Mineral-aqueous fluid partitioning of trace elements at 900–1200°C and 3.0–5.7 GPa: new experimental data for garnet, clinopyroxene, and rutile, and implications for mantle metasomatism

Geochimica et Cosmochimica Acta, Vol. 62, No. 10, pp. 1781–1801, 1998 Copyright © 1998 Elsevier Science Ltd Printed in the USA. All rights reserved 0016-7037/98 $19.00 1 .00

Pergamon

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Mineral-aqueous fluid partitioning of trace elements at 900 –1200°C and 3.0 –5.7 GPa: New experimental data for garnet, clinopyroxene, and rutile, and implications for mantle metasomatism R. STALDER,1,*,† S. F. FOLEY,2 G. P. BREY,1 and I. HORN3,‡ 1

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Institut fu¨r Mineralogie, Universita¨t Frankfurt, Senckenberganlage 28, 60054 Frankfurt/Main, Germany Mineralogisch-Petrologisches Institut, Universita¨t Go¨ttingen, Goldschmidtstrabe 1, 37077 Go¨ttingen, Germany Dept. Earth Sciences, Memorial University of Newfoundland, St. John’s, Newfoundland A1B 3X5, Canada

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(Received July 8, 1997; accepted in revised form February 4, 1998)

Abstract—In order to constrain the role of fluid phases during metasomatic processes in the upper mantle, trace element partition coefficients for Ba, Sr, Pb, Nb, Ta, Zr, Hf, Ti, La, Ce, Sm, Tb, and Yb between aqueous fluids and eclogite assemblage minerals (garnet, clinopyroxene, and rutile) have been determined experimentally at 900 –1200°C and 3.0 –5.7 GPa. Using a new experimental technique in which diamond aggregates are added to the experimental capsule set-up, the fluid was separated from the solid residue so that both quenched solute and residual minerals could be analysed directly. Trace element concentrations were determined in situ by laser ablation microprobe (LAM). The partitioning behaviour is controlled by temperature, pressure, and crystal chemistry; whereas fluid composition is not as crucial. Neither addition of hydrochloric acid nor high silica concentrations in the fluid have strong effects on trace element partitioning. Results indicate that in the presence of garnet or clinopyroxene, Nb and Ta are highly soluble in aqueous fluids, whereas Zr and Hf show variable solubilities. Low field strength elements (LFSE) and light rare earth elements (LREE) are always enriched in the fluid (D(fluid/Min) . 1). Generally, D(fluid/Cpx) is positively correlated with temperature only for high field strength elements (HFSE), but positively correlated with pressure for all other elements. Therefore, the lowest Nb/La is achieved at high pressures and low temperatures. However, even the highest pressures and lowest temperatures examined did not exhibit strong negative HFSE anomalies in the fluid. Garnet retains compatible trace elements at 3 GPa and 1000°C much more effectively (D(fluid/gt)Yb5 0.002) than at 5.7 GPa at the same temperature (D(fluid/gt)Yb 5 0.04). Decreasing temperature results in a lowered D(fluid/gt), particularly for Zr, Hf, and heavy rare earth elements (HREE). At 5 GPa and 900°C a strong intra-REE fractionation is observed (D(fluid/gt)Sm/Yb around 100) and significantly negative anomalies for Hf and Zr, but not for Nb and Ta, are developed. Only residual rutile fractionates all HFSE from all other trace elements. Tantalum and niobium are retained most effectively by rutile, as is the case for rutile/melt partitioning. Fluid/mineral trace element partitioning has important implications for mantle metasomatism in subarc regions. A model is proposed in which HFSE depletions, as observed in island arc volcanic rocks, could originate from a selective enrichment of the mantle wedge in LFSE and LREE by aqueous fluids derived from a rutile-bearing subducted slab. It is shown that melting of the enriched mantle wedge, which had previously been depleted by melt extraction (depleted MORB mantle) can produce magmas with trace element patterns similar to those of subduction-related volcanic rocks. Copyright © 1998 Elsevier Science Ltd a hydrous fluid of the LREE and LFSE relative to the HFSE exerts the most important control over trace element fractionation (Tatsumi et al., 1986, Saunders et al., 1991, McCulloch and Gamble, 1991; Hawkesworth et al., 1993, 1994), although opinions differ as to the necessity of a residual titanite mineral in equilibrium with the fluid. In the second, partitioning of the HFSE into a titanite phase in the mantle wedge and/or in a felsic melt residue during melting in the uppermost part of the slab effectively decouples the element groups (Nicholls and Ringwood, 1973; Marsh, 1976; Green, 1981; Foley and Wheller, 1990; Ringwood, 1990). Fundamental constraints on these models are the partition coefficients, at appropriate conditions, between (1) silicate mantle phases and melt, (2) titanate phases and melt, (3) fluid and silicate mantle phases, and (4) fluid and titanate phases. Of most interest are those phases which are known to incorporate high amounts of trace elements of a specific element group, e.g., clinopyroxene and garnet for MREE and HREE, amphibole for LFSE and LREE, and rutile and sphene for HFSE. Other OH-bearing minerals like serpen-

1. INTRODUCTION

The most consistently observed geochemical difference between arc and nonarc volcanics is the depletion in high field strength elements (HFSE), especially Nb and Ta, relative to low field strength elements (LFSE, e.g., K, Ba) and light rare earth elements (LREE). The behaviour of all of these element groups, but particularly the HFSE, during island arc magma genesis processes is controversial. The decoupling of different groups of trace elements during subduction-related enrichment processes or magma genesis is generally attributed to one of two mechanisms. In the first model, preferential enrichment in *Author to whom correspondence should be addressed (stalder@ erdw.ethz.ch). † Present address: Institut fu¨r Mineralogie und Petrographie, Sonneggstrasse 5, ETH Zentrum, 8092 Zu¨rich, Switzerland. ‡ Present address: Department of Earth and Planetary Sciences, Harvard University, 20 Oxford Street, Cambridge, Massachusetts 02138, USA. 1781

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tine or lawsonite may also be important for trace element partitioning in subduction zones. In contrast, olivine and orthopyroxene generally contain low amounts of trace elements (Green, 1994), and thus it is probably not necessary to take them into consideration. According to the depth of the subducted slab beneath the volcanic front in island arcs (100 –120 km; Gill, 1981; Tatsumi, 1986) appropriate conditions for fluid input in subduction zones may be 3–5 GPa. Experimental studies of fluid/solid trace element partitioning have been done but were limited to pressures at or below 3 GPa (Mysen, 1979; Ayers et al., 1997) or even lower maximum pressures (2 GPa: Brenan et al., 1994, 1995 a,b; Adam et al., 1997; 0.3 GPa: Keppler, 1996). In comparison to other studies, we investigated a higher pressure range (3–5.7 GPa), which corresponds to a depth of 90 –180 km and thus represents more directly the depth where metasomatising fluids are believed to be released. In silicate/fluid systems (e.g., Silica-H2O, Kennedy et al., 1962; Anderson and Burnham, 1965; Diopside-H2O, Eggler and Rosenhauer, 1978; Eggler, 1987; Albite-H2O, Paillat et al., 1992; Shen and Keppler, 1997) a second critical endpoint exists, above which a distinction between a fluid and a melt is no longer possible. Therefore, it seems questionable whether experimentally determined partition coefficients in fluid/silicate systems at subcritical conditions can be extrapolated simply to higher pressures. For a wide range of trace elements, especially for REE, fluid/mineral partitioning data are not well constrained. Thus there is a need for a complete and consistent data set for all trace element groups determined with one method. In order to fill this gap this study provides a data set for thirteen trace element partition coefficients between aqueous fluid and four different solid phases, i.e., two garnets (pyrope, py90gr10), clinopyroxene (diopside), and rutile. Temperature, pressure, and chemistry of the fluid and solid phases have been varied in order to evaluate their influence on D(fluid/min). Starting materials were doped with a more complete element set of LFSE, HFSE, and REE than many previous studies (Mysen, 1979; Brenan et al., 1994, 1995a,b; Keppler, 1996). 2. EXPERIMENTAL AND ANALYTICAL TECHNIQUES

Fig. 1. SEM-picture of a diamond trap after the run (5 GPa, 1000°C). Precipitated material (light) between the diamonds (dark) comprises approximately 15% of the pore space.

inclusions have high concentrations of the most incompatible trace elements and, therefore, inhibit the attainment of equilibrium for these elements. The resulting concentrations in the diamond trap are, therefore, too low. This problem is not a result of the trap technique, but of the glassy starting material used. At 1000°C the glassy starting materials crystallised completely after less than 14 h; residual phase assemblages were determined by XRD for at least one run product of each starting material. These runs were conducted without a diamond trap in Au-capsules with an outer diameter of 4.4 mm and an inner diameter of 4 mm in order to produce higher amounts of run product for XRD. For a more detailed description of the experimental procedure see Stalder et al. (1997).

2.1. Experimental Technique Experiments were carried out in a belt apparatus (Brey et al., 1990). As sample containers we used Au-capsules with an outer diameter of 3.0 mm and an inner diameter of 2.6 mm. Each capsule contained solid starting material (18 –20 mg synthetic silicate glass doped with thirteen trace elements) in the lower half and synthetic diamond crystals in the upper half (grain size 40 –50 mm). 3 mL fluid was added directly as water or hydrochloric acid (1.5–5 mol/L). The pore space between the diamond crystals was preserved at high pressure-temperature conditions, so that the fluid was able to communicate with the solid starting material throughout the experiment. During quenching the dissolved silicate material precipitated into the pore space between the diamond grains (Fig. 1) and thus was physically separated from the residual phases. Each capsule was checked for leakage by weighing, piercing, drying, and weighing again. Usually more than 90% of the initial water could be dried out after the experiment. Experimental run times were 48 h for runs at or above 1000°C and 336 h for runs at 900°C. One reversal experiment at 1000°C was performed by using undoped starting material and 3 mL 1000 ppm Hf-solution, demonstrating that equilibrium for Hf is reached within this time (Stalder et al., 1997). For more incompatible elements (e.g., La, Nb) equilibrium has not been shown; during crystallisation of the silicate glasses (in an early stage of the run), it may be possible that fluid inclusions form and that these

2.2. Starting Material Composition and Experimental Strategy The major element composition of the solid starting materials was obtained by mixing oxides or carbonates (MgO, SiO2, Al2O3, CaCO3) restricted to the CMAS system. Trace elements (Ba, Sr, Pb, Nb, Ta, Zr, Hf, Ti, La, Ce, Sm, Tb, Yb) were added as powdered oxides at a level of 100 ppm (Table 1) and to ensure homogeneous distribution in the starting material and shorter equilibration time during the run, major and trace element oxide mixtures were fused, following sintering, under atmospheric conditions at temperatures 20°C above the liquidus for half an hour. In the experiments involving rutile, natural rutile was added. It was finely ground in an agate mortar to grain sizes ,50 mm. To keep the trace element inventory and amount of dissolved silicate material in the fluid at approximately the same level as in the fluid/ silicate partitioning experiments, rutile bearing experiments also contained various amounts of diopside glass G13 (i.e., 50 –90 wt%). The major element composition of each glass starting material was analysed with an electron microprobe (Table 1). Measurement of ten points on each sample revealed that glasses were homogeneous at a scale of 1 mm. Concentrations of trace elements were determined with LAM-ICP-MS at the Memorial University of Newfoundland (Table 1). Five measured points on each sample confirmed homogeneity at a scale of 20 mm for trace elements.

Mineral-aqueous fluid partitioning of trace elements

The different starting materials furnish different phase assemblages. Glasses G11–G14 are Ol-normative (garnet 1 forsterite or diopside 1 forsterite) to keep the silica activity low, whereas G15 is Qz-normative (diopside 1 quartz). Runs with different silica activities enable direct comparison of fractionation patterns for high and low aSiO2 and provide information on the general dependence of trace element partitioning on the amount of dissolved material. It was expected that at least some of the additional silica dissolves preferentially in the aqueous fluid due to the supercritical state of the system H2O-SiO2 at these pressures (Kennedy et al., 1962). A high silica activity probably exists in the Qz-eclogite of the subducted slab or in veins of solidified melts derived from it. In contrast, low silica activities are more representative of the peridotitic mantle wedge. In the upper mantle, pyrope-grossular solid solutions with significantly variable Ca-contents occur: a depleted harzburgitic mantle source will contain grossular-poor garnets, whereas a Cpx-rich fertile peridotite will contain garnets with considerably higher grossular contents. Thus, the effect of the garnet composition on fluid/mineral partitioning behaviour has been included in our investigations (G11 5 Pyr90Gr10, G14 5 pure pyrope). In order to constrain the effect of complexing ligands at high pressure, sample containers were loaded with HCl (i.e., 1.5 m, 5 m) in two experiments. 2.3. Trace Element Analyses Trace element concentrations of the fluid trapped in the diamonds and the crystalline residues were analysed by laser ablation microprobe-inductively coupled plasma mass spectrometry (LAM-ICP-MS, Jackson et al., 1992, Longerich et al., 1993), using a frequencyquadrupled (266 nm) Nd-YAG laser (Table 2). Due to direct analysis of both quenched fluid and solid residues, partition coefficients could be calculated directly. Traps were analysed with a 0.6 mJ defocused beam, that produced pits 50 –200 mm in diameter (Fig. 2); residues were analysed with a 0.4 mJ beam, pits were usually smaller than 30 mm. Since grain sizes of the silicate phases were usually much smaller than 10 mm (Fig. 2c), measurement of single grains was not possible and concentrations are only given for bulk residues. At the beginning and the end of each analytical run (consisting of twenty measured points) NIST 610 glass standard was analysed. For each trap and residue 5–7 points were measured. Neither in the trap nor in the residue zoning could be observed. Internal standardisation for residues was not critical, since all experiments were run with a high solid/fluid mass ratio of 6 –7, and concentrations for major elements and most trace elements were not changed significantly during the run. Especially the HREE can be considered as compatible; therefore, Yb was used as the internal standard and concentrations in the solids could be analysed without further assumption. As mentioned above, each measured point is a multi-grain analysis due to the small grain size of the silicates. Residua containing diopside showed irregular signals for Zr and Hf in some cases. In experiments where this occurred, the system was probably slightly oversaturated with respect to these elements. To separate the Zr-rich phase from the Cpx and calculate the concentra-

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tions of the saturated Cpx, the irregular parts of the laser signal were cut out during the integration procedure. Thus, mass balance cannot be performed for all data given in Table 2. It should also be emphasised that quenched fluid veins occurred within the residues and that residues may, therefore, also contain quenched fluid. Where this was obviously the case (i.e., irregular signals, unusually high concentrations of LFSE, and/or major element ratios deviative from the stoichiometry of the silicate phase), integration was limited to quench-free parts of the residue. However, we cannot be rule out the possibility that inclusions or small amounts of quench are present. Therefore, calculated D(fluid/min) values for extremely incompatible elements should be regarded as minimum values. Internal standardisation for the trap analyses turned out to be difficult. The traps consisted of (1) diamond crystals, (2) quenched, mostly glassy material from the fluid (between the diamonds) and (3) open pore space, which previously was filled by the fluid phase and remained open after drying out the water. The problem is that the ratio of water/solid in the fluid and the volume of the pore space between the diamonds is unknown. We chose Ni as the internal standard element in the trap since it was a suitable trace element which was easily detectable in the synthetic diamonds (due to the use of a metal alloy as a catalyst for diamond growth) but absent in starting materials. Therefore, it is a reasonable assumption that all analysed Ni is hosted in the trap and relative concentrations of each analysed point could be related to each other. Quantification of trace element concentrations was achieved by performing a mass balance for Ba (the most incompatible element in this study), taking into account the fluid/solid mass ratio and the concentrations of the starting material and the solid residue. Mass balances were successful if Ni in the trap was set to 20 ppm (for further details see Stalder et al., 1997 and Stalder, 1997). 3. RESULTS

XRD analyses confirmed in all cases the expected residual phases based on the normative glass composition of the starting material. In all experiments minor amounts of olivine were present which kept the silica activity of the system low. An exception is run #58 (G15), where additional coesite was present in the residue. All doped trace elements were easily detectable in the trap. Concentrations measured with LAM and calculated partition coefficients are given in Table 2. Note that LAM is not the optimal method for analysing major elements. Furthermore, Al could not be measured, since the sample was ground with corundum paper, and small particles left behind impeded correct analysis. Therefore, sums of major elements in solid residues do not equal 100%. For experiments above 1000°C, sums exceeded 100 wt%, probably due to problems arising from internal standardisation because of the presence of a melt.

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However, partition coefficients relative to each other (within one experiment) and derivative values (see discussion) should not be affected. Results are illustrated in Figs. 3–9. D(fluid/mineral) values for all elements determined are plotted in the order of their incompatibility in the upper mantle using a modified order of Pearce (1983); this facilitates comparison with trends in arc volcanics, especially with respect to HFSE anomalies. In contrast to Pearce (1983), we plot the LFSE in order of decreasing D(fluid/mineral). In this paper the partition coefficient is expressed as D(fluid/mineral), following Mysen (1979), Keppler (1996), and Adam et al. (1997). For comparison with data of Brenan et al. (1994, 1995) and Ayers et al. (1997) reciprocal values have to be taken. The formulation D(fluid/mineral) gives a direct impression of the effect on volcanic trace element patterns: elements with a positive anomaly in the partitioning pattern will also exhibit a positive anomaly in a spidergram for rocks which source was enriched by a fluid. In the following paragraphs, trace element partitioning patterns of all fluid/solid pairs studied are presented. Attention is focussed on (1) decoupling of elements belonging to different trace element groups and (2) fractionation of elements within one group (especially intra REE-fractionation and Nb/Ta variations). 3.1. Diopside-Aqueous Fluid

Fig. 2. SEM-pictures of craters produced by laser ablation. Crater diameters of the traps varied between 50 mm (a) for samples run at 1200°C and 200 mm (b) for samples run at 1000°C. Analyses of crystalline residues are always multi-grain analyses, since the grain size of the crystalline material is much smaller than the laser spot (c).

The results for fluid/diopside partitioning at 1000°C (Fig. 3a) show a slight fractionation of the LFSE, HFSE, and REE; however, there is no marked fractionation of HFSE from the REE or LFSE. Niobium and tantalum partition into the fluid to a similar extent as LREE and some LFSE, and D(fluid/Cpx)Nb/Ta is constantly around 2. At 3 GPa, a HFSE enrichment rather than a HFSE depletion is observed in the fluid and at 5 GPa partition coefficients follow the same order of incompatibility as during melting in the upper mantle, and a smooth fractionation pattern results. The REE show the largest pressure dependence, with D’s increasing by nearly 1 order of magnitude from 3 to 5.7 GPa. Partitioning of HFSE exhibits no significant pressure dependence, whereas all other elements fractionate more strongly into the fluid with increasing pressure. Increasing temperature has the opposite effect to increasing pressure and at 1100 and 1200°C even positive HFSE anomalies develop (Fig. 3b; at 1200°C, as stated above, the quenched material was probably at least partly a melt); only DZr and DHf show positive temperature dependence. However, even under the highest pressures and lowest temperatures investigated no discernible negative anomalies for Nb and Ta, and only small ones for all other HFSE could be observed. In short, low temperature and high pressure favour Zr and Hf anomalies to a certain degree. One experiment was conducted with a high silica activity imposed on the system. Figure 4 shows that additional SiO2 and thus a high silica concentration in the fluid (Table 2) has only minor effects on trace element partitioning. Partition coefficients for some incompatible elements considered (e.g., Sr, Pb, Nb, Ta) are lowered by a factor of 2, whereas Nb/Ta stays unchanged. Thus, it is apparent that silica activity has generally only minor effects on the control of trace element partitioning relative to the effects of pressure, temperature, and the lattice structure of the residual phases.

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Fig. 3. Pressure dependence of partition coefficients D(fluid/Cpx) at 1000°C and various pressures (a) and at 5 GPa and various temperatures (b).

3.2. Garnet-Aqueous Fluid In our study of fluid/garnet partitioning, two different compositions have been investigated: one solid solution consisted of pyr90gr10, in the following called only “garnet,” and the second pure pyrope. For garnet the overall fractionation of the elements at 3 GPa is more pronounced than at 5.7 GPa (Fig. 5a). Generally, incompatible elements become more compatible at higher pressure and compatible elements more incompatible. There may be some fractionation of the HFSE from LREE and LFSE by garnet, particularly for Zr and Hf at 3 GPa.

Increasing temperature generally enhances D(fluid/gt) for REE and HFSE. In particular, D-values for Zr and Hf are increased by more than 1 order of magnitude during an increase of temperature from 900 to 1200°C at 5 GPa. At 900°C Zr and Hf, but none of the other HFSE, are retained more strongly by the garnet residue but still show a much higher concentration in the fluid than HREE such as Yb (Fig. 5b). At all temperatures and pressures an extreme internal REE fractionation is observed. Increasing pressure and particularly increasing temperature generate a less pronounced fractionation. D(fluid/gt)Sm/Yb at 5

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Fig. 4. Influence of silica activity in the fluid at 1000°C and 5 GPa. High silica activity in the fluid (#58, residual diopside 1 coesite) does not reveal significant changes in partition behaviour. For comparison #59 (diopside 1 minor amounts of Ol) is plotted in the same diagram.

GPa decreases from 96 at 900°C to 16 at 1200°C (Fig. 6). In the present study the steepest fractionation between MREE and HREE is developed at 1000°C and 3 GPa ( DSm/Yb 5 106). The chemical composition of the garnet also affects the partitioning of some trace elements. At 5 GPa and 1000°C LREE and HFSE are slightly better retained in pyrope and intra-REE fractionation is weaker (DSm/Yb 5 24) compared to garnet pyr90gr10 (DSm/Yb 5 46). However, this effect is small in comparison to the temperature effect. In contrast to pyr90gr10 at 5 GPa, pyrope exhibits discernible negative anomalies for Zr and Hf at 5 GPa (Fig. 7). 3.3. Garnet-Hydrochloric Acid Results for experiments containing hydrochloric acid are plotted in Fig. 8; for comparison runs with pure water are also shown. Partitioning patterns do not reveal significant differences to experiments with pure water. Neither a decoupling between element groups nor changes in the intra REE fractionation are observed. We found no effect on partitioning as a function of HCl concentration (1.5 m, 5 m HCl). 3.4. Rutile-Aqueous Fluid To study fluid/rutile partitioning, mixtures of rutile and Cpxglass (1:9 and 1:1) were equilibrated with an aqueous fluid at 5 GPa and 1000°C. The resulting fractionation diagrams are shown in Fig. 9. All HFSE are retained by more than 1 order of magnitude stronger than in the rutile-free residue. BulkD(fluid/solid) for other trace elements are slightly higher than for clinopyroxene alone. The presence of high amounts of silicate in the solid residue results in a major element content of dissolved material in the fluid at the same level as in the runs with only Cpx. Further-

more, the investigation of two different proportions , i.e., 1:9 (#71) and 1:1 (#73), also could be regarded as a bracket for partitioning of the HFSE. In #71 38.5% of the total Nb and 4.4% of the total Ta was hosted in rutile in the starting assemblage (Tables 1 and 2); in #73 85% of the total Nb and 29% of the total Ta was hosted in rutile in the starting assemblage. Since it was not possible to analyse single grains in the residual solids, only concentrations of the fluid and the bulk residue, but not the exact concentration in the rutile and the Cpx after the run could be determined. Partition coefficients are bracketed between two values; the first is obtained by assuming that the HFSE in the solid residue are exclusively hosted in the rutile after the experiment (ignoring that a small fraction also is present in Cpx). For the second extreme value it has to be taken into account that elements which are highly incompatible in Cpx and were hosted in the Cpx-glass before the run cannot attain equilibrium due to formation of small fluid inclusions in the Cpx during crystallisation. In this case the trace element content in the residue is only decreased by the fraction of the trace element which is not trapped by fluid inclusions. Taking the measured concentrations for Nb of #71 (Table 2, cfluid 5 10.9 ppm, csolid 5 94.9 ppm), D(fluid/rutile)Nb can be constrained as lying between 0.028 and 0.011. A better estimate of D(fluid/rutile) for Nb and Ta (for which equilibrium between silicate and fluid probably have not been reached) can be achieved if the amount of HFSE mobilised from the silicate starting material is taken into account. The concentration in the rutile after the experiment can be calculated as crt 5 (mfl*(c9fl2cfl)1 mrt*c0,rt)/mrt

where c’fl is the concentration of the given element in the fluid of run #59 (Cpx without rutile, Table 2), cfl the concentration in the fluid of the rutile-containing experiment and c0,rt the start-

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Fig. 5. Pressure dependence of partition coefficient D(fluid/garnet) at 1000°C and various pressures (a) and at 5 GPa and various temperatures (b).

ing concentration of the rutile (Table 1); mfl is the mass of fluid (3.0 mg) and mrt the mass of rutile (1.8 mg in #71 and 9.0 mg in #73). For Zr and Hf (for which silicate and fluid probably are close to equilibrium) a good estimate can be achieved by a mass balance ms*cs 1 mrt*c0,rt 5 mfl*cfl 1 mCpx*cCpx 1 mrt*crt

where cs is the concentration of the HFSE saturated silicate (known from #59, i.e., 26 ppm for Zr and Hf) and cCpx is the concentration in the Cpx after the run. It is assumed that the mass of the silicate starting glass (ms) equals the mass of the

clinopyroxene (mCpx) run product (i.e., 16.2 mg in #71 and 9.0 mg in #73). cCpx can be calculated as cfl/D(fl/Cpx). crt remains as the only unknown, which can then be expressed as crt 5 (ms*cs 1 mrt*c0,rt 2mfl*cfl 1 mCpx*cfl/D(fl/Cpx))/mrt

By dividing the measured cfl by the obtained crt, best estimates for the partition coefficients can be established. Calculated D(fluid/rutile) are given in Table 3. Niobium and tantalum show the highest compatibility in rutile, followed by Zr and Hf. Partition coefficients resemble the three-tier-pattern (Nb,Ta , Zr,Hf , others) already known from rutile/melt partitioning

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Fig. 6. Intra REE fractionation at 5 GPa in the presence of garnet. Most extreme Sm/Yb fractionation is observed at lowest temperatures.

(Jenner et al., 1993). Data for Nb and Ta are in good agreement with the 1000°C and 1100°C results of Brenan et al. (1994). The Nb/Ta fractionations in this study are not highly significant, and it cannot be concluded with certainty whether Ta is retained better than Nb. The stronger negative anomalies for Ta than for Nb in Fig. 9 can be explained by the different concentrations in the starting rutile and Cpx-glass, respectively. DTi was calculated directly from the concentrations cfluid/ rutile c since the system was saturated with respect to TiO2. In both #71 and #73 the Ti-concentration was around 2400 ppm, which is about 1 order of magnitude higher than predicted by Ayers and Watson (1993) for a pure aqueous solution (without dissolved silicate material in the fluid). 4. DISCUSSION

4.1. Comparison to Previous Studies Measured partition coefficients of this study are compared with previously published data in Table 4 and Fig. 10. Some features are confirmed by all these studies, including ours. It is generally accepted that in equilibrium with garnet and clinopyroxene LFSE partition strongly into the fluid; only Sr is compatible in Cpx (Fig. 10a). There is also a consensus that garnet retains HREE much better than LREE (Mysen, 1979; Brenan, 1995b), as in melting experiments. Major discrepancies to Brenan et al. (1995a,b) for D(fluid/Cpx) (Fig. 10a) and D(fluid/Gt) (Fig.10b) occur for the most incompatible elements, which are probably too low in our study (see previous section). Values for D(fluid/Cpx) are similar to those of Keppler (1996) and Adam et al. (1997). Partitioning data for rutile from this study are similar to those of Brenan et al. (1994), also with respect to Nb/Ta fractionation; rutile retains all HFSE. So far the pressure dependence of D(fluid/mineral) for HFSE has only been studied in the presence of rutile (Brenan et al.,

1994; 1–2 GPa) and for some REE in the presence of Cpx (Mysen, 1979; 0.5–3 GPa). The influence of temperature was only examined for HFSE in rutile (Brenan et al., 1994; 900 – 1100°C). Our results confirm the results of Mysen (1979) that there is a considerable pressure effect on trace element partitioning, but D(fluid/Gt) for REE determined by Mysen are by at least 1 order of magnitude higher than in this study. In most of the previous studies only solid phases were analysed, and trace element concentrations in the fluid were calculated by mass balance (Mysen, 1979; Brenan et al., 1994, 1995a). Only in few previous studies were all coexisting phases (fluid and solid or fluid and melt) analysed and mass balance established (Brenan et al., 1995b; Adam et al., 1997). Furthermore, most other studies were restricted to one or two trace element groups. Keppler (1996) determined fluid/melt partition coefficients at 0.3 GPa by analysing fluid and melt and calculated D(fluid/mineral) by dividing D(fluid/melt) by previously published values of D(mineral/melt). By doing so he circumvented the problems of long equilibration times for solid phases and of nonstoichiometric dissolution, but his results have to be extrapolated over several hundreds degrees and an order of magnitude in pressure. The large P-T-extrapolation of obtained values may explain differences to other studies, since the solubility of silicate material in fluids and the solubility of fluids in melts is dependent on pressure. Owing to the possible existence of a second critical endpoint under mantle conditions, it seems questionable whether experimentally determined partition coefficient in fluid/silicate systems at subcritical conditions can be extrapolated reliably to higher pressures. Previous studies on experimental fluid/mineral trace element partitioning have investigated the effect of Cl on partitioning behaviour with contrasting results. Keppler (1996) observed D(fluid/melt) for LFSE and LREE in a 5 M (Na,K)Cl aqueous solution to be enhanced by 1 order of magnitude at 0.3 GPa

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Fig. 7. Trace element partitioning of pyrope (#83) compared to pyr90gr10 (#64) at 1000°C and 5 GPa. Partition coefficients for LFSE, Nb, Ta, and LREE are lowered, but not changed relative to each other. Zirconium and hafnium are retained slightly better in pure pyrope.

compared to pure water. On the other hand, Brenan et al. (1995a) observed only slightly enhanced D(fluid/Cpx) or D(fluid/garnet) but no selective depletion of HFSE in the fluid with additional 0.5 m NaCl at 2 GPa, and our study found no significant effect for a 5 m HCl at 5 GPa. Ayers and Eggler (1995) carried out supraliquidus experiments and found a perceptible influence of complexing ligands like Cl on the amount of total dissolved silicate material in the fluid; this could not be

confirmed in our (subsolidus) study (SiO21MgO1CaO in the fluid at 1000°C and 4.5–5 GPa (Table 2) were 13 wt% for pure water, 19 wt% for 1.5 m HCl, and 10 wt% for 5 m HCl). A main reason for the contrasting results on the effect of Cl may be the differing experimental strategies and chemical systems (addition of different solutions, different alkali concentrations) between these studies. If the substitution 2 Ca21 5 REE31 1 Na1 controls the incorporation of REE, a higher concentration

Fig. 8. Effect of pure water vs. hydrochloric acid at 1000°C and 4 –5 GPa. Addition of 1.5 m HCl (#57) and 5 m HCl (#69) to the experimental charges show nearly no effect in trace element partitioning. For comparison runs at 1000°C with pure water only (#64 : 5 GPa, #55 : 4 GPa) are plotted.

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Fig. 9. Trace element partitioning at 1000°C and 5 GPa for experiments with residual Cpx (#59), 90% Cpx 1 10% rutile (#71) and 50% Cpx 1 50% rutile (#73). Rutile lowers the bulk D(fluid/solid) by more than 1 order of magnitude for all HFSE. D-values for other elements stay nearly unchanged.

of alkalis in the system (and in the Cpx) would promote a partitioning into the crystalline phase. In general, incorporation of more highly charged cations on the eightfold coordinated site can also be achieved by producing vacancies (3 Ca21 ---. 2 REE31 1 vacancy). However, differences in the results are probably more likely due to differences in confining pressure. 4.2. Controlling Factors of Fluid/mineral Trace Element Fractionation at Upper Mantle Pressures Trace element fractionation in igneous systems usually corresponds to CHARAC behaviour (CHarge-and-RAdius-controlled; Bau, 1996). In contrast, aqueous solutions show a non-CHARAC trace element behaviour, which does not mean that charge and radius have no influence, but that additional controls are operating (Bau, 1996). In aqueous systems electron configuration and complexing ligands (e.g., Cl2, SO422 ) cause selective solution and effective fractionation of trace elements (Johannesson et al., 1996), which usually are not fractionated during igneous processes. A famous example for this feature is the tetrad effect for the lanthanides, which results in a discontinuous behaviour for elements with quarter- and half-filled f-orbitals in terms of solubility products (Peppard et al., 1969). Previous studies dealing with fluid/crystal trace element partitioning have proposed both CHARAC and non-CHARAC behaviour as controlling mechanisms. Brenan et al. (1994) favoured major control by the chemistry of crystalline phases. They found that HFSE partitioning for fluid/rutile was not affected by 1m hydrochloric acid, but on the other hand, found a positive correlation between Al-content in the system (and thus in the fluid) and D(rt/fluid); with increasing alumina content HFSE became more compatible in rutile which was explained by coupled substitution in the rutile lattice (2 Ti41 ---. Al31 1 Nb51). Brenan et al. (1995) also suggest CHARAC pro-

cesses; their results exhibit a linear correlation between logD and rcation and a similar Th/U fractionation of Cpx/fluid compared to Cpx/melt, and trace element control by the crystal chemistry of the residual minerals at least for those elements seems to be probable. In contrast, Keppler (1996) explained strong trace element fractionation by calling on the preferred formation of Cl-complexes of some element groups in Cl-bearing systems. The observed non-CHARAC behaviour may be attributed to the considerably lower pressure of his experiments. In our study, neither Cl-concentration in the fluid nor silica content play a significant role for partitioning behaviour (Figs. 4 and 8). The effect of fluorine has not been investigated in this study, but its influence was a small at 2 GPa, where Cl still had a perceptible effect (Adam et al., 1997). Therefore, it seems improbable that trace element partitioning at the pressures and temperatures applied in our study is controlled by the aqueous fluid itself, which implies that crystalline phases exert the dominant control. The physical properties for high pressure aqueous fluids in silicate systems under subduction-relevant conditions probably are between hydrous fluids and silicate melts (Shen and Keppler, 1997), so that controlling mechanisms may lie between CHARAC and non-CHARAC behaviour. At higher pressure, aqueous silicate systems approximate a second critical endpoint, where no distinction between fluid and melt can be made (Eggler, 1987; Ryabchikov, 1993). Thus, during a progressive increase in P and T, an increase in CHARAC behaviour for fluid is likely. However, only in alkali-Al-Si systems are supercritical conditions proven for pressures in excess of 1 GPa; in alkali-poor, Mg-rich systems supercritical conditions are more likely to be reached at much higher pressures (e.g., Eggler and Rosenhauer, 1978).

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A charge- and radius-dependent model to predict mineral/ melt partitioning has recently been proposed by Blundy and Wood (1994); parameters taken into account are the optimum radius of the lattice sites and the Young’s Modulus (E) of the host crystal. All log-D-values for a given ionic charge and lattice site lie on a polynomial which obeys the following equation: D~ p,T, x! 5 D 0~ p,T, x!

5

3 exp

4 p EN A

F

G6

r0 1 ~r 2 r 0! 2 1 ~r i 2 r 0! 3 2 i 3 RT

(1)

NA is the Avogadro constant, r0 the optimum radius of the lattice site, ri the radius of the substituent cation and R the gas constant. If three D-values for a given charge and coordination are known, the curve can be fitted and the three parameters (D0, r0, and E) can be derived. For our data satisfactory results were obtained for REE in Cpx, when D-values of Ce, Sm, Tb, and Yb were fitted (Fig. 11a). The equation of the fitted curve furnishes the parameters r0, D0, and E which are presented in Table 4 for different pressures. For garnet, direct fitting was not successful, because real D0 are very close to zero and values derived from least square fits were negative (but close to zero). An alternative curve following Eqn.1 was constructed by trial and error (Fig. 11b); this shows minor deviation for Sm, Tb, and Yb. The reason for the mismatch observed for La and Ce could have various explanations. Probably, the experimental

technique applied does not furnish accurate results for strongly incompatible elements, since they may be partly trapped in the residue and lead to artificially lower D(fluid/min). If this is the case, D(fluid/min) for elements Ba-Ce would be too low, but the order of incompatibility would probably not be changed significantly. Furthermore, the proposed model might be restricted to concentrations ,100 ppm in the fluid due to the restriction of the Henry’s law region (Brenan et al., 1995) or increasing non-CHARAC behaviour occurs when the radius deviates greatly from D0. Data shown in Fig. 11 demonstrate that the lowest D(fluid/min) are achieved for those elements whose cation radius reveals the smallest deviation from the optimum size. No systematic change was observed between 4 and 5.7 GPa. At 3 GPa, the optimum cation radius r0 and the Young’s modulus E show the highest values and the minimum partition coefficient D0 the lowest (Table 5). At 3.0 GPa Blundy and Wood (1994) obtained r051.02 Å for the M2 site of diopside. This value is confirmed by our result r051.023 – 1.035 Å. The calculated Young’s modulus passes through a minimum at 5 GPa (corresponding to 150 km depth) for Cpx; whether this has important implications for seismic wave velocities (e.g., low velocity zone) cannot be concluded within the scope of this work. Using the proposed model, the differing behaviour between pure pyrope and a pyrope-grossular solid solution could be explained. Pyrope has a slightly smaller A-site which causes an increase of all D(fluid/gt) for elements in eight-fold coordination whose cation radius is larger than the optimum size and a decrease for all elements with a smaller cation radius. In other words, we should expect a stronger retainment for Hf and Zr relative to the REE, as observed (Fig. 7). Thus, we conclude that fluid/mineral trace element partitioning (at least for the P-T-conditions investigated), as is the case for mineral/melt partitioning, is mainly controlled by the crystal chemistry of the solid phase (CHARAC behaviour), and non-CHARAC behaviour is only of minor importance. Uncertainties still remain for temperatures lower than those investigated here. Zr and Hf

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Fig. 10. Comparison of experimentally determined D(fluid/Cpx) (a) and D(fluid/gt) (b). Not shown are data for some LFSE (e.g., K, Rb, Th, U) determined in other studies.

especially seem to become rather insoluble in the fluid at low temperatures. However, strong depletions in the fluid were only observed when garnet, which is known to incorporate high amounts of Zr and Hf in its lattice, is the residual phase. The model could also be applied to the twofold charged cations Ba, Sr, and Ca. Due to the low accuracy of the partition coefficients, the derived parameters E, r0, and D0 will be of low quality. 4.3. Implications for HFSE Contents in IAB Previous hypotheses on subarc trace element enrichment by aqueous fluids were based on less complete trace element partitioning data sets. D-values obtained in this study are used to calculate bulk partition coefficients for fluid/eclogite. Calculated D(fluid/Gt 1 Cpx 1 Rt) can be used to calculate the trace element content of the fluid that is released from the downgoing slab and introduced to the mantle wedge. In the first step we make a simplification. Ionic radii for Ca and Na are similar

(Shannon, 1976), and it can be assumed to a first approximation that both should have similar trace element partitioning behaviour. Therefore, we assume that omphacite behaves similarly to diopside. Similarly, the restriction of our experiments to CMAS should not have a significant influence on D(fluid/mineral) since cation radii for Mg and Fe are nearly identical. The amount of rutile present in the subducted slab depends on the Ti-content of the downgoing slab which we assume to have an average MORB composition, namely 1–1.7 wt% TiO2 (Sun et al., 1979). Thus, the subducted slab in our model contains 1.5% rutile. Modelled fluid/eclogite bulk partition coefficients (Cpx:Gt 5 50:50) at 5 GPa and 1000°C (Table 2) are presented in Fig. 12 for compositions with and without 1.5% rutile. Data are compared to modelled D(fluid/slab) as estimated by McCulloch and Gamble (1991); in their model they estimated D(fluid/slab) for several trace elements which are necessary to explain trace element contents in IAB. An addition to the depleted mantle wedge of 5 wt % fluid flux was taken

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Fig. 11. Crystal chemistry and cationic radii exert control over fluid/mineral trace element partitioning. REE lay on curves as predicted by Blundy and Wood (1994). Ionic radii are taken from Shannon (1976). Parameters, derived from fitting are given in Table 5. Optimum size for the Cpx M2-site is 1.03 Å (a) and for garnet A-site 0.91 Å (b).

into consideration, and this enriched IAB source was subsequently melted to a degree of F 5 0.1. Our model agrees in terms of enrichment in LFSE (Ba, Sr), but a higher amount of fluid is needed to enrich the mantle wedge since absolute D-values for most elements in this study are lower than those of McCulloch and Gamble (1991). We find decoupling of HFSE from REE only if rutile is present. Our model for the generation of IAB can be summarised as follows (only the role of the aqueous fluid has been assessed): (1) During subduction a fluid is liberated from the rutilebearing slab; its source may be hydrous minerals in the former oceanic crust or serpentinites from hydrated oceanic mantle (Tatsumi and Nakamura, 1986; Ringwood, 1990; Ulmer and Trommsdorff, 1995). (2) This fluid infiltrates the overlying mantle wedge, which has previously been depleted by melt

extraction (Gast, 1968; Green, 1972) and may lead to the formation of new hydrous minerals (e.g., mica, Saunders et al., 1991). The mass ratio fluid/(depleted mantle 1 fluid) is considered to be f 5 0.05– 0.1, corresponding to 5–10 wt% fluid in the system. (3) The depleted and re-enriched upper mantle is the source region for island arc magmas. The degree of melting of the garnet lherzolite (residual phases: 70% Ol, 20% Opx, 5% Cpx, 5% Gt) lies between 10 and 15% (F 5 0.1– 0.15). As the new hydrous assemblage is dragged down partial melting is initiated. In our calculation only batch melting is considered, i.e., the initial melt stays in equilibrium with the enriched assemblage until the final degree of melting is reached. This model (Fig. 13) combines the scenarios described by Ryerson and Watson (1987) and Brenan et al. (1994) on the one hand and Ulmer and Trommsdorff (1995) on the other hand.

Mineral-aqueous fluid partitioning of trace elements

Ryerson and Watson (1987) took into account that IAB are undersaturated in respect to TiO2, and thus a Ti-rich accessory phase is unlikely to control HFSE partitioning in the homogeneous peridotitic wedge itself. Residual rutile in a peridotite in equilibrium with a basaltic melt would not survive partial melting. Rutile is more likely to persist in the eclogitic slab. In contrast to many previous models, we favour serpentine breakdown as the principal water source (cf. Ringwood, 1990; Ulmer and Trommsdorff, 1995) instead of amphibole. The input of subducted sediments, which certainly contribute to the trace element inventory of island arc magmas (Morris et al., 1990; Plank and Langmuir, 1993) have not been included in the model in order to evaluate the input of the fluid from the eclogite separately. Regarding the major element concentrations of the fluid (Table 2, column SiO2, MgO, CaO), the major element composition of the mantle wedge will hardly be affected by the fluid input. The fluid contains only 10 –20% dissolved silicate material (which consists of not more than 60% silica and, therefore, is only slightly more siliceous than the parental basaltic system). Therefore, addition of 10% fluid to the mantle lherzolite only contributes 2% silicate material to the system and thus is negligible for the major element chemistry of the

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whole system. It is likely that most of the newly formed accessory phases are totally consumed by the melting process and the initially generated melt thus inherits the trace element inventory of the fluid flux, and only four phase lherzolite has to be considered for the residual mineral assemblage. The metasomatising agent released from the slab is considered to be an aqueous fluid rather than a partial melt (as far as a distinction between both can be made) because temperatures around the subducted slab are generally too low for a partial melting of the slab (Peacock, 1990). Chemical modification of the wedge by partial melting of the Qz-eclogite in the presence of high amounts of water is also possible (Foley and Wheller, 1990; Ringwood, 1990), but not considered in this model. Figure 14 shows the trace element pattern of a modelled IAB. The fluid composition was calculated for a 1.5 wt% rutile bearing slab and the fluid/eclogite mass ratio was estimated to be 0.05. The model calculation for Fig. 14b is outlined in Table 6. Modelled trace element patterns show negative anomalies for Nb and Zr. A mantle wedge which is modified by 5% fluid input (f 5 0.05) and subsequently partially melted to a degree of F 5 0.1 exhibits a discernible decoupling of Nb from La (Fig. 14a), but cannot explain the observed Nb/La ratio in IAB. To illustrate the contribution of the introduced fluid, a trace

Fig. 12. Modelled D(fluid/eclogite) with and without 1.5% rutile. For the sake of simplification the slab is assumed to consist of equal amounts of garnet and Cpx. Furthermore diopside is believed to behave similar to jadeite (see text for discussion). Modelled pattern are compared with D(fluid/slab) as predicted by McCulloch and Gamble (1991). Our modelled rutile-bearing eclogite may be able to produce similar trace element fractionation pattern as needed for selective enrichment in the mantle wedge by fluids released from the subducted slab.

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Fig. 13. Simplified scetch for the proposed model for the enrichment process in island arcs.

element pattern for dry melting (f 5 0) without previous enrichment is also depicted in Fig. 14a. A larger fluid input and subsequently a higher degree of melting (F 5 0.15, f 5 0.1) shows a much better resemblance to IAB trace element pattern (Fig. 14b). The amount of fluid influx in our model might be overestimated even if parental magmas in subduction zones can contain high amounts of water (Anderson, 1980 (up to 4 wt%), Sisson and Layne, 1993; Sisson and Grove, 1993; .8 wt%). Under pressures of 3– 4 Gpa, where melting may occur, the melt is able to dissolve large amounts of water, resulting in H2O-saturation during emplacement to lower pressure regimes and subsequent exsolution. However, if partition coefficients for LFSE, LREE, Nb, and Ta in this study are considerably higher, the amount of water necessary in the model is 1 order of magnitude lower. Furthermore, less rutile is neccessary if it is taken into account that all HFSE partition more strongly into rutile at lower temperatures (Brenan et al., 1994). Uncertainties still remain with respect to Nb/La fractionation; hence additional fractionation mechanisms are not excluded. For further discussion of rutile in subduction zones see Brenan et al. (1994). Even if TiO2 is highly soluble in basaltic melts, this does not mean that rutile is certainly absent in the mantle wedge; it may be stable in enriched veins of micapyroxenites which solidified from a former metasomatising low-degree partial melt (Foley and Wheller, 1990). Since Tisaturation levels of melts derived from these K-rich veins are lower, nonperidotitic veins could host Ti-rich accessory phases even during partial melting, and it is possible that they contribute to the observed negative HFSE-depletion. We do not preclude the survival of rutile in mantle veins of mixed mineralogy after partial melting, but do not take it into account in our model. 4.4. Nb/Ta Fractionation Niobium and tantalum have often been regarded as geochemical twins (Green, 1995) and the reason for their fractionation in CHARAC processes remains enigmatic. Recent experimental work suggests that rutile saturated melts would

increase their Nb/Ta due to more effective retainment of Ta by rutile (Jenner et al., 1993). The opposite effect was observed for rutile in equilibrium with an aqueous fluid (Brenan et al., 1994). Hence, it was concluded that low Nb/Ta values for the crust could be attributed to metasomatism by fluids equilibrated with residual Ti-phases (Green, 1995). Within the limits of our results it remains unsure whether rutile fractionates Nb and Ta; if existent, the effect probably is weak. In the present work experiments with clinopyroxene and garnet show Nb/Ta in the fluid to be enhanced by a factor of 2 compared to the starting composition (Fig. 3–5) thus suggesting that aqueous fluids are likely candidates to produce HFSEenriched hydrous minerals with exceptionally high Nb/Ta. During a metasomatic event fluids will have Nb/Ta ratio well above the mantle value ( ..17.5, Sun and McDonough, 1989). Recently Ionov and Hofmann (1995) described amphiboles and micas from enriched mantle xenoliths with Nb/Ta up to 95, which they proposed to have been crystallised during metasomatism by a fluid originated from a subducted slab.

4.5. Effective Intra-REE Fractionation and Implications for Subcalcic Garnets Fractionation between fluid and garnet is much more pronounced than that between fluid and diopside (Fig. 5a). For the REE this fractionation appears to be much stronger than that accomplished by silicate melt/garnet partitioning. Compared to DSm / DYb 5 3.5 for melt/garnet at 2.5 GPa and 1430°C (Hauri et al., 1994) we determined fluid/garnet DSm / DYb up to 100 at 3 GPa and 1000°C and 5 GPa and 900°C. Still higher DSm/DYb are expected for lower temperatures (Fig. 6). Pressure and temperature dependencies lead us to the conclusion that the most extreme intra-REE fractionation in the mantle is realised in cold regions within the garnet stability field. This is most likely realised in regions with a low geothermal gradient, e.g., in subduction zones and in the upper lithospheric mantle beneath old cratons. Deeper in the Earth a less pronounced fractionation would be generated. The observation that granular garnet lherzolite nodules from kimberlites, which are depleted in major elements, show in bulk a high LREE/HREE (Shimizu, 1975), led to the conclusion that they were enriched by a metasomatic event after major element depletion (e.g., melt extraction). Calculation of the REE composition of the metasomatising agent revealed very high LREE/HREE. Mysen (1979) proposed that due to the high fractionation this metasomatic agent could be an aqueous fluid. Similarly, sinusoidal REE patterns (with low LREE/MREE and high MREE/HREE) of garnets from diamondiferous harzburgitic xenoliths (Shimizu and Richardson, 1987) and from harzburgitic inclusions in diamonds (Stachel and Harris, 1997) would require a strongly fractionated REE pattern (Sm/Yb around 100) in a potential parent liquid and, therefore, cannot be explained by crystallisation of any known melt composition. Thus, it was concluded that the responsible process was a metasomatic enrichment by a fluid (Stachel and Harris, 1997). This fluid was estimated to have introduced high amounts of LFSE (Rb, Sr) and LREE, low amounts of Zr and Hf and basically no HREE. All these characteristics are observed in garnet-equilibrated fluids in this study, so we reaffirm the conclusion that LREE enriched garnets could be

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Fig. 14. Modelled trace element pattern for island arc basalts. The depleted mantle (McCulloch and Bennett, 1994) is replenished with trace elements by an aqueous fluid equilibrated in the eclogitic slab with an approximate composition like MORB (Pearce et al., 1981). Fluid fraction f is given in wt ratio. Subsequent melting to a degree of F 5 0.1 (a) and F 5 0.15 (b) of depleted (f 5 0) and replenished mantle source furnishes the model arc basalt (see text for further discussion). Concentrations are normalised to primitive mantle (Hofmann, 1988) and compared to average IAB (McCulloch and Gamble, 1991).

the product of aqueous fluid metasomatism at depths of 80–100 km, corresponding to pressures of 3 GPa.

5. CONCLUSIONS

Several findings from the present study are in accord with those of Brenan et al. (1995a):

(1) LFSE and LREE may not be easily fractionated from the HFSE simply by calling on a role for fluids. In particular Nb and Ta seem to soluble in these high pressure fluids to a similar extend as LREE. (2) Only in the presence of residual titanate phase (e.g., rutile) will there be significant fractionation of the HFSE from the REE and LFSE. (3) In contrast to Keppler (1996), we do not find evidence to

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support a major effect of chloride on relative HFSE vs. REE/LFSE partitioning. Contrasting results may be due to different run pressures; we do not doubt that complexation with halogens may be very important for lower crustal conditions, but suggest that they are of only minor importance for trace element transport under upper mantle conditions. (4) It is impossible to produce a relatively HFSE-depleted trace element signature in IAB-sources with a metasomatising fluid which is released from a rutile-free source. Thus, it seems probable that a residual Ti-rich accessory phase in the eclogitic slab would specifically retain HFSE during release of a silicate-bearing aqueous fluid. The fluid flux modifies the mantle wedge by a selective enrichment in LFSE and LREE and generates relative HFSE depletions in the source region of island arc volcanics. The subsequently produced IAB would inherit the signature of the modified mantle. In addition we can conclude that (5) Fluid/mineral trace element partitioning is, like mineral/ melt partitioning, mainly controlled by the crystal chemistry of the residual phases, i.e., by charge- and radiuscontrolled processes (CHARAC behaviour). This may be linked to the approaching of a supercritical point, where no distinction can be made between hydrous melts and fluids. Behaviour at temperatures lower than in this study may deviative from this (non-CHARAC behaviour). (6) Extreme REE-fractionation, as observed in subcalcic garnets and inclusions in diamonds, may be accomplished by interaction with a fluid phase. Decreasing temperature reinforces this effect and LREE-enriched assemblages are most likely to be generated in the cold subcratonic lithosphere at depths of around 80 –100 km. Acknowledgements—This project was kindly supported by the Deutsche Forschungsgemeinschaft (BR 1012/6-1). R.S. acknowledges funding by the University of Frankfurt. The manuscript benefitted much from discussion with George Jenner, Thomas Zack and Dorrit Jacob. Anni Bieniok is thanked for helping with XRD and Peter Meier (Uni Heidelberg) directed microprobe analysis. We also thank John Ayers, James Brenan and an anonymous reviewer for their thorough and constructive reviews.

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