Partitioning of U, Pb, Cs, Yb, Hf, Re and Os between chromian diopsidic pyroxene and haplobasaltic liquid

Chemical Geology, 62 (1987) 191-208 Elsevier Science Publishers B.V., Amsterdam - - Printed in The Netherlands

191

[1]

PARTITIONING OF U, Pb, Cs, Yb, Hf, Re AND Os BETWEEN CHROMIAN DIOPSIDIC PYROXENE AND HAPLOBASALTIC LIQUID E. BRUCE WATSON .1, DALILA BEN O T H M A N .2, JEAN-MARC LUCK .2 and A L B R E C H T W. H O F M A N N Max-Planck-Institut fi~r Chemie, Abteilung Geochemie, D-6500 Mainz (Federal Republic of Germany) (Received July 23, 1986; revised and accepted September 11, 1986 )

Abstract Watson, E.B., Ben Othman, D., Luck, J-M. and Hofmann, A.W., 1987. Partitioning of U, Pb, Cs, Yb, Hf, Re and Os between chromian diopsidic pyroxene and haplobasaltic liquid. Chem. Geol., 62: 191-208. Partition coefficients for U, Pb, Cs, Yb, Hf, Re and Os between chromian diopsidic pyroxene and haplobasaltic liquid were experimentally determined at 1275°C and 1-atm. pressure. Unusually large ( ~ 300 mg) charges were utilized in order to facilitate separation of the pyroxenes by hand-picking; subsequent purification by rinsing with dilute HF yielded crystals demonstrably free of glass. Trace-element contents of pyroxene and glass separates were determined either by measuring the gamma activity of doped-in radiotracers ("~7Cs, l~gyb, lSlHf), or by isotopedilution mass spectrometry (U, Pb, Re, Os). The resulting partition coefficients are reproducible at various levels of added trace components, and have the following values: U, 0.0003; Pb, 0.01; Cs, 0.001; Yb, 0.19; Hf, 0.36; Re, ~0.04; Os, ~ 0.08 (the U value may be unrealistically low due to the presence of U 6+ in our charges ). Together with previous experimental studies of U partitioning, our results confirm the long-standing belief that U is less compatible in mantle residues than is Pb, and in doing so reemphasize the paradox surrounding the poor correlation of U/Pb with Rb/Sr and Nd/Sm in mid-ocean-ridge basalts (MORB). Somewhat contrary to expectations, Hf is more compatible in clinopyroxene than are the heavy rare earths. This suggests that the positive correlation of Sm/Nd and Lu/Hf in the suboceanic mantle (implied by the isotopic coherence of Hf and Nd in MORB ) is due not to clinopyroxene control during the pre-MORB depletion event, but to a combination of clinopyroxene and garnet or clinopyroxene and orthopyroxene.

1. I n t r o d u c t i o n

A principal limitation to the further refinement of trace-element and isotopic models of *1Permanent address: Department of Geology, Rensselaer Polytechnic Institute, Troy, NY 12180-3590, U.S.A. *2Present address: Laboratoire de G~ochimie et Cosmochimie, Institut de Physique du Globe et Ddpartement des Sciences de la Terre, Universitds de Paris-6 et Paris-7, 4, Place Jussieu, 75005 Paris, France. 0009-2541/87/$03.50

crust/mantle evolution is knowledge of the mineral/liquid partition coefficients operative during mantle melting. Values for some elements (e.g., Sr) are fairly well known, generally as a result of the combined efforts of experimentalists and natural-rock analysts. The partitioning behavior of other elements (e.g., Pb), however, is poorly understood, in most cases due to difficulties associated with

© 1987 Elsevier Science Publishers B.V.

192

experimental determinations and/or lack of necessary constraints from natural systems. Thus, there is a clear need for information on the partitioning of elements between residual minerals of the mantle and derivative basaltic liquids; this is especially true of some isotopic systems, for which not only the specific values of partition coefficients, but also the general parent/daughter fractionation effects, are poorly known. It is for this reason that we undertook a broadbased experimental study of mineral/liquid partitioning, focusing initially upon diopsidic clinopyroxene in equilibrium with haplobasaltic liquid. The choice of this particular mineral was made because residual clinopyroxene has broad potential to affect trace-element ratios of partial melts in the mantle. The specific elements of interest in this study were U, Pb, Lu, Hf, Re, Os and Cs. The overall objectives were two-fold: (1) to acquire new partitioning data applicable primarily to parent/daughter fractionation in important isotopic systems; and ( 2 ) to develop and routinize experimental and analytical techniques that have been previously but not widely used in partitioning studies. These involve separation of crystals from glass and subsequent analysis of trace elements in both phases by mass- or gamma-ray spectrometry.

Experimental methods 2.1. Choice of system and conditions The base composition for all experiments was selected by reference to the system diopside-albite-anorthite, which is commonly used for basalt analog, clinopyroxene/liquid partitioning studies (e.g., Lindstrom and Weill, 1978; Ray et al., 1983). Although a small measure of geological realism is lost by use of an Fe-free system, the advantages are numerous. Most importantly, the absence of Fe from the charges allows the use of sealed Pt containers, w~i~h

are necessary to confine volatile elements of interest such as Pb, Re, Os and Cs. The Fe-free aspect of the system was the only major-element compositional simplification: From an initial composition similar to "G" of Ray et al. (1983), we 'removed' some SiO2 and added ~ 1% each of TiO~, A120~ and Na20 to bring the concentrations of most major oxides close to those found in natural basalts. The SiO2 content was kept deliberately above 50 wt.% in order to ensure good quenches of the large-volume charges necessitated by our analytical methods ( see p. 195). This difference from most natural basalts is not viewed as serious, because melt composition effects on partition coefficients are not large for Si02 contents < 55 wt.% (Watson, 1977; Watson and Green, 1981; Ray et al., 1983). A final 'additive' included in our base composition was 0.3 wt.% Cr20~. This component was included not only for geological realism (being a trivalent minor cation capable of charge balancing trace-element substitutions ), but also with a view toward simplifying the hand-picking and handling steps in our analytical procedure by coloring the pyroxene crystals. The final bulk composition of starting material used for the partitioning experiments was (in wt.%): SiO2, 51.84; TiO2, 1.00; A12Q, 12.16; Cr2Q, 0.30; MgO, 11.17; CaO, 20.94; and Na20, 2.59. This resembles a basalt in which FeO has been replaced by CaO. Although the data acquired in this study are of potential relevance primarily to mantle proceses, all experiments were conducted at atmospheric pressure. This simplification was made because we needed a combination of large sample volume, .oxidizing conditions, and extremely accurate knowledge and control of temperature. These characteristics are not readily realized in a piston-cylinder apparatus, especially simultaneously. Our bulk composition was such that a suitable proportion of crystals to liquid could be obtained at 1275 ° C; thus, although the pressure of the experiments was low, the temperature was appropriate to that

193

prevailing during partial fusion of the mantle. Given the compositional similarity of the resulting clinopyroxene crystals to natural chromian diopsides (see p. 200 and Table IV), and provided that pressure effects on partition coefficients are not large, our results are relevant to mantle melting despite the discrepancy in pressure. [ In this context, we note that T.H. Gr_een and Pearson (1985) observed an approximately two-fold increase in average clinopyroxene/liquid partition coefficient for the R E E over a pressure change of 7.5-20 kbar. This change is coupled with variations in pyroxene composition, so it is not possible to separate out an intrinsic pressure effect. Watson and Green (1981) saw no change in apatite/liquid R E E partition coefficients over the same pressure increase. ]

2.2. Detailed procedures 2.2.1. Preparation of starting materials A 5-g batch of glass with the composition described above was prepared by mixing and grinding reagent-grade oxides and carbonates under alcohol in an agate mortar. W h e n a grain size of < bl3 z m was obtained, the mixture was dried, placed in a P t crucible, and heated slowly to 1000 ° C in a muffle furnace to drive off CO2. The resulting 'briquet' was re-ground and the powder placed in a P t crucible for a 4-hr. fusion at 1450°C in a molybdenum-wound, verticaltube furnace. The resulting clear, green glass was re-powdered and subjected to an additional fusion at 1450 ° C for 5 hr. A final crushing produced a coarse glass powder ( hereafter referred to as 'base glass') to which all additions of trace elements were made. Trace elements were added independently or in pairs to ~ 300-mg splits of base glass at levels appropriate to the analytical method to be used. U and Pb were added together, and because their concentrations were determined by isotope dilution, fairly high doping levels were required to give sufficient total masses of these elements in the 2-5 mg of clean crystals recoverable from

each charge. U was added as U~Os (N.B.S. standard reference material 950b) and Pb as common PbO. Powders of these oxides (weighed on a Mettler ® ME22 microbalance) were combined with ~ 300 mg of the base glass and pre-mixed thoroughly by grinding under alcohol in an agate mortar. Four separate U - P b experiments were done, with the following bulk concentrations: ~ 80 p p m U + ~ 200 ppm Pb; ~ 250 ppm U + ~ 600 ppm Pb; -~ 900 ppm U + ~ 2000 ppm Pb; and U only at a concentration of ~ 2 5 0 0 ppm. These four starting mixtures also contained ~ 0.01 pCi l:~TCs, added as CsC1 solution, for the purpose of monitoring the purity of crystals separates obtained from the run products (see pp. 196 and 197). Cs partition coefficients were, of course, readily obtained as a by-product. Experiments pertinent to the H f - L u isotopic system were carried out using gamma-emitting radioisotopes of appropriate elements. A 90 : 10 mixture of lSlHf and 17SHf was commercially available (Amersham Buchler G.m.b.H. & Co. K.G.), but because no suitable isotope of Ln exists, lSgYb was used instead (all previous studies of clinopyroxene/liquid partitioning of' heavy rare-earth elements ( H R E E ) indicate that Lu and Yb have indistinguishable partition coefficients). Four separate experiments were done, two each involving 169Yb and ls~ ÷ 17SHf' and with 1 and 10 ppm of stable Yb and Hf. The low-level doping of stable components was accomplished by pre-synthesizing glasses containing 100 ppm of the elements of' interest and diluting these with the undoped base glass. The radioactive isotopes were added in dilute acid solution at a level of ~ 0.01 uCi per charge. Because of the inherent instability of most Os compounds (and to a lesser extent those of R e ) , starting materials containing these elements were the most difficult to prepare. The only practical approach was to mix a weak acid solution (1 N HC1) containing Re and Os as oxide ( R e O j ) and chloride (OsC12 ) anions, respectively. About 50 ml of the acid solution

194

were quantitatively pipetted into a mortar containing ~ 300 mg of the base glass, and the mixture was gently dried in air at 110°C. After thorough mixing, the material was transferred to an open A1 boat and heated slowly, in vacuo, to 200-250°C. In the absence of oxygen, the formation of the extremely volatile oxide of Os ( O s Q ) was suppressed, and hydrate precipitares from the acid solution could be degassed to an extent sufficient for closed-capsule experimentation. Using this method, a preliminary run was made in which no crystals were grown, and the resulting glass was analyzed in order to assess the possibility of Re and/or Os loss to the P t container. About half the initial contents of both elements were retained in the melt, so we proceeded to perform a real experiment with bulk-system concentration of ~ 5 ppm Re and ~ 3 ppm Os. A summary of information on the identities and concentrations of doped-in trace elements and radioactive isotopes is included in Table I.

2.2.2. Execution of experiments The starting material for each experiment consisted of finely-ground glass powder containing added trace components as described

above. About 300 mg of this powder were packed into 2.5-cm lengths of 4.5 I.D. × 5 mm O.D. P t tubing that had been pre-annealed and closed at one end with a three-corner weld. The tubes were then gently flat-crimped to remove as much air space as possible, placed in a drying oven (110°C) for at least 6 hr., and quickly sealed with a single straight weld. This procedure yielded a ~ 90% success rate for capsule integrity during ~ l - w e e k runs. [The R e - O s experiments were the most problematic with respect to bursting of capsules, due to the presence on the glass powder of volatile residues that could not be completely dried off without concurrent loss of Os (see Section 2.2.1 ). ] It should be noted that maintaining capsule integrity is crucial to the success of most of the experiments described in this paper, not only because some of the trace elements of interest (e.g., Pb, Os ) are extremely volatile, but also because loss of Na changes the pyroxene liquidus enough to affect crystal growth. An experimental run was initiated by lowering the P t capsule into a vertical tube furnace already at temperature. The first portion of each run consisted of a constant-temperature (1300 ° or 1310 ° C) superliquidus interval of 0.5-2-day

TABLE I S u m m a r y of i n f o r m a t i o n on a d d e d t r a c e e l e m e n t s a n d r a d i o t r a c e r s Experiment No.

Activity (#Ci)

Element concentration*' (ppm) U

3 4 6 7

255 82

11 12 13 14 15

908

Pb (246) (78)

(884)

2,530 ( 2 , 2 9 7 )

625 202

(585) (173)

2,200 ( 2 , 3 0 5 )

Yb 0.95 9.60 0.95 9.60 _

Hf

Cs

Re

0s

13VCs

l~9yb

ISlHf

(,2) (,2)

_ _ -

_ _

0.01 0.01

-

0.01

0.01 0.01 -

3

0.01 0.01

0.01 0.01 -

0.94 9.50 (,2) 0.94 9.50

(,2) (*~)

5

-

Half-lives: 137Cs = 30 yr.; '69Yb = 31 days; l S ' H f = 45 days; 'VSHf= 70 da ys . Notes: (1) '69yb is c a r r i e r - f r e e ; (2) is, +,VSHf is o n l y ~ 0.1% of t o t a l Hf; c a r r i e r s t i l l a m o u n t s t o o n l y ~ 3 p p b i n c h a r g e . *~First v a l u e is w e i g h e d - i n a m o u n t ; n u m b e r i n p a r e n t h e s e s is c o n c e n t r a t i o n d e t e r m i n e d b y a n a l y s i s of c h a r g e a f t e r e x p e r i m e n t . *2No s t a b l e Cs w a s a d d e d , a l t h o u g h t h e r e m a y h a v e b e e n s o m e Cs i m p u r i t y i n t h e r e a g e n t - g r a d e s t a r t i n g m a t e r i a l s . T h e c o n c e n t r a t i o n of ~37Cs w a s ~ 0.5 ppb.

195

duration, designed to ensure diffusional homogenization of doped-in trace elements and isotopes. Following this step, the temperature was abruptly dropped to 1288 ° C, the predetermined liquidus of the base glass composition. After an additional constant temperature period of 12-16 hr., a cooling interval was initiated (I°C hr. -1) to bring the sample down to 1275°C, which was maintained for 3-4 days before quenching the capsule in water. The time-temperature history of a typical experimental run is shown schematically in Fig. 1, and the actual durations of all intervals are summarized in Table II. The 5-step procedure resulted in good nucleation and growth of elongate, inclusion-free clinopyroxene crystals of generally square cross section, ~ 3 : 1-5 : 1 aspect ratio, and 100-500-~m length (Fig. 2). Due to the presence of Cr, these crystals were bright green and commonly visible to the naked eye, set in a matrix of clear, pale-green glass. The proportion of crystals appeared to vary slightly among runs, presumably because of the sometimes substantial additives (e.g., 2000 ppm Pb). A value of 3-5%, estimated visually and by major-element mass balance (see Section 3.1 ), is typical of all runs.

2.2.3. Separation of crystals from glass An initial separate of nearly-pure clinopyroxene crystals was obtained by simple hand-

picking under a binocular microscope with fine tweezers. Because the crystals spanned a considerable size range, the picking was most effectively done in successive stages, starting with coarsely-crushed sample ( ~ l-mm grain size) and proceeding to progressively finer material. The first round of picking yielded large, nearlyintact crystals with a small amount of attached glass; as the remaining fraction was crushed to finer sizes, the smaller crystals were recoverable. On average, the time required for recovery of most of the crystals present in a charge was 6-10 hr. Although the separation was expedited by the bright green color of the Cr-bearing pyroxenes, it was eventually found even more advantageous to do the picking between crossed polarizers, where even very small crystals were readily visible. Because of the low percentage of' crystals present, it was relatively easy, in the course of 2-3 hr., to obtain 30-80-mg separates of pure glass. Once the 'first-order' separation of crystals from glass had been accomplished by handpicking, all that remained was to remove the small amounts of attached glass by differential acid dissolution [Shimizu (1974) used this technique alone to obtain pure crystal separates in his clinopyroxene/liquid partitioning studies.] Although other types of acid might also be suitable, we found a ~ 5% HF solution to work extremely well, dissolving the glass very

T A B L E II S u m m a r y of t i m e - t e m p e r a t u r e information for p a r t i t i o n i n g experiments ( refer to text a n d Fig. 1 ) Run No.

3 4 6 7 11

12 13 14 15

D u r a t i o n of experiment steps (hr.) homogenization 1300 ° or 1310°C

nucleation 1288°C

crystal growth

constant 1275°C

12 12 41 41 22 46 46 46 45

12 12 16 16 16 16 16 16 16

13 13 13 13 13 13 13 13 13

83 83 67 67 43 41 41 41 48

Total duration (hr.)

120 120 137 137 94 115 115 115 121

196

1310 - - - ~

homogenization

1300

? q}

1290 .......... nucleation ~

1280

V 102 ' 'o Jo

pro-determinedIiqu/dus of base g/ass

ow~

constant T

5060708090100 . . . . . . . .

quench 120

t, /4.0

time, h Fig. 1. Typical time-temperature profile of experimental runs. See text for discussion.

rapidly and the crystals quite slowly. In order to minimize the problem of handling small amounts of tiny grains, we placed each crystal separate in a 'tea-bag' fabricated from ~ 20-/~m Nylon ® mesh. During dissolution, this bag was suspended in a beaker of distilled acid solution, so that the grains were in contact with acid but could readily be separated again for washing, weighing, and/or gamma-ray counting. The acid treatments were done for 1-5 hr. at a time, at which interval the crystals were rinsed ultrasonically in doubly-distilled water, dried at 110°C, and counted to determine the gamma activity of the added radiotracer. This activity dropped very rapidly in the first 3-5 hr. of acid treatment, and only slowly over much longer treatment times, suggesting that the contaminating glass was quickly removed and the subsequent slow decrease in activity was due to dissolution of the crystals themselves. This hypothesis was rigorously tested in the Hf experiments (run Nos. 12 and 13) by weighing the crystal separates after each acid treatment interval, and establishing that the count rate

Fig. 2. Reflected-light photomicrograph of pyroxene crystals from experiment No. 12. Crystals shown have been separated from glass by hand-picking, treated with dilute HF to remove small amounts of attached glass (see text), and mounted in epoxy for polishing. Occasional etch pits are visible along pre-existing parting surfaces.

400

2.77

~ 500 \

~1.96 I~ • ~ . g J

~

3 1__1__

200

k' - - m

440

tre

1.66

l_

#2° ~ - - - e -

4.25 3.95 ~59

2.35 i

0

;

i

I0

i

15 time, h

i

J

20

25

34

Fig. 3. 18~÷]TSHf activity vs. HF rinse time for pyroxene crystal separates of runs 12 and 13. Numbers next to the plotted points denote milligrams of crystals remaining. See text for discussion.

per milligram of remaining sample did not change after the first 5 or so hours of exposure to the HF (Fig. 3). Ideally, this procedure would have been done for all experiments, but was not practical due to the hygroscopic nature of the

197

fine Nylon ® mesh in which the crystals were contained (and consequent error in any weight measurements made during the acid treatments). Periodically removing the crystals from the Nylon ® for weighing was out of the question, because some were inevitably lost in the handling. For the acid-treatment procedure in the Hf experiments, the weighing error was minimized by containing the crystals in a smaller-than-ideal Nylon ® mesh bag. The results of the ~s~+ ~75Hf activity vs. time studies (Fig. 3) are taken as a demonstration that all contaminating glass is removed from the pyroxene separates within 5 or 6 hr. This conclusion is especially significant in view of the fact that the minimum acid treatment time for all other experiments was 24 hr., with some (runs 3 and 4) ranging up to 60 hr. In retrospect, then, it was not really necessary to include a highly incompatible radiotracer (137Cs) to establish the purity of pyroxene separates to be analyzed by mass-spectrometric techniques. It was reassuring, however, to see the ~37Cs count rate drop virtually to zero within the first few hours of acid treatment. Moreover, subsequent long-term counting of this isotope also yielded new values for Cs partitioning that show interexperiment consistency and general agreement with results of previous studies (see Section 3.2.1). As a last note in this section, it should be added that the final weights of crystals and glasses used to compute concentrations were determined by weighing the samples in small polyethylene vials. These vials were counted directly for gamma-activity measurements. In the case of samples to be run by isotope dilution, the material was transferred from the polyethylene vials to Teflon ® beakers for acid dissolution. 2.3. Analysis 2.3.1. Ytterbium, hafnium and cesium Once the pyroxene and glass separates were obtained and the crystals cleaned of all adhering glass ( see Section 2.2.3 ), the determination

of l~Yb, lsl+175Hf and 137Cs partition coefficients was a simple matter of gamma-ray counting. For Hf and Yb, whose concentrations were reasonably high in both phases, a NaI detector and 2K multichannel analyzer were used. The background count rate was reduced to a minimum by enclosing the detector in a Pb box with 10-cm thick walls. 1GgYband the combined isotopes of Hf have gammas of several energies. Four 169Ybpeaks in the ~ 0.1-0.3-MeV range were counted, and for the Hf-doped samples, three ~S~Hf peaks in the ~0.1-0.4-MeV range and the single 175Hf peak at 0.34 MeV were used. Background measurements were made both by interpolation under the peaks in the sample spectra and by counting blanks. Gamma intensities were taken as the integrated areas under the spectral peaks. In the case of 13~Cs, the count rate on the 0.66MeV peak for the pyroxene separates was well below the detection limit of the NaI system, so we resorted to extremely long count times (up to 2 days) using a low-background Ge(Li) detector. The uncertainties reported in Table V for partition coefficients determined by the radiotracer method reflect the counting error associated with gamma intensities. 2.3.2. Lead and uranium As noted previously, the concentrations of U and Pb in the crystal and glass separates were determined by isotope dilution. The samples were first dissolved in a mixture of distilled HF and HNO3 in closed Teflon ® beakers. U and Pb spiking was done at the time of dissolution. All Teflon ® labware was thoroughly pre-cleaned several times with doubly-distilled acids ( H N Q , HC1 and HBr). Pb and U were separated and purified on anion-exchange resins ( A G I × 8 ®, 25 and 10 ml) with HBr (for Pb) and an HF-HC1 mixture (for U). The isotopic analyses were performed on a Finnigan ® MAT 261 mass spectrometer. 2.3.3. Rhenium and osmium Re and Os concentrations were also determined by isotope-dilution mass spectrometry,

198

using the analytical procedure of Luck et al. (1980) adapted for small Re and Os quantities in the study of silicates. Because of the very refractory behavior of both elements in thermal-ionization mass spectrometry, a Cameca ® IMS 3lion microprobe was used for the isotopic analysis. There was no existing basis for making a priori estimates of Re and Os partition coefficients, so the crystal separate from run 15 was split in half to provide a second chance at analysis if a totally inappropriate spike were used in the first attempt. The amounts of pyroxene and glass analyzed were 0.760 and 1.600 mg, respectively. The large uncertainties in the final concentrations, particularly for Os ( see Table V ) do not reflect counting statistics, but 'overspiking' of the runs.

2.4. The case for equilibrium Although none of the experiments reported here has been reversed, there are several lines of evidence supporting the general attainment of equilibrium. First, it should be noted that the proportion of crystals was very small (3-5%) so that systematic zoning of pyroxenes during growth, due to changing trace-element concentration in the melt reservoir, is not likely to have occurred (this conclusion is substantiated below for the case of Hf). Additional lines of evidence suggestive of a close aproach to equilibrium are that: (1) Ray et al. (1983) successfully reversed partitioning experiments of closely similar design; and (2) Lindstrom (1983) showed that, even for diopsides grown at rates much higher than those of the present study, deviations from equilibrium partitioning behavior of the REE are not severe ( i.e. the apparent partition coefficients are at worst within a factor of 2 of the equilibrium values). A final point regarding the attainment of equilibrium in our experiments is that, should deviations from equilibrium exist, these must be such that the apparent partition coefficients are higher than the equilibrium values. This is always true of "crystal-growth" partition coef-

ficients for incompatible elements (see, e.g., Albar~de and Bottinga, 1972). In this respect, our partition coefficients can be viewed as maximum values. 3. R e s u l t s and discussion

3.1. Major elements The concentrations of major and minor elements in the clinopyroxenes and glasses were determined by energy-dispersive electron microprobe analysis, using the ETEC facility at the State University of New York at Albany. Because the major-element bulk compositions and temperatures of all experiments were the same, there is relatively little variation in the compositions of the run products. Intracharge variation in glass composition, as gauged by comparison of 5- or 6-analysis spots, was not detectable. Intercharge variation among the nine reported experiments can be judged from the standard deviations given with the overall average glass (melt) composition listed in Table III. Because the compositions of the clinopyroxenes appear to vary slightly from one charge to the next, the mean composition for each run is listed individually in Table III. The apparent interexperiment variability shows up primarily in the A1203 and Cr2Q contents, which also show, relative to other major oxides, significant variability among crystals in a single charge ( see standard deviations given in Table II1). In an attempt to ascertain whether this variation is systematic within individual crystals, we traversed a large grain both parallel and perpendicular to the long axis. Scans consisting of 8 - 1 0 spots revealed no radial heterogeneity whatsoever. Additional tests on other grains led to the conclusion that the smaller crystals in a given charge sometimes have higher A1 and Cr contents than is typical of the overall population. This particular non-uniformity apears to be the primary contributing factor to the large standard deviations associated with analyses of

199

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d d d d ~ d d d d e4 M ~ ,,~ d

o

o

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o

o

o

o

o

o

o

o

~ d d o d d

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~ d d ~ d d ~

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d d M M ~ d

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d d ~ d d d d

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0

200 pyroxenes from a single experiment. Although the average compositions of crystals from the nine runs have not been shown to be identical, the apparent differences are not outside of the 2a variation attributable to intracharge heterogeneity. The average clinopyroxene and glass analyses can be used in mass-balance calculations to estimate typical proportions of crystals and liquid in the experiments. Calculations based upon A1203, Cr203, CaO and Na20 yield crystal contents of 3-7%, while TiO2 indicates 12%, MgO 19% and SiO2 - 6 % . Values of 3-7% are in good agreement with visual approximations, but 12% and 19% are too high and - 6 % is of course impossible. The discrepancies in the T i Q - and SiQ-based calculations are not surprising in view of the relative and absolute concentrations of these oxides in the two phases. The discrepancy in the MgO-based value is probably due to slightly inaccurate microprobe analyses for this oxide. Despite the Fe-free aspect of our experimental system and the low pressure at which the runs were made, the crystals obtained bear considerable resemblance to natural diopsidic clinopyroxenes. There is, for example, general similarity, especially in terms of A1 and Cr content, to clinopyroxenes from mantle peridotites (Table IV). Pyroxenes found in some Ca-,Mgrich oceanic basalts are even more closely similar. Perhaps the most important point to make in establishing relevance of the experimentally-determined partition coefficients is that our synthetic pyroxenes contain the same minor elements (e.g., Ti, Cr, A1, Na) that might be expected to participate in charge-balanced substitutions involving trace elements. 3.2. Trace elements 3.2.1. Cesium Four different experiments (Nos. 3, 4, 11 and 15) yielded clinopyroxene/liquid partition coefficients (Kd's) for Cs (Table V). Although the values from runs 3 a n d 4 ( ~4-10 -3) are in

good agreement with previous experimental determinations on similar pyroxenes (Shimizu, 1974) and with numbers inferred from analysis of crystal separates from basaltic rocks (Hart and Brooks, 1974), it should be noted that the uncertainties on these two values are extremely large. This is due strictly to gammacounting errors: When longer counting sessions became available on the Ge (Li) detector system, it was possible to greatly reduce the counting uncertainty, with the consequence of decreasing the Cs partition coefficient by a factor of ~ 3 (to 1.5" 10 -3 and 0.9" 10 ~ for runs 11 and 15, respectively). When the error estimates are taken into account, our data set is internally consistent, and indicates a clinopyroxene/liquid partition coefficient somewhat lower than previous estimates. 3.2.2. Ytterbium The partition coefficient for 169Yb in the experiments containing both i and 10 ppm stable Yb is 0.19 [runs 6 and 7 yield values that are the same within uncertainty ( see Table V) ]. This is significantly lower than many previously-determined clinopyroxene/liquid values for the HREE. In a study centered principally upon crystal separates from volcanic rocks, Fujimaki et al. (1984), for example, report numbers in the 0.3-0.7 range. Grutzeck et al. (1974) carried out experiments in the diopside-albite-anorthite system and obtained HREE partition coefficients of ~ 0.3 for diopside in equilibrium (at 1265°C) with a liquid similar to that of the present study. The best HREE values from the experimental work of Tanaka and Nishisawa (1975) on a natural tholeiite composition are 0.7-0.8, while Mysen (1978) measured a value for Tm in a simple, hydrous system of ~ 0.5, similar to the HREE experimental determinations of Nicholls and Harris (1980) for clinopyroxene in equilibrium with hydrous basaltic liquid at 1080°C. T.H. Green and Pearson (1985) report HREE values close to 1, but these pertain to quite silicic melts ( ~ 60 wt.% SiO2 on a hydrous basis ) and

201 T A B L E IV Selected analyses of natural diopsides resembling those of the present experimental study

SiO2 Ti02 A120~ Cr~O:~ FeO* MgO CaO Na20

1

2

3

4

5

6

53.7 0.3 3.1 0.7 3.1 17.8 21.5 0.2

55.4 < 0.03 2.19 1.57 1.44 16.80 22.00 1.26

54.7 0.26 1.69 0.78 4.78 18.20 17.10 1.46

55.2 0.11 1.28 0.98 2.84 19.50 18.70 0.98

53.5 0.54 5.43 2.05 1.51 14.00 18.70 3.32

54.3 < 0.03 2.13 1.96 1.30 16.70 21.20 2.00

Number of cations (based on 6 oxygens): Si Ti A1 Cr Fe Mg Ca Na

1.935 0.008 0.132 0.020 0.093 0.956 0.830 0.014

1.983 0.001 0.092 0.044 0.043 0.896 0.844 0.087

1.994 0.007 0.073 0.022 0.146 0.989 0.668 0.103

1.990 0.003 0.054 0.028 0.086 1.048 0.722 0.069

1.943 0.015 0.232 0.059 0.046 0.758 0.728 0.234

1.970 0.001 0.091 0.056 0.039 0.903 0.824 0.141

1 =diopside, mid-ocean-ridge basalt (Watson, 1976); 2=diopside, granular peridotite nodule in kimberlite (Boyd et al., 1976) ; 3,4 = diopsides from sheared lherzolite nodules in kimberlite (Boyd et al., 1976); 5 = diopside, granular garnet pyroxenite (Boyd et al., 1976) ; 6=diopside in lherzolite (Boyd et al., 1976).

generally low temperatures (900-1100 ° C ). Partition coefficients from other published studies compare quite favorably with our Yb value of 0.19. If one takes into account the fact that light REE (LREE) partition coefficients are uniformly lower than those for middle REE (MREE) and HREE (see Irving, 1978; Fujimaki et al., 1984), the D.H. Green et al. (1971) Gd value of 0.16 for augite coexisting with marebasalt liquid is consistent with the present Yb number. The simple-system study by Ray et al. (1983) is possibly more relevant - these workers obtained a Sm partition coefficient of 0.13 for diopside coexisting with their liquid "G", which does not differ greatly from our melt composition. Obviously, all comparisons of partition coefficient data sets should be made with the realization that inter-study differences in intensive and extensive parameters will affect the extent of. agreement. If the systems of interest are

restricted to generally basaltic or haplobasaltic melt compositions, temperature probably has the greatest effect on clinopyroxene/liquid partition coefficients for the REE (see Ray et al., 1983). Inasmuch as REE partition coefficients increase with decreasing temperature, it is possible that the natural rock cyrstal/matrix values of Fujimaki et al. (1984) are higher than ours because they reflect magmatic temperatures somewhat lower than our 1275 °C experiment temperature. Alternatively, the differences may derive simply from the fact our crystals are diopsidic and the Fujimaki et al. natural samples are all augites. It is unlikely that the discrepancy is due to differences in pressure of crystallization because: (1) most of the pyroxenes of Fujimaki et al. coexist with plagioclase and thus appear to be low-pressure minerals; and (2)'the work of T.H. Green and Pearson (1985) indicates that HREE partitioning is insensitive to changes in pressure.

202 TABLE V Summary of analytical information and partition coefficients (A) Gamma-ray spectrometry Experiment No.

Tracer

BG (c.p.m.)

3 4 6 7 11

l:~VCs 1:~7Cs 169Yb ~gYb 137Cs

12,2

,v~ ~lS~Hf ~7.,,+ ~S~Hf 17~t ~S~Hf ~w+ ~SlHf

27 33 37 28

13..~

~w+tSlHf ~w +~S~Hf 17~'+~SlHf ~7~~~S~Hf

27 33 37 28

15

~37Cs

0.46 0.53 109 109 0.5

0.5

Crystals

K,~ +_2a

Glass

(c.p.m. mg 1)

a ( , l l (%)

(c.p.m. mg-1)

lai,l~(%)

0.0426 0.0801 29.43 20.53 0.0415

58 33 2.6 4.9 16.3

10.22 21.97 146.55 115.08 28.05

1.8 0.9 0.44 0.41 0.4

(0.0042 ,+ 0.0040) *;~ (0.0036,+ 0.0024 ) *~ 0.201 _+0.010 0.178 -,-_0.018 0.0015 ,+ 0.0005

43.32 31.85 74.70 15.86

2.4 3.2 1.5 6.2

120.08 84.86 201.45 43.91

0.8 1.0 0.6 1.4

0.361 0.375 0.371 0.361

_+0.018 ,+ 0.026 _+0.012 + 0.046

56.89 42.18 99.05 21.27

4.2 5.5 2.5 10.7

173.78 122.75 292.22 62.39

0.4 0.5 0.5 0.7

0.327 0.344 0.339 0.341

,+ 0.028 _+0.038 ,+0.018 _+0.074

20.0

25.52

0.5

0.0009 ,+ 0.0004

0.0223

(B) Isotope-dilution mass spectrometry Experiment No.

Elements

3 3 4 4 11 11

U Pb U Pb U Pb U Re Os

14 15 15

Content (ppm)

Kd

crystals

glass

(.4) 5.29 *5 (.4) 1.54 *~ 0.437 *5 28.8 *5 0.543 *5 0.088 0.11

246.0 585 78.4 172.5 883.7 2,305 2,297 2.5 1.4

0.00904 0.00892 0.000495 0.0125 0.000236 0.035 _+0.005 0.075 ,+ 0.040

BG = background. * ~Gamma counting error. *2Four gammas of the Hf radiotracers were counted separately (see text) ; results are reported for all four in order of increasing gamma energy. *:~Due to large counting error on clinopyroxene separate, Kd-value (listed in parentheses) is marginally meaningful (see text and Kd'S for experiments 11 and 15). *4No determination possible. *~Blank correction applied.

3.2.3. Ha[nium Runs 12 and 13, containing I and 10 p p m stable Hf, respectively, yield an average Hf partition coefficient of 0.36 (Table V ) . Relatively few previous H f measurements are available for comparison, but it is probably valid to consider Zr as indistinguishable in its partitioning behavior. Even so, the data are limited, and

existing values show considerable variation. McCallum and Charette (1978), for example, obtained Zr partition coefficients ranging from 0.05 to 0.22 for clinopyroxene coexisting with a Ti-rich mare-basalt liquid analog at ~ 1100 ° C. In this case, the substantial range was attributed to major heterogeneity of the pyroxenes with respect to Ti and A1, and it was noted that

203 Zr content is correlated with these elements. Fujimaki et al. (1984) report clinopyroxene/ basaltic rock matrix concentration ratios for Zr and Hf ranging from 0.1 to 0.6, but give no compositional information on the pyroxenes other than that they are augites. (The highest concentration ratios are for a peralkaline basalt. ) Although our number is conveniently centered in the range given by Fujimaki et al., the agreement may be fortuitous considering the differences in composition (and probably in equilibration temperature) of the pyroxenes being compared. It should also be noted that the ratio of Hf to HREE partition coefficients is quite different: the highest ratio suggested by Fujimaki et al.'s data is ~ 1, while our data indicate a value of 1.9. 3.2.4. Uranium Although four experiments were doped with U, only two yielded pyroxene/liquid partition coefficients. The U contents of the crystals in experiments 3 and 4 could not be determined because 23~U/238U was equal to the spike ratio. Experiments 11 and 14 yielded Kd's of 4.95.10 -4 and 2.36-10 -4, respectively (Table V). The factor-of-2 difference between these two values can be explained simply on the basis of counting statistics, because a very small amount of 'overspiked' U from the crystal separates was analyzed. Because the partition coefficient is so small, a 'real' difference of this order could also conceivably result from a tiny amount of glass contaminant in one of the crystal separates, but this possibility is considered unlikely inasmuch as the Cs partition coefficients determined from identical separates were indistinguishable for runs 11 and 14. Under the circumstances, the agreement between the two measured U partition coefficients is considered quite good. Previous determinations of clinopyroxene/ liquid partition coefficients for U, all made by fission-track techniques, are considerably higher than ours. Seitz (1973) reported a range of 0.0008 to 0.0025, while Benjamin et al. (1978) obtained values between 0.0016 and 0.0047.

Neither of the previous studies is closely comparable to the present one in terms of experimental system or conditions. Perhaps more importantly, both previous sets of experiments were done in a piston-cylinder apparatus and therefore at lower oxygen fugacity than the present work, so the differences in partitioning behavior could be ascribed to differences in the oxidation state of U (Calas, 1979). Hexavalent ions may exist in glasses of the present study, which casts some doubt on the significance of our measured partition coefficients to natural melting processes. 3.2.5. Lead Partition coefficients for Pb were obtainable from all three experiments in which Pb was included, ie. runs 3, 4 and 1I. The values of 0.904"10 -2 , 0.892-10 2 and 1.250.10 2 (pertaining to crystals with 5.3, 1.5 and 28.8 ppm Pb, respectively) are in very good agreement. The partition coefficient for run 11 is high relative to the other two by an amount outside of the estimated analytical uncertainty; this slight discrepancy may be due to the difference in AlzO3 and Cr203 contents of the pyroxenes (which, in turn, may be attributed to the high Pb content of the melt in this experiment). To our knowledge, no previous measurements of clinopyroxene/liquid partitioning of Pb have been published, so there exists no basis for comparison and evaluation of the present numbers. In the context of simple ionic charge/ radius considerations, our numbers appear to be quite low - - Sr e+, for example, which is of similar size and charge to Pb 2+, has a considerably higher partition coefficient (~0.1; Shimizu, 1974; Ray et al., 1983). It should be borne in mind, however, that the nature of SrO and PbO bonds is quite different, the latter being much more covalent in character, so some difference in partitioning behavior of the two elements should be anticipated. A similar factor-of-10 difference exists between apatite/liquid partition coefficients for Sr ( ~ 1.5; Watson and Green, 1981) and for Pb (0.1-0.4; Watson

204

et al., 1985 ). The unexpectedly low clinopyroxene/liquid partition coefficient for Pb is definitely not attributable (as the case may be for U) to the presence of a higher oxidation state than is present in nature. Tetravalent Pb is unstable in air above 605°C (White and Ray, 1964). Most clinopyroxenes separated from natural mafic and ultramafic rocks have Pb contents considerably higher than would be predicted on the basis of our experimental determinations (Zartman and Tera, 1973; Kramers, 1977, 1979; GSpel et al., 1984). However, inasmuch as the analyses of natural samples were performed on clinopyroxenes separated from either ophiolite harzburgites (GSpel et al., 1984) or mantlederived peridotites, the concentrations do not necessarily result from equilibrium with basaltic liquid.

3.2.6. Rhenium and osmium With respect to Re and Os, this study is of a strictly reconnaissance nature, the main idea being to test the feasibility of performing the experiments and analyzing the separated phases by ion microprobe mass spectrometry. As noted in Section 3.1, neither element was quantitatively retained in the melt during the preliminary, superliquidus heating step. We consider it reasonable to assume, however, that most or all of the loss of the Pt container occurred before the pyroxene crystals were nucleated and grown, so no large error in the measured partition coefficients is introduced. In this respect, experimental studies of Re and Os partitioning that involve the use of Pt containers are not unlike the case of Ni - that is, there are obvious complications introduced because the element (s) of interest is soluble in the container, but this does not preclude the acquisition of reliable data (e.g., Hart and Davis, 1978). Our partition coefficients of ( 3.5 + 0.5 ) - 10- 2 for Re and ( 7.5 _+4). 10 -2 for Os are believed to be accurate to within a factor of 2.

4. Geochemical implications All seven elements treated in this study are moderately to highly incompatible. Cs was included primarily as a monitor of our experimental procedures and with the idea that the results might set a new lower limit to the partition coefficient determined in previous experiments (Shimizu, 1974). We have indeed confirmed the highly incompatible character anticipated on the basis both of Shimizu's experimental work and the behavior of Cs relative to other highly incompatible elements such as Rb and Ba in natural basaltic systems (e.g., Hofmann and White, 1983). Re and Os have also been shown to be moderately to highly incompatible in clinopyroxene, and it is now clear that this mineral is not responsible for the major fractionation of Re/Os observed between the Earth's crust and mantle or within a single komatiite flow (Luck and Arndt, 1985). Additional work on other minerals, especially olivine, is needed before the geochemical significance of the present results on Re and Os can be discussed. We therefore focus our discussion on the element pairs U-Pb and Lu-Hf. Of the major mineral phases believed to exist in the mantle - olivine, orthopyroxene, clinopyroxene and garnet - clinopyroxene is arguably the most likely to have caused the incompatible trace-element fractionation that accompanied the differentiation of the primitive mantle into the continental crust and depleted mantle. Clinopyroxene is the principal suspect because the characteristic REE patterns of mid-ocean-ridge basalts ( MORB ) - flat for the HREE and increasingly depleted from the middle to the LREE - mimic the pattern of clinopyroxene/liquid partition coefficients. If the MORB source was depleted by a partial melting event of small enough proportion to leave residual clinopyroxene, and if that clinopyroxene dominated the REE pattern of the bulk residue, then the pattern of the residue would be similar to those of MORB. Accordingly, the following discussion will test (and

205 ultimately discredit) the hypothesis that clinopyroxene has played a dominant role as a residual phase during the mantle melting events leading to the formation of either continental or oceanic crust. 4. I. U - P b

The fractionation of U/Pb ratios in the suboceanic mantle that is necessitated by the isotopic composition of Pb in different oceanic basalts was first discovered by Gast et al. (1964) and by Tatsumoto (1966). It has been abundantly borne out by more recent data (e.g., Sun, 1980 ), and presents a major geochemical puzzle (All~gre, 1969) because it is poorly correlated with the chemical fractionation inferred from the isotopic data of Sr, Nd and Hf. The most perplexing aspect of the Pb isotopic data is that they require increases of the U/Pb source ratios of both oceanic island basalts (OIB) and most MORB. This is contrary to the widespread expectation that U is more highly incompatible than Pb and that the MORB source should therefore have a lower-than-primitive U/Pb ratio. [This expectation has been based on the similarity in ionic parameters of Pb 2+ and Sr 2÷ and the known, only moderately incompatible behavior of Sr (Shimizu, 1974; Ray et al., 1983; see discussion in Section 3.2.5). The large size of U 4+ relative to moderately incompatible ions of Ti and Zr has suggested highly incompatible behavior for U]. Our results cannot directly confirm or reject this assessment because of the uncertainty in the oxidation state of U in our experiments, so any definitive discussion of the "Pb paradox" on the basis of experimental work must await additional measurements on U. It may be worth noting, however, that our partition coefficient for Pb, combined with those of Seitz (1973) and Benjamin et al. (1978) for U, suggests that U is indeed much less compatible than Pb. Even in the absence of a U partition coefficient directly comparable with our value for Pb, it is worth briefly exploring the question of whether residual clinopyroxene can affect the

abundance of Pb in oceanic basalts. A mantle peridotite containing ~ 10% clinopyroxene will have a bulk partition coefficient for Pb of K pb - 10 -3, provided there is no other residual mineral with a similar or larger partition coefficient. This would require melt fractions of 0.01 or less in order for the residual clinopyroxene to retain an appreciable amount of Pb and thus affect the Pb abundance in the melt relative to other incompatible elements. Such small melt fractions are generally believed to be implausibly small for MORB and tholeiitic OIB. However, complex melting models involving source regions with variable melt fractions (e.g., Magaritz and Hofmann, 1978; O'Hara, 1985; Galer and O'Nions, 1986) can be conceived such models would predict that the relative trace-element abundances in the melt are controlled by those portions of the source that contain very small melt fractions. A stronger argument against clinopyroxene control of Pb abundances in oceanic basalts derives from the correlation of Pb with Ce in these rocks (Hofmann et al., 1986; Newsom et al., 1986). This correlation indicates that Pb and Ce have similar bulk partition coefficients in natural systems producing oceanic basalts. This cannot b'e explained by clinopyroxene-melt equilibria because the Kd for Ce is at least a factor of 5 larger than our value for Pb [ cf. overall REE partitioning patterns of Fujimaki et al. (1984) with our value of 0.19 for Yb]. As discussed by Hofmann et al. (1986), the correlation of Pb and Ce does not extend to continental rocks, where Pb/Ce is dramatically higher than in oceanic basalts. This indicates that during the (presumably complex) processes forming the oceanic crust, Pb (in contrast with Ce ) behaved as a highly incompatible element; this is completely consistent with the low partition coefficient for Pb in the clinopyroxene-melt system. 4.2. L u - H [

The isotopic correlation of Hf and Nd in oceanic basalts (Patchett et al., 1984) requires

206 that the S m / N d and L u / H f element ratios have been positively correlated in the suboceanic mantle for a significant part of the Earth's history. There is every reason to believe that, in clinopyroxene/melt systems, the partition coefficient of Lu is very similar to that of Yb (e.g., Fujimaki et al., 1984). From this observation and from our experimental results follows the conclusion that K Lu < K w. There is also abundant evidence that, in clinopyroxene-melt systems, K~ m > KdNd. Therefore, during any kind of magmatic fractionation dominated by clinopyroxene, the S m / N d ratio (and therefore 14~Nd/144Nd) should be negatively correlated with the L u / H f ratio (and therefore with 176Hf/177Hf). As mentioned above, the correlation actually observed in basaltic rocks is positive. It follows that the L u / H f ratio has been fractionated by a phase other than clinopyroxene. Two minerals with the appropriate relative partition coefficients for Lu and H f are orthopyroxene and garnet, one or both of which must have been residual mantle phases during the melting event ( s ) that produced the depletion of the mantle and presumably also the complementary enrichment of the continental crust. The above-mentioned conclusion should be regarded with the additional knowledge that the abundances of H f and Sm are very well correlated in oceanic basalts, and their abundance ratio is approximately chondritic (Bougault and Treuil, 1980; J o c h u m et al., 1984). While this might suggest that the concentrations of the two elements are controlled by a single mineral for which K w ~ K ~ m , it now seems more likely that the roughly chondritic ratio has resulted from a fortuitous balance of equilibria involving a combination of minerals such as clinopyroxene and garnet. The abundance of H f in oceanic basalts and in the suboceanic mantle might indeed be controlled by clinopyroxene, which causes H f to be only moderately incompatible. The observed fractionation of L u / H f indicates that Lu is still less incompatible, perhaps by virtue of residual garnet. The presence of gar-

net during the mantle depletion event, for example, would 'buffer' Lu in the residue and keep its abundance nearly unchanged. The presence of clinopyroxene would partially buffer H f in the residue and prevent it from being depleted too much. These qualitative conclusions are borne out by quantitative model calculations, the presentation of which is beyond the scope of the present paper.

Acknowledgements This work was carried out during the first author's 5-month stay as a visiting scientist at the Max-Planck-Institut; he is grateful to have had the chance to work in so stimulating an environment. The third author expresses his sincere thanks to Professor F. Begemann for providing the opportunity to work in his Department (Cosmochemie) and to use the Cameca ® ion microprobe. The ion probe measurements were greatly expedited by the attention and assistance of Mr. Specht. The expertise and time devoted by Dr. H. Palme to set-up and operation of the NaI detector is greatly appreciated, as is the willingness of Dr. B. Spettel and Professor H. W~inke to provide time on their Ge (Li) detector systems. Bill White provided insightful discussions as well as indispensable technical and computer assistance in relation to the mass spectrometry. Constructive comments by Rick Ryerson, Fred Frey, Mike Roden, Bernard Duprd, Claude All~gre, Dave Lindstrom, and an anonymous reviewer are much appreciated.

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