Experimental and natural partitioning of Th, U, Pb and other trace elements between garnet, clinopyroxene and basaltic melts

CHEMICAL GEOLOGY IS;t)TOPE

ELSEVIER

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ChemicalGeology117 (1994) 149-166

Experimental and natural partitioning of Th, U, Pb and other trace elements between garnet, clinopyroxene and basaltic melts Erik H. Hauria'% T h o m a s P. W a g n e r b, T i m o t h y L. G r o v e b aDepartment of Geology and Geophysics, WoodsHole OceanographicInstitution, WoodsHole, MA 02543, USA bDepartment of Earth, Atmospheric and Planetary Sciences, Massachusetts Institute of Technology, Cambridge, .VIA 02139, USA

Received 16 September 1993; revisionaccepted 11 April 1994

Abstract Partition coefficients for Th, U, Pb, rare-earth elements (REE), high field strength elements (HFSE), alkalineearth elements, Sc, Cr, V and K were measured by ion microprobe techniques in two experiments on a natural high-alumina basalt composition from Medicine Lake, California. All elements were measured at natural abundance levels except Th, U and Pb, which were each present in the starting mix at 1-wt% levels. The results show that garnet retains U preferentially over Th ( D ~ t/mett = 0.0059, D~-th/rne~t= 0.0014), while clinopyroxene shows the opposite sense of partitioning (D~x/melt = 0.0127, D - ~ x/melt =0.014). The experimental Th, U and Pb partition coefficients for garnet-melt and cpx-melt are consistent with garnet-cpx pairs from garnet-bearing ultramafic rocks which exhibit U-Pb isochrons, thus demonstrating equilibrium (D~)/cpx=0.30, D~C°x=0.072, Dt~,/cvx= 0.016). The partition coefficients for Th and U between clinopyroxene and basaltic melt vary systematically as a function of the tetrahedral A1 content of clinopyroxene. Garnet/melt values for Th, U and Pb agree with previous determinations, indicating that mid-ocean ridge basalt (MORB) generation begins in the stability field of garnet lherzolite. However, high 226Ra//3°Th ratios in MORB require very small porosities near the region where the melts lose chemical equilibrium with the mantle. Partitioning data for HFSE and REE suggest that this region of melt segregation is not in the spinel lherzolite field. This requires either rapid transport of MORB magmas from >/70 km, or some degree of disequilibrium during melt generation and/or transport.

I. Introduction Knowledge of the partitioning behavior of trace elements is crucial to understanding the generation of basaltic rocks. Inferences on magma source compositions and the depth and degree of melting of basaltic magmas are limited, to some degree, by our understanding how elePresent address: Department of Terrestrial Magnetism, Carnegie Institutiton of Washington, 5241 Broad Branch Road, N.W., Washington,D.C. 20015, USA.

ments partition between mantle minerals and basaltic melts. For example, there is a growing recognition that mid-ocean ridge basalt (MORB) generation begins in the stability field of garnet lherzolite (Salters and Hart, 1989, 1990; Spiegelman and Elliot, 1993). This observation has only come about because of the constraints offered by high-quality partitioning data for trace elements such as Sm, Nd, Lu, Hf, Th and U (E.B. Watson et al., 1987; Hart and Dunn, 1992; LaTourrette and Burnett, 1992; Beattie, 1993a, b; LaTourrette et al., 1993 ). The best estimates for

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the depth of magma generation at hotspots, beginning in the stability field of garnet lherzolite (Hofmann et al., 1984; Frey and Roden, 1987; S. Watson and McKenzie, 1991 ; Eggins, 1992) are only lower limits, since we have little knowledge of the mineral/melt partitioning of trace elements in mantle phases deeper than ~ 200 km. In addition, much emphasis in trace-element geochemistry is placed on the ability to model trace-element ratios, especially high field strength element/rare-earth element (HFSE/REE) ratios, in mantle-derived magmas (All6gre et al., 1977; Hofmann and White, 1983; Hofmann et al., 1984, 1986; Salters and Hart, 1989; Williams and Gill, 1989; McKenzie and O'Nions, 1991; Spiegelman and Elliot, 1993) and ultramafic rocks (Vannucci et al., 1993; Hauri and Hart, 1994). These increasingly sophisticated methods require very accurate knowledge of the partitioning behavior of trace elements relative to each other (e.g., HFSE/REE, Ce/Pb, N b / U , K / Ti, etc. ). In this regard, it is difficult to construct a useful, self-consistent partitioning database based on many experiments in different systems with limited numbers of elements. This study reports the results of two partitioning experiments which attempt to overcome these limitations. The experiments were conducted on a natural basalt doped with U, Th and Pb, with all other trace elements at natural levels. The experiments produced clinopyroxene+melt and garnet+clinopyroxene+melt, and partition coefficients were measured by ion microprobe techniques. These experiments are the first to provide partitioning data for U, Th and Pb relative to a large number of other trace elements, and in garnet and clinopyroxene grown in the same experiment. In addition, the U - T h - P b results are compared to a unique set ofgarnet/cpx partitioning data in equilibrated garnet-bearing ultramafic rocks determined by isotope dilution.

2. Experimental and analytical techniques

2.1. Experimental conditions The starting composition was a primitive high-

alumina basalt from Medicine Lake, California, U.S.A. (sample number 79-35g). Phase relations of this sample up to 1.5 GPa have been studied by Bartels et al. ( 1991 ). Initial majorand trace-element compositions of the starting mix are given in Table I. To the powdered sample (grain size ~ 10 pm) was added 1 wt% each U, Th and Pb as oxides, and the mixture was homogenized by grinding in ethanol in an agate mortar, then dried at 100°C. This mixture was packed into graphite capsules, enclosed in MgO spacers, graphite heater and BaCO3 sleeve wrapped with Pb foil, and run in a ½-in ( 1.27 cm ) piston-cylinder apparatus with W97Re3-W75Re25 thermocouples. The samples were pressurized to 10 kbar cold, and the temperature increased to 865°C, where it was held for 10 rain while the pressure was increased to run pressure. The sample was then ramped up to run temperature while maintaining run pressure. The cpx/melt experiment was run at 1405°C and 1.7 GPa for 22 hr, and the garnet + cpx/melt experiment was run at 1430°C and 2.5 GPa for 69 hr. Experiments were quenched by abruptly dropping the pressure to 1 GPa and shutting off power to the heater. This decompression quench technique minimized the growth of quench crystals in the charge. No pressure correction was applied to thermocouple electromotive force (EMF), and temperatures are thought to be accurate to _+ 15 ° C. The temperature gradient across the graphite capsule was < 20 ° C based on measurements with paired thermocouples. Oxygen fugacity was maintained at or slightly below the C-CO-CO2 buffer by the graphite capsule. Pressure was monitored with a bourdon tube gauge, and is accurate to within 0.1 GPa based on the melting point of gold at 1 GPa (Akella and Kennedy, 1971). Each charge consisted of ~ 10% crystals ( 10200-#m clinopyroxenes, 5-50-~tm garnets) and ~ 90% glass. Thin rims (2-10 ~tm ) of quench clinopyroxene were observed in backscattered mode around cpx and garnet. Glass regions for ~ 50/~m around garnet crystals were rich in garnet-forming components, and contained tiny quench spinels (1-3 /~m). These regions may have formed by melting of garnet during de-

E.H. Hauri et al. / Chemical Geology 117 (1994) 149- I66

151

Table 1 Compositions of starting mix and constituent phases from trace-element partitioning experiments Major-element compositions 1430cC, 2.5 GPa (69 hr)

1405 C. 1 . 7 G P a ( 2 2 h r )

79-35 ga starting mix

cpx (n = 20)

garnet ( n = 30 )

glass ( n = 10)

SiO2 TiO2 A1203 Cr203 CaO MgO FeO MnO Na20 K2O P:Os

46.7 0.50 18.0 12.0 9.52 8.03 0.15 2.05 0.07 0.05

48.84 0.37 13.23 0.05 17.22 12.07 4.99 0.12 2.26 -

40.4 0.5 21.84 0.03 9.15 14.43 11.64 0.28 -

48.26 0.8 17.33 0.06 9,76 6.54 8,43 0,2 2,85 0, l 7 O. 18

49.3 0.33 9.16 0.11 15 18.8 5.64 0.12 1.87 -

48.4 0.69 18.11 0.03 10,77 8,69 9,12 0,18 2.88 0, l 5 0.11

Sum

97.07

99.15

98.27

94.4

100.J~i

991 i3

cpx (n = 20

glass (n=10)

Trace-element compositions 1430°C, 2.5 GPa (69 hr)

Ba Th U K Pb Nb La Ce Sr Nd Zr Hf Sm Eu Ti Dy Er Yb Lu Sc Cr V

1405:C, 1.7 GPa (22 hr)

starting mix a (n=3)

cpx

garnet

glass

cpx

glass

( n = 16)

(n=6)

(n=5

( n = 12)

(n=5)

28.3 10,000 10,000 l, 165 10,000 2.65 1.18 3.80 335 3.91 88.3 2.13 1.44 0.55 4,366 2.51 1.73 1.77 0.26 21.1 121 109

0.17 152 140 8.84 113 0.026 0.067 0.49 58.1 1.17 20.1 0.53 0.78 0.27 2,174 1.99 1.21 1.21 0.19 19.5 232 232

0.021 14.9 64.6 0.23 1.33 0.14 0.021 0.29 3.66 1.54 218 2.92 1.85 1.19 3,316 11.6 7.23 7.41 1.14 63.1 281 189

29.8 10,840 10,990 1,320 11,100 3.15 1.30 4.51 370 4.24 103 2.39 1.68 0.59 4,820 2.80 1.83 1.91 0.30 24.1 140 128

0.019 133 116 10.9 101 0.031 0.074 0.42 47.9 0.59 19.5 0.72 0.62 0.22 1,760 1.19 0.77 0.81 0.13 39.7 399 251

31.1 11,100 11,300 1,350 11,600 3.36 1.45 4.66 386 4.36 119 2.51 1.87 0.59 4,920 2.90 1.89 1.95 0.30 23.9 145 125

Major-element data in wt%, errors are _+ 1-2% for major oxides, _+5-10% for minor oxides. Trace-element data in ppm by weight, errors are _+0.3-3% for glass, _+3-15% for crystal phases. - = not determined. aAnalysis on super-liquidus glass ( 1 GPa, 1400 ° C. 5 hr).

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E.H. Hauri et al. / Chemical Geology 117 (1994) I49-166

compression quenching of the experiments. Outside of these Al-rich zones around garnet, the glass was homogeneous to nearly analytical uncertainty levels for both major (1-2%) and trace (5-10%) elements. Clinopyroxene was generally mildly concentrically zoned in A1 and Ti, but the total variation was within _+5% for major elements. Typically, only 2 or 3 spots on each clinopyroxene crystal could be analyzed for trace elements without overlapping previous spots or glass. The variability within crystals was similar to the variability between crystals for trace elements ( _+3-10%). Garnets exhibited complicated forms, often containing inclusions of glass, and tended to show complex zoning in pyrope and almandine components (extremes of Py57AI22Gr2~ to PYasA126Gr26). Because of the small size and complex forms of garnet crystals, only 1 or 2 spots per crystal could be analyzed for trace elements. Garnet trace-element variations were also within _+ 10%.

2.2. Microbeam techniques Major-element compositions were determined with the JEOL ® 733 Superprobe at MIT using Bence-Albee matrix corrections modified by A1bee and Ray (1970), using a focussed beam (23-#m activated diameter) for crystals and a defocussed beam (15-#m activated diameter) for glass. For standards, we used clinopyroxene, garnet and glass standards of similar composition as the run products. Trace-element concentrations were determined with the Camrca ® IMS 3fion microprobe at the Woods Hole Oceanographic Institution. For all trace elements, a primary O - beam of 0.2 nA was focussed to create a sputtered crater 710/~m in diameter. The analyzed spot was sputtered for 5 min before analysis in order to allow the sputtering process to reach steady state. Thus, for analyses of pure phases (e.g., crystals without glass inclusions), there was no systematic time dependence of intensity ratios relative to silicon. Secondary ions were accelerated to 4.5 kV and energy filtered at - 90 V offset ( ___10-V energy window) to suppress molecular isobaric interferences. Counting times were increased from 30

to 60 s to improve precision. Because U, Th and Pb were doped to percent levels, it was possible to use a small primary ion beam without sacrificing counting statistics. In this way, we could analyze crystals as small as 10/tm in diameter, and measure several spots on larger crystals. Traceelement concentrations were calculated from linear calibration curves determined as a function of isotopic intensities relative to a major isotope (e.g., 139La/3°Si vs. La in ppm ). All trace elements except Ba, Th, U, Pb, K and Nb were standardized using well-determined garnet (Monastery megacryst, South Africa) and clinopyroxene (Kilbourne Hole megacryst, New Mexico, U.S.A.) standards; the remaining elements were standardized using basaltic glasses. Using these standards, matrix effects are unresolvable at the 10-15% level for these elements. When measuring mineral trace-element compositions, only the central parts of crystals were analyzed to avoid quench growth around crystal rims. Since the amount of crystallization was small, no corrections to partition coefficients were necessary for zoning. In each ion probe analysis, the ion current for K was used as a monitor for glass contamination in the ion probe spot, since this element has a very low partition coefficient in cpx and garnet, and is present at > 1000 ppm in the glass. The distribution of K concentrations in clinopyroxene and garnet was peaked at 8 and 0.3 ppm, respectively, with a tail toward higher concentrations. Thus, points with K concentrations higher than ~ 10 ppm (clinopyroxene) or ~ 1 ppm (garnet) indicated the presence of glass in the analyzed ion probe spot, and these analyses were eliminated. Glass inclusions were also suspected when there were large ( > 20%) fluctuations in the secondary ion intensity ratios of highly incompatible elements (e.g., 39K/3°Si, t38Ba/3°Si), since these fluctuations were usually correlated with higher crystal concentrations for these elements. In hindsight, it was found that crystal analyses with K as high as 50 ppm in clinopyroxene and 5 ppm in garnet (indicative of ~ 5% glass in the spot) could be "corrected", by a preliminary procedure, to give results similar to analyses of pure crystals. This could be done by subtracting ion counts in the

E.H. Hauri et al, / Chemical Geology I 17 (1994) 149-166

proportions in which they were measured in the glass, until "corrected" K concentrations in clinopyroxene and garnet reached 8 and 0.3 ppm, respectively. However, since these "corrections" are preliminary and more needs to be learned about preferential ion-beam sputtering of mixed solid phases, no attempt was made to incorporate "corrected" results. K content was only used to screen the crystal analyses. The errors on partition coefficients given in Table 1 are calculated from the propagated uncertainty in the average glass composition and variability in individual crystal analyses.

2.3. Isotope techniques Natural samples are spinel peridotites (R238,

R324), garnet peridotites (R498, R521, R618), olivine-bearing garnet pyroxenites (R256, R740) from the Ronda Ultramafic Complex, Spain (Obata, 1980), and a fertile garnet lherzolite (LS-33) from Pali-Aike, Chile (Skewes and Stern, 1979). To avoid contamination during processing, samples were wrapped in plastic wrap and cotton towels, and crushed to 50-100 mesh size with a plastic mallet on a wooden block. Minerals were concentrated magnetically and purified by repeated crushing (with an agate mortar and pestle) and hand-picking following procedures similar to Zindler and Jagoutz (1988). Great care was taken to avoid inclusions, films and subgrain boundaries known to host incompatible elements. Mineral separates were leached using HF-HC1-HNO3 acids. The leaching procedure was developed to ensure the solution of U and Th during the leaching procedure, though repeat analyses of R238 cpx indicate that the different leaching procedures did not fractionate U, Th and Pb. All leaching steps used reagents cleaned by sub-boiling distillation in Teflon ® stills, and took place in clean, filtered air. Mineral separates of 17-300 mg were dissolved on a hotplate at 125°C for 5-10 days in 3- or 7-ml Savillex ® screw cap beakers in a mixed-acid solution consisting of equal parts concentrated HNO3-6.2 N HCl-concentrated HF. A mixed spike containing 235U, 23°Th and

153

2°spb was added before dissolution. Pb was separated in 0.5 N HBr on 20 #1 of Dowex 1 ® using a scaled down version of the technique of Manh~s et al. ( 1978 ). The fluoride residue after dissolution was dissolved in aqua regia at 150°C, evaporated to dryness, and redissolved in 8 N HNO3. Radiotracer studies have shown that U and especially Th partition strongly into Ca-fluoride formed during the dissolution step (A. Zindler, pers. commun., 1989). The potential for sample-spike disequilibrium for U and Th may thus exist during dissolution, and it is important to make every effort to achieve dissolution of the fluoride residue. The sample was centrifuged, and U and Th were separated together on AGIX8 ® anion-exchange resin in 8 N HNO3, and eluted with 6.2 NHC1 and 0.5 NHBr. One drop of0.015 N H2SO 4 was added to the U + Th fraction, and the solution was evaporated down to a tiny drop for loading. Sm and Nd were separated from the remaining solution using procedures described by Hart and Brooks (1977) and Zindler (1980). Pb was run on single Re filaments (0.001 in×0.030 in or 0.025 m m × 0 . 7 5 m m ) using the silica gel-phosphoric acid activation technique (Akishin et al., 1957; Cameron et al., 1969). Due to the very small amounts of Pb analyzed ( 1502000 pg), data were collected using an electron multiplier operated in ana'-g mode at a gain of 2" 10 4. Data for 2°Spb+ for isotope dilution measurements were always taken after the data for Pb isotopic composition, in order to burn off TI. Typical ion currents for 2°8pb+ were (0.53)- 10 -12 A for 1 ng ofPb. A correction factor of 0.375% amu-1 was employed based on repeated analysis of NBS 981 common Pb, using the values ofTodt et al. (1984). Precision for most runs is less than +_0.3% on all ratios, and the reproducibility of NBS 981 runs (300-2000 pg) is + 0.07% a m u - ~. Reproducibility of duplicate samples is in general also on the order of + 0.07% a m u - i. U and Th were loaded together on single Re filaments with colloidal graphite, and also run using the electron multiplier. The zone-refined Re filament sometimes had a small Th blank, which was generally eliminated by outgassing the Re filaments at 5.5-6.0 A (for ~ 2 0 rain), and

154

E.H. Hauri et al. / Chemical Geology 117 (1994) I49-166

checking each filament for Th before loading. A correction factor of 0.4% a m u - 1 was applied to the U and Th data based on repeated analyses of SRM 960 U. Blanks for the total procedure ranged from 9 to 15 pg Pb, and were < 1 pg for Th and U. Reagents used during the separation procedures typically accounted for 75-100% of the measured Pb blank. The isotopic composition of the Pb blank was consistently MORB-like (2°6pb/2°apb= 18.5-19.5, 2°7pb/2°npb= 15.415.6, 2°8pb/z°4pb= 37-39 ), which resulted in minimal blank corrections for most analyses. Isotope dilution measurements of U, Th and Pb are accurate to ~ + 5%, and this error is reported except for concentration data for olivine, opx and garnet, where uncertainties approach + 10%.

3. Partitioning results 3.1. REE, HFSE and other trace elements The measured cpx/melt and garnet/melt partition coefficients are given in Table 2 and displayed graphically in Fig. 1. There is an enormous literature on cpx/melt partitioning, and a full review will not be attempted here. The measured E L,/RE cpx/melt ~ of the cpx + melt experiment are similar to the data of McKay et al. (1986) and Hart and Dunn (1992) in that the partitioning pattern is slightly steeper than most other data (Irving, 1978). The measured D ~ x/melt in the garnet + cpx + melt experiment are ~ 50% higher than in the cpx + melt experiment, and these data are thus on the high side compared to most other cpx/melt data. It is likely that this is due to the higher CaO contents of these clinopyroxenes, as REE partitioning has been shown to be a function of Wo (wollastonite)or CaTs (Ca-Tschermak molecule) content of clinopyroxene (Jones and McKay, 1992 ). The values xutro~'eaJHFSEF~ c p x / r n e l t between the two experiments are different by <25%, with no apparent systematic offset as a function of Wo or CaTs in clinopyroxene. Relative to the REE, the cpx/melt partitioning patterns in Fig. 1A show negative anomalies for Ti and Zr, whereas H f is generally consistent with its nearest REE neighbor Sm, and the partition

coefficient for Nb is very low relative to La. These results are consistent with other partitioning studies incorporating HFSE and REE in the same experiments (Johnson and Kinzler, 1989; Hart and Dunn, 1992 ). The measured garnet/melt partitioning pattern is shown in Fig. lB. The measured values for O~ta/mell and D~tu/melt are similar to results obtained by Shimizu and Kushiro (1975), Irving and Frey ( 1978 ) and Nicholls and Harris ( 1980 ) on pyrope-rich garnets. However, our values for the intermediate REE (MREE) in our almandine-rich garnets are somewhat higher than the pyrope results. Similar variation axl '~ E/3gt/melt .t..,RE with garnet composition was observed by Nicholls and Harris (1980). As a result, our garnet/melt partitioning pattern for the REE is considerably flatter in the heavy REE (HREE) than most other experimental results. Alternatively, the flatness of the HREE partitioning in garnet could be the result of boundary layer depletion of compatible elements if diffusion through the melt did not keep pace with garnet growth. However, we would then expect higher values for the more incompatible elements, which is inconsistent with the measured results for D~ta/melt. Fig. IB shows substantial anomalies for the HFSE relative to the REE in the garnet/melt partitioning pattern. The partition coefficient for Ti is quite low relative to Dy, H f and especially Zr are high relative to Sm, and Nb is high relative to La. The data for Zr and H f are unexpectedly high, and appear to be compatible elements in garnet in this bulk composition. The only data with which to compare these results are for an almandine-rich garnet megacryst-host pair given by Irving and Frey ( 1978 ), which shows a value for O~/mett of 0.43, similar to D ~[me, determined experimentally for more pyropic garnets (Fujimaki et al., 1984).

3.2. U-Th-Pb partitioning The measured cpx/melt partition coefficients are shown in Table 2 and displayed in Fig. 1 relative to partition coefficients for the other elements. These data can be compared to the recent experimental data of LaTourrette and Burnett

155

E.H. Hauri et al. / Chemical Geology 117 (I 994) 149-166 Table 2 Trace-element partition coefficients from experiments and natural garnet-bearing rocks Experimental partition coefficients 1430°C, 2.5 GPa (69 hr)

Ba Th U K Pb Nb La Ce Sr Nd Zr Hf Sm Eu Ti Dy Er Yb Lu Sc Cr V

1405°C, 1.7 GPa (22 hr)

cpx/melt (n=7)

_+20. (%)a

garnet/melt (n=5)

_+20 (%)a

cpx/l (n=5)

+ 2or (%)a

0.0058 0.014 0.0127 0.0067 0.0102 0.0081 0.0515 0.108 0.157 0.277 0.195 0.223 0.462 0.458 0.451 0.711 0.66 0.633 0.623 0.808 1.66 1.81

31 16 22 11 14 20 8.9 6.2 5.5 7.8 3.3 21 13 16 9.9 12 8.4 6.3 16 14 9.4 14

0.0007 0.00137 0.00588 0.0002 0.00012 0.0538 0.0164 0.065 0.0099 0.363 2.12 1.22 1.1 2.02 0.688 4.13 3.95 3.88 3.79 2.62 2.01 1.48

63 20 14 42 39 16 23 18 20 15 4.3 26 10 13 8.6 6.3 7.8 4.3 8.8 6.3 9.5 4.8

0.0006 0.012 0.0103 0.0081 0.0087 0.0093 0.0512 0.089 0.124 0.136 0.164 0.288 0.331 0.373 0.358 0.412 0.41 0.413 0.425 1.66 2.75 2.01

43 31 38 12 31 27 16 8.4 4.3 7.6 6.1 33 6.6 12 8.1 10 11 8.9 15 12 6.6 8.7

Natural and experimental mineral/mineral partition coefficients

Olivine/cpx Opx/cpx Garnet/cpx

Sample

Dxh

Du

Dpb

R324 R238 * 1 R238 # 2 R256 R498 R521 R618A R618B R740 LS-33

< 0.002 < 0.008 < 0.008 0.069 0.061 0.072 0.038 0.032 0.052 0.18

< 0.002 < 0.017 < 0.020 0.25 0.33 0.30 0.16 0.098 0.69

0.008 0.018 0.019 0.012 0.0098 0.016 0.0085 0.0079 0.027 0.034

0.072

0.30

0.016

0.097

0.46

0.012

Mean Experimental garnet/cpx

1,430°C, 2.5 GPa

Partition coefficients are calculated from concentration ratios. Errors are calculated from variations in measured partition coefficients, propagated through uncertainties in glass compositions. Uncertainties on mineral/mineral partition coefficients are estimated at _+20-30% for the natural samples and + 30-60% for the 2.5-GPa experiment.

E.H. Hauri et al. / Chemical Geology I 17 (1994) 149-166

156

1:

.1

.01

.001 c x/melt

.0001

2.5 (SPa

. . . . . . . . . . . . . . . . . . Ba Pb Th U Nb La Ce Sr Nd Zr HI Sm Eu Ti Dy Er Yb Lu

10:

B

1 ca

.1' Qa

.01 ¸

.001

0 0 0 1 ,Ba (b T L 6

' Sr' d d La Ce

6 ' Sm Eu' Ti' 6 y ' Er

' Lu

Fig. 1. Cpx/melt (A) and garnet/melt (B) partition coefficient patterns from experiments on high-alumina basalt 79-35g. Cpx was the only crystalline phase in the experiment at 1.7 GPa, while garnet and cpx were grown together in the 2.5-GPa experiment. Cpx/melt results for the REE at 2.5 GPa are ~ 50% higher than at 1.7 GPa. The cpx/melt pattern shows negative anomalies for HFSE relative to the REE. The garnet/melt pattern shows a negative anomaly for Ti, but positive anomalies for Hf, Zr and Nb relative to the REE.

( 1992 ) and Beattie (1993a). Between these different studies, the cpx/melt partition coefficients for U, Th and Pb show substantial variability. The valence state of U is a function of oxygen fugacity. Very low U partition coefficients have been measured obtained under less reducing conditions (above C-CO-CO2 buffer) by E.B. Watson et al. (1987), LaTourrette and Burnett ( 1992 ) and Beattie (1993a), and are due

to the presence of U 6÷. In the graphite-buffered high-pressure experiments of Beattie (1993b) and this study, U 4÷ is present in higher proportion. The actual distribution of U among different valence states is also dependent on temperature and composition. For U and Th, our results are substantially higher than the results of LaTourrette and Burnett (1992) and Beattie (1993a) at similar oxygen fugacities (near the

E.H. Hauri et al. / Chemical Geology 117 (I 994) 149-166

quartz-fayalite-magnetite oxygen buffer, QFM); the total variation in the data is about a factor of +_ 12. However, when comparing U and Th partitioning data in the vicinity of the QFM buffer, it is apparent that much of the variation is the result of systematic crystallographic control over the partitioning of U and Th. Fig. 2 shows the U and Th results plotted against the proportion of tetrahedral AI in the experimental clinopyroxene compositions reported by LaTourrette and Burnett (1992), Beattie (1993a) and this study. The positive correlation with A1~xv) is the expected result for + 4 cations substituting into clinopyroxene. Most mantle diopsides have ~ 10-15% of their A1 as A1
e p x / m e l t

0.020-

0.015"

0.005"

l

'~ A

0.000 0.00

•

ii

0.05

,

B LT HWG / • • • D(Th) A [] 0 D(U)

I

i

0.10

•

i

0.15

Fe ratios exert little influence on the partitioning of + 2 cations. Our results show that Df,g"/mel' is similar to, or slightly lower than O ~ ×/melt , and in this respect are contradictory to the results of Beattie (1993a). However, from Fig. 2, it is apparent that this is due to the low values for U and Th due to low-A1 °v) clinopyroxenes in the Beattie (1993a) experiments. Measured garnet/melt partition coefficients for U, Th and Pb are compared with the results ofBeattie (1993b) and LaTourrette et al. ( 1993 ) in Table 3. The results for O~)/meh and D ~ meh show good agreement, and do not vary as a function of garnet composition. There is substantial noise in the data for D~/me~', with a range of 10 -5 to 10- 3, with our value in the middle of the range. Since there is no systematic relationship of O~/melt with garnet composition, it is likely that this noise reflects measurement errors associated with very low Pb concentrations in the experimental garnets. Since the present experiments were doped with 1 wt% Pb, our garnets have the highest Pb concentrations and thus the smallest errors on D~/melt. However, it is clear that D~/melt is very small, and that garnet preferentially rejects Pb relative to U and Th.

3.3. Garnet/cpx partitioning in garnet peridotites and pyroxenites

036I

0.010"

157

,

i

0.20

,

i

0.25

,

0.30

A1 (IV) in Clinopyroxene Fig. 2. D ~ ~/melt and D ~ ~/'~l~ vs. AltlV) in clinopyroxenes from the experiments of Beattie (1993a) ( n = 5 ) , LaTourrette and Burnett ( 1992 ) (n = 3 ) and the present study ( n = 2 ), in order of increasing AI
Table 4 shows U, Th and Pb concentration data and Pb isotopic compositions for clinopyroxene and garnet mineral separates from garnet peridotites and garnet pyroxenites from the Ronda Ultramafic Complex, Spain (Obata, 1980), and a single garnet lherzolite xenolith from Pali-Aike, Chile (Skewes and Stern, 1979; Stern et al., 1989). The data on these mineral pairs are unique for mantle samples, as demonstrated by t h e 2 3 8 U - 2 ° 4 p b isochron diagram in Fig. 3. The garnet-cpx pairs from Ronda give isochron ages of 17-44 Ma with substantial errors. These ages are in agreement with Sm-Nd ages of two of the mineral pairs (Fig. 3 ), Sm-Nd ages obtained on different samples from Ronda (Zindler et al., 1983; Reisberg et al., 1989), and Rb-Sr dating of anatectic granites intruding the ultramafic body (Priem et al., 1979). In addition, the gar-

E.H. Hauri et al. / Chemical Geology 117 (1994) 149-166

158

Table 3 Comparison of natural and experimental partition coefficients for U, Th and Pb A v g . E. C h i n a peridotites a

Average peridotite data (this study )

1430°C, 2.5 G P a ( this study )

D~}~x/mdt ( X 10 - 3 )

-

-

10.3-12.7

0.94-3.61

O{Phx/melt ( X 10 - 3 )

-

-

12.0-14.0

1.25-5.31

D~,~x/melt ( X 10 - 3 )

-

-

8.7-10.2

4.89-9.68

-

D ~ rnm/mel' ( X I 0

-3 )

-

-

5.88

9.18-17.9

15

D f ~ rnel/melt ( X | 0 - 3 )

-

-

1.37

1.50-3.26

1.7

D ~ met/melt ( X I 0

_

_

0.12

0.01-0.67

-

D~/cpX

0.419

0.306

0.269

D~th/cpx

0.066

0.0715

0.0593

1.15

0.086

Dlg~/cox

1.97

0.0166

0.011

2.94

-

3)

Beattie ( 1993a, b)

LaTourrette and B u r n e t t ( 1 9 9 2 ), L a T o u r r e l t e et al. (1993) c 4.7-11.3 II . 3 - 3 6

10.7

2.05

= no data. a G a r n e t - c p x p a i r s f r o m g a r n e t l h e r z o l i t e s a n d g a r n e t c l i n o p y r o x e n i t e s f r o m e a s t e r n C h i n a , d a t a f r o m T a t s u m o t o el al. ( 1 9 9 2 ). ~ C a l c u l a t e d f r o m a v e r a g e c p x / m e l t a n d g a r n e t / m e l t d a t a in s e p a r a t e h i g h - p r e s s u r e e x p e r i m e n l s . C C a l c u l a t e d f r o m a v e r a g e c p x / m e l t a n d g a r n e t / m e l t d a t a in s e p a r a t e e x p e r i m e n t s . -

Table 4 U-Th-Pb

R238

R324

R256 R498 R521 R618

R 740

LS-33

-

and Sm-Nd

cpxl cpx2 opx cpx opx olivine cpx gt cpx gt cpx gt cpx gtA gtB cpx gt cpx gt

isotope data for coexisting minerals from mantle peridotites and mafic rocks

U

Th

Pb

(ppb)

(ppb)

(ppb)

11.6 9.86 <0.20 21.4 <0.10 <0.05 7.85 1.95 8.53 2.83 8.89 2.70 18.6 3.06 43.8 4.27 12.7 8.77

24.6 24.4 <0.20 22.3 <0.10 <0.05 24.2 1.68 25.6 1.55 26.1 1.87 45. I 1.71 1.44 t00 5.17 57.6 10.2

42.9 41.2 0.79 122 1.90 0.98 227 2.79 149 1.46 173 2.83 276 2.35 2.18 357 9.65 27.2 0.93

2°6pb/2°4pb

17.047 17.077 18.259 18.984 18.784 18.544 18.797 18.927 17.715 . 18.292 18.740 18.289 18.531 . 18.500 18.659 18.072 18.260

.

.

2°7pb/2°4pb

15.203 15.206 15.693 15.649 15.655 15.406 15.648 15.344 . 15.454 15.524 . 15.627 15.400 15.390

2°8pb/2°4pb

.

36.277 36.292 38.188 38.98I 38.512 38.128 38.612 38.595 37.021 . 37.988 38.252 37.866 37.952 . 38.813 38.582 37.614 37.770

Sm

Nd

(ppm)

(ppm)

. -

143Nd/144Nd

.

.

.

-

-

1.618 0.650 2.196 0.749

4.345 0.342 7.280 0.445

0.513185 0.513299 0.513236 0.513362

-

-

-

-

-

-

.

.

= not determined.

net lherzolite xenolith LS-33 from Pali-Aike gives a U - P b date of 2.16_+ 2o31.4~Ma, which is in agreement with the eruption ages of volcanoes on the Patagonia Plateau, southern Argentinia (C.R.

Stern, pers. commun., 1990). While two of the U / P b isochrons (R521, R740) are not in agreement, this may be due to small fluctuations in analytical Pb blank, which would shift the garnet

E.H. Hauri et al. / Chemical Geology 117 (1994) 149-166

159

.51340

.51335

.51330"

Z

garnet .51325-

Z .51320

cpx

.51315-

.51310 0.0

o.~

o.~

0

0

0

o:8o 147Srn/ l~4Nd

o.'8

1.8

1.2

0

0

0

19,00

2v

18.80-

-'f .

J3

/

/

/

gai,o,/ I,'~82~

40

18.60.

g

/

X

¢q

g

v

185_+3.6~

U

18.40

18.20

~ 18.00-

~ ..... ......

1 7 . 9 0j l

o

. 290

,

•

..

+36. - , 1A~Ma

. . . . . .

4.0,0

I

I

I

,

•

i

20

40

. . . . . .

600 ......

I

I

68

. 2 ....... 800

!

I

8b

...... "

oo

238U/ 2°4pb Fig. 3. S m - N d (A) and U - P b (B) isochron plots for garnet-cpx pairs from Ronda garnet peridotites from southern Spain ( R521, R 618, R 740 ) and a garnet-bearing mafic layer ( R256 ). Uncertainties (2a) are estimated at +_0.0035% and _+0.5% on 143Nd/144Nd and t~TSm/144Nd, respectively; +_0.14% and + 0.5% on 2°6pb/2°4pb for cpx and garnet, respectively; and +_20% on z38U/z°4pb. The U - P b ages are consistent with the above S m - N d ages and S m - N d ages of other garnet peridotites from Ronda ( Reisberg et al., 1989 ) and Rb-Sr dates of anatectic country rocks (Priem et al., 1979 ). Also shown on a compressed U / P b scale is a similar U - P b isochron for a garnet lherzo)ite xenolith from Pali-Aike, Chile; this age is within error of the eruption age of the host alkali basalt. These isochronous relationships demonstrate isotopic and U - P b partitioning equilibrium.

160

E.H. Hauri et al. / Chemical Geology 117 (1994) 149-166

isotopic compositions to higher 2 ° 6 p b / 2 ° 4 p b , and thus older ages. In addition, there is some evidence that the U - P b system has a higher closure temperature than the S m - N d system (Mezger et al., 1992), which would result in U - P b ages which are older than S m - N d ages in slowly cooled rocks. The 2 3 2 T h - 2 ° 8 p b data for LS-33, while burdened by excessive errors relative to the very young age, are also consistent with this age. These garnet-clinopyroxene pairs provide a unique constraint on the U - T h - P b partitioning data, because these isochrons are proof of isotopic and U - T h - P b equilibrium at the time of U - P b closure. Such convincing independent evidence of equilibrium is rare or non-existent in partitioning experiments. Equivalent data on natural garnet-cpx pairs with which to compare the U - T h - P b data are rare. Only the recent study of Chinese xenoliths by Tatsumoto et al. (1992) provides four garnet-cpx pairs from peridotites with U, Th and Pb concentration and isotope data. Their results are in excellent agreement for garnet/cpx partitioning of Th and U, but the Pb data are in the opposite direction of ours, i.e. Pb is preferentially partitioned into garnet relative to cpx in their samples (Table 3). However, the Chinese peridotite data do not form U - P b isochrons, and in fact U / P b is negatively correlated with 2 ° 6 p b / 2°4pb in their data. This suggests intermineral isotopic disequilibrium in the Chinese samples, or possibly some contamination problem associated with Pb in garnet (since the U and Th partitioning data are in perfect agreement). The natural garnet/cpx partitioning data are shown in Table 3 relative to the experimental results of the present study and those of LaTourrette and Burnett (1992) and Beattie (1993a, b). These results are also displayed in Fig. 4. The garnet/cpx data measured in the 2.5GPa experiment are in excellent agreement with the natural garnet/cpx data from Table 2. However, when we compare the preferred garnet/melt and cpx/melt data of LaTourrette and Burnett (1992), Beattie ( 1993a, b) and LaTourrette et al. (1993) to calculate garnet/cpx partitioning, it is apparent that there are problems in the extrapolation of these mineral/melt data for differ-

100~

] 10

cj

0,... "-+.,...,....

. . . Q . ----C~.--

I Natural g~/cpxIthtsstud}] Hauri et al (expts. tins stud},) LaTourretteet al gt/cpx Beattie gt/'cpx

1

.1

,01

%+ .+ D

.001

U

Th

Pb

Fig. 4. Comparison ofgarnet/cpx partitioning data for U, Th and Pb. Data are calculated from cpx/melt and garnet/melt data of LaTourrette and Burnetl ( 1992 ), LaTourrette et al. (1993) and Beattie ( 1993a, b). Measured garnet/cpx data are from natural garnet-cpx pairs in peridotites and pyroxenites from the Ronda Ultramafic Complex (Spain) and PaliAike (Chile), and the garnet + c p x + m e l t experiment on 7935gat 2.5 GPa (this study).

ent phases in different experiments (Table 3). The calculated value of D~th/cpx from the data of LaTourrette agrees very well with the present value, but D~/cpX is high by a factor of 3-10. D g'/cPX-values calculated from the data of Beattie (1993a, b) are even more problematic, with /-'1 g t / c p x and Jt.,pb/Tgt/cpxlow by a factor of 10, and ,--u Dg~ cpx high by a factor of 10-50. While the Pb data are quite noisy, the difference in D gl/cp× for the U and Th data is due to the fact that the garnets and clinopyroxenes produced in the separate experiments of LaTourrette and Beattie were not in equilibrium with each other. The clinopyroxenes of Beattie (1993a) simply have too little A1203 to be in equilibrium with garnet. This is emphasized by the observation that the experimental results prior to this study would suggest that U is preferentially partitioned into garnet relative to clinopyroxene, in direct conflict with all of the natural data to date. As a result, combining the partition coefficients for different phases from separate experiments can lead to erroneous results, especially in the relative values of partition coefficients for elements with quite different charges and ionic radii.

E.H. Hauri el al. / Chemical Geology 117 (1994) 149-166

161

4. Discussion

[] Spinel Lherzolite i

4.1. Implications for trace-element ratios in mantle-derived basalts A clear limitation of our data is that the experimental system is essentially a basalt, not an ultramafic composition, and as such, the resulting clinopyroxenes and garnets are considerably richer in iron than mantle phases. Nonetheless, results for + 2 and + 3 cations are consistent with other experiments of different M g / F e ratios. While there are too little data to be able to say this for + 4 cations (HFSE), the relative sense of fractionation between elements of different charge and ionic radii is likely to be similar in ultramafic compositions. Because of the large number of trace elements determined, and considering the above discussion, the present partitioning data set has implications for trace-element fractionation in the mantle. Shown in Fig. 5 are ratios of partition coefficients for cpx/melt, garnet/melt and a hypothetical bulk partition coefficient for garnet lherzolite/melt ( 10% garnet+ 10% cpx), based on the results of this study. These particular trace-element ratios have been examined because they show limited variations in MORB and oceanic island basalts (OIB), except for K/Ti. The D ratios for spinel lherzolite/melt would not be substantially different from the cpx values, whereas the D ratios for garnet lherzolite represent the combined effects of cpx and garnet. D ratios close to unity would result in very small variations of these trace-element ratios in mantle-derived magmas and residual peridotites. D ratios substantially different from unity would result in fractionation of these trace-element ratios during melting or melt migration if the degree of melting or the porosity is on the same order as the bulk D. With the exception o f T h / U , Ti/Eu, Zr/Hf, and possibly N b / U , most of the D ratios for spinel lherzolite do not approach unity. As a result, substantial fractionation of these traceelement ratios could be expected upon melting of spinel lherzolite, both in magmas and residual solids. These data clearly point out the fallacy of using incompatible trace-element ratios such as

001-

Th/U

K/L

K/T:

Nb/tT

Nb/Ia

C¢,Pb

Zr/bm

HI/Sin

Zr/Ht

Ti/Eu

Partition Coefficient Ratios (e.g. D~/D v) Fig. 5. This figure shows the values of ratios of bulk partition coefficients (e.g., Dcc/Dpb) for spinel lherzolite ( open boxes) and garnet lherzolite Cfi/led boxes), calculated from the results of this study. A garnet/cpx ratio of unity was assumed for garnet lherzolite. The boxes enclose the possible range in partition coefficients, based on propogated analytical uncertainties reported in Table 2. With the exception of K/Ti, these trace- element ratios show limited variability in mid-ocean ridge basalls (MORB) and oceanic island basalts (OIB), which would predict partition coefficient ratios near unity. The D ratio fox C c / P b assumes no contribution from peridotitic sulfide, which would probably lower this ratio substantially. In detail, the constancy of the ratios Zr/Hf, Zr/ Sm and Hf/Sm in MORB and OIB supports initiation of MORB melting in the garnet Iherzolite stability field (Salters and Hart, 1989). The constancy of N b / U ratios in MORB and OIB ( Hofmann et al., 1986 ) would seem to favor melting of spinel lherzolite: howe~er, REE systematics of most OIB are the result of melting garnet lherzolite (Hofmann el al., 1984), which would result in marked Nb/l~ fractionation according to our results.

K/Ti as proxies for mantle heterogeneity in mantle-derived magmas (Shen and Forsyth, 1993). In contrast, the D ratios for Nb/La, Zr/ Sm, H f / S m and Z r / H f are all moderately close to unity for a garnet lherzolite assemblage, and as such the range of these trace-element ratios would be limited in melts in equilibrium with garnet lherzolite. The ratios Zr/Sm, H f / S m and Z r / H f show limited variation in most oceanic basalts, including MORB and OIB. In fact, the H f / S m ratio is significantly less variable in oceanic basalts than either Z r / S m or Z r / H f (Salters and Hart, 1989; 1990), which is entirely consistent with our partitioning results. Relative

162

E.H. Hauri et al. / Chemical Geology I 17 (1994) 149-166

to these observations, the present set of partitioning data provides further support from HFSE/REE ratios that MORB melting begins in the stability field of garnet lherzolite (Salters and Hart, 1989). Surprisingly, the D ratios for Nb/U, Ce/Pb and K / U are significantly different from unity for both spinel and garnet lherzolite assemblages (Fig. 5 ). This is in conflict with the uniformity of these ratios in most oceanic basalts (Hofmann, 1986; Hofmann et al., 1986). For Ce/Pb, it is possible that residual sulfide in the mantle controls the partitioning of Pb to some degree (Meijer et al., 1990), which could lower this D ratio substantially. However, this would require a very specific amount of sulfide to be residual after melting, for both MORB and OIB, a scenario which seems unlikely. For Nb/U, K / U and Th/U, the relative constancy of these ratios in oceanic basalts may be suggesting that degrees of melting, or effective porosities (or both), are sufficiently large ( > 1-2%) that these elements are not fractionated significantly from each other during melt generation and transport. However, porosity larger than ~ 2% presents problems for the generation of U - T h - R a disequilibria, as will be shown below.

4.2. Generation of U-series disequilibria in MORB The general sense of U and Th partitioning between cpx, garnet and basaltic melts observed in this study is consistent with the results of LaTourrette and Burnett (1992), Beattie (1993a, b), and LaTourrette et al. ( 1993 ) in demonstrating that 23°Th-238U disequilibria can be derived by initiation of MORB melting in the garnet lherzolite stability field. Garnet lherzolite with a 1:1 garnet/cpx ratio will have DTh/Du <0.8, SO that melting and melt transport will result in 23°Th excess. Generation of MORB entirely within the spinel lherzolite field will result in 23°Th deficiencies, since cpx has DTh/Du> 1.1 (Table 2), and even higher iffo2 is above the graphite-CO2 buffer (LaTourrette and Burnett, 1992; Beattie, 1993a, b). In either the spinel- or garnet-lherzolite stability field, the fractiona-

tion of highly incompatible elements such as U, Th and Ra, can only take place where the mantle porosity is similar to the bulk equilibrium partition coefficient for these elements. Generating 23°Th/238U and 226Ra/23°Th ratios observed in MORB by equilibrium melting thus requires melting and/or melt transport to take place in the garnet lherzolite field at porosities of < 10- 3. Fig. 6 shows the evolution of (23°Th/238U) and (226Ra/23°Th) in a one-dimensional melt column (parentheses denote activity ratios), where the magma maintains equilibrium with the solid during melt transport, after Spiegelman and Elliot (1993). The column consists of a layer of spinel lherzolite (0-60 km) overlying a layer of garnet lherzolite (60-80 km). Melting begins at 80-km depth and continues up to 0 km. The maximum degree of melting of 25% and the maximum porosity is 1%. The partition coefficients are from the present study, assuming 18% garnet+ 12% cpx in the garnet peridotite, and 18% cpx in spinel peridotite. This results in D gl/melt =0.00193, D~/m~'l' = 0.00258, ,_.r~ _sl/mel~ f h -Th

0.00252 and D~J/m~qt=0.00229, and a value of Dgl/m,qt =0.000l is assumed for the entire CO1Ra umn, based on the partition coefficients for Ba. These partition coefficients are constant throughout the column. Using these partition coefficients, 34% fractionation in ( 23°Th/238U ) can be created in the first instantaneous melt at the bottom of the column. With continued migration through garnet lherzolite, the excess 23°Th will be diluted as more 23sU is partitioned into the melt at higher degrees of melting. However, within the garnet lherzolite field, (23°Th/238U) does not decay to equilibrium ( = 1 ). This is because 23°Th is being produced in the solid residue and is subsequently partitioned into the melt; the preferential partitioning of U over Th into garnet results in retention of 238U in the matrix, producing 23°Th which is then partitioned into the melt phase, resulting in an excess of 23°Th over 238U in the melt (Spiegelman and Elliot, 1993). Since the porosity is < 10 -3 throughout the column, 226Ra excesses are also sustained through melt transport, since DRa < DTh. However, migration of the melt into the spinel

E.H. Hauri el al. / Chemical Geology I 17 (1994) 149-166

163

1.30

Garnet I Spinel Lherzolite, i Lherzolite

A

i

1.20

, zc

""

1,10

|

00

0.90

I

L

Garnet Lherzolite [-..

I

I

I

I

Spinel Lherzolite

B

4 3 2

0 80

70

60

50

40

30

20

10

0

Depth (km) Fig. 6. Evolution of ( 23°Th/238U ) (A) and ( 226Ra/23°Th ) ( B ) in a basaltic melt, as a function of height in a melting column o f garnet lherzolite ( 6 0 - 8 0 kin) below spinel lherzolite ( 0 - 6 0 km), according to the model of Spiegelman and Elliot (1993). Partition coefficients are taken from this study. Garnet lherzolite is assumed to contain 18% garnet and 12% cpx, and spinel lherzolite contains 18% cpx. Melting begins at 80 kin, with an adiabatic ascent rate of 5 cm yr- ~, a m a x i m u m porosity of 1%, and a m a x i m u m degree of melting o f 25%. The model shows that excesses in 2~°Th can be generated in the garnet lherzolite field. However, at the small porosities required to maintain 226Ra excesses, migration o f melt through spinel lherzolite results in 23°Th deficiencies under these melting conditions.

lherzolite field will alter the Th isotope signature. Spiegelman and Elliot (1993) have shown that excesses in 23°Th and 226Ra can be maintained in a melt which migrates through spinel lherzolite, even though the relative partitioning of Th and U in spinel lherzolite will drive the melt to 23°Th deficiencies (DTh/Du=I.1). A very small porosity, on the order of 10 -3 , must be maintained in the spinel lherzolite field in order to maintain 226Ra excesses in this model. However, this small porosity will allow effective T h / U fractionation, resulting in 23°Th deficiencies in

the melt as it interacts with spinel lherzolite. This small porosity will also induce significant changes in trace-element ratios such as Nb/La, Zr/Sm and Hf/Sm, which have bulk partition coefficients much higher than 10-3 and will thus be fractionated in the spinel lherzolite field. Fig. 5 shows that these partition coefficient ratios ( DNb/ DLa , D z r / Dsm, DHJ Dsm ) are very close to unity for garnet lherzolite, but significantly below unity for spinel lherzolite. Since it is unlikely (though not impossible ) that MORB generation always occurs under the same conditions of up-

164

E.H. Hauri et al. / Chemical Geology 117 (I 994) 149-166

welling rate, porosity and degree of melting, more variability is expected for Nb/La, Z r / S m and Hf/Sm ratios in melts of spinel lherzolite than in melts of garnet lherzolite. The limited variability in these ratios both MORB and OIB (Salters and Hart, 1990), and the presence of excess 23°Th in MORB, favor the latter scenario. Alternatively, the abundances of key minerals like garnet and cpx will decrease with height in the melting column, and as a result the bulk peridotite/melt partition coefficients will decrease with height, and the column will lose its ability to fractionate incompatible trace elements at high levels. Harzburgite probably has bulk partition coefficients on the order of 10-5 for Th and U (Table 3). This might help to maintain high ( 23°Th/238U ) ratios in the spinel lherzolite field, but if transport takes place by grain boundary melt percolation at low porosity ( ~ 1-2%), the melt velocities will be too small to maintain high 226Ra/23°Th ratios, as 226Ra will decay back toward equilibrium with 23°Th due to its short halflife ( 1600 yr). This situation would seem to favor a rapid rise of MORB from the garnet stability field, rapid enough not only to maintain ( 23°Th/238U ) through the spinel peridotite field, but also to preserve high (226Ra/23°Th) ratios. The most straightforward way to do this would be through a vein or fracture network (Spiegelman and Kenyon, 1992; Hart, 1993). If trace-element fractionation during melting and melt transport were effected, to some degree, by volume diffusion of trace elements in the solid, as advocated by Iwamori ( 1993 ), then the result would be to increase the effective partition coefficients of the most incompatible elements. These higher, diffusion-controlled effective partition coefficients would relax the requirements of very small porosities in the source region of MORB, and thus go part of the way toward resolving the paradoxical requirements of small porosities and rapid melt transport to maintain 23°Th and 226Ra excesses. It is apparent from Fig. 2 that AI is involved in the substitution mechanism for incorporating Th and U in clinopyroxene. Recent experiments on the diffusion of AI into diopside (Grove and Wagner, 1993) suggest that diffusion of Th and U, whether it is con-

trolled by diffusion of (Mg,Ca) (Th,U)AlzO 6 or coupled substitution with A1<~v), could be quite slow relative to the best estimates for melting rates (McKenzie and Bickle, 1988; Iwamori, 1993). In addition, two- and three-dimensional melt migration effects such as melt focusing could be important in generating U-series isotopic disequilibrium. Clearly, more work is needed, on both experimental and dynamic fronts, to resolve these issues.

5. Conclusions

We have measured the partition coefficients of a large number of trace elements, at natural and doped abundance levels by ion microprobe techniques, on cpx and coexisting cpx and garnet in single charges. This experimental strategy minimizes the errors in relative values of partition coefficients. In addition, we have shown that measurements of garnet/melt and cpx/melt partition coefficients in charges with coexisting garnet and cpx provide the most accurate method of determining these parameters, as shown by the comparison of U - T h - P b garnet/cpx partitioning in experiments and in natural garnet-bearing ultramafic rocks where equilibrium is demonstrated. The results show that garnet retains U preferentially over Th (D~/m~l~=0.0059, D~/meh = 0.0014), while clinopyroxene shows the opposite sense of partitioning (D~:px/melt =0.0127, "-'Thr~¢px/m¢lt=0.014). The T h / U partitioning results support previous suggestions that MORB melting begins in the stability field of garnet lherzolite, which is required in order to generate 23°Th excesses. However, the small porosities required to maintain excess 226Ra in MORB magmas would result in substantial fractionation of other incompatible trace elements, especially HFSE/REE ratios. If uranium-series disequilibrium in MORB is due to melting and melt transport effects, then rapid extraction of MORB melts from below the garnetspinel transition at 60-70-km depth, or some degree of diffusion-controlled melting and melt transport, is required.

E.H. Hauri et al. / Chemical Geology 117 (1994) 149-166

Acknowledgements Thanks go out to Nobu Shimizu, Glenn Gaetani, Stan Hart, Marc Spiegelman, Tim Elliot and Peter Kelemen for discussions on these issues. Assistance provided by Mike Jercinovic on the electron probe and Ken Burrhus on the ion probe is sincerely appreciated. Thanks to Charles Stern for collecting and providing the garnet lherzolite LS-33. This work was supported by NSF EAR9096194 to S.R. Hart and EAR-9115901 to T.L.G.

References Akella, J. and Kennedy, G.C., 1971. Melting of gold, silver and copper: proposal for a new high pressure calibration scale. J. Geophys. Res., 76: 4969-4977. Akishin, P.A., Nikitin, O.T. and Panchenkov, G.M., 1957. A new effective ion emitter for the isotopic analysis of lead. Geokhimiya, 5: 429-434. Albee, A.L. and Ray, L., 1970. Correction factors for electron probe microanalysis of silicates, oxides, carbonates, phosphates and sulfates. Anal. Chem., 42: 1408-1414. All~gre, C.J., Treuil, M., Minster, J.F., Minster, B. and AIbar6de, F., 1977. Sytematic use of trace element in igneous process, Part I. Fractional crystallization processes in volcanic suites. Contrib. Mineral. Petrol., 60: 57-75. Bartels, K.S., Kinzler, R.J. and Grove, T.L., 1991. High pressure phase relations of primitive high-alumina basalts from Medicine Lake volcano, northern California. Contrib. Mineral. Petrol., 108: 253-270. Beattie, P., 1993a. The generation of uranium series disequilibria by partial melting of spinel peridotite: constraints from partitioning studies. Earth Planet. Sci. Left., 117: 379-391. Beattie, P., 1993b. Uranium-thorium disequilibria and partitioning on melting of garnet peridotite. Nature (London), 363: 63-65. Cameron, A.E., Smith, D.H. and Walker, R.L., 1969. Mass spectrometry of nanogram size samples of lead. Anal. Chem., 41: 525-526. Eggins, S.M., 1992. Petrogenesis of Hawaiian tholeiites, 2. Aspects of dynamic melt segregation, Contrib. Mineral. Petrol., 110: 398-410. Frey, F.A. and Roden, M.F., 1987. The mantle source for the Hawaiian islands: constraints from the lavas and ultramafic inclusions. In: M.A. Menzies and C.J. Hawkesworth (Editors), Mantle Metasomatism. Academic Press, London, pp. 423-463. Fujimaki, H., Tatsumoto, M. and Aoki, K., 1984. Partition coefficients of Hf, Zr and REE between phenocrysts and groundmasses. J. Geophys. Res., 89: 662-672.

165

Grove, T.L. and Wagner, T.P., 1993. Is adiabatic melting of oceanic mantle a disequilibrium process? Constraints from experimental measurement of element diffusion rates in high-Ca pyroxene. Eos (Trans. Am. Geophys. Union), 74:284 (abstract). Hart, S.R., 1993. Fractal tree melting networks. Eos (Trans. Am. Geophys. Union I, 74:283 (abstract). Hart, S.R. and Brooks, C., 1977. The geochemistry and evolution of the early precambrian mantle. Contrib. Mineral Petrol., 61: 109-128. Hart, S.R. and Dunn, T., 1992. Experimental clinopyroxene/melt partitioning for 24 trace elements. Contrib. Mineral. Petrol., 113: 1-8. Hauri, E.H. and Hart, S.R., 1994. Constraints on melt segregation from mantle plumes: a trace element study of peridotite xenoliths from Savai'i, Western Samoa. J. Geophys. Res. (in press). Hofmann, A.W., 1986. Nb in Hawaiian magmas: constaints on source composition and evolution. In: S. Deutsch and A.W. Hofmann (Editors), Isotopes in Geology - - Picciotto Volume. Chem. Geol., 57:17-30 (special issue). Hofmann, A.W. and White, W.M., 1983. Ba, Rb and Cs in the Earth's mantle. Z. Naturforsch., 38a: 256-266. Hofmann, A.W., Feigenson, M.D. and Raczek, I., 1984. Case studies on the orgin of basalt, III. Petrogenesis of the Mauna Ulu eruption, Kilauea, 1969-1971. Contrib. Mineral. Petrol., 88: 24-35. Hofmann, A.W., Jochum, K.P., Seufert, M. and White, W.M., 1986. Nb and Pb in oceanic basalts: new constraints on mantle evolution. Earth Planet. Sci. Lett., 79: 33-45. Irving, A.J., 1978. A review of experimental studies of crystal/liquid trace element partitioning. Geochim. Cosmochim. Acta, 42: 743-770. Irving, A.J. and Frey, F.A., 1978. Distribution of trace elements between garnet megacrysts and host volcanic liquids of kimberlitic to rhyolitic composition. Geochim. Cosmochim. Acta, 42:771-787. Iwamori, H., 1993. Dynamic disequilibrium melting model with porous flow and diffusion-controlled chemical equilibration. Earth Planet. Sci. Lett., 114: 301-313. Johnson, K.T.M. and Kinzler, R.J., 1989. Partitioning of REE, Ti, Zr, Hfand Nb between clinopyroxene and basaltic liquid: an ion microprobe study. Eos (Trans. Am. Geophys. Union), 70:1388 (abstract). Jones, J.H. and McKay, G.A., 1992. REE partitioning between pyroxene/liquid and garnet/liquid: parameterization using Dca. Eos (Trans. Am. Geophys. Union), 73: 607 (abstract). LaTourrette, T.Z. and Burnett, D.S., 1992. Experimental determination of U and Th partitioning between clinopyroxene and natural and synthetic basaltic liquid. Earth Planet. Set. Lett., 110: 227-244. LaTourrette, T.Z., Kennedy, A.K. and Wasserburg, G.J., 1993. Thorium-uranium fractionation by garnet: evidence for a deep source and rapid rise of oceanic basalts. Science. 261: 739-742. Manhbs, G., Minster, J.-P. and Allbgre, C.J., 1978. Compar-

166

E.H. Hauri et al. / Chemical Geology 117 (1994) 149-166

alive uranium-thorium-lead and rubidium-strontium of St. Severin amphoterite: concequences for early solar system chronology. Earth Planet. Sci. Lett., 39: 14-24. McKay, G., Wagstaff. J. and Yang, S.-R., 1986. Clinopyroxene REE distribution coefficients for shergottites: the REE content of the Shergotty melt. Geochim. Cosmochim. Acta. 50: 927-937. McKenzie, D. and Bickle, M.J., 1988. The volume and composition of melt generated by extension of the lithosphere. J. Petrol.. 29: 625-679. McKenzie, D. and O'Nions, R.K., 1991. Partial melt distributions from inversion of rare earth element concentralions. J. Petrol., 32: 1021-1091. Meijer, A., Kwon, T.-T. and Tilton, G.R., 1990. U - T h - P b partitioning behavior during partial melting in the upper mantle: implications for the origin of high-,u components and the " P b paradox". J. Geophys. Res., 95: 433-448. Mezger, K., Essene, E.J. and Halliday, A.N., 1992. Closure temperatures of the S m - N d system in metamorphic garnets. Earth Planet. Sci. Lett., 113: 397-409. Nicholls, I.A. and Harris, K.L, 1980. Experimental rare earth element partition coefficients for garnet, clinopyroxene and amphibole coexisting with andesitic and basaltic liquids. Geochim. Cosmochim. Act& 44: 287-308. Obata, M., 1980. The Ronda peridotite: garnet-, spinel-, and plagioclase-lherzolite facies and the P - T trajectories of a high-temperature mantle intrusion. J. Petrol., 21: 533-572. Priem. H.N.A., Boelrijk, N.A.I.M., Hebeda, E.H., Oen, l.S., Verdurmen, E.A.T. and Verschure, R.H., 1979. Isotopic dating of the emplacement of ultramafic masses in the Serrania de Ronda, southern Spain. Contrib. Mineral. Petrol., 70: 103-109. Reisberg, L.. Zindler, A. and Jagoutz, E., 1989. Further Sr and Nd isotopic results from peridotites of the Ronda U1tramafic Complex. Earth Planet. Sci. Lett., 96:161 - 180. Salters, V.J.M. and Hart, S.R., 1989. The Hf-paradox, and the role of garnet in the MORB source. Nature (London), 342: 420-422. Salters, V.J.M. and Hart, S.R., 1990. The mantle sources of ocean islands and arc basalts: the Hf isotope connection. Earth Planet. Sci. Lett., 104: 364-380. Shen, Y. and Forsyth, D.W., 1993. The initial and final depths of melting beneath mid-ocean ridges. Eos (Trans. Am. Geophys. Union ), 74:283 (abstract). Shimizu, N. and Kushiro, I.. 1975. The partitioning of rareearth elements and barium between garnet and liquid at high pressures: preliminary, experiments. Geophys. Res. Lett., 2:413-416.

Skewes, M.A. and Stern, C.R., 1979. Petrology and geochemistry of alkali basalts and ultramafic inclusions from the Pali-Aike volcanic field in southern Chile and the origin of Patagonian plateau lavas. J. Volcanol. Geotherm. Res., 6: 3-25. Spiegelman. M. and Elliot, T., 1993. Consequences of melt transport for uranium series disequilibrium in young lavas. Earth Planet. Sci. Lett., 118: 1-20. Spiegelman, M. and Kenyon, P., 1992. The requirements for chemical disequilibrium during magma migration. Earth Planet. Sci. Lett., 109:611-620. Stern, C.R., Saul, S., Skewes, M,A. and Futa, K., 1989. Garnet peridotite xenoliths from the Pali-Aike alkali basalts of southernmost South America. In: Kimberlites and Related Rocks. Geol. Soc. Aust.. Spec. Pubt. No. 14. Tatsumoto, M.. Basu. A.R., Wankang, H., Junwen, W. and Guanghong, X.. 1992. Sr. Nd and Pb isotopes of ultramafic xenoliths in volcanic rocks of Eastern China: enriched components EM l and E M I I in subcontinental lithosphere. Earth Planet. Sci. Lett.. I13: 107-128. Todt, W., Cliff, R.A., Hanser, A. and Hofmann, A.W., 1984. 2°2Pb + 2°5pb double spike for lead isotope analysis. Terra Cognita, 4:209 (abstract). Vannucci, R., Shimizu, N., Piccardo, G.B., Ottolini, L. and Boltazzi, P., 1993. Distribution of trace elements during breakdown of mantle garnet: an example from Zabargad. Contrib. Mineral. Petrol., I I 3: 437-449. Watson. E.B., Ben Othman, D., Luck, J.M. and Hofmann, A.W., 1987. Partitioning of U, Pb, Hf, Yb, Cs, Re, Os between chromian diopsidic pyroxene and haplobasaltic liquid. Chem. Geol., 62:191-208. Watson, S. and McKenzie, D., 1991. Melt generation by plumes: a study of Hawaiian volcanism. J. Petrol., 32:501 537. Williams, R.W. and Gill, J.B., 1989. Effects of partial melting on the 238U decay series. Geochim. Cosmochim. Acta, 53: 1607-1619. Zindler, A., 1980. Geochemical processes in the earth's mantle and the nature of crust-mantle interactions: Evidence from studies of Nd and Sr isotope ratios in mantle derived igneous rocks and lherzolite nodules. Ph.D. Thesis, Massachusetts Institute of Technology, Cambridge, Mass. Zindler, A. and Jagoutz, E., 1988. Mantle cryptology. Geochim. Cosmochim. Acta, 52: 319-333. Zindler, A., Staudigel, H., Hart, S.R., Endres, R. and Goldstein, S., 1983. Nd and Sr isotopic study o f a mafic layer from Ronda ultramafic complex. Nature ( London ), 304: 226-230.