Journal of Petrology

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Journal of Petrology Volume 41 Number 2 Pages 229-256 2000
© Oxford University Press 2000

Oxygen Isotope Geochemistry of Oceanic-Arc Lavas

JOHN M. EILER^1,*, ANTHONY CRAWFORD², TIM ELLIOTT³, KENNETH A. FARLEY¹, JOHN W. VALLEY⁴ and EDWARD M. STOLPER¹

¹DIVISION OF GEOLOGICAL AND PLANETARY SCIENCES, CALIFORNIA INSTITUTE OF TECHNOLOGY, PASADENA, CA 91125, USA
²DEPARTMENT OF GEOLOGY, UNIVERSITY OF TASMANIA, GPO BOX 252C, HOBART, TAS. 7001, AUSTRALIA
³FACULTEIT AARDWETENSCHAPPEN, VRIJE UNIVERSITEIT, DE BOELELAAN 1085, 1081 HV AMSTERDAM, NETHERLANDS
⁴DEPARTMENT OF GEOLOGY AND GEOPHYSICS, UNIVERSITY OF WISCONSIN, MADISON, WI 53706, USA

Received November 2, 1998; Revised typescript accepted July 15, 1999

	ABSTRACT

TOP ABSTRACT INTRODUCTION SAMPLES AND ANALYTICAL METHODS RESULTS DISCUSSION SUMMARY AND CONCLUSIONS APPENDIX REFERENCES

 
Variations of oxygen isotope ratios in arc-related lavas can constrain the contributions of subducted crustal igneous rocks, sediments, and fluids to the sub-arc mantle. We have measured oxygen isotope ratios in 72 arc and back-arc lavas from five ocean–ocean subduction zone systems using laser-fluorination analyses of olivine and other phenocrysts and glass. Eighty percent of our samples have ¹⁸O values for any given phase (olivine, plagioclase, glass, or biotite) within 0·2 of the average value for that phase in upper-mantle peridotites and mid-ocean ridge basalt (MORB); the range for each phase is 1·0. This result contrasts with previous studies of whole-rock samples (which are significantly more variable even after exclusion of samples believed to be altered or fractionated by magmatic differentiation) and demonstrates that most arc-related lavas contain 1–2% of ¹⁸O-enriched crustal oxygen from any source (i.e. assimilation or subducted contributions). Elevations in ¹⁸O that do occur in these basic, arc-derived magmas relative to values most common for mantle-derived lavas are associated both with ‘enriched’ radiogenic isotope signatures and, even more strongly, with chemical indices consistent with high integrated extents of melting of their peridotite sources. We interpret these relationships as evidence that melting in the sources of the high-¹⁸O lavas we have studied was fluxed by addition of high-¹⁸O aqueous fluid (or perhaps a hydrous melt) from the subducted slab, such that sources that contain relatively large components of slab-derived fluid or melt are both relatively ¹⁸O enriched and also experienced relatively large amounts of melting. We have developed a quantitative model linking the amount of melting to the extents of ¹⁸O, radiogenic isotope, and trace-element enrichment in a mantle source being fluxed by addition of aqueous fluid. Comparison of this model with observed variations in the geochemistry of lavas from the Vanuatu–Fiji–New Caledonia region (the suite of related samples showing the greatest range in ¹⁸O observed in this study) constrains the amounts and chemical and isotopic compositions of slab-derived phases in the sources of these arc-related lavas. Assuming a ¹⁸O value of 20 for the slab-derived fluid, 0·5–1·0 wt % is added to the sources of most mantle-derived arc magmas; the maximum amount of slab-derived flux in the sources of arc magmas according to our results is 2·5 wt %.

KEY WORDS: oxygen isotopes; arc volcanism

	INTRODUCTION

TOP ABSTRACT INTRODUCTION SAMPLES AND ANALYTICAL METHODS RESULTS DISCUSSION SUMMARY AND CONCLUSIONS APPENDIX REFERENCES

 
There is considerable evidence that igneous rocks, sediments, and fluids from subducted oceanic crust contribute to the formation of convergent margin magmas (e.g. Morris & Tera, 1989; McCulloch & Gamble, 1991; Stolper & Newman, 1994; Elliott et al., 1997; Hawkesworth et al., 1997). These contributions are generally assumed to be introduced into the overlying mantle wedge as aqueous fluids and/or silicate melts released from the downgoing slab as it heats up. Mixing of these water-rich fluids and melts with peridotite in the overlying mantle leads to the production of magmas with the distinctive petrology and geochemistry of convergent margin igneous rocks. Given the likely thermal structure of the subducting slab and overlying mantle, melting in the mantle wedge may in many cases require and/or be controlled by the introduction of these water- and incompatible-element-rich fluids and/or melts because they significantly lower the temperature of the peridotite solidus (Kushiro et al., 1968).

Efforts to identify and characterize the contributions of subducted materials to arc volcanism have emphasized the abundances and isotopic ratios of minor and trace elements that are relatively abundant in altered basaltic rocks and sediments and are concentrated into aqueous fluids and silicate melts relative to residual solids (e.g. K, B, Sr, U; Morris & Tera, 1989; Plank & Langmuir, 1993; Hawkesworth et al., 1994; Leeman et al., 1994; Ryan et al., 1995). These elements can be sensitive indicators of the presence of slab-derived components in the sources of arc lavas; they can also be used to discriminate among and characterize various slab-derived components (e.g. aqueous fluids vs silicate melts; Elliott et al., 1997) and to estimate the time required for these components to travel from the slab to the source regions of magmas in the mantle wedge (Hawkesworth et al., 1997; Regelous et al., 1997). However, such elements are poor monitors of the absolute amounts of slab-derived components in magmatic sources because the concentrations of these elements vary widely among subducted crustal rocks and sediments (Plank & Langmuir, 1998) and because the partitioning of these elements among minerals, melts, and aqueous fluids is strong, highly variable, and only partially constrained by experiment (Green, 1994; Brenan et al., 1995; Keppler, 1996; Johnson & Plank, 1999).

Oxygen isotopes can provide constraints on the sources of arc volcanism that are complementary to those of trace element abundances and their isotope ratios. Most ultramafic rocks and inferred sources of basaltic lavas have ¹⁸O [¹⁸O=(¹⁸O/¹⁶O_sample/¹⁸O/¹⁶O_VSMOW - 1) x 1000, where ¹⁸O/¹⁶O_VSMOW = 0·0020052] values of 5·5 ± 0·2 (Mattey et al., 1994). In contrast, ¹⁸O values of altered basalts in the uppermost oceanic crust are typically +10–20 (compared with values of 0–7 for the lower oceanic crust; Muehlenbachs, 1986; Staudigel et al., 1995); values for coarse clastic sediments are also typically +10–20 (Arthur et al., 1983); values for oceanic clays are typically +15–20 (Arthur et al., 1983); and typical values for carbonate-rich and siliceous pelagic sediments are +30–35 (Kolodny & Epstein, 1976; Arthur et al., 1983). The oxygen isotopic contrast between upper oceanic crust (i.e. sediments and altered basalts) on the one hand and typical upper mantle on the other therefore provides a potential indicator of the presence of subducted upper oceanic crustal materials in the sources of subduction-zone lavas (note, however, that the contrast in oxygen isotope ratios between the mantle and lower-crustal materials may be more subtle and in the opposite direction; Muehlenbachs, 1986). It is known that the distinctive ¹⁸O values of the oceanic crust can be preserved to depths of tens of kilometers in subduction zones (Bebout & Barton, 1989), can persist for long times in deeply recycled materials (Garlick et al., 1971), and can survive transport through the sub-arc mantle in a melt or fluid phase (Eiler et al., 1998). Moreover, there are only modest variations in the concentration of oxygen in most geological solids and fluids, and the isotopic composition of oxygen is only weakly and predictably fractionated by high-temperature exchange among solids, melts, and fluids (e.g. Chiba et al., 1989; Palin et al., 1996); therefore, variations in oxygen isotope ratios in arc-related magmas may more strongly constrain the absolute amounts of slab-derived components contributing to the mantle sources of arc-related magmas than do incompatible trace elements and their isotopes, i.e. uncertainties in such estimates should stem only from uncertainties of a factor of 2 in the ¹⁸O values of subducted upper-crustal materials, rather than from uncertainties of an order of magnitude or more in the abundances of incompatible trace elements. Such constraints on the amounts of fluid and/or melt added to the sources of arc lavas could in turn help to constrain the concentrations of incompatible trace elements in the slab-derived components, to distinguish between metasomatic fluids and melts (which may differ substantially in their abundances of certain elements relative to oxygen), and to define the processes by which slab-derived components interact with the mantle (e.g. bulk mixing, chromatographic exchange, fluid-driven melting).

The oxygen isotope geochemistry of arc-related lavas is known principally through measurements of whole-rock samples of andesite and more evolved rocks [see review by Harmon & Hoefs (1995)]. Such data are the basis of several early and fundamental constraints on subduction-related volcanism, including the discovery of the approximate (±1–2) correspondence between arc-related andesites on the one hand and peridotites and basalts from other tectonic settings on the other (suggesting an origin of the parental liquids of andesitic magmas by melting of peridotite rather than by direct melting of subducted crustal rocks—a significant finding at the time it was made; e.g. Matsuhisa et al., 1973) and of the importance of crustal assimilation and large-scale crustal melting in arc magmatism (e.g. Hildreth & Moorbath, 1988; Davidson & Harmon, 1989). However, the oxygen isotope ratios of the abundant evolved lavas from arcs have been modified from primitive, mantle-derived values by magmatic processes (e.g. fractional crystallization and crystal accumulation; Woodhead et al., 1987; Harmon & Hoefs, 1995), and fine-grained and glassy lavas are particularly susceptible to subsolidus alteration (Davidson & Harmon, 1989; Harmon & Hoefs, 1995). These high-level fractionation and alteration processes can produce shifts in ¹⁸O of 1 or greater in lavas and thus hinder the study of the smaller (<1) variations in oxygen isotope ratios known to characterize the mantle sources of ocean-island basalts (OIB) (Eiler et al., 1996) and that might be more typical consequences of the interactions of the sub-arc mantle with fluids or melts from the subducted slab. Previous efforts to circumvent these difficulties in applying oxygen isotope geochemistry to the study of arc magmatism have included correction or exclusion of measured ¹⁸O values of whole-rock specimens inferred to be significantly altered and/or unusually evolved (Woodhead et al., 1987; Davidson & Harmon, 1989; Harmon & Hoefs, 1995), or analysis of glass (Ito & Stern, 1985; Macpherson & Mattey, 1997) or phenocrysts (Singer et al., 1992; Smith et al., 1996; Thirlwall et al., 1996; Macpherson & Mattey, 1997; Macpherson et al., 1998) rather than whole rocks.

We describe in this paper the results of a study of oxygen isotope variations in relatively magnesian lavas from oceanic arcs (and associated back-arc basins) based primarily on measurements on olivine phenocrysts, supported with less extensive measurements on other phenocryst phases and glass. Olivine has a low rate of oxygen exchange in the absence of replacement by alteration minerals (Reddy et al., 1980); it is an early phase in the low-pressure crystallization sequences of primitive basaltic liquids; and it is thought to be present as a major phase in the mantle sources of basalts. It is therefore an obvious target for providing insights into the upper-mantle sources of basic magmas. We note, however, that our emphasis on olivine-bearing lavas may have prejudiced our sampling against magmas (and their fractionation products) produced by single-stage melting of non-ultramafic sources (e.g. partial melts of subducted crust unmodified by reaction with the mantle wedge in transit to the surface), if such magmas exist (Defant & Drummond, 1990). Approximately 20 laser-fluorination measurements of the ¹⁸O of olivines from arc-related magmas have been previously published, distributed among four studies of the Lau basin, the Lesser Antilles, and the Kermadec–Hikurangi margin (Smith et al., 1996; Thirlwall et al., 1996; Macpherson & Mattey, 1997; Macpherson et al., 1998). We have not integrated these data with our study because of both the small number and wide distribution of these previous measurements and the difficulty in interpreting small (several tenths of per mil) variations in ¹⁸O measurements from different laboratories without careful cross-standardization. However, we note that data from those studies are generally within the range of results presented here. Results for other phases (glass and clinopyroxene) from these recent studies will be discussed further below.

	SAMPLES AND ANALYTICAL METHODS

TOP ABSTRACT INTRODUCTION SAMPLES AND ANALYTICAL METHODS RESULTS DISCUSSION SUMMARY AND CONCLUSIONS APPENDIX REFERENCES

 
The suite of samples we have studied includes 72 basalts, basaltic andesites, shoshonites, boninites, and alkali basalts from five ocean–ocean subduction systems: the Mariana arc and back-arc trough, the South Sandwich arc and associated Scotia Sea back arc, the Vanuatu–Fiji arc and nearby New Caledonia arc (these are treated as a single suite in much of the subsequent discussion), and Papua New Guinea [alkaline lavas of the Tabar, Lihir, Tanga, and Feni (‘TLTF’) islands, derived from melting of the mantle wedge of a recently active subduction zone]. The sample suite was restricted to lavas from oceanic (as opposed to continental) arcs and to olivine-phyric lavas that are more primitive than those typically considered in previous studies of the oxygen isotope geochemistry of arc lavas (the mean MgO and SiO₂ contents of samples for which major element data are available are approximately 7 and 51 wt %), although samples were also chosen so as to include a significant range in minor and trace element geochemistry and radiogenic isotope ratios. Analyses are available for 58 of the 72 samples for major element concentrations, trace element concentrations, mineral chemistry, and/or Sr, Nd, Pb, and/or U-series isotope geochemistry, providing a petrological and geochemical context in which to interpret our data. The supplemental chemical data are reported using the same sample labels as in the literature (Hawkesworth et al., 1977; Saunders & Tarney, 1979; Hawkins & Melchior, 1985; Maillet et al., 1986; Kennedy et al., 1990; Eggins, 1993; Monzier et al., 1993; Sigurdsson et al., 1993; Rogers & Setterfield, 1994; Crawford et al., 1995; Pearce et al., 1995; Elliott et al., 1997; Peate et al., 1997) or are from unpublished studies of A. Crawford or R. Stern; a table including the supporting geochemical data is available on request from the authors. The 14 samples that have not been previously characterized are related to characterized samples and come from previously studied sequences of lavas, such that aspects of their geochemistry can be confidently inferred; these samples are indicated in Table 1 and in the figures. All of the significant variations and trends in ¹⁸O observed in this study have extremes exhibited by well-characterized samples.

All samples were crushed in a steel percussion mortar and dry-sieved to size fractions of 600–300 µm and 300–150 µm for hand picking of olivine, plagioclase, glass, and/or biotite separates. All separates were cleaned of dust by briefly blowing filtered compressed air over them while holding them in a 100 µm sieve. Olivine separates were prepared for 67 of the 72 samples; 24 plagioclase separates, 11 glass separates, and two biotite separates were also prepared.

Most mineral and glass separates were analyzed for ¹⁸O at the CO₂-laser-fluorination laboratory at the University of Wisconsin, Madison [‘UW’, described in detail by Valley et al. (1995)]. A subset of the data was measured by the same methods at the University of Southern California (‘USC’). Analyses at UW were made on 12 days over a period of 7 months, whereas those at USC were made on 3 days over a period of 1 month. Measurements of unknowns on each day in both laboratories were accompanied by 4–8 measurements of UWG-2 garnet (Valley et al., 1995) and 2–4 measurements of the SCO-1 olivine working standard (Eiler et al., 1995). The 1 uncertainty for a given standard on any given day averaged 0·07. In addition, 60 of the unknown samples were measured 2–3 times, and the mean of the 1 uncertainties of these multiple analyses is 0·06. On 12 of the 15 days on which analyses were performed, measurements of each standard averaged within 0·1 of the accepted value and no systematic correction was applied to the data. On 3 days, however, the averages for both UWG-2 and SCO-1 differed from their accepted values by 0·1–0·2, and a correction was applied to the data equal to the average of the differences between the mean-measured value and the accepted value for each of the two standards. Such corrections are required on 25% of the working days at the UW laboratory and have been inferred to be caused by small variations in the condition of the gas-purification and gas-transfer apparatus (Valley et al., 1995). Four olivine samples (i.e. unknowns) that had been measured on a day on which a correction was applied based on the analyses of the standards were reanalyzed on another day on which no correction was required; the average difference between the corrected and uncorrected measurements was 0·06. One analysis of an olivine separate (GUG-3) yielded an anomalous result (3·5) that was not reproduced by two further measurements of a new mineral separate from this sample. The anomalous value was discarded, and only the average of two additional measurements is listed in Table 1a.

View this table: [in this window] [in a new window]   Table 1a: Oxygen isotope data and select major elements

 

The accuracy of our analyses was assessed by measurement of seven silicates (three olivines, two basaltic glasses, one biotite, and one plagioclase; Table 1b), six of which had been previously analyzed multiple times by other laboratories using conventional fluorination procedures. The average difference between the mean value of ¹⁸O for these materials determined by laser fluorination at UW and the accepted values was 0·09; that is, comparable with our precision based on repeated analyses of the UWG-2 working standard. Moreover, the UWG-2 garnet standard and three other materials (SCO-1, KHO-1, and AH95-22 glass) were analyzed in both the UW and USC laser fluorination laboratories, and the differences between the average values for these materials from these two laboratories averaged 0·06. Overall, these demonstrations of reproducibility and of agreement between laser fluorination analyses at UW and USC and with conventional fluorination analyses allow us to infer that our data are accurate and precise to approximately ±0·07 (1) for the phases analyzed in this study.

View this table: [in this window] [in a new window]   Table 1b: Oxygen isotope standards

 

	RESULTS

TOP ABSTRACT INTRODUCTION SAMPLES AND ANALYTICAL METHODS RESULTS DISCUSSION SUMMARY AND CONCLUSIONS APPENDIX REFERENCES

 
Values of ¹⁸O determined in this study vary over a range of 1 or less for each of the phases analyzed: the ranges are 0·98 for the 67 olivine analyses; 0·38 for the 24 plagioclase analyses; and 0·14 for the 11 glasses (Table 1a). Average values of ¹⁸O for olivine, plagioclase, and glass are within 0·1 of those for fresh mid-ocean ridge basalts (MORBs) or their inferred sources (Ito et al., 1987; Eiler et al., 1996); likewise, the average for olivine in our sample suite is within 0·1 of the average for olivine from mantle peridotite xenoliths (Mattey et al., 1994). Even the largest range in ¹⁸O for a single phase (olivine) is small relative to expectations based on most previous studies of arc-related lavas (Harmon & Hoefs, 1995). The relatively small range in ¹⁸O is particularly striking given the number and diversity of samples that we have studied: our suite of samples is taken from five different arcs and three back-arc regions and includes much of the range in radiogenic isotope ratios (e.g. ⁸⁷Sr/⁸⁶Sr = 0·7026–0·7044), trace element compositions (e.g. Ba/Nb = 6–400), U-series radioisotope activity ratios (e.g. [U²³⁸/Th²³⁰] = 1·0–1·6), and petrologic diversity (i.e. island-arc basalts and basaltic andesites, back-arc-basin basalts, shoshonites, boninites, and alkali basalts) known from oceanic-arc lavas. It should be noted, however, that we have not included any samples characterized by exceptionally high ⁸⁷Sr/⁸⁶Sr ratios and Th concentrations, such as those common to the Philippines, Indonesia, and the Aeolian islands [i.e. the ‘high-Ce/Yb’ arcs of Hawkesworth et al. (1994)].

Approximately 80% of the ¹⁸O values for olivine are within the range for typical upper-mantle olivine based on peridotite xenoliths (5·0–5·4; Mattey et al., 1994). Only one sample (Charlotte 22 from Vanuatu) is significantly below this range (and this one by only 0·15); the remaining 20% of the samples are enriched in ¹⁸O relative to this range by up to 0·43 (Fig. 1). Relatively ¹⁸O-enriched samples include three island-arc basalts (all from western Vanuatu), three shoshonites (also from the Vanuatu region), three boninites (from Vanuatu and New Caledonia), and six alkali basalts from the TLTF islands (Papua New Guinea). In the following paragraphs we assess some of the systematics of our data, comparisons with previous work, and correlations of our data with other geochemical characteristics of the samples.

View larger version (46K): [in this window] [in a new window]   Fig. 1. Values of 18O for olivine from oceanic-arc lavas analyzed in this study, organized into groups defined by geographic location and petrologic type. The gray field indicates the range typical of olivine from upper-mantle peridotite xenoliths (Mattey et al., 1994) and inferred for olivine in the MORB source (5·2 ± 0·2; Ito et al., 1987; Eiler et al., 1996). Short vertical lines mark average values for each suite of related samples. Samples are labeled by location and by petrologic type: island-arc basalt (IAB, circles), back-arc-basin basalt (BABB, triangles), shoshonite (squares), boninite (stars), and alkali basalt (Alk B, diamonds). TLTF indicates lavas from the Tabar–Lihir–Tanga–Feni islands, Papua New Guinea. Symbols for Mariana arc and back-arc samples are open; for South Sandwich–Scotia Sea samples they are gray; for the Vanuatu–Fiji–New Caledonia suite they are filled; for TLTF samples they are open.

 

Distribution of oxygen isotopes among coexisting phases
We have made nine determinations of basaltic glass–olivine oxygen isotope fractionation (_{glass–olivine}; Fig. 2a) and 21 determinations of plagioclase–olivine fractionation (_{plagioclase–olivine}; Fig. 2b; where _i–j = ¹⁸O_i – ¹⁸O_j). Both of these fractionations are essentially constant given the precision of our measurements: _{plagioclase–olivine} = 0·65 ± 0·14 (1) and _{glass–olivine} = 0·36 ± 0·11 (1). The expected precision of a determination of _i–j is 0·10 (1), based on independent errors of 0·07 for a determination of ¹⁸O on each phase.

View larger version (42K): [in this window] [in a new window]   Fig. 2. Histograms illustrating the variability in glass–olivine (a) and plagioclase–olivine (b) measured in this study (dark gray boxes) and from previous studies (diagonally ruled boxes; data from Onuma et al., 1970; Clayton et al., 1971, 1972; Kyser et al., 1981, 1982; Macpherson & Mattey, 1997). Vertical dashed lines and light gray fields illustrate the range of fractionations between olivine and either An50–90 plagioclase or basaltic melt at 1300°C based on experimental studies of isotope partitioning and models of isotope partitioning among silicate phases (Muehlenbachs & Kushiro, 1974; Chiba et al., 1989; Rosenbaum, 1994; Matthews et al., 1998).

 
The average fractionation we determine between plagioclase and olivine phenocrysts compares favorably with that expected at magmatic temperatures based upon exchange experiments (e.g. 0·58 for An₉₀ plagioclase to 0·76 for An₅₀ plagioclase at 1300°C; Chiba et al., 1989). The fractionation between basaltic liquid and olivine is not known directly from experiment, but a value of 0·32 is predicted at 1300°C based upon the combination of experimentally determined melt–CO₂, forsterite–calcite, and calcite–CO₂ fractionations (Muehlenbachs & Kushiro, 1974; Chiba et al., 1989; Rosenbaum, 1994). However, the accumulated error in such an estimate is relatively large (±0·2–0·3, 1), so the close correspondence between this predicted value and our measured value may not be significant. An empirical model for oxygen isotope partitioning between minerals and melts (Matthews et al., 1998) predicts a value of 0·4–0·5 for _{melt–olivine} for the lava compositions considered in this study. Our results compare favorably with the range of values expected from these two independent estimates.

The fractionations of oxygen isotopes we observe among coexisting phases are within the ranges of but more restricted than fractionations observed in previous studies of phenocrysts and glass or groundmass in basic lavas (Fig. 2). Values of _{glass–olivine}, _{groundmass–olivine}, or _{whole rock–olivine} from previous studies range from 0·2 to 2·0, and values for equilibrium melt–olivine fractionation have been suggested to be as high as 0·9–1·0 (Onuma et al., 1970; Clayton et al., 1971, 1972; Hoernes & Friedrichsen, 1977; Kyser et al., 1981, 1982; Macpherson & Mattey, 1997). Our results indicate a value of 0·4 ± 0·1, significantly lower than the upper end of the range of previous data. Previous measurements of _{plagioclase–olivine} are similarly more widely distributed (0·4–1·3) and on average larger than observed in this study (Onuma et al., 1970; Clayton et al., 1972; Hoernes & Friedrichsen, 1977; Kyser et al., 1981, 1982). These differences may be due to the fact that ours is the first study to make a significant number of determinations of _{glass–olivine} and _{plagioclase–olivine} using laser-based measurements, and therefore we may have been more successful at avoiding sample alteration and the well-known analytical difficulties associated with conventional oxygen isotope measurements of olivine (Mattey et al., 1994).

Comparison with previous work
The range of ¹⁸O values we find for olivine phenocrysts (0·98) is a factor of 10 lower than that reported previously based on conventional fluorination measurements of whole-rock samples from oceanic arcs [summarized by Harmon & Hoefs (1995)]. Some of this difference could reflect our emphasis on olivine-bearing samples, which are primitive relative to the full range of arc-related lavas represented in the conventional data base. However, this is not likely to be the principal cause because much of the reported range in whole-rock measurements can be found in basalts, basaltic andesites, and andesites in the same arcs we have studied (Pineau et al., 1976; Ito & Stern, 1985; Woodhead et al., 1987; Harmon & Hoefs, 1995). We infer, as has previously been concluded from study of mantle peridotites and ocean-island basalts (Mattey et al., 1994; Eiler et al., 1995, 1996), that the oxygen isotope variations in olivine phenocrysts are more restricted than in whole rocks primarily because analysis of fresh phenocrysts (particularly olivines) avoids or minimizes the effects of low-temperature alteration; nevertheless, an additional factor may be that the increases in ¹⁸O that are known to build up in residual liquids from fractional crystallization (Matsuhisa et al., 1973) are minimized by focusing on relatively primitive, olivine-bearing samples.

Our data contradict a conclusion reached by some previous studies of oxygen isotope variations in oceanic-arc basalts that the average value of ¹⁸O in oceanic-arc magmas, even after correction for low-temperature alteration and magmatic fractionations, is 0·5 higher than for MORB and back-arc-basin basalts (Woodhead et al., 1987; Harmon & Hoefs, 1995) and that this difference reflects a common and relatively large amount of subducted crustal oxygen in the sources of these arc magmas. In contrast, we find that elevations in ¹⁸O of 0·5 with respect to MORB are an extreme found only in a minority of arc samples and that the sources of both arc and back-arc lavas are on average closely similar to estimates for the typical upper mantle based on measurements on MORBs and peridotite xenoliths (Fig. 1). We note that in this respect our results confirm the work of Ito & Stern (1985), who found that glasses from both the Mariana arc and trough are similar in ¹⁸O to MORB. Similarly, recent results for glasses from the Lau basin (Macpherson & Mattey, 1997) span a small range in ¹⁸O near the average value for MORB and display correlations with chemistry that are analogous to correlations displayed by our data (discussed below). Recent study of lavas from the Lesser Antilles and the Kermadec–Hikurangi margin have found larger ranges and higher average values of ¹⁸O than both MORB and the results of this study, but in both cases these differences were interpreted to reflect assimilation of continental material (Smith et al., 1996; Thirlwall et al., 1996; Macpherson et al., 1998); these studies confirm the conclusions of Harmon et al. (1981), Hildreth & Moorbath (1988), and Davidson & Harmon (1989) that variations in ¹⁸O of arc-related lavas erupted through thick sequences of continental rocks or sediments dominantly reflect assimilation–fractional crystallization processes.

Correlations between ¹⁸O and sample chemistry
When comparing our oxygen isotope data with other chemical indices, it is useful to focus on a single phase to remove the effects of oxygen isotope fractionations among different crystalline phases and glass. We have selected olivine for this purpose because it was measured in most of our samples and, for the reasons discussed above, we believe it to be a reliable monitor of ¹⁸O in unaltered and relatively primitive basaltic lavas. For the five out of 72 samples in which olivine was not analyzed, we calculated an ‘equivalent olivine’ value equal to the value measured in plagioclase or glass in that sample minus the average fractionation between olivine and plagioclase or glass observed in this study. Although these fractionations are well defined by our data, this correction introduces additional uncertainty for these samples; they are distinguished from the other data by parentheses in Table 1. None of the oxygen isotope variations or correlations observed in this study are significantly dependent on these five values, but they are provided for completeness. We emphasize at the outset of this discussion that all values of ¹⁸O in excess of 5·4 occur in lavas from the Vanuatu and New Caledonia arcs or the TLTF islands (i.e. Papua New Guinea), and we reiterate that such ¹⁸O-enriched samples are atypical (20% of our survey sampling); thus, the observed trends and our interpretations of them are biased toward processes relevant to these samples. However, the other studied suites are consistent with the trends defined by these extreme samples and therefore are consistent with the interpretation that such suites were influenced, though not as strongly, by similar processes.

Values of ¹⁸O in olivine phenocrysts in our samples are not correlated with the whole-rock SiO₂ or MgO contents of the host lavas when the data set is viewed as a whole (Table 1). We infer that the effects on ¹⁸O of magmatic fractionation and accumulation leading to such correlations for whole-rock data from arc settings (Matsuhisa et al., 1973; Woodhead et al., 1987) have been minimized by our emphasis on relatively primitive olivine-bearing lavas.

A key question is whether ¹⁸O values for our samples correlate with conventional chemical and isotopic monitors of slab components in the sources of arc-related lavas. To answer this question, our results for the oxygen isotope composition of olivine in arc-related lavas are compared in Fig. 3 with ⁸⁷Sr/⁸⁶Sr and various trace- and minor-element indices that plausibly monitor the contribution of subducted materials to the sources of arc lavas.

View larger version (30K): [in this window] [in a new window]   Fig. 3. Comparison of oxygen isotope ratios in olivine phenocrysts from arc and back-arc lavas with other geochemical indicators in the host lavas: (a) 87Sr/86Sr; (b–f) various minor and trace element abundance ratios (by weight). Sr isotope and element-abundance ratios are either measured on the same samples (small symbols) or estimated based on data for closely related samples (large symbols). (See text for sources of trace element and Sr isotope data.) Symbols are as in Fig. 1. Gray boxes mark typical NMORB compositions (Sun & Nesbitt, 1979; Ito et al., 1987; Michael, 1995). It should be noted that only lavas with MgO contents between 6 and 12 wt % are plotted in (c), to minimize the effects of fractionation on the K2O/TiO2 ratio. This range was selected based on the observed relationship between this ratio and MgO in Vanuatu and Mariana arc lavas.

 

Values of ¹⁸O in olivine span a restricted range near the average upper-mantle value in lavas with low-⁸⁷Sr/⁸⁶Sr ratios; elevated ¹⁸O values are restricted to samples at the high end of the distribution of ⁸⁷Sr/⁸⁶Sr (Fig. 3a). Radiogenic strontium in arc-related lavas is generally interpreted as a monitor of slab-derived contributions to their sources (Gill, 1981), and thus the relationship shown in Fig. 3a suggests that the elevated oxygen isotope ratios we observe are related to slab-derived components. The observed relationship between ¹⁸O and ⁸⁷Sr/⁸⁶Sr is consistent with a hyperbola that results from mixing between an Sr-poor, low-¹⁸O, low-⁸⁷Sr/⁸⁶Sr component (i.e. the approximately MORB-like mantle wedge) and an Sr-rich, high-¹⁸O, high-⁸⁷Sr/⁸⁶Sr component (i.e. slab-derived fluid and/or melt). It should be noted that the overall hyperbolic trend shown in Fig. 3a, although defined by the data set as a whole, is best defined by the single suite of island-arc basalts from the western Vanuatu arc (filled circles); this suite is the only group of related samples considered in this study that covers a sufficient range in ¹⁸O and ⁸⁷Sr/⁸⁶Sr to display the full trend. It should be noted also that correlations between ¹⁸O and ⁸⁷Sr/⁸⁶Sr as a result of assimilation of pre-existing crust are generally linear rather than hyperbolic, because of the similarity in Sr concentrations of arc-related lavas and older rocks in the arc crust (e.g. Davidson & Harmon, 1989); therefore, the hyperbolic relationship in Fig. 3a suggests processes other than crustal assimilation.

Arc-related lavas are characterized by distinctive enrichments and depletions in certain trace elements relative to MORB and OIB that are associated with (or defined as characteristics of) slab-derived components (Gill, 1981). Certain chemical characteristics are common to nearly all arc-related lavas and are rare elsewhere [e.g. enrichment of large ion lithophile elements (LILE) such as K, Ba, and Sr with respect to high field strength elements (HFSE) such as Ti and Nb], whereas others are thought to be diagnostic of the presence in the sub-arc mantle of distinctive components such as melts or aqueous fluids that may sample a specific part of the subducting slab (e.g. the basaltic crust vs the overlying sediments; Kay, 1980; Ellam & Hawkesworth, 1988; Elliott et al., 1997; Turner et al., 1997) or pre-existing ‘OIB-like’ enrichments in the mantle (Lin et al., 1989). The number and identity of such components differ according to the set of elements considered; for example, the classification of Hawkesworth et al. (1997) would define all lavas considered in this study as falling within the compositional field (Sr/Th > 200; Th < 5 ppm) interpreted to reflect dominantly addition of a slab-derived aqueous fluid; in contrast, Elliot et al. (1997) concluded that sediment-derived melt controls the budgets of many trace elements in the sources of samples we have studied from the Marianas. However, certain common themes have emerged in recent studies: uranium excesses (i.e. [²³⁸U /²³⁰Th] > 1) and enrichments of fluid-soluble elements (Ba, Sr, U) with respect to non-fluid-soluble but incompatible elements (Th, Nb, Zr) are generally identified with slab-derived aqueous fluids dominated by contributions from the basaltic portions of the subducted oceanic crust (Ellam & Hawkesworth, 1988; Elliott et al., 1997; Turner et al., 1997); in contrast, high concentrations of Th, elevated ratios of Th and rare earth elements (REE) to Nb, strong light rare earth element (LREE) enrichments, negative Ce anomalies, and/or U-series isotopic compositions near the equiline (i.e. [²³⁸U /²³⁰Th] 1·0) are associated with a component of sediment-derived melt (Ellam & Hawkesworth, 1988; Elliott et al., 1997; Turner et al., 1997), or possibly a pre-existing ‘OIB-like’ enrichment in the mantle (Lin et al., 1989).

Elevated values of ¹⁸O (= 5·4) observed in this study are generally associated with elevated ratios of highly incompatible LILE to moderately incompatible trace elements not believed to be highly fluid soluble (e.g. Sr/Yb and K₂O/TiO₂; Fig. 3b and c). These trends, although somewhat scattered, are consistent with the conclusion drawn from Fig. 3a that ¹⁸O enrichments in lavas are associated with distinctive, ‘enriched’ components in their sources. Relationships between ¹⁸O and trace-element ratios thought to discriminate among different varieties of enriched components are complex and fail to reveal a unique association between ¹⁸O enrichments and any one of the several distinctive slab-derived components currently believed to contribute to the sub-arc mantle (Fig. 3d–f). For example, arc lavas with high-¹⁸O olivines have Ba/La ratios that are high relative to NMORB, but they span a considerable range and are not significantly elevated in Ba/La relative to the lower-¹⁸O arc-related lavas; in particular, the most extreme elevations of Ba/La among lavas examined in this study are in samples from the South Sandwich arc, yet these have ¹⁸O values indistinguishable from average upper mantle (Fig. 3d). It should be noted that enrichment of Ba relative to other LILE such as La is one of the most widely used indicators of contributions of fluid from the basaltic portions of the subducted slab to the sub-arc mantle (e.g. Gill, 1981; McCulloch & Gamble, 1991; Hawkesworth et al., 1994), so our data do not support an exclusive association of ¹⁸O enrichments with such a component. Similarly, lavas from Guguan island in the Mariana arc have ¹⁸O values for olivine within the range of olivine in typical upper-mantle peridotites (Table 1), despite spanning a significant range in [²³⁸U/²³⁰Th] ratio from 1·29 to 1·56; such large U excesses are taken as evidence that fluids from the basaltic portions of the subducting slab dominate the U and Th budget of the mantle sources of Guguan lavas (Elliott et al., 1997).

High-¹⁸O lavas are generally LREE enriched relative to NMORB and to most lower-¹⁸O arc-related lavas (Fig. 3e). However, shoshonites from the northern Mariana arc have ¹⁸O values similar to average upper mantle, yet they display extreme LREE enrichments, demonstrating that LREE and ¹⁸O enrichments are not uniquely linked in arc environments. Similarly, enrichments in LREE with respect to HFSE (e.g. La/Zr; Fig. 3f) are generally positively correlated with ¹⁸O among our samples, but the Mariana shoshonites have extreme enrichments of La with respect to Zr without any elevation in ¹⁸O. Enrichments of LREE relative to heavy rare earth elements (HREE) and HFSE in arc lavas have been proposed as evidence that sediment-derived melts dominate the trace-element characteristics of enriched sub-arc mantle (Brenan et al., 1995; Elliot et al., 1997); the results shown in Fig. 3e and f indicate that ¹⁸O elevations are generally but not universally associated with such a signature.

Figure 3d–f thus demonstrates an important result of our study: extreme ‘slab-derived fluid’ signatures (e.g. Ba/La and U/Th enrichment usually identified with fluids from the basaltic component of the subducting crust) and ‘sediment-derived melt’ and/or ‘OIB-like’ signatures (e.g. LREE/HREE or LREE/HFSE enrichment) can each be present in the sub-arc mantle sources of arc-related magmas without a distinctive oxygen isotope signature. Collectively, the relationships illustrated in Fig. 3 suggest that, although elevated ¹⁸O values of arc-derived magmas are associated in general terms with commonly used trace element indicators of slab-derived components in their sources (Fig. 3a–c), the abundances of such components as measured by oxygen isotopes are not related one-to-one to the proportions of distinguishable ‘slab fluid’ and ‘sediment melt’ components based on diagnostic trace element ratios (Fig. 3d–f).

Although the relationships between ¹⁸O and commonly used trace element monitors of slab-derived components are scattered, there are several well-defined relationships between the ¹⁸O values from this study and other geochemical indices, which rather than being direct monitors of slab contributions to the sources of arc lavas are useful as monitors of the extent of melting and/or prior depletion of the peridotitic sources of basaltic melts. These relationships are illustrated in Fig. 4 and described in the following paragraphs.

View larger version (28K): [in this window] [in a new window]   Fig. 4. Comparisons of oxygen isotope ratios in olivine phenocrysts from arc and back-arc lavas with other geochemical indicators: (a) wt % TiO2 in the host lava; (b) TiO2(8·0) values of the host lava [calculated following methods described by Klein & Langmuir (1987); see text for details]; (c) the cr-numbers of coexisting chromite phenocrysts [i.e. the Cr/(Cr + Al) ratio on a molar basis]; (d) the Yb/Sc ratio (by weight) of the host lavas. Gray boxes mark typical NMORB compositions. Chemical and mineral-composition data are either measured on the same samples (small symbols) or estimated based on data for closely related samples (large symbols). (See text for data sources.) Symbols are as in Fig. 1.

 

TiO₂ contents
Values of ¹⁸O are compared with concentrations of TiO₂ (wt %) measured in whole-rock specimens of host lavas in Fig. 4a; samples plotted in this panel are restricted to those with 6–12 wt % MgO to minimize variations related to fractional crystallization. This figure shows that the ¹⁸O-enriched samples all have low TiO₂ contents, generally less than 0·8 wt %—approximately half the average concentration in NMORB (Sun & McDonough, 1989) and at the lower end of the range observed in arc lavas (Plank & Langmuir, 1988). The overall trend shows both systematically higher TiO₂ contents for back-arc magmas (triangles) than for arc magmas with comparable ¹⁸O values and a negative correlation between TiO₂ and ¹⁸O when only the arc lavas are considered. All other things (e.g. source composition, mineralogy of the residue, and extent of fractionation) being equal, lower values of TiO₂ in basalts indicate higher extents of single-stage melting and/or melting of a more depleted (i.e. previously melted) peridotite source.

Although Fig. 4a is restricted to lavas with 6–12 wt % MgO, some of the variation in TiO₂ could still reflect superposition of the effects of fractionation on source and degree-of-melting effects. We have evaluated this by calculating values of TiO_2(8·0) for suites of related samples to correct for low-pressure fractionation (Klein & Langmuir, 1987). Eleven suites of lavas were selected that were sufficiently well sampled to permit an estimation of TiO_2(8·0) values: two islands from the Mariana arc (Guguan and Agrigan), the Mariana trough as a whole, the Scotia Sea as a whole, one island from the South Sandwich arc (Zavodovski), three islands from the Vanuatu arc (Ambrym, Tanna, and Tongoa), boninites from New Caledonia and the Hunter fracture zone (Vanuatu), and one island from the ‘TLTF’ arc (Lihir). TiO₂ contents of individual samples from each suite and related samples from that suite not analyzed for ¹⁸O in this study were regressed with MgO content, and the TiO₂ value of the resulting line at 8·0 wt % MgO was taken as the TiO_2(8·0) value for that suite; this value applies to all members of that suite and the ¹⁸O value is the average for one or more members of that suite analyzed for oxygen isotope composition in this study. Figure 4b confirms the key feature of Fig. 4a: elevated ¹⁸O values are associated with low values of TiO_2(8·0) in the host lavas, generating an overall negative correlation. Scatter about the trend in Fig. 4a is reduced in Fig. 4b. The trend in Fig. 4b is particularly well defined by data other than alkali basalts from Papua New Guinea (the open diamond). Alkali basalts, including those in arc settings, are systematically higher in Ti than related tholeiitic lavas (e.g. Shimizu & Arculus, 1975), so this discrepancy is not unexpected.

Macpherson & Mattey (1997) presented data for oxygen isotope variations in Lau basin back-arc lavas that show a subtle trend of increasing ¹⁸O with decreasing Na₂O_(8·0). The behavior of Na during partial melting of peridotite is, to first order, similar to that of Ti, and thus this observation is analogous to that described above. Lavas examined in this study and found to have relatively high values of ¹⁸O (¹⁸O_olivine = 5·4) also generally have low concentrations of Na relative to the sample suite as a whole (2·1 vs 2·5 wt % Na₂O), although the overall correlation between ¹⁸O and Na₂O is not as well defined as those illustrated for TiO₂ in Fig. 4a and b.

Chrome number
Among samples containing chromite phenocrysts (approximately half of our samples; mostly from the Vanuatu suite), high ¹⁸O values are restricted to samples with chromites having chrome numbers (i.e. cr-number = Cr/[Cr + Al]) of 0·80 or greater (Fig. 4c). Such values are high compared with chromites in abyssal peridotites (0·1 < cr-number < 0·6) and most basaltic lavas (0·2 < cr-number < 0·7); comparable values are common only in alpine peridotites, boninites, high-Mg andesites, and, less commonly, oceanic plateau basalts (Dick & Bullen, 1984; Arai, 1994). Relatively high cr-number values in spinels from peridotites are interpreted as an indication that the rocks are residual to high integrated extents of melting (i.e. the total amount of melt extracted from the source, whether in single or multiple melting events, is high) or high extents of melt–peridotite reaction (Arai & Matsukage, 1996); this is because Cr is compatible and Al incompatible in peridotitic residues equilibrated with basaltic melts (Dick & Bullen, 1984). When found in basaltic phenocrysts, such spinels are similarly thought to reflect derivation of those lavas from sources that have undergone unusually high integrated extents of melting (Dick & Bullen, 1984; Arai, 1994). Thus, the association of high ¹⁸O values with high cr-number suggests that the residues remaining in the sources of lavas with elevated ¹⁸O are highly depleted.

Yb/Sc
There is a negative correlation between ¹⁸O and Yb/Sc among samples from this study, regardless of geographic location and petrologic type (Fig. 4d). The Yb/Sc ratio of unfractionated, mantle-derived magmas is predicted to decrease continuously with increasing integrated extents of melting of peridotitic sources because Yb is somewhat more incompatible than Sc during peridotite melting (Klein & Langmuir, 1987). Clinopyroxene fractionation will lead to progressive increases in this ratio in residual liquids, but it should be insensitive to fractional crystallization of other phases and particularly during fractionation of olivine from relatively primitive magmas. Although care must be taken to exclude highly fractionated magmas when examining variations in Yb/Sc, observed relationships between Sc and Yb concentrations and other indices of extent of melting in both MORB and arc-related magmas (Klein & Langmuir, 1987; Plank & Langmuir, 1993) confirm that they can be a valuable measure of integrated degree of melting. We have examined the Yb/Sc relationship shown in Fig. 4d for the effects of clinopyroxene fractionation; although this process probably influenced the two highest Yb/Sc ratios, the trend is unchanged if only samples with 8% MgO and/or 25 ppm Sc are plotted (a compositional range over which the effects of clinopyroxene fractionation are not expected to be significant). The correlation in Fig. 4d is therefore consistent with the relationships exhibited in Fig. 4a–c; that is, higher ¹⁸O is associated with higher extents of single-stage melting and/or higher degrees of depletion of the peridotitic sources of arc-related basalts as a result of previous melting events.

	DISCUSSION

TOP ABSTRACT INTRODUCTION SAMPLES AND ANALYTICAL METHODS RESULTS DISCUSSION SUMMARY AND CONCLUSIONS APPENDIX REFERENCES

 
Key results of this study are that variations in ¹⁸O among arc-related lavas, although unusual, are associated with known trace element and radiogenic isotopic tracers of slab input to their sources and, more importantly, they are correlated with chemical indices that are sensitive to the integrated extent of melting of those sources. In the following discussion we develop a model for the petrogenesis of arc-related lavas that both explains this observation and is consistent with the other known characteristics of these lavas. Much of this discussion is biased toward explanation of the relationship between ¹⁸O and other characteristics of lavas from the Vanuatu–Fiji–New Caledonia region because this is the suite of samples covering the full range of ¹⁸O observed in this study. However, the samples from the other suites fall on the same overall trends, suggesting that the model we develop could apply more broadly; more work will be needed to define whether these other suites (as well as suites not considered in this study) can be understood in the framework of the same model.

Oxygen isotope shifts associated with melting and magmatic evolution
Fractionations of oxygen isotopes among coexisting minerals and melts are small at the temperatures of peridotite melting and basaltic volcanism (Chiba et al., 1989; Palin et al., 1996; Fig. 2), but it is possible that even small fractionations could produce the 1 variations in ¹⁸O observed for olivine in this study by Rayleigh fractionation during melting or crystallization. This possibility can be tested using experimentally known and empirically estimated high-temperature fractionations of oxygen isotopes among minerals and melts.

Crystallization of olivine ± pyroxene will raise both the ¹⁸O values (Chiba et al., 1989; Matthews et al., 1998; Fig. 2) and TiO₂ contents of residual liquids, and therefore should produce trends opposite to that observed in Fig. 4a. Moreover, the correction of TiO₂ contents to a common MgO content should minimize the effects of fractional crystallization, and thus the variations in TiO_2(8·0) and the correlation with ¹⁸O shown in Fig. 4b would be difficult to account for by this process. Crystallization of Cr-rich spinel as part of a fractionating assemblage is expected to lead to progressive decreases in the cr-number of spinel phenocrysts in most circumstances (Arai, 1994); therefore, fractional crystallization can produce negative correlations between ¹⁸O and the cr-number of spinels but not trends with positive slope such as that shown in Fig. 4c for arc-related lavas. Fractionation of clinopyroxene rapidly removes Sc from magmas, such that Yb/Sc ratios in low-MgO lavas can be high (>0·1, based on data reported in the references in the ‘Samples and analytical methods’ section for evolved lavas in the arcs we have studied). However, the MgO and Sc concentrations of the samples defining the trend in Fig. 4d (average 8 wt % and 35 ppm) are substantially higher than those typical of such fractionated magmas (<5 wt % and <25 ppm). In any case, fractionation of clinopyroxene is expected to produce a positive trend in Fig. 4d rather than the negative one that is observed (i.e. because it raises both the ¹⁸O and Yb/Sc of residual magma). The relationships observed in Fig. 4 are therefore at odds with the expected effects of fractional crystallization, and we conclude that such processes are not major factors in the observed covariations between ¹⁸O and other geochemical and petrological indicators.

The difference in ¹⁸O between basaltic melt and a residual peridotitic mineral assemblage during mantle melting is expected to be a small positive number (0·2; Chiba et al., 1989; Matthews et al., 1998; Fig. 2), and thus even extensive fractional melting of peridotite will produce depleted residues that are only slightly lower in ¹⁸O than initial fertile sources (e.g. 20% fractional melting lowers ¹⁸O of the residue by only 0·04). This expectation is confirmed by analysis of olivine and pyroxene from highly depleted harzburgites of the Voykar massif, Polar Urals: these rocks are thought to be residual to 15% fractional melting (Sharma et al., 1997), yet they contain olivines that are indistinguishable in ¹⁸O from those in average fertile mantle peridotite (Eiler et al., 1996). Based on this and the predicted effects of melting on elemental abundances described above (i.e. residues will have lower TiO₂ and Yb/Sc, and their spinels will have higher cr-number), melting alone could therefore only produce correlations in Fig. 4 that are opposite in slope from those observed (and in any case so steep so as to appear vertical at the scale plotted). We therefore conclude that equilibrium partitioning of oxygen isotopes during simple batch or fractional fusion of peridotite cannot produce the trends in Fig. 4.

Given that fractionation of oxygen isotopes during melting or crystallization cannot explain the observed relationships between oxygen isotopes and sample chemistry, variations in ¹⁸O observed in this study are most simply interpreted in terms of mixing of two or more isotopically distinct reservoirs. High-level assimilation of oceanic crust and/or the island-arc volcanic edifice has been suggested as a way to introduce ¹⁸O-enriched material into mantle-derived magmas (Harmon et al., 1981; Hildreth & Moorbath, 1988; Davidson & Harmon, 1989). Although the effects of assimilation are widespread in arc-related environments, particularly in evolved magmas from arcs with thick crustal sections of continental rocks or sediments, it is difficult to explain by such a mechanism the relationships between ¹⁸O and sample chemistry observed in this study. In particular, this explanation would require preferential addition of crustal assimilants to lavas that are relatively high-degree mantle melts and/or are derived from relatively depleted mantle sources; moreover, such a process is not expected to generate the relationship between ¹⁸O and ⁸⁷Sr/⁸⁶Sr shown in Fig. 3a (i.e. assimilation of crustal rocks generally leads to more linear trends in these dimensions; Davidson & Harmon, 1989); finally, assimilation–fractional crystallization (AFC) processes generally produce the largest deviations in ¹⁸O in the most fractionated samples (DePaolo, 1981; Taylor, 1986) and therefore would produce trends generally opposite those observed in Fig. 4 (for reasons similar to those discussed above in reference to fractional crystallization). We conclude that although it would be possible to construct models of assimilation processes consistent with the trends in Figs 3 and 4, such models would contrast with the typical behavior of systems influenced by AFC processes.

Fluxed melting—a model for the interactions between slab-derived fluid or melt and the sub-arc mantle
A possible explanation for the correlations in Figs 3 and 4 is that peridotites in the sub-arc mantle span a restricted range in ¹⁸O (5·0–5·2 for olivine; 5·3–5·5 for the bulk peridotite), comparable with the average value of ¹⁸O elsewhere in the upper mantle (Ito et al., 1987; Mattey et al., 1994; Eiler et al., 1996), and that addition of a high-¹⁸O, slab-derived component is closely linked to enhanced melting of these sources. This explanation is attractive for several reasons:

1. as described in the Introduction, ‘slab’ signatures in the geochemistry of arc-related lavas are plausibly attributed to the transfer of aqueous fluids and/or hydrous silicate melts from the slab to the mantle wedge. Both of these metasomatic agents would probably be derived from sediments and/or hydrated basalts in the upper portions of the subducted crust because these regions either have high initial water contents or become water rich through low-temperature sea-floor alteration; as there are several high-¹⁸O components in the top half of the subducted crust, it is thus further expected that addition of either of these slab-derived metasomatic agents to mantle peridotite will raise its ¹⁸O.
   
2. Addition of water-rich components (whether aqueous fluids or hydrous silicate melts) decreases significantly the solidus temperature of peridotite at the pressures and temperatures of the sub-arc mantle (Kushiro et al., 1968; Hirose & Kawamoto, 1995; Gaetani & Grove, 1998), and thus metasomatism and ¹⁸O enrichment are expected a priori to be accompanied by enhanced melting of the mantle wedge.
   
3. An additional feature of this explanation is that it relates variations in ¹⁸O and chemical composition among our studied samples by a relatively simple common mechanism, and it is therefore suitable for examination with a detailed forward model. In this section we develop such a model to elaborate and test this hypothesis. The model is constructed with specific reference to variations in ¹⁸O, chemical composition, and Sr isotope ratio among lavas from the Vanuatu–Fiji–New Caledonia region, but we also examine its consistency with the presence or absence of high ¹⁸O lavas in other suites.
   

Before detailing this model, we note a possible alternative is that the sources of arc lavas are variably depleted before addition of a volatile-rich slab component and that the extent to which such a component is added is a function of the extent of prior depletion. This model reverses the causality of the model detailed in this discussion by requiring that the extent of depletion in the sub-arc mantle has some means of limiting or controlling the amount of slab-derived component that can be added to it. This might be possible if, for example, the permeability of depleted peridotite with respect to flow of slab-derived fluids or melts were higher than that of fertile sources (Toramaru & Fujii, 1984) or due to differences in the melt productivity of depleted vs fertile peridotites having equal abundances of slab-derived components. We have not developed a forward model for this hypothesis, but improved understanding of these phenomena may permit detailed examination of such models in the future.

General features of the model
Our model of the mantle sources of arc lavas relates quantitatively the abundance of slab-derived constituents in these sources to oxygen isotope ratios and other geochemical signatures of arc lavas and to the productivity of melting in the mantle wedge. We accomplish this by integrating what is known of the productivity of isobaric, isothermal, wet peridotite melting (Hirose & Kawamoto, 1995; Gaetani & Grove, 1998; Hirschmann et al., 1999), trace element partitioning between basaltic melts and residual peridotitic assemblages (see the Appendix), and approximations of the geochemical properties of slab-derived fluids and melts. We acknowledge at the outset that we have made a number of simplifying assumptions that must be elaborated upon and/or modified if this geochemical and petrological model is to be explicitly integrated with dynamical models of subduction zone processes.

We assume that pre-metasomatic sources in the mantle wedge consist of peridotite having concentrations of most major and minor elements based on the ‘mm3’ synthetic peridotitic mix of Baker & Stolper (1995), but with approximately half the Cr₂O₃ and somewhat less Al₂O₃ to be comparable with other estimates of the composition of relatively fertile upper-mantle peridotite. We also assume that these sources have abundances of selected minor and trace elements that will, on 5% batch melting (using partition coefficients detailed in the Appendix), produce a melt with a composition near the low-¹⁸O ends of the data arrays in Figs 3 and 4. This composition is listed in Table 2 and is broadly similar to (but somewhat more depleted than) the inferred sources of NMORB (Sun & McDonough, 1989; McKenzie & O’Nions, 1995; Hirschmann & Stolper, 1996). The initial value of ¹⁸O for olivine in this source is assumed to be 5·0 (approximately the low-¹⁸O end member to trends in Figs 3 and 4), implying a whole-rock value of 5·3 (Chiba et al., 1989; Mattey et al., 1994).

View this table: [in this window] [in a new window]   Table 2: Composition of components for fluxed melting models

 

The source is assumed to be held at a constant temperature and pressure (detailed below) while either a high-¹⁸O aqueous fluid or a water-rich silicate melt is added to it. We first describe the assumed composition of the model aqueous fluid; we estimate below the composition of the model hydrous silicate melt. The fluid is assumed to be 90 wt % H₂O, 6 wt % SiO₂, 2 wt % Na₂O, 1 wt % K₂O and 1 wt % Al₂O₃, and to contain no Ti, Cr, Yb, or Sc. This composition is based upon the solubility of major elements in water in equilibrium with silicate minerals at 15–20 kbar and 600–1100°C (Schneider & Eggler, 1986), plausible concentrations of Ti, Yb, and Sc in upper-crustal sediments and rocks in the slab (Albarède & Michard, 1989; Plank & Langmuir, 1998), the results of trace-element partitioning experiments between aqueous solutions and silicate minerals (Brenan et al., 1995; Keppler, 1996), and the assumption that Cr has a low concentration in the fluid phase. The abundances of other minor and trace elements in this model fluid will be considered in more detail in the section on concentration of solutes, but we note here that we do not assume that the composition of this fluid corresponds in these respects to any of the several previously proposed compositions of slab-derived components [e.g. the ‘slab fluid’ vs ‘slab melt’ of Elliott et al. (1997)].

We assume a ¹⁸O value of 20 for our model fluid, intermediate between observed values in siliceous glasses from arc xenoliths (Eiler et al., 1998; P. Schiano & J. M. Eiler, unpublished data, 1999) and the highest values expected in subducted sediments and volcanoclastic rocks in the upper kilometer of the subducted crust (Kolodny & Epstein, 1976; Arthur et al., 1983; Muehlenbachs, 1986; Plank & Langmuir, 1998). Oxygen isotope fractionation between water and silicate and carbonate minerals at temperatures higher than 500°C is negligible (2; O’Neil & Taylor, 1967) and has not been considered in our estimate. If a slab-derived fluid or melt moved through the mantle wedge by percolation such that it exchanged efficiently with surrounding peridotite, its oxygen isotope composition would come into equilibrium with the peridotite through reaction, solution–reprecipitation, and diffusive exchange. Therefore, implicit in our model is the assumption that slab-derived components can be introduced into the melting region without having fully exchanged with a ‘chromatographic column’ of normal mantle. Movement of fluids or melts by fracture propagation or through veins is a plausible physical mechanism for satisfying this condition.

A parallel set of calculations was made assuming that the fluxing agent is a water-rich silicate melt with the same ¹⁸O value (20) as the model aqueous fluid described above. The model melt composition is given in Table 2 and is based on experimental melts of red clay at 15–20 kbar and 900–950°C (Nichols et al., 1994; Johnson & Plank, 1999). We assumed that the model silicate melt contains 1 ppm Yb and 30 ppm Sc based on measurements of hydrous, siliceous glasses in mantle xenoliths from above subduction zones (Schiano et al., 1995).

Description of wet melting
If the starting peridotite composition is held at a temperature and pressure between its dry and wet solidi and outside the stability limits of hydrous minerals in a peridotitic assemblage, addition of the aqueous fluid or hydrous melt will flux melting. Under these conditions, the increment in the melt fraction (F) per increment of the mass fraction of fluid in the source (+X_fl; note that both F and X_fl are fractions relative to the total mass of the system, which changes with addition of fluid or melt) is a function of pressure, temperature, and the compositions of the fluid and peridotite. We used the MELTS algorithm (Ghiorso et al., 1994) to estimate F/X_fl (which we refer to as the ‘productivity’ for isobaric, isothermal, fluid-enhanced melting) for addition of the assumed aqueous fluid or hydrous melt to the model peridotite source. Using this algorithm, the fluid or melt phase was added to the initial solid source in increments of 0·1 wt % at a pressure of 10 kbar and a temperature of 1200°C; these conditions were chosen to be within the pressure range at which MELTS accurately describes the equation of state of water (10 kbar; Hirschmann et al., 1999) and 50°C below the MELTS-calculated dry solidus of the model peridotite at that pressure. We assume that the calculated productivity of wet melting under these conditions approximates the productivity under similar conditions at the higher pressures at which melting in the mantle wedge is believed to take place (e.g. at 20 kbar and 1300°C or 30 kbar and 1400°C; Plank & Langmuir, 1988). A calculation of this type at the more relevant, higher pressure conditions is precluded by the absence of data on the productivity of hydrous melting and the inaccuracy of the model for water used by MELTS under these conditions. In our view, a ‘homologous’ calculation at lower pressure of the sort we have performed is at this time the best and most quantitative compromise for modeling these processes.

MELTS calculations of productivity were made for two types of models of melt generation: (1) nearly fractional melting (i.e. extraction of each increment of melt generated after addition of each increment of 0·1 wt % fluid; bold curves in Fig. 5), and (2) batch fusion (i.e. the melt generated after each increment of fluid added to the source remains in the source in equilibrium with the residue; fine curves in Fig. 5). Calculations for three fluxing agents are shown in Fig. 5: results for addition of model aqueous fluid described above are shown as continous black curves; results for fluxing with the model silicate melt described above are shown as dashed black curves; and results for addition of pure water are shown as continuous gray curves for comparison with more compositionally complex fluxing agents. Figure 5a shows the melt fraction as a function of the amount of flux added to the source; Fig. 5b shows the melt fraction as a function of the amount of water added to the source by the fluxing agent.

View larger version (31K): [in this window] [in a new window]   Fig. 5. (a) MELTS-calculated relationships between the extent of melting (F, in wt %) of model peridotite (Table 2) and the amount of a slab-derived water-rich phase (Xfl, in wt %) added to that peridotite at a temperature of 1200°C (unless otherwise marked) and a pressure of 10 kbar. The temperature was chosen to be 50°C below the MELTS-calculated dry solidus for the peridotite composition at 10 kbar. The continuous black curves are for addition of a solute-rich aqueous fluid (Table 2); the dashed black curves are for addition of hydrous silicate melt (Table 2); and the continuous gray curves are for addition of pure water. Bold curves are for nearly fractional melting (i.e. removal of melt from the residue following addition of each small increment of slab phase); fine curves are for batch melting (retention of melt in equilibrium with the residue over full range in extent of melting). Small open circles indicate the conditions of clinopyroxene exhaustion from the residue. Truncation of a curve with a short perpendicular line indicates that further addition of fluid or melt does not drive further melting under the conditions of the calculation (i.e. because the wet solidus is higher than the assumed temperature). Heavy black curves marked ‘+50°C’ and ‘+100°C’ indicate models of fractional melting fluxed with model aqueous fluid at higher temperatures than assumed for all other calculations. The ‘+50°C’ calculation is approximately at the dry solidus of the initial peridotite; in the case of the ‘+100°C’ calculation, an increment of 5% batch melting takes place before addition of any fluid or hydrous melt. (b) reproduces the results in (a) but uses the water content of the system as the abscissa, illustrating that components of the fluxing agent other than water have little effect on the productivity of fluxed melting. The shaded field in (b) spans a range of estimated curves independently calculated based on hydrous melting experiments (Gaetani & Grove, 1998); filled circles in (b) illustrate results for wet batch melting experiments near the dry peridotite solidus (Hirose & Kawamoto, 1995), and are expected to be comparable with our ‘+50°C’ curves.

 

The models of productivity for both fractional and batch fusion fluxed by the model aqueous fluid yield a similar result: F/X_fl decreases steadily from 20 wt % melt per wt % fluid in the mixed source at the first introduction of fluid, to 15 wt % melt per wt % fluid at 15% melting and 10 wt % melt per wt % fluid at 30% melting. The model results for addition of pure water and the more complex aqueous fluid are not significantly different; this is well illustrated by the comparison in Fig. 5b, which normalizes out the difference in water content between these two fluxes (see below). Figure 5b shows that the MELTS model predicts a relationship between X_fl and F for pure H₂O similar to the results of Gaetani & Grove (1998) for a small range of temperatures near the dry solidus and less productive—by about a factor of two—compared with the results of Hirose & Kawamoto (1995) for melting slightly above the dry solidus (i.e. at relatively higher temperatures than assumed for most of our model calculations) fluxed by low total amounts of water; the comparison with experiments suggests that, in agreement with the analysis of Hirschmann et al. (1999), the MELTS algorithms provide a reasonable quantitative basis for modeling the productivity of hydrous melting. It should be noted that several previous treatments of the productivity of fluid-enhanced melting have considered the effects of flux addition well above the temperature of the dry solidus, and therefore produced higher estimates of the productivity (F/X_fl) (Stolper & Newman, 1994; Hirose & Kawamoto, 1995; Gaetani & Grove, 1998; Hirschmann et al., 1999). Calculated curves for fractional fusion on addition of the model aqueous fluid are shown at 50 and 100°C higher temperatures (i.e. at the dry solidus and 50°C above the dry solidus; labeled continuous lines, Fig. 5a) and to illustrate the sensitivity of productivity to temperature according to the MELTS model.

Figure 5a shows that the calculated productivity of melting fluxed by hydrous silicate melt is substantially lower on a weight-fraction basis than is that for aqueous fluid or pure water. This is a consequence of the lower concentration of water (10 wt %; Table 2) in the model silicate melt relative to the model aqueous fluid. Our calculations suggest that, regardless of the metasomatic agent, it is primarily the water content of the system that dictates the productivity of fluxed melting (although other incompatible elements have similar but less extreme effects; Kushiro, 1975; Hirschmann et al., 1999). This is illustrated in Fig. 5b, in which X_fl (i.e. the mass fraction of the fluxing agent in the system, whether it is pure water, solute-rich aqueous fluid, or hydrous melt) has been replaced on the abscissa by X_H2O, the mass fraction of water in the system. In these coordinates, melting fluxed by all the water-rich phases we have considered has a comparable productivity of 12–16 wt % melt per wt % H₂O added to the system. This is important for understanding our efforts described in the next section in this discussion to discriminate aqueous fluid from silicate melt as the metasomatic agent that could have produced the trends in Figs 3 and 4. Figure 5 also shows that according to the MELTS calculations, variations in the composition of the metasomatic phase or the mode of melting (i.e. batch vs fractional) have only small effects on the productivity of fluxed melting and on the extent of melting at which clinopyroxene is exhausted from the residue (marked by circles in Fig. 5a and b). In addition, it should be noted that fractional melting fluxed by pure H₂O (or by aqueous fluid for the higher temperature calculations) does not proceed beyond 22–24 wt % melting under the conditions of our calculations because the temperature of the wet solidus rises to higher than the assumed temperature, reflecting the highly refractory nature of residues of pure fractional melting that have not been sufficiently refertilized by Na and Al in the fluxing agent.

Element partitioning during melting
To maintain consistency with the relationships between F and X_fl shown in Fig. 5, we also used MELTS to calculate the phase proportions of the residue as a function of progressive fluid- or melt-fluxed melting. Again, we carried out these calculations for batch fusion and for incremental fractional fusion. Batch and incremental fractional fusion fluxed by the model aqueous fluid or by pure H₂O yielded melting reactions up to the exhaustion of clinopyroxene (at 25–30% melting) approximated by the reaction (given in units of mass)

The melting reaction for fluxing by hydrous silicate melt differs substantially from this:

The partitioning of minor and trace elements among coexisting phases was calculated for each melting increment using the phase proportions output by the MELTS calculations plus solid–melt partition coefficients. For each element, the partition coefficient between melt and each solid phase was taken from published experimental studies (referenced in the Appendix) and was assumed to be constant. Residue–melt partition coefficients were then calculated after each increment of melt production based on the phase proportions in the residue. The cr-number of chromite was described using the observed relationships among cr-number, F, and the Al₂O₃ content of the residue from the wet peridotite melting experiments of Hirose & Kawamoto (1995; see the Appendix). The MELTS calculations produced similar trends of cr-number vs F, but these results were not used because MELTS does not allow for Cr partitioning into pyroxenes.

Model algorithm
Given the assumed compositions of the metasomatizing aqueous fluid or hydrous melt and of the peridotite source (Table 2), the productivity and stoichiometry of fluxed melting output by MELTS [illustrated in Fig. 5a and approximated by the melting reactions given by equations (1) and (2)], and values for the solid–melt partition coefficients (see the Appendix), modeling of fluxed melting of the mantle wedge is straightforward. The bulk composition of the source (including its isotopic composition) is recalculated after the addition of every increment of fluid to the peridotite source. The phase proportions of this modified bulk composition are then calculated using the MELTS algorithm, and the trace elements are partitioned between the melt and solid phases according to the phase proportions and the assumed constant mineral–melt partition coefficients. Oxygen isotopes are assumed not to be fractionated between the residue and melt (i.e. we argued above that such effects are insignificant), and the ¹⁸O of the source increases nearly linearly with X_fl. For the batch melting case, the calculation is made for a range of amounts of fluid added to the initial source; for the fractional fusion calculation, the bulk composition of the source at each step is taken to be that of the solid residue before the next increment of fluid is added (although F and X_fl are, as indicated above, normalized to the total mass of the system, including the removed melt).

Results
The results of these calculations are compared with our data in Fig. 6. In Fig. 6b, we compare the output of our model with the trend plotted in Fig. 4b rather than with that in Fig. 4a because TiO_2(8·0) is more resistant to the effects of crystallization differentiation than is raw TiO₂ content. Each figure contains fine and bold curves representing the batch and fractional melting models. Faint symbols reproduce the data from Fig. 4. Continuous curves illustrate the results of our model for addition of aqueous fluid; dashed curves are based on a model of melting fluxed by hydrous melt. It is clear that the forward models of fluxed melting driven by aqueous fluid capture the overall trends of the correlations we observed and do so simultaneously for correlations among ¹⁸O, crystal chemistry (cr-number), major element composition (TiO_2(8·0)), and trace element ratios (Yb/Sc). In contrast, models of fluxed melting driven by our model silicate melt fail to describe, or even approach, the observed trends; the possible significance of this discrepancy will be discussed in the following section.

View larger version (22K): [in this window] [in a new window]   Fig. 6. Comparison of data from this study (gray symbols, reproduced from Fig. 4b–d) with calculated trends for fluxed batch fusion (fine continuous lines) and fluxed fractional fusion (bold continuous lines) models assuming an aqueous fluid as the flux; fine and bold dashed lines show the equivalent models assuming a hydrous melt as the flux. The batch and fractional models of slab-melt-fluxed melting (dashed fine and bold curves) overlap in (c). All calculations assume an initial peridotite source given in Table 2. Tick-marks indicate the total amounts of fluid or melt introduced to the source in 1% increments. (See text for further details.) It should be noted that in (a) the calculated relationship between cr-number and F is based on the observed relationships among cr-number, F, and bulk Al content in the residue during wet peridotite melting (Hirose & Kawamoto, 1995). For (b), the calculated TiO2 contents of model melts were normalized to TiO2(8·0) values by calculating the low-pressure differentiation of model melts using the MELTS algorithm.

 

Fractional melting calculations of fluid-fluxed melting are particularly successful at matching the data trends, whereas the batch melting calculations of fluid-fluxed melting produce less depleted lavas (i.e. higher TiO₂ and Yb/Sc) for a given amount of addition of aqueous fluid. In detail, the fractional fusion model predicts: (1) a gently curved trend of increasing ¹⁸O with decreasing TiO_2(8·0); (2) a four-fold rise in the cr-number of residual chromite with a relatively subtle rise in ¹⁸O (i.e. remaining within the range of ‘normal’ mantle peridotite), followed by a sharper rise in ¹⁸O at nearly constant cr-number; (3) a slightly curved, almost linear, trend of increasing ¹⁸O with decreasing Yb/Sc. The agreement between model and data is particularly good for the plot of Yb/Sc ratios vs ¹⁸O (Fig. 6c). As pointed out previously, the alkali basalts from the TLTF volcanoes fall off the trend defined by the rest of the data in Fig. 4b, and thus they also deviate from the model calculations in Fig. 6b. Given that these samples are not members of the Vanuatu–Fiji–New Caledonia suite for which our model was constructed, we do not try to explain this deviation in detail; however, as noted above, these lavas are the only alkali basalts considered in this study, and elevated TiO₂ contents are a characteristic of alkalic lavas that may be due to differences in the initial fertility and/or mineralogy of the source and/or their fractionation history.

Quantitative comparison of the data with the model calculations suggests that most of the studied lavas from the Vanuatu–Fiji–New Caledonia suite could be derived from portions of the mantle fluxed by up to 1·0 wt % of aqueous fluid from the slab, but that some lavas (boninites, low-TiO₂ shoshonites, and low-TiO₂ island-arc basalts) require fluxing of the mantle by up to 2·5 wt % of such a fluid. Formation of the back-arc-basin basalt from the Hunter ridge (Vanuatu sample 123/3) requires fluxing by no more than 0·25 wt % of this fluid based on the low integrated extent of melting of its source and its ¹⁸O value; however, we note that the sources of back-arc magmas could be initially less depleted than those of arc lavas, which may typically have been previously depleted by the extraction of back-arc magmas (Woodhead et al., 1993; Stolper & Newman, 1994), so the two groups of magmas may not be simply related by a continuous model of the sort we have developed here. Inspection of Fig. 6 demonstrates that samples other than members of the Vanuatu–Fiji–New Caledonia suite are also consistent with our model curves and therefore could reflect similar processes. In particular, the alkaline TLTF (Papua New Guinea) suite conforms to the model trend in Fig. 6c and the less ‘depleted’ chemical compositions and low ¹⁸O values in Mariana and South Sandwich arc lavas are consistent with a lower abundance of a slab-derived fluxing agent in their sources than is present in higher-¹⁸O, more depleted lavas from the Vanuatu–Fiju–New Caledonia suite. More confident generalization of our model will require discovery of new suites that vary substantially in both ¹⁸O and apparent integrated extent of melting.

Slab melt or slab fluid?
Comparison of our observations with model curves for fluxing with hydrous silicate melt (dashed curves, Fig. 6) suggests that addition of a slab-derived melt to the mantle wedge cannot reproduce the trends of our observations. In particular, melt-fluxed models generate curves that, although similar in shape to those generated by fluid-fluxed models, are too shallow to explain the data. To first order, this simply reflects the facts that water is the major fluxing agent (Fig. 5b) and that the concentration of water is much higher in the aqueous fluid than in the hydrous melt. These calculated trends also reflect the high concentrations of TiO₂ and Al₂O₃ in the model silicate melt when compared with the model aqueous fluid (i.e. the model melt actually enriches the source in these elements during fluxing). An important limitation of the melt-fluxed model as we have developed it is that given the TiO₂, Al₂O₃, and Yb contents of the model silicate melt (Table 2), the low extents of melting per increment of fluxing (Fig. 5a), and the partition coefficients for these elements during peridotite melting (given in the Appendix), melting fluxed by sediment-derived melt cannot produce lavas at the extremes of the observed trends (TiO₂ < 0·8 wt %; cr-number > 0·3, Yb/Sc < 0·05).

However, it is important to note that our model contains many variables (e.g. the composition of model hydrous silicate melt, the conditions under which it interacts with the mantle wedge, etc.), changes in which might permit fluxing of the mantle wedge by hydrous melt to produce trends such as those in Fig. 4. We have explored a number of such alternative models; although full description of these alternatives would require a digression from our discussion, the following are examples of conditions under which models similar to those described above can match our results using slab-derived melt as a fluxing agent. For example, we can match the observed trends if the metasomatizing hydrous melt is similar to model aqueous fluid in having near-zero abundances of Al, Ti, and Yb and if it has a relatively low ¹⁸O value of 10 (similar in ¹⁸O and Ti content to hydrous silicate glass in mantle nodules from Lihir island, but much lower in Al and Yb content than them; Eiler et al., 1998). Alternatively, we can generate successful models if the slab-derived melt is water rich (30 wt % H₂O), low in ¹⁸O, and intermediate between our standard model melt and the model aqueous fluid (Table 2) in Al, Ti, and Yb. Finally, successful models can result if slab-derived melt is the fluxing agent provided that it is poor in Al, Ti, and Yb and interacts with the mantle wedge at temperatures on the order of 100°C higher than the dry peridotite solidus. An important characteristic of all these successful models is that they require addition of up to 20 wt % of the metasomatizing melt to the mantle wedge to generate the observed trends; this simply reflects the need to bring in sufficient H₂O to generate the required levels of depletion in the source. Although not a priori unphysical, the requirement of such large amounts of slab-derived melt [almost two-thirds of the total amount of melt produced by the melting reaction; see equation (2)] appears to be a necessary feature of successful models of melt-fluxed melting and thus may lead to tests of the plausibility of this process. Overall, although we can develop successful melt-fluxed models, all require choices of slab melt properties (e.g. low Al, Ti, Yb, and ¹⁸O) and/or temperatures in the mantle wedge and quantities of flux that are restrictive and not obviously compatible with experimental constraints on the compositions of slab-derived melts (Nichols et al., 1994; Johnson & Plank, 1999) or with thermal models of the sub-arc mantle (Peacock, 1996). Nevertheless, although we do not favor these versions of our model, we emphasize that neither can we rule them out.

The overall results of our modeling are that fluxing of a relatively cool mantle wedge by solute-rich aqueous fluids describes the characteristics of high-¹⁸O, arc-related lavas without the need for adjustments (i.e. beyond the assumed constraints taken from previous studies cited above and the general characteristics assumed for our model), but that the relationship between indices of extent of melting and ¹⁸O can only be described by fluxing the mantle wedge with hydrous slab-derived melts if large adjustments are made to our best estimates for the properties of such melts and/or to the conditions of their interaction with the mantle wedge. Our preference, based on this modeling, is either that slab-derived melts are absent from the mantle sources of the lavas we have studied or, more plausibly, that they are present in such low abundance that they do not contribute significantly either to the bulk ¹⁸O of melt-generating regions in the sub-arc mantle or to the productivity of melting. This may appear to be inconsistent with recent studies that have concluded that sediment-derived melts (rather than sediment- or basalt-derived hydrous fluids) dominate aspects of the trace element budget of some arc lavas (e.g. Hawkesworth et al., 1997). However, the results of our modeling can be reconciled with trace-element evidence for the importance of slab-derived melts in metasomatism of the mantle wedge if such melts are a minor component of a fluid-dominated slab enrichment of the sub-arc mantle but are sufficiently rich in certain incompatible trace elements that they nevertheless dominate the budgets of those elements.

Concentrations of solutes
To this point, most of the details of the trace-element chemistry of slab-derived metasomatic fluids and/or melts have been unspecified and relatively unimportant to our discussion; that is, the metasomatizing phase has been assumed to be water rich, high in ¹⁸O, and poor in Ti, Cr, HREE, and Sc, and this allowed us to construct models that match our observations. A key feature of this modeling is that it provides estimates of the amounts of slab-derived fluid or melt required in the source regions of studied lavas from the Vanuatu–Fiji–New Caledonia region. These estimates in turn allow us to solve for the abundances of other trace elements in the slab-derived metasomatizing phase that, in the context of our models, are required explain their concentrations in these lavas. In this section, we assume that the metasomatizing phase is the aqueous fluid described previously and solve for its concentrations of a set of minor and trace elements. We emphasize that these calculations are simply logical consequences of the modeling described above and are not framed with the intent of relating our results to recent models of multicomponent slab additions to the mantle wedge (e.g. Elliott et al., 1997).

These calculations are straightforward: a sample’s ¹⁸O value can be used to constrain in the context of our model the total amount of fluid added to its source; that amount of fluid is taken to have driven a total extent of melting given by the batch or fractional curve for solute-rich aqueous fluid in Fig. 5a; partitioning of each element between residue and melt for each increment of melting is calculated using the phase proportions from MELTS, the chosen set of partition coefficients (see the Appendix), and mass balance; and finally, a series of forward calculations are made in which the concentration of each element in the fluid is varied until a match is obtained with the measured abundance in the lava (see the Appendix). The initial mantle source was assumed to have characteristics given in Table 2, as explained above. It should be noted that we solve for only a single slab-derived component for each sample, even though there is evidence for at least two such components in some arcs (Elliott et al., 1997; Fig. 3e and f); if more than one slab-derived component contributes to the trace element budget of the sub-arc mantle, our results can be viewed as a sum of the contributions of all such components divided by the abundance of the aqueous fluid component (i.e. our calculation would systematically overestimate the abundances in the fluid of elements highly concentrated in other components such as slab-derived melts).

We have solved for the required trace element composition of the model aqueous fluid for several samples from the Vanuatu–Fiji–New Caledonia suite. These samples include two island-arc basalts from Vanuatu (one a relatively high-¹⁸O lava from Ambrym, Ambrym 28; the other a lower-¹⁸O lava from Tongoa, Tongoa-6), one shoshonite from Fiji (E2c), and one boninite from New Caledonia (NC-52). For comparison, we also computed trace element contents of model aqueous flux for an island-arc basalt from Uracas in the Mariana arc (URA-5), although we emphasize that this sample is not part of the Vanuatu–Fiji–New Caledonia suite and thus may not be directly comparable with the fluxed melting model examined in detail in the preceding discussion. The results of these calculations are presented in Fig. 7 as NMORB-normalized compositions of slab-derived fluids calculated for each sample; fluxed-batch-fusion results are presented in Fig. 7a, and fluxed-fractional-fusion results are shown in Fig. 7b. For comparison, we have plotted in Fig. 7c the compositions of the average ‘slab’ component of McCulloch & Gamble (1991), the ‘water-rich component’ for the Mariana back-arc from Stolper & Newman (1994), the expected compositions of slab-derived fluids calculated by Brenan et al. (1995) based on experimental study of partitioning between solids and fluids and estimated compositions of subducted sediments and altered basalts, and the expected composition of sediment-derived melt and sediment-derived fluid based on the experiments of Johnson & Plank (1999) and the average composition of sediments subducting into the western Vanuatu trench (Peate et al., 1997; Plank & Langmuir, 1998). It should be noted that absolute concentrations of trace elements in some or all of the samples we have considered may have been modified by fractional crystallization; this could contribute systematic offsets (i.e. shifting the entire pattern up or down in Fig. 7) of tens of percent in the estimated composition of the slab-derived fluid based on any given sample. It also should be noted that to satisfy the concentrations of some HREE of some lavas, calculated concentrations of these elements in the aqueous fluid are required to be negligible; in these cases, reflecting the insensitivity of our model to the slab contribution for these elements, no results are shown in Fig. 7.

View larger version (67K): [in this window] [in a new window]   Fig. 7. Calculated concentrations of trace elements in slab-derived aqueous fluid required to satisfy simultaneously the restrictions of fluxed-batch- and fluxed-fractional-fusion models (Fig. 6) and the trace element chemistry of selected lavas considered in this study. (a) Results for fluxed-batch-fusion model; (b) results for fluxed-fractional-fusion model; (c) comparison of ranges of fluid compositions calculated for fluxed-batch- and fluxed-fractional-fusion models with previous estimates of the composition of slab-derived metasomatic components (McCulloch & Gamble, 1991; Stolper & Newman, 1994) or calculated slab-derived melts or fluids based on partition coefficients from Brenan et al. (1995) or Johnson & Plank (1999) and assuming equilibrium of an infinitesimal amount of fluid or melt with sediment subducting beneath the Vanuatu arc (Peate et al., 1997). Error bars associated with the estimates of Stolper & Newman [open circles in (c)] are approximately equal to the size of the symbol. (See text and the Appendix for details of the calculations and sources of elemental data on lavas.)

 

The results of our calculations show that the aqueous fluids required by our model are characterized by enrichments in minor and trace elements relative to NMORB that are typically higher for more incompatible elements (i.e. for elements toward the left side of Fig. 7). Superimposed on this general trend are positive anomalies in Rb, Ba (except for the boninite sample), K, and Sr (again nearly absent in the boninite), and a negative anomaly in Nb. These features are typical of arc-related lavas and have been previously inferred to reflect high solubilities of alkali- and alkali-earth elements and low solubilities of HFSE in aqueous fluids (Gill, 1981; McCulloch & Gamble, 1991).

The compositions of fluids calculated assuming a fluxed-fractional-fusion model (Fig. 7b) are similar to one another in their relative abundances of most elements (results for fluxed-batch fusion are only somewhat more scattered). This is particularly true for the four non-boninitic samples. This relatively small range in trace element concentrations of the model fluids, when combined with the success of the model in reproducing the trends in Fig. 4, demonstrates that much of the geochemistry of this group of arc-related lavas can be explained by fluxing the sub-arc mantle with fluids having a small range in oxygen isotope ratios and trace element compositions. The mean composition of these calculated fluids (including the boninite, NC-52) is presented in Table 2 under the heading ‘Vanuatu model aqueous fluid’. The fluids calculated assuming a fluxed-batch-fusion model are broadly similar to this composition, the only significant differences being in Y and the middle REE to HREE (see comparison in Fig. 7c).

The calculated trace element composition of the aqueous fluid is similar to that of the H₂O-rich component previously identified as a source of geochemical variability in the source regions of Mariana back-arc lavas [Fig. 7c; note, however, that the H₂O-rich component of Stolper & Newman (1994) is richer in Na and poorer in H₂O than assumed here]. Despite similarities between the models used in this study and by Stolper & Newman (1994) (i.e. both link the abundance of a water-rich component in the source to the extent of melting of that source), the estimates of trace element concentrations in the water-rich components of these two studies are independent, so this is not a necessary or circular result and is therefore encouraging with respect to the robustness of the results of both studies. The success of the fluxed melting model in both cases and the correspondence between the inferred trace-element compositions of the water-rich components in the two studies suggest a similarity between the hydrous component in the Mariana back-arc mantle and the sub-arc mantle sources of several types of lavas from the Vanuatu–Fiji–New Caledonia region. This correspondence is surprising given the expectation that there is not a single, universal slab-derived component (e.g. its composition is expected to vary among different slab-derived components and with the age and sediment stratigraphy of the subducted plate; Elliott et al., 1997; Plank & Langmuir, 1998). It should be emphasized, however, that the patterns compared in Fig. 7 are shown on a logarithmic scale, and there are, in detail, significant differences among the fluid compositions (e.g. the Rb/Ba ratio varies by a factor of 12 among the calculated fluid compositions for different samples shown in Fig. 7).

Although the calculated fluid compositions from this study are similar in terms of trace elements to that inferred by Stolper & Newman (1994), they differ from some other estimates of the composition of such aqueous fluids. For example, the calculated concentrations of Th and U in the slab-derived fluid from this study (for either the batch or fractional melting models) are significantly higher than predicted using known aqueous-fluid–mineral partition coefficients and the estimated composition of typical subducted crust (Brenan et al., 1995). This may be due to natural slab-derived fluids having more ‘melt-like’ properties with respect to trace element partitioning than do aqueous fluids that have been investigated experimentally (e.g. Shen & Keppler, 1997). Alternatively, the slab-derived components that produced the correlations in Fig. 4 may in fact include contributions from silicate melts that are minor in absolute abundance compared with aqueous fluids but still dominate the budgets of U and Th; such a view of metasomatism of the mantle wedge dominated by aqueous fluids but with contributions from melt has been presented previously (e.g. Moriguti & Nakamura, 1998). Another difference between our results and previous work is that the composition of the slab-derived component given by McCulloch & Gamble (1991), although overlapping for most elements with the range of estimates from this study, is consistently near the low end of the range given here (and some elements, such as Rb, Th, U, and Zr, are more than two times lower than the low end of our range; Fig. 7c). These differences probably reflect a number of factors, but an important one is that McCulloch & Gamble selected a range of 1–5% for the abundance of the slab-derived component in the sub-arc mantle. Our results suggest that the average value is closer to 0·5–1·0 wt %. Modification of the McCulloch & Gamble model to account for this difference would result in concentrations of most elements in the slab-derived fluid that more closely approach our average estimates. It should be noted that our model slab-derived fluid (and the estimates of McCulloch & Gamble and Stolper & Newman) has higher concentrations of nearly all but the most compatible of trace elements shown in Fig. 7 than is predicted by combining recent experimental results on trace element partitioning for melting or dehydration of sediments (Johnson & Plank, 1999) with the known compositions of sediments subducting beneath the western Vanuatu arc (Peate et al., 1997; filled symbols, Fig. 7c).

The isotopic composition of Sr provides a consistency check on our estimated concentration of Sr in the slab-derived fluid; i.e. if the concentration of Sr arrived at through the preceding exercise is correct, then given reasonable values for the Sr isotope ratios of the slab and mantle, our model should predict the observed relationship between ¹⁸O and ⁸⁷Sr/⁸⁶Sr (Fig. 3a). We have made such a prediction for the suite of island-arc lavas from western Vanuatu (i.e. that subset of Vanuatu lavas erupted adjacent to a portion of the adjacent trench having a well-characterized sedimentary section; Peate et al., 1997). Figure 8 compares our data (black filled circles and triangle for Vanuatu data; gray symbols show data from other locations for comparison) with the trends calculated by our fluid-fluxed melting model (i.e. the mantle and fluid have Sr concentrations listed in Table 2; fluid-fluxed melting follows our model as described above; and the ¹⁸O of the fluid is as previously assumed for the calculations presented in Fig. 6). We assumed that the mantle wedge initially has an ⁸⁷Sr/⁸⁶Sr ratio of 0·7027 (comparable with the NMORB source; Ito et al., 1987), that the initial ¹⁸O of olivine in the source is 5·1 (i.e. slightly higher than the value of 5·0 in our general model, so that the predicted curve passes through the low-⁸⁷Sr/⁸⁶Sr end of the subset of the data being considered), and that the fluid has an ⁸⁷Sr/⁸⁶Sr ratio between 0·7044 (for altered volcanic rocks) and 0·7072 (for pelagic sediment; both taken from measured values in samples of the slab subducting beneath the western Vanuatu arc; Peate et al., 1997). The resulting calculated curves show that our model slab-derived fluid predicts a strongly hyperbolic covariation of ¹⁸O and ⁸⁷Sr/⁸⁶Sr, consistent with the observed trend. The value of ⁸⁷Sr/⁸⁶Sr at which ¹⁸O values rise sharply is substantially lower than the value of pelagic sediment subducting beneath Vanuatu, requiring in the context of our model that a significant fraction of the Sr in the slab phase is derived from altered volcanic material.

View larger version (21K): [in this window] [in a new window]   Fig. 8. Comparison of data from Vanuatu (black symbols; gray symbols reproduce other data from Fig. 3a) with calculated trends for our fluxed fusion model assuming that the flux is an aqueous fluid having an Sr concentration equal to the average of the results shown in Fig. 7b (4572 ppm; see Table 2). Two calculated curves are shown representing the extremes in the isotopic composition expected for Sr released from the top of the slab subducting beneath western Vanuatu (Pearce et al., 1998): one assuming that the fluid acquires its Sr from altered volcanoclastic sediments (87Sr/86Sr = 0·7044), the other assuming that Sr is from pelagic sediment (87Sr/86Sr = 0·7072). Batch and fractional models coincide to within the width of the plotted lines. In combination with calculations presented in Figs 6 and 7, this calculation is consistent with the source of Vanuatu lavas being a peridotite similar to the sources of NMORB that has been fluxed by 0–2 wt % of an aqueous fluid that is high in 18O, highly enriched in Sr compared with peridotite in the mantle wedge, and in which the Sr is derived principally from altered volcanoclastic sediment (or perhaps from the underlying altered basalt of the oceanic plate, which probably has a similar 87Sr/86Sr ratio; Staudigel et al., 1995).

 

	SUMMARY AND CONCLUSIONS

TOP ABSTRACT INTRODUCTION SAMPLES AND ANALYTICAL METHODS RESULTS DISCUSSION SUMMARY AND CONCLUSIONS APPENDIX REFERENCES

 
The key observations in this study are the low abundance (1%) of crustal oxygen in most (80%) oceanic arc-related lavas and the relationship between those ¹⁸O enrichments that do exist and the integrated extent of melting in the sources of arc-related lavas. This second observation suggests that the extent of melting in the sources of the studied lavas is related to the supply of ¹⁸O-enriched, slab-derived oxygen. In our view, the most plausible mechanism for such a relationship is that slab-derived oxygen is introduced as a metasomatic agent into the peridotitic mantle wedge in amounts up to a few percent in the form of a slab-derived aqueous fluid, the abundance or availability of which controls the progress of peridotite melting. This explanation of depleted chemical signatures in arc lavas differs from the widely held view that they reflect a template of variable depletion related to melting in the back-arc environment onto which slab signatures are superimposed in the sub-arc mantle (e.g. Woodhead et al., 1993). The forward model we propose to describe this process can explain quantitatively observed variations in chemistry and ¹⁸O of lavas from the Vanuatu–Fiji–New Caledonia region, which display the largest range in ¹⁸O and apparent source fertility of any single suite of related samples examined in this study. We cannot rule out the possibility that this process is driven by hydrous slab melts, but this would require addition of large amounts of slab-derived melt to the mantle wedge, unexpected chemical and isotopic compositions for the hydrous melts, and/or unexpectedly high temperatures in the metasomatized sub-arc mantle wedge (i.e. well above the dry peridotite solidus).

Our interpretation is similar in several of its details to conclusions previously reached through study of the abundances of major and minor elements and water in back-arc-basin and arc basalts and more generally to the widely held notion that slab-derived fluids contribute to both the extent of melting and the chemistry of subduction zone lavas. In particular, Stolper & Newman (1994) and Gribble et al. (1998) observed a relationship between the water contents of back-arc-basin basalts in the Mariana trough and the inferred extent of melting of the sources of those basalts. They concluded that melting in those sources was fluxed by addition to a peridotite similar to the NMORB source of small amounts (typically up to 0·5 wt %—just below the limit at which an oxygen isotope signal is expected to be seen by our measurements) of a water-rich component having trace element contents similar to those we infer for the agent of fluxed melting in the Vanuatu–Fiji–New Caledonia sub-arc mantle. Crawford et al. (1989) proposed that the low CaO/Al₂O₃ ratios and Cr-rich spinels characteristic of boninites reflect derivation from a depleted peridotitic source in the sub-arc mantle that has been melted as a result of fluxing by 2–3 wt % water. Stern et al. (1991) similarly proposed that many aspects of boninite chemistry can be explained by melting a water-rich, highly depleted mantle source. Our results for lavas from the Vanuatu–Fiji–New Caledonia region agree quantitatively with these estimates and serve to link the petrogenesis of these diverse arc-related lavas through a common mechanism. Comparison of our model with samples other than the Vanuatu–Fiji–New Caledonia suite and similar, but more subtle, observations of a relationship between ¹⁸O and Na₂O_(8·0) in a previous study of back-arc lavas from the Lau basin (Macpherson & Mattey, 1997) suggest that this mechanism may be consistent with a wider spectrum of arc-related magmas. However, more specific generalizations must await the discovery of more arc-related suites demonstrating a large range in ¹⁸O and chemistry.

	APPENDIX

TOP ABSTRACT INTRODUCTION SAMPLES AND ANALYTICAL METHODS RESULTS DISCUSSION SUMMARY AND CONCLUSIONS APPENDIX REFERENCES

 
Calculated concentrations of trace- and minor-element solutes in the slab-derived fluid were made as follows. The concentration of a given element in a lava produced by fluxed batch fusion is described by the equation

where X_fl is based on the ¹⁸O of olivine in the lava, the estimated ¹⁸O of our slab and mantle end members, and the contrast in oxygen concentration between peridotite and water. Values of F are based on estimates of X_fl and the MELTS-derived relationship between F and X_fl (Fig. 5a). C_s is the initial concentration of the element of interest in peridotite before any addition of a slab phase (Table 2), and C_l is the concentration of the element of interest measured in the lava. Equation (A1) is for batch melting, modified for addition of the aqueous slab component to the initial source. Solid–melt distribution coefficients were taken from several sources (Table A1). The bulk partition coefficient is based on the mineralogy of the residue as generated by our MELTS calculations. Fractional melting calculations were made similarly, but the process was broken into steps in which a small amount of slab fluid (X_fl = 0·0005) was added to the initial peridotite source and a small extent of melting (F = 0·005–0·01) proceeded as modeled by MELTS (Fig. 5a). The increment of melt was then removed and the residue used as the initial source for the next step. This calculation closely approaches the analytical solution for perfect fractional fusion for elements not supplied by the slab-derived fluid phase.

View this table: [in this window] [in a new window]   Table A1: Mineral–melt distribution coefficients used in this study

 

The calculation of the cr-number of chromite was treated somewhat differently. The experiments of Hirose & Kawamoto (1995) were used to constrain the dependence of the distribution coefficient for Al on the extent of melting and the relationship between the cr-number of the residue and the cr-number of chromite in the residue. Cr was assumed to be a compatible element (D_bulk = 10; Dick & Bullen, 1984). Results of the MELTS algorithm include calculated cr-numbers of residual spinel and agree qualitatively with the results of this treatment, but are systematically offset from the plotted model curves in Fig. 6a, reflecting the absence of Cr in pyroxene in the MELTS model.

	ACKNOWLEDGEMENTS

 
We thank Bob Stern for his gracious supply of samples and unpublished data for shoshonites from the Mariana arc. Glen Gaetani, Mike Baker, Ronit Kessel, Peter Wyllie, Bob Stern, and Terry Plank improved upon this work through informal reviews of an early draft of this manuscript and/or helpful discussions on subjects related to this work. We thank Jean Morrison for allowing us to use her laboratory facilities at USC, and Mike Spicuzza and Nami Kitchen for laboratory assistance at the UW and at USC, respectively. A portion of the data presented in this manuscript was collected by Ronit Kessel as part of her graduate studies at Caltech. We gratefully acknowledge the improvements to this manuscript resulting from reviews by Chris Hawkesworth, Peter Kelemen, and Colin Macpherson. This work was supported by NSF Grant EAR-9805101. This is Contribution 8556 of the Division of Geological and Planetary Sciences at the California Institute of Technology.

	FOOTNOTES

 
^*Corresponding author.

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