Behaviour of boron, beryllium, and lithium during melting and crystallization: constraints from mineral-melt partitioning experiments

Geochimica et Cosmochimica Acta, Vol. 62, No. 12, pp. 2129 –2141, 1998 Copyright © 1998 Elsevier Science Ltd Printed in the USA. All rights reserved 0016-7037/98 $19.00 1 .00

Pergamon

PII S0016-7037(98)00131-8

Behaviour of boron, beryllium, and lithium during melting and crystallization: Constraints from mineral-melt partitioning experiments J. M. BRENAN,1 E. NERODA,2 C. C. LUNDSTROM,3 H. F. SHAW,4 F. J RYERSON,5 and D. L. PHINNEY6 1 Department of Geology, University of Toronto, Toronto, Ontario M5S 3B1, Canada Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, California 91125, USA 3 Earth Sciences Board, University of California, Santa Cruz, California 95064, USA 4 Geoscience and Environmental Technology Division, L-201, Lawrence Livermore National Laboratory, Livermore, California 94551, USA 5 Institute of Geophysics and Planetary Physics, L-202, Lawrence Livermore National Laboratory, Livermore, California 94551, USA 6 Nuclear Chemistry Division, L-237, Lawrence Livermore National Laboratory, Livermore, California 94551, USA 2

(Received August 25, 1997; accepted in revised form March 10, 1998)

Abstract—In order to provide a more substantial foundation for interpreting the behaviour of B, Be, and Li during the production and early crystallization of primitive igneous rocks, we have measured olivine-, clinopyroxene-, orthopyroxene-, and amphibole-melt partition coefficients for these elements involving broadly basaltic-andesitic melt compositions. Experiments were conducted at both one atmosphere and 1.0 –1.5 GPa and employed a time-temperature history that yielded large crystals with minimal compositional zoning. Experiment temperatures ranged from 1000 to 1350°C and were selected to minimize the total crystal fraction in a given experiment. Partition coefficients for olivine and clinopyroxene were found to be independent of run duration or total concentration of B, Be, or Li suggesting that crystal-liquid equilibrium was closely approached. Olivine-, orthopyroxene-, and clinopyroxene-melt partition coefficients decrease in the order: Li (0.1– 0.2) @ Be ; B (0.002– 0.03), whereas amphibole-melt partition coefficients for Be and Li are similar (;0.2) and larger than those for B (;0.02). Comparison of partition coefficients measured in this study with previous determinations yields good agreement, with the exception of some of our mineral-melt values for B, which are uniformly lower (up to 10 times) than values determined at similar conditions of pressure and temperature. The latter discrepancy could be due to mineral or melt compositional effects, but this hypothesis is currently untestable owing to the absence of reported mineral compositions in previous studies. Partition coefficients for olivine and clinopyroxene have been found to vary as a function of mineral and melt composition, and with the exception of B partitioning into clinopyroxene, this variation can be modeled using simple exchange reactions involving the trace element and a substituent element, such as Na, Mg, or Al. Partition coefficients measured in this study were combined with simple models of melting and crystallization to evaluate how accurately element ratios such as B/Be, B/K, B/Nb, Be/Nd, Li/V, and Li/Yb in primitive magmas reflect that of their source. These models further confirm that the source regions of IAB magmas are enriched in B/Be, B/Nb, and Li/Yb relative to the MORB source, thus lending further support to the notion of metasomatic enrichment of the IAB source by slab-derived fluids. Moreover, our modeling also indicates that the low B/Be and B/Nb in primitive OIB magmas is indicative of similarly low values in OIB sources, which is consistent with the hypothesis that OIB sources contain a B-depleted component, such as subducted, dehydrated oceanic crust. Partial melting models have also been constructed to explore the possibility of using the Li/V ratio in MORB and IAB as a monitor of redox conditions in their source-regions. Models indicate that this ratio does not uniquely constrain source fO2 without a priori knowledge of the degree of melting. However, the small amount of dispersion in MORB Li/V is consistent with (1) the small variation in source-region fO2 inferred for MORB by independent means and (2) degrees of melting close to clinopyroxene exhaustion. The very large dispersion in Li/V ratios in the IAB suite can be reconciled by melt generation under more oxidising conditions than that for MORB, in addition to variation in source composition resulting from metasomatism involving a Li-rich component. Copyright © 1998 Elsevier Science Ltd concentrations (i.e., B and 10Be) and elevated concentration ratios (i.e., Li/Yb, B/Nb, B/Be) of these elements in island arc basalts (IABs), relative to either mid-ocean ridge basalts (MORBs) or ocean island basalts (OIBs), provides a strong indication for a component of subducted oceanic crust (fluid or melt) in the IAB source region (Ryan and Langmuir, 1987, 1988, 1993; Ryan et al., 1996; Tera et al., 1986; Morris et al., 1990). Much of the current insight regarding the behaviour of B, Be, and Li during igneous processes has been derived from studies of specific magma suites in which the variation of B, Be, or Li is compared with other elements for which the experimental/

1. INTRODUCTION

The behaviour of B, Be, and Li in igneous rocks has received much attention owing to the potential capacity for these elements to trace the chemical flux from subducted oceanic crust to the source region for convergent margin magmas. The utility of B, Be (as the cosmogenic isotope 10Be) and Li as geochemical tracers stems from the fact that each of these elements (or isotopes) are relatively enriched in altered seafloor basalt and/or pelagic sediment (cf. Leeman and Sisson, 1996; Tera et al., 1986; Ryan and Langmuir, 1987), yet depleted in the oceanic upper mantle. As such, both the observed elevated 2129

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J. M. Brenan et al.

observational database is more substantial (Ryan and Langmuir, 1987, 1988, 1993; Ryan et al., 1996). Although valuable insight can be gained from this exercise, the general utility of such observations in developing quantitative models for B, Be, and Li variation during melting and crystallization is limited by an inadequate knowledge of the partitioning behaviour of these elements between common igneous minerals and silicate melts. Previous experiments to measure partition coefficients for these elements has largely focused on B (Chaussidon and Libourel, 1993; Seitz, 1973), and to date, only Ryan (1989) has provided partition coefficients for B, Be, and Li for some of the common igneous minerals at low pressure conditions. Inasmuch as such partitioning data is crucial to separating the effects of melting and crystallization from variations in source composition and to establishing the efficiency of 10Be transfer during melting, a comprehensive set of partition coefficients for these elements is clearly required. It is in this context that we have measured partition coefficients for B, Be, and Li in experiments conducted at both ambient pressure and 1–1.5 GPa involving the minerals olivine, clinopyroxene, orthpyroxene, and pargasitic amphibole coexisting with broadly basaltic melt compositions. 2. EXPERIMENTAL METHODS

2.1. Partitioning Experiments Starting materials for partitioning experiments were synthesized by repeated grinding and fusion of reagent grade oxides and carbonates. The exception to this procedure was the single experiment involving an Fe-bearing composition (the Baker/ Stolper partial melt, see below). This composition was synthesized using the standard oxide/carbonate mixture, decarbonated at 1200°C, reground, then sintered at 1000°C in a CO2-CO mixture at an oxygen fugacity equivalent to the fayalite-magnetite-quartz buffer. Beryllium was added to the starting materials as a dilute HNO3 solution, Li as Li2CO3 in dilute HCl, and B was added as H2BO3 in H2O. Subsequent to adding the Be and Li, the starting material was thoroughly mixed under ethanol, dried at 400°C; the B solution was then added and the sample dried at 100°C. The concentrations of B, Be, and Li added to our experiments were dictated by the minimum detection limits for these elements (see below) and guided by the previously mentioned partitioning and observational studies. For the minerals under consideration, we anticipated that B and Be would be moderately to highly incompatible, which warranted the use of element concentrations in our experiments (20 –190 ppm B or Be) that exceeded levels typically found in mafic to intermediate igneous rocks (,1–20 ppm B and ,1–2 ppm Be; Ryan and Langmuir, 1988, 1993). On the other hand, because of the more compatible behaviour expected for Li, we employed Li concentrations in our experiments (17–115 ppm) that overlap natural levels (2–50 ppm; Ryan and Langmuir, 1987; Leeman and Sisson, 1996). Bulk compositions for partitioning experiments were selected based on their relevance to basaltic-andesitic magmatism, whether they afforded the opportunity to study bulk compositional effects and if the phase equilibria was reasonably well known. Bulk compositions for the one-atmosphere olivine/melt partitioning experiments consist of three compositions along the 1350°C isotherm in the system forsterite-

albite-anorthite (Watson, 1977; Schairer and Yoder, 1966). In terms of the proportions of nonbridging oxygens, CaO and Na2O contents, composition 1b (Fo17.3Ab82.7An0 by weight), and 8c (Fo30Ab23.3An47.8) represent extremes along this isotherm, without crossing into the field of spinel saturation. A third composition consisted of a 1:1 mixture (by weight) of the latter endmembers. Compositions employed for the one-atmosphere clinopyroxene/melt partitioning experiments consist of two endmembers along the 1275°C isotherm in the system diopside-albite-anorthite (Bowen, 1915; Schairer and Yoder, 1960; Osborn and Tait, 1952) doped with ;1 wt% of TiO2 and ;0.2 wt% Cr2O3 to allow for additional charge balance possibilities. Composition Di-An plots along the diopside-anorthite join (;Di65An35), whereas composition Di-Ab plots along the diopside-albite join (;Di55Ab45). Amphibole/melt partitioning experiments were conducted using the same composition employed previously by Brenan et al. (1995). The composition is roughly andesitic and was selected because pargasitic amphibole is the primary near-liquidus phase at 1.5 GPa, 1000°C under water-bearing conditions and the coexisting melt readily quenches to a homogeneous glass. We also performed one experiment at 1.0 GPa and 1320°C using the liquid composition from lherzolite melting experiment #15 (1280°C, 1.0 GPa) of Baker and Stolper (1994). Approximately 2 wt% H2O was added to this experiment and crystallization products consist of olivine, orthopyroxene, and clinopyroxene. In all partitioning experiments the degree of crystallization was ,15% (as estimated visually). All one-atmosphere experiments were conducted in vertical tube furnaces with MoSi2 heating elements and temperature was monitored with Pt-Pt10%Rh thermocouples calibrated against the melting point of pure gold. All high pressure experiments were conducted using a piston-cylinder pressure apparatus employing a crushable ceramic-pyrex-NaCl pressure cell with samples immediately surrounded by either MgO (i.e., the Baker/Stolper composition) or calcium carbonate (amphibole-melt partitioning). Temperature was monitored in the piston-cylinder experiments using either W3%Re–W25%Re or W5%Re–W26%Re thermocouples with NIST-traceable calibrations. With the exception of the Baker/Stolper composition, all other compositions are nominally Fe-free and were, therefore, encapsulated in Pt and sealed with an arc welder. The Baker/Stolper composition was welded into a graphite-lined platinum capsule to eliminate Fe loss. The time-temperature history for each partitioning experiment was designed to promote the growth of a few, large, equant, and compositionally homogeneous crystals. To this end, a typical experiment included (1) an initial superliquidus step to reduce or eliminate pre-existing nuclei and promote melt homogenization, (2) a temperature drop (typically ;100°/ min) to the predetermined liquidus with or without a subsequent isothermal soak at this temperature (to promote crystal nucleation), and (3) a slow cooling step (typically 0.2–1°/h to promote crystal growth) over a temperature interval of 4 – 6°C, followed by an isothermal soak lasting from 1 to 100 h. Samples were then quenched by either immersion in cold water (one-atmosphere experiments) or by shutting the power off to the sample heater (piston-cylinder experiments). A summary of experiments and their time-temperature histories is provided in Table 1.

Mineral-melt partitioning of boron, beryllium, and lithium

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Table 1. Summary of Mineral/Melt Partitioning Experiments Experiment I.D.

Mineral

Pressure

initial T(°C)1

initial time (hrs)2

liquidus T(°C)3

liquidus time (hrs)4

dT/dt (deg/hr)5

final T(°C)6

final time (hrs)7

8c-2a/low 8c-2/low 8c-2/high lb Baker/Stolper Di-An Di-Ab BBeLi5 BBeLi50

olivine olivine olivine olivine olivine, opx.cpx cpx cpx amphibole amphibole

1 atm 1 atm 1 atm 1 atm 10 kb 1 atm 1 atm 15 kb 15 kb

1450 1450 1450 1450 1400 1293 1293 1200 1200

1 1 1 1 1 40 40 2 2

1353 1353 1353 1353 1325 1281 1281 1005 1005

0 0 0 0 1 24 24 2 2

0.2 0.2 0.2 0.2 1 0.2 0.2 0.3 0.3

1349 1349 1349 1349 1320 1275 1275 1000 1000

1 100 63 63 24 3 3 72 72

notes: 1) Initial homogenisation temperature. 2) Duration of initial homogenisation step. 3) Predetermined liquidus temperature. 4) Duration of nucleation step. 5) Cooling rate for crystal growth step. 6) Final “soaking” temperature. 7) Duration of final “soak”.

2.2. Analysis of Run-Products Operating conditions for electron microprobe analyses of run products were optimized to minimize Na loss. Stable emission of Na x-rays (i.e., for at least 60 s) from samples was achieved using a 20 mm, defocused beam and a 10 –20 nA or 2–5 nA beam current on crystals and glass, respectively. Synthetic mineral standards were used, and raw x-ray intensities were converted to oxide weight percent using the Bence-Albee method (Bence and Albee, 1968). Reported errors for the electron probe analyses are one standard deviation in the variation of multiple spots. The presence of ;13 wt% water in the glass coexisting with amphibole has been confirmed in a previous study (Brenan et al., 1995) using FT-IR spectroscopy and accounts for the low totals obtained in the analyses reported in this work. Concentrations of B, Be, and Li were measured using a Cameca IMS-3F ion microprobe at LLNL using a 15–50 nA Oprimary beam with a net energy of ;17 KeV and a 20 mm spot diameter. Positive secondary ions were extracted and accelerated to ;4.5 KeV. For each sample, 3– 4 spots on each phase were analyzed, and the results were averaged. Isobaric molecular interferences were minimized through the use of an energy-filtering technique (Shimizu et al., 1978; Zinner and Crozaz, 1986); only secondary ions having energies lying within a 32.5 eV window centered on an 80 eV offset from the peak of the secondary ion energy distribution were collected. As previously shown by Ottolini et al. (1993), the magnitude of this voltage offset is sufficient to minimize the differences in ion yields between a variety of different silicate minerals and glasses. Reported trace element concentrations were determined by comparing the 30Si-normalized isotopic ratios of the samples to those of NIST glass standards. The quenched melt from the partitioning experiment involving amphibole contains ;13% H2O, and we found that ion yields for B, Be, and Li relative to 30 Si were lower in this hydrous glass than in an anhydrous glass of the same composition. The reduction in yield is largest for B (wet/dry 5 0.69), intermediate for Li (wet/dry 5 0.87) and only

minor for Be (wet/dry 5 0.97). Based on analyses of NIST glass standards, the estimated combined error in the accuracy of trace element concentrations is 620%. As a result of surface contamination, the count rate for 11B usually decayed from somewhat elevated values to a constant level only after the first 10 –15 counting cycles (i.e., 10 –15 min on most mineral specimens); only the data obtained in subsequent cycles was considered to reflect the intrinsic abundance within the sample. Partition coefficients were calculated from the ratio of the average 30Si-normalized count rate of an isotope in the mineral to the average 30Si-normalized count rate in the glass, correcting for the differences in Si contents and the difference in ion yields (for hydrous glasses) due to the effect noted above. Reported 1s errors reflect propagation of the larger of the counting statistics errors or the variability of the spot analyses. In most cases, analyses of quenched melt yielded a standard deviation from separate spots that was similar to errors based on counting statistics (i.e., ;1%). However, analyses of different spots on crystals from run products were in most cases reproducible to within 3–30%, which exceeds counting statistics (i.e., 1–5%). This latter variation is certainly due to some heterogeneity in run product crystals (i.e., zoning, small melt inclusions), although the poorest reproducibility usually involved B analyses at concentrations that were close to the minimum detection limit. To assess minimum detection limits we have analysed synthetic crystals of forsterite and diopside, both of which are materials that we believe to be nominally free of B, Be, and Li. In each case, the count rates at masses 7 (Li), 9 (Be), and 11 (B) were indistinguishable from instrumental background (determined at a nonintegral mass) after the first 10 –15 min of sputtering. Minimum detection limits, calculated as two standard deviations above the background count rate are ;180 ppb for B, ;0.1 ppb for Be, and ;15 ppb for Li. For all of the glasses and minerals analysed in this study we obtained count rates for B, Be, and Li that were above background levels. In most cases, measured concentrations were at least five times

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J. M. Brenan et al. Table 2. Melt Compositions Produced in Experiments1

oxide

8c-2 (n 5 20)2

lb (n 5 20)

50:50 (n 5 20)

Di-Ab (n 5 10)

Di-An (n 5 8)

Baker/Stolper (n 5 20)

BBeLi (n 5 10)

SiO2 TiO2 Al2O3 Cr2O3 FeO MgO CaO Na2O Total

49.21(0.41)3 NA 22.22(0.33) NA 0.06(0.04) 14.78(0.27) 9.96(0.15) 2.90(0.06) 99.14

65.47(0.56) NA 16.88(0.23) NA 0.07(0.02) 6.60(0.14) 0.04(0.02) 9.83(0.17) 98.90

57.30(0.47) NA 19.64(0.24) NA 0.03(0.03) 11.02(0.21) 4.91(0.07) 6.59(0.11) 99.49

61.34(0.45) 1.05(0.04) 9.63(0.12) NA NA 8.48(0.19) 12.54(0.15) 5.64(0.07) 98.88

49.50(0.30) 1.03(0.03) 14.68(0.24) 0.15(0.04) NA 10.13(0.10) 23.10(0.19) 0.05(0.02) 98.64

49.05(0.81) 0.64(0.04) 16.60(0.32) 0.03(0.03) 5.95(0.21) 10.79(0.16) 11.31(0.25) 2.49(0.05) 96.854

51.28(0.79) 0.18(0.05) 20.74(0.35) NA NA 2.33(0.09) 6.64(0.26) 5.55(0.15) 86.724

notes: 1) all analyses are by electron microprobe, see text for analytical details 2) n is the number of analyses 3) number in parentheses refers to error based on one standard deviation of n analyses 4) sample contains H2O not analysed by emp NA 5 not analysed

background levels with the exception of the B content of olivine from experiments involving the 8c-2 compositions, which are within a factor of two of the detection limit. 3. RESULTS AND DISCUSSION

3.1. General Observations Tables 2 and 3 list the major element compositions of coexisting crystals and melt from the partitioning experiments. In all experiments, large (100 –500 microns) equant crystals were grown using the protocol described previously. The reader is referred to Fig. 2 of Brenan et al. (1995) and Fig. 1a of Lundstrom et al. (1994) for representative examples of runproduct textures. Both quenched melt and coexisting crystals were found to be homogeneous within the precision of the electron microprobe analyses, with the general exception of clinopyroxene in which CaO, MgO, and Al2O3 typically show 5– 6% relative variation. The major element compositions of olivine, clinopyroxene, and orthopyroxene produced in our experiments are within the range reported by Hervig et al (1986) for minerals in spinel

lherzolite xenoliths from geographically diverse localities (Table 4). Amphiboles produced in our experiments are very similar in composition to the low Ti pargasites found in spinel lherzolite xenoliths (Table 4). The most notable distinctions between the amphiboles synthesized in this study and those found in nature are the lower Ti and the absence of Cr, K and Fe in the synthetic material. Owing to the general compositional similarity between the minerals produced in our experiments and those present in lherzolite xenoliths, we expect that our partitioning data will be particularly useful to modeling the processes of partial melting and the initial stages of fractional crystallization of primitive magmas. 3.2. Trace Element Partition Coefficients 3.2.1. General observations Boron, beryillium, and lithium are moderately to strongly incompatible in the minerals considered for this study, as summarized in Table 5. For olivine, clinopyroxene, and orthopyroxene, mineral/melt partition coefficients decrease in the

Table 3. Mineral Compositions Produced in Experiments1 oxide

olivine: 8c-2 (n 5 28)1

olivine: lb (n 5 10)

olivine 50:50 (n 5 20)

olivine Baker/Stolper (n 5 20)

cpx: Di-Ab (n 5 10)

cpx: Di-An (n 5 10)

cpx: Baker/Stolper (n 5 20)

opx: Baker/Stolper (n 5 20)

ampibole: BBeLi (n 5 45)

SiO2 TiO2 Al2O3 Cr2O3 FeO MgO CaO Na2O Total

42.64(0.41)3 NA 0.11(0.05) NA 0.04(0.02) 56.50(0.51) 0.18(0.02) NA 99.47

43.36(0.58) NA 0.04(0.01) NA 0.04(0.02) 56.63(0.32) 0.01(0.01) NA 100.08

42.27(0.48) NA 0.00(0.00) NA 0.01(0.01) 57.01(0.48) 0.11(0.01) NA 99.40

40.70(0.24) 0.04(0.04) 0.07(0.01) 0.09(0.01) 8.89(0.15) 50.65(0.36) 0.26(0.03) 0.02(0.01) 100.69

55.24(0.33) 0.22(0.02) 0.75(0.04) 1.59(0.33) NA 18.20(0.20) 23.22(0.33) 0.54(0.09) 99.76

51.55(0.41) 0.40(0.03) 4.76(0.27) 2.13(0.19) NA 16.40(0.22) 24.53(0.20) 0.00(0.00) 99.78

50.88(0.39) 0.29(0.03) 7.94(0.42) 0.89(0.11) 4.10(0.19) 19.40(0.58) 15.86(1.03) 0.47(0.02) 99.83

54.46(0.34) 0.09(0.02) 5.08(0.12) 0.85(0.04) 5.21(0.07) 31.85(0.33) 2.06(0.05) 0.07(0.01) 99.68

43.54(0.76) 0.34(0.03) 19.64(0.49) NA NA 18.13(0.43) 11.14(0.32) 3.12(0.07) 95.92

notes: 1) all analyses are by electron microprobe, see text for analytical details 2) n is the number of analyses (typical of 4 –5 different grains) 3) number in parentheses refers to error based on one standard deviation of n analyses NA 5 not analysed

Mineral-melt partitioning of boron, beryllium, and lithium

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Table 4. Average and representative analyses of olivine, orthopyroxene, clinopyroxene and amphibole from spinal lherzolite xenoliths oxide

olivine1

orthopyroxene1

clinopyroxene1

amphibole 100062

amphibole WGBM153

amphibole PHN40224

SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO CaO Na2O K2O Total

41.1(4) 0.003(2) 0.027(8) 0.016(4) 9.8(1.1) 0.143(17) 49.3(9) 0.074(13) 0.05(3) 0.0020 100.85

55.3(8) 0.093(55) 4.1(0.9) 0.46(10) 6.2(7) 0.144{16) 33.0(8) 0.82(14) 0.091(33) 0.0046(33) 100.34

52.0(1.0) 0.35(23) 5.6(1.4) 0.91(19) 2.90(41) 0.091(14) 15.8(7) 20.5(1.) 1.31(41) 0.0023(12) 99.58

45.99 0.08 12.91 3.14 3.42 0.07 18.45 9.07 4.79 0.92 98.84

43.98 1.60 13.26 1.91 4.15 0.07 17.32 10.95 3.11 1.28 97.63

43.2 1.08 15.9 0.81 3.98 0.03 17.6 11.0 4.07 0.02 97.69

notes: 1) Average analyses from Hervig et al. (1986), number in parentheses is one standard deviation of 3–21 analyses (see Hervig et al. [1986] for details). 2) Representative analysis of amphibole from Nunivak Island xenolith in Francis (1976). 3) Representative analysis of amphibole from southeastern Australia xenolith in O’Reilly et al. (1991). 4) Representative analysis of amphibole from Malaita, Solomon Islands xenolith in Neal (1988).

order: Li (0.1– 0.2) @Be ; B (0.002– 0.03). Amphibole/melt partition coefficients for Be and Li are similar (;0.2) and larger than those for B (;0.02). Results for partitioning experiments that involved different concentration levels (BBeLi5, BBeLi50, and 8c-2/low, 8c-2/high) and duration (8c-2a/low, 8c-2/low, and 8c-2/high) are reproducible at the 2s level or better. This observation, combined with the generally homogeneous nature of the phases produced in our experiments, suggests a close approach to crystal-melt equilibrium and Henry’s Law behaviour. Moreover, although some workers have reported apparent B transport into or outof experiments employing sealed platinum capsules (Hart and Dunn, 1994), others have not documented such behaviour (Chaussidon and Libourel, 1993; Ryan, 1989). The excellent reproducibility of the B content of glasses from experiments containing identical starting materials (i.e., 8c-2a/low and 8c-2/low), but run for vastly different soak durations (1 and 100 h, respectively), suggests our experiments are closed with respect to B gain or loss.

3.2.2. Comparison to previous studies A summary of mineral/melt partition coefficients for B, Be, and Li determined in previous studies is provided in Table 6. In addition to values for the minerals considered in the present work, for completeness, we have also included partition coefficients involving spinel, phlogopite, and plagioclase. With the exception of some of the previously published values for B, partition coefficients determined in the present study are in good agreement with past work. Significantly, previous measurements of partition coefficients for Li substantiate our findings that it is at least an order of magnitude more compatible than B or Be in anhydrous mafic silicates. It is also of note that, like amphibole, plagioclase shows a similar propensity to reject B relative to either Be or Li, both of which show essentially identical partitioning behaviour. A prominent exception to the relative compatibility of Li in the aformentioned minerals is the markedly more compatible behaviour of Be (D 5 0.95) relative

Table 5. Summary of Mineral/Melt Partition Coefficients Experiment I.D.

Mineral

ppm B1

Dmineral melt B2

ppm Be1

Dmineral melt Be2

ppm Li1

Dmineral melt Li2

8c-2a/low 8c-2/low 8c-2/high lb 50–50 Baker/Stolper Baker/Stolper Baker/Stolper Di-An Di-Ab BBeLi5 BBeLi50

olivine olivine olivine olivine olivine olivine opx cpx cpx cpx amphibole amphibole

22.2(0.2) 21.8(0.4) 28.9(0.4) 171.9(3.2) 90.0(1.5) 90.0(6.2) 90.0(6.2) 90.0(6.2) 65.8(4.3) 34.6(1.2) 12.0(0.9) 16.2(0.9)

0.007(0.001) 0.007(0.002) 0.004(0.001) 0.004(0.002) 0.003(0.001) 0.008(0.003) 0.018(0.002) 0.025(0.003) 0.016(0.002) 0.017(0.003) 0.022(0.003) 0.015(0.001)

61.0(0.4) 65.1(0.7) 79.7(0.9) 59.2(2.4) 68.9(2.4) 46.6(2.4) 46.6(2.4) 46.6(2.4) 48.9(0.6) 54.6(0.5) 8.0(0.1) 24.1(0.4)

0.0018(0.0002) 0.0018(0.0003) 0.0018(0.0001) 0.0029(0.0004) 0.0021(0.0001) 0.0015(0.0001) 0.016(0.001) 0.021(0.002) 0.011(0.001) 0.0025(0.0004) 0.26(0.02) 0.21(0.01)

26.5(0.3) 28.2(0.7) 54.2(0.9) 25.7(0.8) 19.0(0.6) 17.5(1.2) 17.5(1.2) 17.5(1.2) 49.9(3.5) 115.4(2.4) 48.0(0.3) 80.6(0.8)

0.15(0.02) 0.13(0.01) 0.15(0.01) 0.15(0.02) 0.15(0.02) 0.35(0.03) 0.20(0.02) 0.27(0.05) 0.20(0.03) 0.14(0.02) 0.19(0.02) 0.13(0.01)

notes: 1) concentration of element measured in glass by SIMS, number in parentheses is the one sigma uncertainty based on the larger of error from counting statistics or the variability of 8 –10 spot analyses 2) error on individual partition coefficients (in parentheses) is the one sigma uncertainty based on the larger of counting statistics of the variability of 8 –10 spot analyses on glass and 3– 4 spot analyses on crystals

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J. M. Brenan et al. Table 6. Summary of B, Be and Li partition coefficients from literature sources

Mineral

Data Source

Experimental Conditions

Dmineral melt boron

Dmineral melt beryllium

Dmineral melt lithium

olivine

1

0.019–0.044

—

—

olivine

2

0.028

—

—

olivine

3

0.01–0.02

,0.02

0.20–0.43

cpx cpx

1 2

0.117 0.67

— —

— —

cpx cpx

3 4

0.01–0.17 0.036*

0.03–0.13 0.047

0.11–0.18 —

opx opx spinel spinel amphibole

1 2 1 2 5

0.027 0.046 0.08 0.08 —

— — — — 0.15

— — — — 0.124

phlogopite plagioclase

5 3

1-1.5 GPa: 1325–1450°C: peridotite partial melts 1 GPa: 1375°C: peridotite partial melts 1 atm: 1150-1245°C: natural basalts as per olivine above 2.5 GPa: 1150°C: synthetic, hydrous, Fe-free peridotite partial melt as per olivine above 3.0 GPa: 1380°C, natural basalt as per olivine above " " " 1.5 GPa, 1092°C, natural basanite " as per olivine above

— 0.02–0.07

0.95 0.21–0.33

0.064 0.22–0.24

Data sources and notes: 1) Chaussidon and Libourd (1993) 2) Seitz (1973) 3) Ryan (1989) 4) Hart and Dunn (1993) 5) LaTourrette et al. (1995) * Value considered suspect by authors owing to evidence of open system behaviour for boron during their experiment

to Li (D 5 0.064) in phlogopite (Table 6; LaTourrette et al., 1995). All of the olivine/melt partition coefficients for B reported in the literature are uniformly higher than those determined in this study. Similarly, the clinopyroxene/melt partition coefficient for B measured by Chaussidon and Libourel (1993) is ;10x the values we have measured, and the former value is generally larger than other values previously reported (Table 6). Various possibilities exist to explain these discrepancies, including mineral and melt composition effects, differences in pressure and temperature, and surface contamination of B during analysis. In terms of mineral or melt composition effects, as outlined below, simple relations exist that describe the exchange of B between olivine and silicate melt, such that the partition coefficient for B should correlate with that for Al (the clinopyroxene-melt systematics for B are somewhat less understood). Unfortunately, no mineral composition data are provided in the previous studies of B partitioning and, therefore, the possibility of high Al partition coefficients correlating with those for B cannot be tested. Differences in pressure or temperature are unlikely to account for the higher partition coefficients in the previous studies, inasmuch as the range of pressure and temperature previously investigated overlaps that of the current study. Boron contamination during surface analysis could lead to anomalously high B concentrations, and this effect would be proportionally more acute for analyses that involved phases that are very low in B. Seitz (1973) used autoradiography to measure the B contents of coexisting phases and this technique has been shown to be particularly sensitive to trace amounts of surface contamination (Shaw et al., 1988). Partition coefficients

determined in the Seitz (1973) study could thus be erroneously high. Both Chaussidon and Libourel (1993) and Ryan (1989), however, used SIMS to measure the B content of coexisting phases from their experiments. As in our study, both groups subjected sample surfaces to extensive cleaning by the ion beam prior to accumulating analytical data, and Chaussidon and Libourel (1993) assessed the analytical blank for B by analysing a piece of vacuum-degassed rhenium foil. Accepting these precautions as adequate to avoid surface contamination, the lower mineral/melt partition coefficients we have obtained for B could be the result of compositional effects, although more extensive work would be required to thoroughly address this issue. 3.2.3. Mineral-melt exchange reactions Partition coefficients for olivine and clinopyroxene were found to vary with mineral and melt composition, and we have attempted to account for this variation in the context of likely mineral-melt exchange reactions. Owing to similarity of both ionic radius and charge, we anticipate that olivine-melt partitioning would involve the exchange of B for Al (or ferric iron), Be for Mg (or ferrous iron), and Li for Na. Similarly, likely exchange reactions involving the major cations in clinopyroxene would be B and Be for Al (in either the tetrahedral or M1 site) or ferric iron in the M1 site, Be for Mg or ferrous iron in the M1 site and Li for Na in the M2 site. The reaction that describes an isovalent exchange reaction between a major cation X having formal charge n with a trace element Y takes the form:

Mineral-melt partitioning of boron, beryllium, and lithium

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Xnmin 1 Ynliq 5 Xnliq 1 Ynmin where the subscripts liq and min denote melt and mineral, respectively. The equilibrium constant for this reaction can be expressed as: [Xnliq] [Ynmin] [Xnmin] [Ynliq] where the square brackets denote activites. Equating activities to mole fractions, the equilibrium constant can be written as the ratio of two mineral/melt molar partition coefficients, Ds, for the elements Y and X: DY DX which is a constant at any given temperature and pressure and invariant with composition of either the mineral or melt, provided activity coefficients do not vary. Thus, provided the appropriate exchange reactions have been chosen, plots of molar mineral/melt Ds involving a trace element and its substituent element should produce linear relations, whose slope is equal to the equilibrium constant and whose intercept is the origin. Such information can then be used to describe partitioning of B, Be, and Li for compositions different than those explored experimentally, provided substituent element partition coefficients are available (cf. Jones, 1995). Figure 1 portrays the results of these calculations for the case of olivine/melt partitioning. Equilibrium constants for each regression are provided in Table 7, along with the assumed exchange reaction. In each case, the regression curve was forced through the origin. Owing to the low abundances of Al and Na in run-product olivines, precise concentrations could not be determined by electron microprobe. Instead, ratios of 23 Na/30Si and 27Al/30Si in run-product olivines and glasses were determined by SIMS and converted to absolute concentrations using the electron microprobe analyses of the glass from each experiment. In Figure 1a, the B partitioning data for the Fe-free experiments are reasonably-well correlated with the partition coefficient for Al and all of those values have been regressed to calculate a single equilibrium constant. The partitioning data for the high pressure Fe-bearing experiment is distinctly above the Fe-free trend, which suggests that either a decrease in temperature or an increase in pressure favours B, relative to Al, substitution in olivine, or that there is an additional exchange mechanism that occurs in the Fe-bearing system, possibly involving ferric iron. As shown in Fig. 1b and 1c, excellent correlations were found for partition coefficients involving the element pairs Be-Mg and Li-Na. Moreover, because all of the partitioning data for each element pair describe coherent trends extrapolating to the origin, all of the data in each figure were regressed with a single curve. This result implies that (1) small differences in temperature (30°C) do not appreciably affect the equilibrium constants for the reactions assumed and (2) activity coefficients are constant over the compositional range of our experiments. Figure 2 portrays the results for clinopyroxene-melt partitioning. To define which site in the clinopyroxene structure B resides, we have attempted to correlate the D for B with the D for Al calculated as either the ratio of tetrahedral or octahdral

Fig. 1. Plots of olivine-melt partition coefficients for B, Be, and Li (molar basis) as a function of partition coefficients for cations likely to be involved in trace element exchange. In each case, the best fit curve through the data is constrained to pass through the origin. Slopes of the regression curves, which define the eqilibrium constant for the exchange, are provided in Table 7. (a) Dboron vs. Daluminum; (b) Dberyllium vs. Dmagnesium; (c) Dlithium vs. Dsodium.

Al in clinopyroxene relative to the total amount of Al in the liquid. In either case, distinct linear relations do not exist between the D for B and the D for Al calculated in this fashion, as is illustrated in Fig. 2a for the exchange involving six-fold Al. Indeed, the Ds for B for the Fe-free compositions are essentially identical, despite distinct differences in the Ds for Al. Although the Ds for both B and octahedral Al are distinctly higher in the Fe-bearing experiment compared to those involving the Fe-free compositions, it is not clear whether this is due to the presence of ferric iron or possibly the higher pressure or lower temperature at which the experiment was conducted. Other reactions, involving the coupled exchange of Na1 1 B31

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J. M. Brenan et al. Table 7. Summary of olivine- and clinopyroxene-melt exchange reactions for boron, beryllium and lithium and equilibrium constants derived from partitioning data Exchange Reaction

Equivalent Ratio of Molar Partition Coefficients

Equilibrium Constant

31 31 31 Al31 ol 1 B liq 5 Al liq 1 B ol

DB ol DAl ol

7.46(0.56) 3 1021

21 21 21 Mg21 ol 1 Be liq 5 Mg liq 1 Be ol

DBe ol DMg ol

3.75(0.13) 3 10-4

1 1 1 Na1 ol 1 Liliq 5 Naliq 1 Liol

21 21 31 Na1 cpx 1 Al (VI) cpx 1 Ca liq 1 Be liq

5 Na

1 liq

1 Al

31 liq

1 Ca

21 cpx

1 Be

21 cpx

1 1 1 Na1 cpx 1 Liliq 5 Naliq 1 Licpx

for Ca21 1 Mg21 or Al(IV)31 1 B31 for Si41 1 Mg21, also did not produce significant correlations in the associated partition coefficients. It is likely, therefore, that clinopyroxene/melt partitioning of B is complicated by the action of several substitution mechanisms that change in response to changes in melt composition, pressure, and temperature. Clinopyroxene/melt partitioning of Be and Li appear to follow much simpler compositional systematics than does the partitioning of B. Like B, Be partitioning cannot be described by simple, isovalent exchange reactions, such as those involving Mg or Ca. We found, however, that the partitioning of Be could be very well described by a coupled exchange involving Ca21 1 Be21 for Na1 1 Al(VI)31 (Fig. 2b), which is consistent with the overall similarity in the ionic radii of Al (0.535 angstroms) and Be (0.45 angstroms) in six-fold coordination (Shannon and Prewitt, 1969, 1970). The value of the apparent equilibrium constant for the aforementioned exchange is ;0.7 (Table 7), which is higher than the value of ;0.2 calculated using the Be partitioning data of Hart and Dunn (1993). However, these latter experiments produced clinopyroxenes having much lower CaO (;15 wt%), and higher Na2O and Al2O3 (;3 and 14 wt%, respectively) than run-products from this study. Thus it seems likely that changes in the activity coefficients for either Ca, Na, or Al in clinopyroxene could account for this difference in apparent equilibrium constant. As in the case for olivine, clinopyroxene/melt partition coefficients for Li are linearly correlated with those for Na, indicating that Na is the most probable substituent element (Fig. 2c). As for Be partitioning, our apparent equilibrium constant for Li exchange (1.3, Table 7) is larger than the value calculated using the Hart and Dunn (1993) partitioning data (;0.8). Again, it seems likely that the substantial differences in clinopyroxene composition between the two studies accounts for this difference. We have also compared our partition coefficients for Li and Na to those predicted using the elastic strain model of Blundy and Wood (1994). According to this model, because Li is smaller (0.92 angstroms) and Na is larger (1.18 angstroms) than the optimal ionic radius for the clinopyroxene M2 site (1.05 angstroms),

DLi ol DNa ol

8.30(0.23) 3 101

Be DCa cpx D cpx Na Al (VI) Dcpx Dcpx

6.71(0.07) 3 1021

DLi cpx DNa cpx

1.30(0.01) 3 100

there is a small difference in the amount of elastic strain energy required to substitute Li or Na into the M2 site. This difference is, in turn, manifested by slightly different partition coefficients for these elements. The Blundy/Wood model predicts that the (wt%) partition coefficient for Li should be only slightly larger than that for Na (DLi/DNa 5 1.04), whereas regression of our partitioning data yields a value of DLi/DNa of 1.30. This difference in predicted vs. observed relative partitioning (;20% relative) is beyond the error of our measurements and could reflect the effect of pressure, temperature, or pyroxene composition on the values of the optimal radius for the M2 site and/or the bulk modulus for singly charged cations determined by Blundy and Wood (1994). 3.3. Variation in Element Ratios During Melting and Crystallization It has now become well established that, under certain conditions, elemental ratios in primitive igneous rocks can be used to uniquely discern the chemical variability inherent to their respective source regions (cf. Hofmann et al., 1986; Miller et al., 1994; Sims and DePaolo, 1997), in essence allowing element ratios to be used in a similar manner to isotope ratios. Except for the case of very high degrees of melting, the necessary condition for the ratio of two elements in a derivative melt to be identical to that of its source is for the bulk solid/ liquid partition coefficients for the elements to be the same. For the few elements for which this latter requisite is satisfied (i.e., Ce/Pb, Nb/U), this realization has provided geochemists with a tool for defining specific geochemical reservoirs independent of isotopic data. Past studies of the behaviour of B, Be, and Li in volcanic rocks have similarly sought to define other elements which exhibit like partitioning behaviour during the melting or crystallization process (Ryan and Langmuir, 1987, 1988, 1993; Ryan et al., 1996). Such observations have subsequently been used to further substantiate the notion that the sources of mid-ocean ridge basalt (MORB), ocean island basalt (OIB), and island arc basalt (IAB; e.g., Ryan et al., 1995, 1996;

Mineral-melt partitioning of boron, beryllium, and lithium

Fig. 2. Plots of clinopyroxene-melt partition coefficients (molar basis), as in Fig. 1. In contrast to the data for olivine, note the general constancy of the partition coefficient for B and the lack of correlation between Dboron and Daluminum. (a) Dboron vs. Daluminum; (b) Dcalcium*Dberyllium vs. Dsodium * Daluminum (see text for details); (c) Dlithium vs. Dsodium.

Ishikawa and Nakamura, 1994) are geochemically distinct. Combining the partitioning data obtained in this study with simple models of melting and crystallization allows one to assess a priori which elements are likely to behave similar to B, Be, and Li during igneous processes. As described below, this information allows for a more complete interpretation of the abundance data for these elements in volcanic rocks, particularly with regard to variations in source-region chemistry. 3.3.1. Variations in boron, beryllium, and lithium abundances in IAB and OIB Studies of the behaviour of B, Be, and Li in primitive MORB glasses (i.e., with greater than ;7 wt% MgO) by Ryan and

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coworkers (Ryan and Langmuir, 1987; 1988; 1993; Ryan et al., 1996) have indicated that the element pairs B-Be, B-Nb, B-K, Be-Nd, Li-V, and Li-Yb appear to have similar partitioning behaviour during the partial melting process. As such, provided the residual mineralogy is similar to that during MORB petrogenesis, ratios of these elements in primitive rocks from other igneous suites should also reflect that of their source. This conclusion has been used to suggest that the high B/Be, B/Nb, and Li/Yb ratios in island arc lavas relative to MORB reflect source regions that are products of metasomatism by B and Li-enriched slab-derived fluids (Ryan and Langmuir, 1987, 1988, 1993; Ryan et al., 1995; Morris, et al., 1990; Ishikawa and Nakamura, 1994). In addition, owing to the perceived mobility of B in fluids produced during subduction of oceanic crust, it is expected that this element will be selectively depleted in the subducted material that descends into the convecting portion of the mantle. As a consequence, the low B/Be and B/Nb ratios in oceanic island basalts relative to MORB have been used to provide additional support for the hypothesis that OIB sources contain a component of subducted oceanic crust (Ryan et al., 1996). In order to further substantiate the inferred similarity of the aforementioned element pairs during the melting process, we have used the partition coefficients measured in this study (selecting those measured from the Baker/Stolper composition as most appropriate), combined with values for K, Nd, Nb, and Yb from Keleman et al. (1994), Brenan et al. (1995), and LaTourrette et al. (1995) to calculate the variation in B/Be, B/Nb, B/K, Be/Nd, and Li/Yb ratios in a melt as a function of the degree of melting of a lherzolite mineralogy (Li/V ratios are considered in the following section). The trace element composition of model melts was calculated assuming equilibrium (batch) melting and that melting was nonmodal. We have applied the variation in mineral mode as a function of melt fraction determined in the experiments of Baker and Stolper (1994) and use their starting mineral mode of 51 wt% olivine, 30 wt% orthopyroxene, 17 wt% clinopyroxene, and 2 wt% spinel. We have also assessed the effect of residual amphibole on element fractionation in separate calculations by assigning half the starting clinopyroxene mode to amphibole and assuming that amphibole and clinopyroxene contribute to the melt in the same proportions. We acknowledge that a simple batch melting model is likely to be an overly simplistic rendition of that which occurs during MORB genesis. However, Williams and Gill (1989) have shown that trace element abundance ratios in the accumulated melt fractions from more complex models (i.e., continuous or dynamic melting) do not differ significantly from that produced in the batch case. Results of the melting calculations, plotted as element ratio relative to that of the source as a function of melt fraction, are shown in Fig. 3a and b. For the case of no residual amphibole, the ratios B/Be, B/Nb, and B/K are essentially unfractionated relative to the source provided the degree of melting exceeds ;10%. For smaller degrees of melting, B/K can be up to three times higher and B/Nb and Be/Nd up to two times lower than the initial source composition. In contrast to ratios involving B or Be, the ratio of Li/Yb in derivative melts is up to a factor of two lower than that of the source over much of the melting interval. This is consistent with the interpretation of Ryan and Langmuir (1987) that DLi is slightly larger than DYb during

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Fig. 4. Abundances of lithium as a function of vanadium for MORB glasses with greater than 7 wt% MgO and lherzolite xenoliths from the data of Ryan and Langmuir (1987) and Jagoutz et al (1979), respectively. Note the similarity in Li/V ratio between mantle lherzolites and the glass array, implying that DLi 5 DV during the melting process that produces MORB.

source ratios involving B, Be, and Li from the compositions of OIB lavas (cf. Sims and DePaolo, 1997). It is important to note, however, that the calculated two-fold (maximum) reduction in B/Be and B/Nb in small degree melts relative to their source is insufficient to explain the three- to ten-fold lower ratio of these elements in OIBs. Thus, unless there is a residual phase in OIB sources that preferentially sequesters B (phlogopite?), the low B/Be and B/Nb in OIBs relative to MORB is likely to reflect similarly low ratios in their sources. Fig. 3. Plots of element ratio relative to source ratio as a function of melt fraction for the ratios (a) B/Nb and B/K and (b) Be/Nd, B/Be and Li/Yb. Curves were calculated for each element pair using the partitioning data determined in this and previous studies, combined with a simple batch melting model that assumes phase proportions change during melting as documented by Baker and Stolper (1994; see text). Solid curves correspond to melts coexisting with an anhydrous residue containing olivine, orthopyroxene, and clinopyroxene whereas the dashed curves correspond to the case when one half of the clinopyroxene is replaced by amphibole.

MORB genesis. Because amphibole rejects B relative to Be, potassium and niobium, B/Be, B/K, and B/Nb ratios are elevated to varying extents relative to source ratios so long as amphibole is residual (in this model, amphibole is consumed at a melt fraction of ;0.12). For the case of island arc magmatism, the degrees of melting are typically assumed to be high (i.e., greater than 10%, cf. Stolper and Newman, 1994), and thus our calculations would suggest that the elevated ratios of B/Be and B/Nb (relative to MORB) in primitive arc magmas are likely to accurately reflect similarly elevated ratios in their sources. Inasmuch as Li/Yb ratios in partial melts are anticipated to be about half as large as the source value, the elevated values of Li/Yb in IABs relative to MORB are certain to reflect elevated values of this ratio in the IAB relative to the MORB source. In contrast to the relatively high degrees of melting inferred for arc magmas, somewhat lower degrees of melting could be inferred for some members of oceanic island suites, owing to their often alkalic character. It would seem prudent, therefore, to consider the effects of melting when estimating

3.3.2. Variations in lithium/vanadium ratios in MORB and IAB In their study of the Li systematics of young volcanic rocks, Ryan and Langmuir (1987) noted that the Li/V ratio in primitive MORB glasses (.7 wt% MgO) was surprisingly constant at 0.020 6 0.003 (1s; Fig. 4) and that this value is essentially identical to the average Li/V found in mantle xenoliths (Jagoutz et al., 1979). As described previously, in order for the average Li/V ratio in MORB to be the same as that of likely sourcerocks, the bulk partition coefficients for Li and V during melting must be similar. It is well documented that mineral/melt partition coefficients for vanadium vary within the range of oxygen fugacities likely for terrestrial magmagenesis, and this is due to the existence of predominantly V31 at low fO2 and V51 at higher fO2. Because the partitioning of V is a function of oxygen fugacity, whereas partitioning of Li is not, the inferred similarity in DLi and DV in MORB could provide a means to estimate the ambient oxygen fugacity during MORB genesis. We have explored the possibility of using Li/V in MORB as a redox indicator by applying a simple nonmodal batch melting model as described in the previous section. In this case, however, we have incorporated the dependence of DV on fO2 using the mineral/melt partitioning data of Lindstrom (1976), Horn et al. (1994), and Canil (1997). These partitioning relations are shown in Fig. 5a along with the best-fit curves used to determine the appropriate DV at any given fO2 with respect to the fayalite-magnetite-quartz (FMQ) buffer; the equations for these curves are provided in the figure caption. In light of the com-

Mineral-melt partitioning of boron, beryllium, and lithium

Fig. 5. (a) Plot of mineral-melt partition coefficients (log scale) for vanadium as a function of oxygen fugacity (log scale) with respect to the fayalite-magnetite-quartz buffer. All data are from experiments performed at one atmosphere and the temperature range indicated. The curves through the data are best-fit linear or polynomial regressions. The equations for the linear regressions through the olivine and orthopyroxene data are: log D 5 21.376 2 0.251 Dlog fO2 FMQ (olivine) and log D 5 20.387 2 0.232 Dlog fO2 FMQ (orthopyroxene). The data for spinel and clinopyroxene were fit with third order polynomials of the form log D 5 a 1 b*(Dlog fO2 FMQ) 1 c*(Dlog fO2 FMQ)2 1 d*(Dlog fO2 FMQ)3. For spinel, the values of a, b, c, and d are: 1.051, 20.360, 20.0262, and 0.00432, respectively. For clinopyroxene, the values of a, b, c, and d are: 0.335, 20.158, 20.00290, and 0.00146, respectively. (b) Plot of Li/V ratio in the melt relative to the source Li/V ratio as a function of melt fraction. Curves were calculated at different values of oxygen fugacity, labelled in log units with respect to the fayalite-magnetite-quartz buffer, using the vanadium partitioning relations shown in Fig. 5a, combined with a simple batch melting model that assumes phase proportions change during melting as documented by Baker and Stolper (1994; see text). Mineral/melt partition coefficients for Li are assumed to be independent of fO2 and DLi for spinel is assumed to be zero. The melt fractions corresponding to spinel and clinopyroxene exhaustion from the residue are demarcated on the diagram.

plete lack of garnet-melt partitioning data as a function of fO2, we cannot evaluate the effect of this mineral on Li/V fractionation during melting in the garnet lherzolite facies. However, because of the relatively large spinel-melt partition coefficients for V over a wide range of fO2, we expect that the signature of garnet-induced Li/V fractionation will be largely subdued during the passage of melts through the spinel lherzolite facies. In

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terms of Li partitioning, we have assumed that DLi is independent of fO2 and used the partitioning data from the experiment involving the Baker/Stolper melt composition. DLi for spinel was assumed to be zero. Owing to the overall similarity in the Ds for Li for the constituent minerals of the lherzolite sourcerock, the bulk partition coefficient for this element is essentially constant (;0.3) up to very high degrees of melting. In contrast, the bulk D for V varies quite strongly with melt fraction, largely as a result of the control exerted by clinopyroxene and spinel. As a consequence, the Li/V ratio in a derived melt is equivalent to that of its source rock at different oxygen fugacites, depending on the extent of melting, which directly corresponds to the abundance of clinopyroxene and spinel in the residue. Specifically, the Li/V ratio in a melt is equal to that of its source at melt fractions ranging from 0.01 (FMQ 1 2) to 0.23 (FMQ-1; Fig. 5b), thus, without a priori knowledge of the degree of melting intrinsic to MORB, the observed constancy of the Li/V ratio in MORB does not uniquely constrain the MORB-source fO2. Independent estimates of the fO2 intrinsic to the MORB source, based on Fe21/Fe31 ratios in MORB glasses (Christie et al., 1986) and spinel barometry in abyssal peridotites (Bryndzia and Wood, 1990), indicate that values of around FMQ-1 (61) are appropriate. Additional modeling of the entire MORB glass array indicates that the relatively small amount of dispersion in MORB Li/V is likely to be a consequence of two factors: (1) the absolute variation in fO2 inferred for MORB sources being small and (2) the total range of melting being relatively limited, but more importantly that melts be produced from residues close to spinel and/or clinopyroxene exhaustion. The second factor arises because, at the fO2s indicated for MORB sources, the difference between DLi and DV is smallest and deviates the least from unity over this melting interval (Fig. 5b). Possible support for MORB melt production being close to spinel and/or clinopyroxene exhaustion would be the low modal abundances of these phases in abyssal peridotites (Dick et al., 1984). It should be noted, however, that the absolute degree of melting required to produce a specific Li/V ratio is somewhat dependent on the starting mineral modes. For example, with less initial clinopyroxene (i.e., ,17%), the same Li/V ratio would be produced at a lower degree of melting. In contrast to the rather limited range in the Li/V ratio of MORB glasses, values exhibited by whole-rock analyses of IABs show considerable dispersion (Fig. 6a). Li/V ratios in IABs range from ;0.01 (Izu suite) to ;0.1 (Aleutian suite) and are uncorrelated with other indicators of crystallization, such as TiO2, NiO, or MgO. Excellent positive correlations exist, however, between Li/V and Li/Yb (Fig. 6b), and all of the arc suites have Li/Yb that exceed MORB values, including samples having Li/V ratios # 0.02. Inasmuch as elevated Li/Yb is considered to be an indicator of the introduction of a Li-rich, slabderived component to the sources of IABs (cf. Ryan and Langmuir, 1987), the observed dispersion in Li/V ratios can be interpretted in a similar manner. However, it is notable that Li/V in some IAB suites is actually lower than the canonical MORB value of ;0.02 (i.e., the Izu and Bonin suites). Two obvious possibilities exist to explain this behaviour, the first being that the Li/V ratios of the Izu and Bonin sources are inherently lower than that for the MORB source or second that the fO2 during melting in the Izu or Bonin sources is higher

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at a higher fO2. Thus, with the assumption that the Li/V ratios in IAB sources are at least as high as in the MORB source, it can be concluded that IAB suites with Li/V ratios lower than 0.02 are formed under more oxidising conditions than MORB. The process of Li introduction into the Izu and Bonin arc sources (metasomatism by slab-derived fluid?) may, therefore, be considered to be coupled to oxidation of these sources as well. Arc suites with Li/V that exceed MORB values may also be produced under relatively oxidising conditions, but the reduction in Li/V to below MORB values that would allow us to distinguish this effect is overwhelmed by the addition of Li to those sources. Our inferences regarding IAB source oxidation state are also consistent with fO2s calculated from Fe21/ Fe31 ratios in calc-alkaline volcanic rocks summarized by Carmichael (1991), with values ranging from ;FMQ to FMQ13. 4. SUMMARY AND CONCLUSIONS

Fig. 6. (a) Abundances of Li as a function of vanadium for IAB whole-rocks with greater than 5 wt% MgO from the data of Ryan and Langmuir (1987). Note the large dispersion in the IAB dataset compared to MORB (Li/V 5 0.02) and the presence of arc lavas with Li/V ratios lower than then MORB array. Lines denoting Li/V 5 0.30, 0.20, and 0.10 are shown for reference. (b) Ratio of Li/Yb as a function of Li/V for the same data-set portrayed in Fig. 6a. Li/Yb and Li/V in MORB are essentially constant and have values of 1.7 and 0.02, respectively (Ryan and Langmuir, 1987). Elevated Li/Yb ratios in IAB relative to MORB are interpretted to be the result of the influx of a Li-rich component into the IAB source region (Ryan and Langmuir, 1987). The positive correlation between Li/Yb and Li/V suggests the dispersion in Li/V in the IAB dataset can also be accounted for by variations in the Li content of IAB sources.

than that of the MORB source. In terms of the first possibility, the MORB glass array suggests that DLi 5 DV during MORB genesis, thus Li/V in the MORB-source will not change during the melting process. One may expect, therefore, that if the mantle wedge at the onset of subduction is a product of MORB genesis, then Li/V would not differ greatly from ;0.02. Moreover, the observation that the Li/Yb ratio exceeds MORBvalues in all of the IAB suites considered suggests that these arc sources have become Li-rich with respect to the MORB-source. Therefore, it seems unlikely that the low Li/V ratio in the Izu and Bonin arc suites are a result of inherently low ratios in their sources. The effect of fO2 on melt Li/V is best illustrated by considering the model results presented in Fig. 5b. Such results indicate that, at any degree of melting of the same source composition, the Li/V ratio will be lower in the melt produced

A suite of partition coefficients has been measured for B, Be, and Li between silicate melts and the major silicate minerals present during the production and initial crystallization of primitive basaltic magmas. Partition coefficients for olivine and clinopyroxene have been found to vary as a function of mineral and melt composition, and, with the exception of B partitioning into clinopyroxene, this variation can be modeled using simple exchange reactions involving the trace element and a substituent element, such as Na, Mg, or Al. Combining our partitioning data with simple models of melting and solidification has allowed us to evaluate how accurately element ratios such as B/Be, B/K, B/Nb, Be/Nd, Li/V, and Li/Yb in primitive magmas reflect that of their source. These models confirm that the library of element ratios that define the distinct nature of IAB and OIB sources relative to MORB should be expanded to include B/Be, B/Nb, and Li/Yb. We have also shown that the observed constancy of the Li/V ratio in MORB does not uniquely constrain the fO2 of the MORB source without a priori knowledge of the degree of melting. However, the small amount of dispersion in MORB Li/V is consistent with the small variation in source-region fO2 inferred for MORB and is likely to also reflect degrees of melting close to clinopyroxene exhaustion. Finally, the large dispersion in Li/V in IAB suites is distinct from MORB and is likely to reflect the open system nature of IAB sources. Specifically, Li/V ratios lower than MORB-values may reflect more oxidising conditions in IAB sources with similar or higher Li/V as the MORB source, thus indicating that the addition of Li to IAB sources may also be coupled to oxidation. Acknowledgments—Ian Hutcheon and Adam Kent are thanked for their advice and assistance concerning SIMS analysis. Reviews by Jeff Ryan and Tom Latourrette helped to improve the clarity of the manuscript, especially Ryan’s comments regarding Li-V systematics. Support for this research was provided by the Geosciences Research Program of the U.S. Department of Energy’s Office of Basic Energy Research and by LLNLs branch of the Institute for Geophysics and Planetary Physics. E. Neroda’s research at LLNL was made possible by the USDOEs Science and Engineering Research Semester administered by LLNL. Work at LLNL was performed under the auspices of the U.S. Department of Energy under contract No. W-7405-Eng-48. J. Brenan also acknowledges funding from the Natural Sciences and Engineering Research Council of Canada under operating grant OGP 0194228.

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