Determining the composition of high-pressure mantle melts using diamond aggregates

Geochimieaet Cosm~himica Acta, Vol. 58, No. 13,pp. 28 I I-2827, 1994 Copyright0 1994EkevierScienceLtd Printed in theUSA. All rights reserved 0016-7037/94$6.00 + 00

Pergamon

0016-7037(94)00080-8

Determining the composition of high-pressure mantle melts using diamond aggregates MICHAEL B. BAKER and EDWARD M. STOLPER Division of Geologicaland PlanetarySciences, California Institute of Technology, Pasadena. CA 9 I 125,USA

(Revived February 16, 1993; accepfed in revisedJ&n Feb~ia~~ 10, 1994)

Abstract-We present a new experimental technique for circumventing the quenching problems that have plagued high-pressure peridotite melting studies. A thin layer of -50 pm diamonds is placed above a layer of peridotite powder. Partial melt extracted from the peridotite layer collects in the pore spaces between the diamonds and equilibrates diffusively with the residual peridotite mineralogy, Isolated from the crystalline residue, the melt quenches to a glass that records the composition of the liquid coexisting with the residual crystalline phases under the conditions of the experiment. We have used this technique to investigate partial melting of a fertile mantle composition at 10 kbar and a temperature range of 12701390°C. Oxide concentrations in the liquids from the longest duration runs (up to 15 1 hours) vary systematic~ly with increasing tempemture: TiOz, A&03, and NazO decrease monotonically, while CrZOj, FeO*, and MgO increase steadily. CaO shows more complicated behavior, first increasing and then decreasing, with the crest in the temperature-CaO trend approximately coincident with the disappearance of clinopyroxene from the residue between 1330 and 1350°C. Overall variation in silica content with temperature is small, and there appears to be a minimum at about 12% melting. The compositions of liquids produced in time series, temperature reversal, and two-stage experiments (conducted to test the technique) all indicate that our experimentally determined liquid compositions represent close approaches to equilibrium. Calculated melt fractions (F) also vary systematically with temperature. The slope of the T(Y)-F curve is not constant over the spine1 lherzolite melting interval, but decreases as temperature increases from 1270 to 1330°C. Extrapolating the curve back to zero melt suggests that the anhydrous solidus temperature for our peridotite starting composition is - 1240°C. At temperatures below the cpxout curve, melt generation occurs via the reaction, 0.38 opx + 0.7 I cpx + 0.13 sp -+ 0.22 oliv + I .O liq, and the proportions of minerals that enter the melt appear to be independent of temperature. At temperatures above cpx-out, the less well constrained melting reaction is: 1.06 opx + 0.04 sp = 0.1 oliv + 1 Iiq. The fact that all of the 10 kbar melts have FeO* contents that are substantially lower than those reported in any primitive MORB glasses further strengthens the conclusions that these glasses are not 10 kbar primary melts, that they involve a component of higher pressure partial melting, and that they have evolved by significant olivine fractionation from more primitive liquids. Our experimental data also provide an independent check of the results of recent peridotite partial melting calculations. Efforts to parameterize the experimental database on peridotite melting, and to calculate melt compositions as a function of P, T, and F are partially successful in reproducing the compositional trends determined in this study. ments on Hawaiian pyrolite minus 40% olivine and Tinaquill0 lherzolite minus 40% olivine, JAQUES and GREEN ( 1980) were able to reconstruct the liquid compositions by visually estimating the mode of each experiment, determining the compositions of the solid phases, and then calculating the liquid composition by mass balance. Although subject to large uncertainties, their dataset has been used as a basis for two modeling studies of mid-ocean ridge basalt (MORB) petrogenesis (KLEIN and LANGMUIR, 1987; NW and BATIZA, 199 1). The “sandwich” technique was intr~uced to circumvent the quenching problems and has provided insight into the compositions of basaltic liquids saturated with ohvine (oliv), orthopyroxene (opx), clinopyroxene (cpx), and spine1 (sp) at high pressures (STOLPER, 1980; TAKAHASHI and KIJSHIRO, 1983; FUJIIand SCARFE, 1985; FALLOON andGREEN, 1987; FALLOON et al., 1988). However, it is difficult to use directly the results of these experiments to follow the melting ofspecific peridotite compositions, because the experimental bulk compositions (i.e., basalt + peridotite) are substantially different from representative mantle peridotite. Moreover, as noted by PRESNALLand HOOVER (1987) and GROVE et al.

DETERMINING THE COMPOSITIONand fraction of melt near the solidus of mantle peridotite as a function of temperature, pressure, and bulk composition is a difficult problem in experimental petrology, but one whose solution is fundamental for understanding the petrogenesis of basaltic magmas. In principle, such experiments are easy to design and carry out; a peridotite sample is loaded directly into a capsule, melted, allowed to equilibrate at pressure and temperature, and then quenched. The difficulty is that low-melt-fraction liquids undergo extensive crystallization on quenching, and thus recognizing what was liquid under the conditions of the experiment and reconstructing its composition are often impossible (CAWTHORN et al., 1973; NICHOLLS, 1974; JAQUES and GREEN, 1979). For this reason, there are few reliable data that provide information on the composition and percentage of melt coexisting with mantle peridotite over a range of pressures and temperatures. This lack of data is especially striking near the peridotite solidus, where the partial melt coexists with a lherzolitic assemblage. In a series of experi2811

2812

M. 8. Baker and E. M. Stolper

(1990), many of the multiply saturated liquids from these experiments may not be directly analogous to mantle partial melts (e.g., many are too rich in FeO*, total Fe as FeO, relative to MgO to be in equilibrium with any reasonable mantle residue), and in some cases, the experimental glasses may not actually be saturated with all of the minerals present in the capsule (KINZLER and GROVE, 1992a). In an effort to circumvent these problems, KINZLER and GROVE (1992a) ex~rimentally produced a suite of basaltic melts saturated with oliv + opx + cpx + sp + plagioclase over a range of temperatures and pressures. They parameterized the multiply saturated liquid surfaces using temperature, pressure, and compositional terms and calculated liquid compositions for both batch and fractional mantle melting paths. Other parameterizations used to model mantle melting include those of MCKENZIE and BKKLE (1988). NIu and BATIZA (1991). WATSON and MCKENZIE (1991), and LANGMUIR et al. ( 1992). In this paper, we describe a technique inde~ndently developed by us (BAKER et al., 1992) and by KUSHIRO and his coworkers (HIROSEand KUSHIRO, 1992, 1993: JOHNSONand KUSHIRO, 1992; KUSHJROand HIROSE, 1992) that circumvents the quenching problems, thus making it possible to determine the compositions of liquids present in peridotite melting experiments, even near the peridotite solidus. The experiments are done by placing a thin layer of diamond powder above a sample of peridotite; the technique is similar to that used by RYABCHIKOV et al. (1989) to investigate highpressure ~~dotite-~uid interactions. When the diamond + peridotite are brought to run conditions, the peridotite partially melts and some of the liquid flows into the open spaces between the diamonds. These pores remain open due to the high yield strength of the diamond. The liquid is probably driven into these pores by a large pressure gradient between the silicate (at the pressure of the experiment) and the air-filled pores (at -atmospheric pressure). Although separated, the liquid in the diamond layer communicates by diffusion with the liquid and crystals in the adjacent peridotite and, given sufficient time, equilibrates with the crystals in the peridotite layer. On quenching, the melt within the diamond layer converts to glass without interacting with the residual peridotite crystals. Even if the melt interstitial to the diamonds were to crystallize on quenching, which may occur under some conditions, the bulk composition of the quenched aggregate would be unchanged from the liquid present under the conditions of the experiment. We have applied this new technique to the melting of an anhydrous, model upper mantle peridotite, and in this paper we report the results of our 10 kbar experiments at I2701390°C. We have concentrated on the temperature interval where clinopyroxene coexists with meit, because clinopyroxene controls the abundances of several major and trace elements in the liquid during partial melting. We show how data of this sort can be used to quantify melting reactions of mantle peridotite (i.e., the proportions of solid phases that enter the melt with increasing temperature). Our goal is to illustrate the potential of this technique for generating a definitive dataset on melting of peridotite and other rock compositions, and thus direct petrogenetic insights are limited. Nevertheless, we note that our data and the data of HIROSE

and KUSHIRO (1993) are inconsistent with a direct, shallow partial melting origin for MORBs; these data also preclude any substantial reequilibration between the mantle and the parental liquids of MORBs at relatively low pressures (ie 1 5 10 kbar). Finally, while some recent parameterizations of peridotite melting predict trends in liquid composition that are broadly consistent with our data, (NIU and BATIZA, 199 I ; WATSON and MCKENZIE, 199 1: KINZI.EK and G~ovt:. t992ab, 1993; LANGMUIRet al., 1992). we expect

the dia-

liquid compositions and provide critical new constraints on the compositions of near-solidus mantle melts over a range of temperatures and pressures. mond

technique

to provide

direct

determinations

EXPERIMENTAL AND ANALYTICAL

of

TECHNIQLJES

Starting Materials and Bulk Composition The synthetic mantle composition used in this study WCS COW strutted using minerals separated from a disaggregated Kilbourne Hole spinet lherzolite nodule. Compositions of the olivine. ortho-

pyroxene. efjnopyroxene, and spine1 were determined by eiectron microprobe as described below and are listed in Table I. Coarse mineral separates were handpicked under a binocular microscope, washed for 20 min in warm 2.4 M HCI, rinsed in distilled water, and then crushed and sieved. For each mineral. the 16-28 mesh size fraction was handpicked again to improve its purity, reduced to 200-325 mesh, and then washed again in warm 2.4 M HCI and distilled water. The starting material (designated as MM3) was produced by combining the olivine, orthopyroxene, chnopyroxene, and spine1 in the proportions 0.50:0.3&O. 1?:0.03 by weight and then grinding for an hour under ethanot. After this final grinding. the majority of the grains were IO-30 jrm in size. The bulk composition listed in Table I is based on the individual mineral analyses and their proportions in the mix. Within the constraints of the mineral compositions. our bulk composition approximates closelythe various estimates of “primitive” upper mantle (e.g.. RINGWOOD,1979: HART and ZINDLER, 1986; MCDONOIJGI-~ and FREY, 1989). Compared to the three primitive mantle estimates listed in Table I. MM3 is slightly lower in FeO* (by 6- 12%relative), but is in the range of all other major elements. Based on charge balance calculations for each mineral composition, we estimate that our peridotite starting material has a molar F$‘/Fe*+ of -0.08. With respect to minor elements, MM3 is higher in CrZOr (bv 4%75% relative,. lower in TiOz (by 40-50s relative), and has negli~ble K,O compared to 0.03% in the primitive mantle estimates. Experimental Methods

Experiments were run in a 1.27 cm piston-cylinder apparatus using inner pieces of crushable MgO that had been dried at 1COOT,straightwalled graphite furnaces, and outer sleeves of calcium fluoride. All runs were done using graphite crucibles sealed inside Pt capsules. Except for #55T and #62T (described below), capsules were constructed as follows: the starting mixture. dried for -30 min in a vacuum oven at I iO”C, was loaded into the graphite capsule after the layer of diamond powder: the capsule was then placed into an open Pt tube that had been welded shut at one end; graphite powder was packed atop the graphite capsule and the open end of the Pt tube was crimped shut: the capsule was returned to the vacuum oven for another 4-6 h, after which the crimped end was welded shut. The diamond-size fraction was 44-54 or 40-50 pm. in most experiments. diamonds constituted IO-20 wt% of the material loaded into the graphite capsule. The exceptions were the experiment at 1390°C. the two reversal runs at 1310°C. and the one hour experiment at 1330°C: in these cases, the diamond constituted 23-35% of the total mass. All capsules were run with the diamond layer above the silicate powder. Figure I shows a cross-section of a typical capsule following an experiment. Runs #55T and #62T involved a two-stage procedure designed to minimize the length of the diffusive path between the

Experimental

study of partial

melting

2813

in the mantle

Table 1. Mineral compositions, starting mix and other upper mantle compositionst SiO2

TioZ

4203

CrzO3

FeO*

MnO

MgO

CaO

Oliv(4)§

40.4(l)

-

0.03( 1)

9.20(9)

0.14(l)

49.80(8)

0.08(l)

OPX(8)§

54.7(2)

0.11(l)

4.73(7)

0.49(l)

5.88(4)

0.14( 1)

32.9(2)

0.86(2)

CPX(9)B

51.8(l)

0.43(S)

6.8(l)

1.09(2)

2.85( 11)

0.08(2)

15.7( 1)

54.99(3)

13.00(S)

0.10(l)

21.01(3)

0.13(l)

38.3(2)

SP(5)S MM3 M&F

0.04( 1) 0.12(l) 45.5(2) 45.08

0.11(l) 0.2 1

3.98(10) 4.48

0.68(l)

10.39(6) 7.18(14) 8.05

0.38

0.14

37.44

19.6(2)

Na20

K20

NiO

Mg#*

0.37( 1)

90.6

0.12(l)

0.10(l)

90.9

1.62(6)

0.05(Z)

90.9

0.35(l)

78.3

0.31(2)

0.23

90.5

3.62

0.34

0.03

0.22

90.5

0.03

0.2

89.2

0.03

0.28

89.5

3.57(4)

R

45.1

0.2

3.3

0.4

8.0

0.15

38.1

3.1

0.4

H&Z

45.96

0.18

4.06

0.47

7.54

0.13

37.78

3.21

0.33

J&G HPy

45.2

0.7 1

3.54

0.43

8.47

0.14

37.5

3.08

0.57

0.13

0.20

89.9

J&G HPy-40

47.9

1.18

5.91

0.72

8.81

0.13

28.8

5.14

0.95

0.22

0.13

85.4

M&B

45.53

0.17

3.09

8.50

-

39.20

2.59

0.44

0.06

W&M

45.16

0.17

3.31

-

8.5

-

39.2

2.78

0.42

0.07

N&B

44.74

0.17

4.37

0.45

7.55

0.11

38.57

3.38

0.40

0.0

H&K HK-66

48.02

0.22

4.88

0.25

9.90

0.14

32.35

2.97

0.66

0.07

H&K KLB-1

44.48

0.16

3.59

0.31

8.10

0.12

39.22

3.44

0.30

0.02

89.2 89.2 0.26

90.1 85.3

0.25

89.6

t Abbreviations: oliv=olivine, opx=orthopyroxene, cpx=clinopyroxene, sp=spinel. MM3 = bulk composition used in this study; other estimated primitive mantle compositions: M&F = McDonough and Frey (1989); R = Ringwood (1979); H&Z = Hart and Zindler (1986). Mantle compositions used in other experimental and numerical melting studies: J&G = Jaques and Green (1980), Hawaiian pyrolite and Hawaiian pyrolite minus 40% olivine; M&B = McKenzie and Bickle (1988); W&M = Watson and McKenzie (1991); N&B = Niu and Batiza (1991); H&K = Hirose and Kushiro (1993). HK-66 and KLB-1. Numbers in parentheses adjacent to the analyses are 10 in terms of the least units cited; e.g., 40.4(l) represents 40.4M.l. 5 Minerals separated from a Kilboume Hole nodule; number of electron microprobe analyses in parentheses, with each analysis taken on a separate grain. i Mg# = lOOMg/(Mg+Fe) on a molar basis with all iron as Fe2+.

partial melt and the residual peridotite minerals. Capsules were loaded with peridotite powder without diamond and run for either 120 h at 1270°C or 96 h at 1330°C; after quenching each run, the coherent peridotite plug was extracted from the capsule, reloaded along with a layer of 44-54 pm diamond into a new graphite and Pt assembly, and rerun at the same pressure and temperature for 26-31 h. The results of these experiments are discussed below. Pressure was applied using the hot piston-out technique and has not been corrected for possible friction effects. Run temperature was monitored and controlled to within I “C using a W,Re/W2,Re thermocouple, and no pressure correction was applied to thermocouple EMF. Based on previous determinations of the thermal profile of our run assembly, we estimate that temperatures are accurate to +510°C and that the thermal gradient over the region of the capsule containing the peridotite and diamond layers is <5”C. Total run times ranged from I - I5 1 h. Power consumption increased slightly during the initial IO-20 h of a given run, and then remained nearly constant. Inspection of the thermocouple wires, especially near the

Pt

2mm FIG. 1. Cross-section of run #I8 drawn diamonds/(diamonds + silicate) = 0.17.

to scale. The mass ratio

steel base plug, revealed only minor oxidation in the longest duration experiments. This and the near constant power readings suggest minimal thermocouple drift. Experiments were terminated by shutting off the power and initial quench rates were on the order of 102”C/s. Although the oxygen fugacity is uncontrolled in these experiments, the presence of graphite 1imitsf02 to below the graphite-C-O vapor buffer @CO), which lies in the wiistite stability field under the conditions of our experiments (TAYLOR and GREEN, 1989; ULMER and LUTH, I99 1). If metastable equilibria involving diamond influence the oxygen fugacity of the experiment, then the upper limit toJO will be lower than GCO. After quenching, each capsule was sliced down its long axis with a diamond wafering blade and one-half was mounted in epoxy. The diamond layer and its interstitial glass were scraped out of the other half of the capsule, and this material was mounted in a separate epoxy plug. Both plugs were polished with 1 and 0.25 pm diamond powder. In the case of the disaggregated material, the -5- 15 pm glass shards generally polished well with little relief. In contrast, polishing the half of the capsule that contained the coherent diamond and peridotite layers did not produce a flat surface in the region of the diamonds. The surfaces of the glass within the diamond matrix were usually concave upward and were several tens of microns below the irregular surface defined by the “tops” of the diamonds.

Analytical

Techniques

Crystalline phases were analyzed with the Caltech 5-spectrometer JEOL 733 microprobe using a I5 keV accelerating voltage, a 10 nA beam current, and a I pm spot. Data were reduced using a modified ZAF procedure (CITZAF, ARMSTRONG, 1988). Interstitial glass within the diamond and peridotite layers as well as the glass separated from the diamond layers and mounted as glass shards were analyzed with a Camscan scanning electron microscope (SEM) fitted with an energy dispersive system. The beam current was 0. I5 nA measured in a Faraday cup in the sample holder and spectra were collected for 200 s. The raw counts were reduced using CITZAF after first normalizing the unknown k-ratios to those of a basaltic glass standard (VG-2; see Table 2) that was analyzed during each SEM session. Multiple analyses of a fused sample of BCR- I (I 5OO”C, 10 kbar,

2814

M. B. Baker and E. M. Stolper

Table 2. Experimental results Run#t Temp Time Silicate (“C) Ous) (mg)

a Tatal

Diamonds (microns)

Run pmductss

Phase proportions~ 27.4(1.0),55.1(1.2),

24

1390

47.7

4.89

0.68

44-54

liq. oliv. opx. sp

13

1360

24.1

8.84

0.82

44-54

liq. oliv, opx. tr cpx. sp

22

1360

58.0

6.48

0.81

44-54

liq. oliv. opx. sp

23.9(1.2). 54.q1.4). 21.2(1.0).0.33(8)

21

1350

53.7

7.63

0.82

40-50

liq. oliv, opx. sp

22.4(1.4). 54.7(1.3). 22.4(1.4). 0.47(9)

28

1330

1.0

4.59

0.73

44-54

liq. oliv, opx. cpx, sp

18

1330

26.4

5.70

0.83

40-50

liq. oliv, opx. cpx. sp

19

1330

42.3

6.68

0.77

40-50

liq. oliv. opx. cpx. sp

24

1330

71.9

6.07

0.79

44-54

liq. oliv, opx, cpx. sp

62T

1330

122.5

6.26

0.84

44-54

liq. oliv, opx. cpx. sp

44R

1310

32172.7 6.32

0.77

44-54

liq. oliv. opx. cpx, sp

43R

1310

4Ol73.5

5.34

0.77

44-54

liq, oliv, opx. cpx. sp

17

13w

24.1

7.45

0.88

44-54

liq, oliv, opx. cpx. sp

14

1300

46.7

8.43

0.81

44-54

liq, oliv. opx. cpx. sp

17.2(1.3),0.26(11)

17.6(1.2). 54.5(1.3). 23.4(1.2).4.0(1.1). 0.48(12)

16

13M)

72.0

4.91

0.84

44-54

liq. oliv. opx. cpx. sp

12

1280

12.0

7.98

0.81

44-54

liq. oliv. opx. cpx. sp

15

1280

51.6

8.30

0.81

44-54

liq. oliv, opx. cpx. sp

10

1270

72.0

5.11

0.88

40-50

liq. oliv. opx. cpx. sp

20

1270

97.9

8.5 1

0.84

40-50

liq. aliv. opx. cpx, sp

7.411.0), 51.9(1.0),27.2(1.1),

55T

1270

151.3

7.02

0.87

44-54

liq. oliv. opx. cpx. sp

8.4(1.1), 53.2(1.2), 26.2(1.0), 11.4(9), 0.9(2)

13.4(1.4).54.2(1.5).24.4(1.1).7.3(1.2).0.67(121 9.2(1.1),54.8(1.3).23.5(9).

11.2(1.4). M(3) 12.2(8). 1.2(4)

addition 26utd 3i.6hrs.

mspectively. 1; tr=tmce. t phase prnpmtions only calculated for longerdurationexperiments.Values in parenthesesare uncertainties; read27.4( 1.O) as 27.4fl .O.

0Abbreviations asin Table

0.5 h) collected over a one month period provide an estimate of the accuracy and precision of the EDS system (Table 3). Based on thirtytwo analyses, relative errors (100 X 1o/mean value) for the major oxides are 0.43% SiO*, 3.0% Ti02, 0.57% A1203, 1.8% FeO*, 2.8% MgO, 1.7% CaO, and 6.6% NazO. Since the gla$sesin our experiments have about one-halfthe FeO* content of BCR- 11relative FeO* errors due to machine precision may be slightly greater than 2% (perhaps as high as 4%). Analyses of glass in place within the diamond layer were generally more variable than analyses of the disaggregatedglass shards. Thisgreater variability is probably produced by the concavity of the glass surfaces within the diamond layer, and by partial interference of the diamonds with the emitted X-rays. For these reasons, we report only glass data collected on the disaggregatedshards. Although we did not analyze for C in the glasses,the experimqntal work of PAN et al. (I 99 1) and HOLLOWAY et al. (1992).coupled with the low Fe203/Fe0 in our bulk composition, suggestthat dissolved carbonate abundances in our partial melts would be at most several tenths of a percent CO* by weight. We do not expect these levels of dissolved carbonate to influence peridotite melting relationships. EXPERIMENTAL

RESULTS

AND

DISCUSSION

Experimental details and the phase assemblage in each charge are reported in Table 2. Compositions of glasses and residual crystals from selected runs are listed in Table 3, and glass compositions from long duration experiments and the reversals are displayed graphically in Fig. 4. Figure 2 compares for an experiment at 1360°C the composition of the glass interstitial to the diamond with that of interstitial glass from the underlying silicate layer. Figure 2 shows clearly that quenching significantly modifies the composition of the melt interstitial to the residual peridotite crystals. The glass pockets within the silicate layer are depleted in MgO and enriched in the other major oxides relative to the glass within the diamond layer. It. is this compositional modification (discussed in the Introduction) that led us to develop the technique described in this paper.

Attainment of Equilibrium

Because partial melt moves into the diamond interstices on a short timescale (< 1 h), the liquid initially present in the diamond layer is out of equilibrium with the residual peridotite. At least two factors are responsible for this disequilibrium: the failure of the first formed melt to interact with the cores of residual crystals by diffusion and/or dissolution and reprecipitation, and the failure of the different minerals in the solid assemblage to melt in equilibrium proportions due to variable melting kinetics among the solid phases. Compositional zonatiqn jn residual crystals (see below) and the high values of K$$& in hour long experiments by us and JOHNSONand KUSHIRO (1992) demonstrate that the first of these factors is important. Other variations in glass composition as a function of time in our experiments and those of JOHNSONand KUSHIRO(1992) suggest that the second factor also plays a role. For example, one hour runs produced liquids with anomalously high silica contents relative to one day and longer runs, suggesting that on short timescales the amount of orthopyroxene dissolving into the melt, and/or the amount of olivine precipitating from the melt, were above the equilibrium values. We depend on long run times to overcome both of these effects, and it is thus essential to determine the timescale required for the segregated melt and the residual crystals to approach equilibrium closely. We have used time series, reversal, and two-stage experiments to address this question. Figure 3 shows concentrations of the major oxides in quenched glass from a series of our IO kbar experiments at 133O“C. Run durations varied from l-72 hours and all charges contained olivine, orthopyroxene, clinopyroxene, and spinel, in addition to melt. Glass compositions in experiments

2815

Experimental study of partial melting in the mantle Table3.Compasitionsofexptimentalnm RunX

Phases

26

Wl)

51a(2)

oliv(7)

40.7(2)

OPXO)

55.2(3)

Si@

Tit&

prcductst

d2%

0.35(6) 12.1(l) 0.06(l) 0.01(l)

Cr203

FeO'

MnO

MgO

O.l7(5)

l5.7(2)

0.58(6)

7.4(3)

0.37(4)

7.90(11) 0.14(4)

2.69(13) 1.42(13) 4.96(12) 0.12(2)

0.42(4) 0.15(3) 23.3(3)

Na20 1.0(l)

0.28(2) -

33.99(14) 1.94(4) 0.05(l)

10.2O(6)

0.18(3)

l8.2(2)

13

iiq(12)

50.9(2)

0.37(Z) 12.8(l)

0.52(E)

7.0(2))

0.13(7)

15.0(2)

12.0(2)

1.34(V)

22

liq(7)

50.5(3)

0.42(2) 13.0(l)

0.50(6)

7.3(3)

0.15(6)

l4.5(2)

12.7(4)

1.12(8)

oliv(8)

40.7(3)

o.zvr3)

8.44(12) O.OV(2)

51.1(4)

OPX(8)

54.7(3)

5.1(2)

O.ll(2)

33.1(3)

2.3(2)

10.22(10) 0.19(l)

18.0(2)

O.ll(5)

SP(4)

SP(5) 21

0.01(l)

3.0(4)

0.24(Z) O.lO(3) 24.0(4)

liq(l4)

50.4(3)

OliV(9)

40.4(2)

OPHll)

55.4(4)

SP(5)

0.06(l)

0.42(7) 13.4(l) 0.05(l) O.o4(4)

2.8(3)

0.23(4) 0.14(6) 24.5(2)

49.1(E)

51.6(3)

CaO 11.5(2)

1.3700) 47.6(4) 0.44(7)

7.1(3)

0.14(6)

13.9(2)

0.2X4)

8.5(2)

O.ll(3)

50.8(2)

1.30(14) 5.20(14) O.li(3) 46.8(4)

0.07(7) -

0.30(2)

13.1(2)

0.06(2) 1.28(13)

0.32(3) -

33.04(14) 2.55(10)0.07(l)

10.9(2)

0.14(2)

18.0(2)

0.15(8) -

28

liq(8)

51.1(5)

0.42(3) 12.8(3)

0.48(4)

7.0(2)

O.lO(8)

13.8(2)

12.9(2)

l.5(2)

18

liq(ll)

50.3(3)

0.43(5) 14.8(2)

0.32(7)

6.7(2)

0.13(2)

12.9(2)

12.6(3)

l.66(14)

IV

liq(ll)

50.2(2)

0.48(6) 15.1(l)

0.33(6)

6.8(Z)

0.13(7)

12.6(2)

12.8(2)

1.62(11)

24

liq(l2)

50.3(3)

0.47(6) 14.7(l)

0.34(6)

6.7(2)

0.13(V)

13.1(2)

12.8(2)

1.5(2)

oliv(l2) 40.5(l)

0.07(4)

0.24(3)

8.6(2)

0.10(2)

50.5(2)

opx(8)

54.4(3)

O.OV(2)

3.3(4)

1.22(6)

5.1(2)

0.11(l)

32.70(14) 2.4(2)

0.052(V)

CPX(8)

52.4(4)

0.05(3)

4.23(E)

1.67(V)

3.5(2)

O.lO(2)

21.3(4)

0.30(2)

18.16(10) 0.07(l)

0.32(2) 0.17(2) 28.5(3)

l5.7(5)

9.9(2)

0.17(2)

44R

liq(l0)

50.0(4)

0.54(6) 16.3(2)

0.18(V)

6.4(2)

0.16(S)

12.0(2)

12.3(4)

43R

liq(6)

49.9(2)

0.48(7) 15.7(2)

0.27(8)

6.6(2)

0.18(6)

12.4(2)

12.6(2)

1.9(l)

16

liq(7)

50.0(2)

0.54(6) 16.1(l)

0.24(8)

6.5(2)

0.12(7)

12.1(l)

12.5(2)

1.9(2)

oliv(6)

40.8(3)

O.O6(3)

0.14(2)

9.1(2)

0.14(2)

51.0(2)

OPX(4)

54.7(3)

0.08(S)

4.4(3)

1.16(12) 5.56(V)

0.12(3)

32.6(2)

cpx(6)

52.1(4)

0.13(6)

5.3(4)

1.50(8)

3.7(2)

O.ll(2)

20.6(2)

10.0(2)

0.12(3)

1X78(7)

SP(3)

sP(5) liq(v)

50.9(2)

oliv(4)

40.8(2)

opx(6)

54.4(4)

CPX(5)

51.8(4)

SP@)

0.69(8) l7.0(2)

32.9(6) O.l6(3)

6.1(2)

0.09(P)

11.1(l)

0.03(2)

0.13(6)

8.98(E)

0.13(l)

49.8(2)

0.07(4)

4.5(4)

l.oO(l3) 5.5(2)

0.14(2)

32.2(2)

0.14(7)

5.0(7)

1.46(8)

3.6(2)

0.07(2)

l9.6(6)

9.3(2)

O.ll(4)

20.0(l)

0.18(3) 0.17(3) 44.2(7)

25.0(E)

2.0(2)

0.28(l) 2.36(6) 0.05(l) l6.6(3)

0.36(3)

O.OV(2) ll.6(2)

2.4(2)

0.22(l) 2.06(5) 0.05(2) 17.2(2)

0.44(5)

0.08(3) -

liq(l0)

51.2(2)

0.64(6) 18.6(l)

O.ll(6)

5.7(2)

0.12(6)

9.5(l)

11.0(2)

3.0(2)

liq(l0)

50.9(2)

0.66(6) 17.9(l)

0.12(6)

5.9(2)

O.l4(6)

10.4(2)

ll.2(3)

2.7(2)

oliv(6)

40.2(4)

0.04(3)

0.12(7)

9.2(2)

0.16(2)

50.4(4)

OPX(4)

53.4(l)

O.ll(6)

5.1(2)

0.8(3)

5.74(13) 0.14(2)

32.7(E)

1.5(6)

0.07(2)

CPX(4)

51.6(2)

0.22(E)

5.6(4)

l.25(10) 3.24(5)

0.12(3)

19.0(4)

17.7(4)

0.52(3)

25.6(1.0) 9.50(13) O.ll(4)

20.4(3)

SP(4) 55T

0.05(4) 0.09(4) 37.6(6)

42.1(4)

0.34(4)

O.lV(7) O.OV(2) 44.2(S)

liq(l0)

50.2(2)

oliv(7)

40.7(2)

opx(6)

54.0(3)

0.10(4)

CPX(8)

51.3(2)

0.18(2)

SP(5)

0.67(6) 18.1(2)

O.ll(8)

6.1(2)

0.12(7)

10.5(2)

0.13(l)

9.14(10) O.ll(l)

50.5(2)

4.8(2)

1.1(l)

5.60(5)

0.16(4)

32.341)

5.8(2)

1.30(V)

3.60(7)

0.09(2)

l9.4(3)

9.80(14) 0.09(l)

20.2(2)

0.27(3) 0.14(4) 43.4(5)

27.0(4)

0.25(3) -

0.05(l) il.4(2)

2.8(2)

0.22(2) 2.11(6) 0.07(l) 17.3(3)

0.47(3)

0.09(2) -

BCR-It

55.6(2)

2.36(7) 14.05(S)

-

12.3(2)

0.23(6)

3.26(10) 7.11(12) 3.2(2)

VG-2*

5l.oo

1.86

-

11.88

0.22

6.74

14.11

11.16

2.63

tSolidphasecompositionsareonly~ponedforthelongerdunuionexperimentateachtempnaturc;sceTable2. Numbersin parentheses nexttoeachanalysisare lointem~softheleastunitscitcd; e.g..5l.O(2)repnsents 51.0f 0.2. 5 phase abbreviationsasinTable 1;numberofanalysesaregiveninparentbeses. t Meanand l~of32analysesofh1sedBCR-1(10kbar.15Oo~C,0.5 hrs)cooaed wera~ne month period.Accepted compositionofstandardglassVG-2,nonnalizcdtoloo wt%,fromJarosewichetal.(l979).

of 226 h do not change systematically with time. Although we demonstrate below that longer times are necessary togenerate homogeneous residual crystals, the results shown in Fig. 3 suggest that in our experiments at 1330°C the liquid begins to approach a constant composition on a timescale of tens of hours. Based on these results, the compositional changes in the liquid associated with the final homogenization of the crystals appear to be subtle and minor compared to the changes occurring over the first day or two. The small, apparently unsystematic variations in glass composition in the longer duration experiments shown in Fig. 3 could reflect the 59°C uncertainty in temperature in our experiments, since the concentrations of the major oxides in the quenched glasses vary strongly with temperature (Fig. 4).

Since time-invariant liquid compositions in long duration experiments do not demonstrate conclusively that these runs have approached equilibrium, we have also conducted a pair of reversal experiments, where we have approached a final temperature of 13 10°C from both higher and lower temperatures. A charge was initially held at either 129O“C (40 h) or 1330°C (32 h), and then the temperature was either raised or lowered to 13 lO”C, and both runs were continued for an additional 73 h. The liquid compositions from these two experiments are plotted in Fig. 4, and within the analytical uncertainty, the quenched melts bracket the temperaturecomposition trends for each oxide. Based on the results from our time series experiments, run times of 40 and 32 h should have been sufficient to generate partial melts in the diamond

M. B. Baker and E. M. Stoiper

2816 Run #Xi, Ill kbar, 1360*F

16

. 12 4

’ -I

’

1 10

’

’ 13

’

16

wt% MgO

FIG. 2. Com~~son of the com~sition of interstitial glass from the diamond layer with the range of interstitial glass compositions from the peridotite layer of run #22. The filled cirete represents the mean of 7 analyses; lo values are smaller than the symbol. Each open circle represents a single glass analysis: errors based on counting statistics are smaller than the symbols.

layers that were similar to the 1290 and 1330°C compositions defined by the trends in Fig. 4. Clearly, over the final 73 h at 13 IO’C, the composition of the liquid in each diamond layer changed in response to the change in temperature and approached very similar compositions that bracket the results of our single-stage melting experiments. Although the reversal experiments indicate that liquids in the diamond layers can “reequilibrate” with the residual peridotite minerals over a period of 73 h, low diffusion rates for components such as silica may still limit the approach to equilibrium in both the melting experiments and the reversal runs. In order to evaluate the importance of this effect and to limit the diffusive length scale within the capsule, we conducted two-stage experiments at 1270 and 1330°C. As discussed above, two capsules containing only peridotite were run at 10 kbar for either 120 h at 1270°C or -96 h at 1330°C. During the first stage of each experiment, all of the liquid remains in intimate contact with the residual peridotite and solid phase dissolution/reprecipitation and solid-state diffusion are the rate-limiting steps for equilibmtion between crystals and liquid (i.e., diffusion of components in the liquid is not a rate-limiting step). After quenching each sample, each peridotite plug was removed, placed inside another capsule with a layer of diamonds, and run at the same temperature and pressure for either 32 h (1270°C) or 26 h ( 1330°C). During these second experiments, the glass (and quench phases) within each peridotite layer remelts and liquid migrates into the diamond layer. In contrast to the single-stage experiments, the melt that migrates into the diamond layer should be close to the desired equilibrium liquid composition due to the lengthy first stage of each experiment. The additional day or so of ~uilibration during the second stage should be sufhcient for any fine-tuning that may be required due to slight differences in the pressure and temperature between the two experimental stages. Phases in run #62T have not yet been analyzed; the main reason for conducting this experiment was to verify the presence of clinopyroxene in the run product (see below). Crystals and glass from run #55T (1270°C) have, however, been analyzed and the glass composition is plotted in Fig. 4. With the exception of SiOZ, the oxide concentrations overlap the melt composition from the

single-stage run #20 ( 1270°C) at the 1(r level. Calculated phase proportions in these two runs are also similar (Table 2; Figs. 9 and 10). The slightly lower SiOz value produced in the run #55T partial melt relative to the melts in runs #20 and #IS ( 1280°C) may suggest that, at the lowest melt fractions in the single-stage experiments, silica contents had not yet reached equilibrium values in both melt reservoirs (i.e.. the diamond and peridotite layers). We stress, however, that equilibrating partial melt and peridotite either in intimate contact or via connected melt channels produces very similar melt compositions; the differences that are present appear to be subtle. We are continuing these multistage experiments at temperatures near the ~~dotite solidus to evaluate more fully these subtle effects. We have also examined compositional heterogeneity of the residual peridotite mineralogy in our experiments. sin&! achievement of equilibrium would require homogeneous crystals and no dependence of mineral composition on distance from the diamond layer. At temperatures i 1300°C. run times ~72 h are required to produce olivines that are unzoned and homogeneous throughout the capsule. In contrast, unreacted orthopyroxene cores exist in all of our experiments. In the longest duration runs, the radius or longest dimension of these cores are generally less than half the combined core and rim dimension, and thus, the cores comprise at most a few tens of percent of the total orthopyroxene in these charges. Orthopyroxene rim compositions are nearly independent of location in the silicate layers at these long run times. At temperatures below cpx-out, all the experiments contain unreacted clinopyroxene cores, which in the longest duration runs are 520 Mm in size. Even in the 9X- I5 I h runs. clinopyroxene rim compositions show more relative variation in CaO and A&O3 contents than the orthopyroxene rims. although this compositional variation does not appear to correlate with position within the capsule. The clinopyroxene

491

6i 14

1

1

1

i’

L-,_.

j

l

a

1

iz

‘A-l__

/

/

2

ei

1ck-----

..^ ~~~~~~~-_~~._~, &_

l_.-II-.--L_“-.

‘7__.--__.._.-.

~+--$-.-~---i--A--

E+o

4,

Time (hrs)

(i,,

*ii,

Time jhrh)

FIG. 3. Mean concentrations of selected oxides in interstitial glass from the diamond layer from 1330°C experiments as a function of run duration (see Tables 2 and 3). Error bars represent AII o of’ the distribution of analyses.

Experimental study of partial melting in the mantle

L.8 1260 ,280

.I

,I

,300

.I.,

,320

1340

.I

1360 1380

.I 1400

Temperature(“C)

I. 1260

I.

1.

,280 ,300

8.

I.

,320 1340

2817

1.

I

1360

1380

.I 1400

Temperatme(YZ)

FIG. 4. Mean oxide concentrations in interstitial glass from the diamond layer vs. run temperature. The circles represent liquid compositions from the longest single-stage experiments at each temperature, horizontally pointing triangles represent liquid compositions from the two reversal runs, and the diamond symbol shows the liquid composition from the two-stage experiment (see Tables 2 and 3). For the reversals, the direction that the triangles point indicates whether the final run temperature (I 3 10°C) was approached from the high (left-pointing) or low temperature (rightpointing) side. Error bars for oxide concentrations are la; if not shown, error bars are smaller than the symbols. Uncertainties in temperature are all taken as +9”C. Here and in the following figures: closed circles indicate results from this study in which the residual mineralogy is olivine, orthopyroxene, clinopyroxene, and spinel; open circles indicate experiments in which clinopyroxene was exhausted from the crystalline residue, i.e., the residual mineralogy is olivine, orthopyroxene, and spinel.

rims are irregularly shaped in two dimensions, and their volume fraction may be up to 50% of the total clinopyroxene present. Although we have not quantified the difference, the ratio of clinopyroxene rim to unreacted core was greater in the two-stage run (#55T; 127O”Q than in the single-stage run (#20; 1270°C) based on backscattered electron images. Individual spine1 grains are generally smaller than 15 Wm. Core and rim compositions determined on some of the rare lo- 15 pm grains show that Mg# decreases and Cr# increases outward from the core (Mg# = lOOMg/(Mg + Fe*) and Cr# = lOOCr/(Cr + Al), where abundances are mole fractions). Single microprobe analyses were collected from the interiors of the 5- 10 pm spine1 grains. In all the experiments, interiors of spine1 grains near the diamond layer have higher Cr# and lower Mg# values than those at the bottom of the charge, although the magnitude of the variation decreases with increasing run time. For example, at 1330°C and -42 h, Cr# values range from 44.8(4) (top) to 35.1(6) (bottom), while after 72 h the variation is 49.8(4)-44.0(1.0) (the values in parentheses here and throughout the paper are lg in terms of the least units cited). In the two-stage run (#55T), spine1 composition is nearly independent of distance from the diamond layer, and is similar to the composition of the spinels nearest the diamond layer in run #20 (Table 3). The greater reequilibration of pyroxene and spine1 in the two-stage ex-

periment may reflect both the longer run time and the larger amount of melt present in the peridotite layer. However, despite the variation in spine1 composition with position and time in the single-stage experiments, spine1 constitutes < 1.5 wt% of the residue in these capsules, and we expect that these heterogeneities do not significantly affect melt composition. Based on our results to-date, run durations of at least 4872 h are required to achieve a nearly time-independent liquid composition, to minimize gradients in residual spine1 composition with distance from the diamond layer, and to reduce compositional gradients in individual residual silicate crystals to acceptable levels. Only results from experiments of this duration or longer are included in the Discussion. In contrast to our results, HIROSE and KUSHIRO (1993) report completely homogeneous residual crystals at 1250- 1300°C after only 48 h. The reason for the convergence in their experiments to near-equilibrium results on more rapid timescales than in our experiments is unknown, but could reflect some difference in the starting materials. Although JOHNSON and KUSHIRO (1992) noted that 24 h was insufficient to produce a timeinvariant liquid composition at 1425°C and 15 kbar, they nevertheless believed that the quenched melt was sufficiently close to equilibrium to provide information on peridotite melting. Based on our results, we think that the crystalline residue does not closely approach equilibrium on this time-

2818 scale at lower, near-solidus temperatures, longer run times if possible.

M. B. Baker and E. M. Stolper and we recommend

tern; i.e., clinopyroxene may be stable to higher temperatures in ultramafic systems that contain CrzOX compared to compositions that are Cr-free (T. FAI.L~ON, pers. commun.).

Phase Relations Phase Compositions Olivine, orthopyroxene, and spine] coexist with melt over the entire investigated temperature interval (1270- 1390°C). while clinopyroxene disappears from the residual peridotite layer between 1330 and 1350°C. Our temperature bracket for cpx-out (T,,,.,,,) is consistent with the results of TAKAHASHI (1986) and TAKAHASHI et al. (1993) on KLB-I, a bulk composition very similar to MM3. The cpx-out curve drawn by TAKAHASHI (1986) for KLB-1 is between 1340 and 1350°C at 10 kbar, within our experimental bracket of 1330- 1350°C. The difference between the 10 kbar spine]-out temperature for KLB-I (1325-1350°C; TAKAHASHI, 1986) and MM3 (> 1390°C; this study) probably reflects MM3’s higher Cr content. At 10 kbar, experimentally determined temperature brackets for T,,,.,,, for naturally-occurring and model mantle peridotites range from 1250- 1300°C for Hawaiian pyrolite minus 40% olivine (HPy-40; JAQUES and GREEN, 1980) to 1330- 1350°C for MM3 and KLB-1 (this study; TAKAHASHI, 1986; TAKAHASHI et al., 1993; see HESS, 1992, for a compilation). The nominally anhydrous cpx-out temperatures reported in the literature and in this study are positively correlated with peridotite Mg# and negatively correlated with total alkali content. These observations are qualitatively consistent with inferences drawn from simple systems. For example, in the system MgSi03-FeSi03-CaSi03, decreasing the Mg# of a bulk composition at constant temperature can move it from inside the oliv + cpx + liq three-phase triangle into the oliv + liq two-phase region (YODER et al., 1964). Likewise. increasing the amount of jadeite (i.e., increasing the Na,O content) lowers the liquidus temperature for clinopyroxene on the diopside-jadeite join at 30 kbar (BELL and DAVIS, 1965). The univariant melting reaction involving forsterite, enstatite, diopside, and spine1 in the CaO-MgO-A120z-Si02 system (CMAS) provides a further check of the consistency of our temperature bracket for cpx-out. For MM3, r,,,.,,, is comparable to the range of experimentally determined temperatures for the melting reaction enstatite + diopside + spine] = forsterite + liq at 10 kbar (- 1350°C. KUSHIRO, 1972: 1330 + 10°C Fig. 2, PRESNALL, 1976; 1319 ? 10°C Fig. 3, PRESNALL et al., 1979). For a diopside-poor synthetic mantle composition in CMAS, the univariant melting curve should be nearly coincident with the diopside-out curve. Based on the discussion above, we would have expected 7,,,.,,, to be higher in CMAS than in the natural system because of the effect of Fe and Na, and yet this does not seem to be the case. For this reason we did a two-stage melting experiment at 1330°C (run #62T; total time = 122.5 h) to confirm that clinopyroxene is present in the MM3 residue at this temperature (Table 2). The abundance ofclinopyroxene is quite low in both of the 1330°C runs (#24, calculated mode and #62T, visual estimate), suggesting that r,,,.,,, is only slightly higher than 1330°C. The unexpectedly small difference between r,,,.,,, for MM3 and for synthetic mantle compositions in CMAS may reflect the influence of Crz03 in the natural sys-

Figure 4 shows how oxide concentrations In the experimental glasses from the longest duration runs vary as a function of temperature. Compositions vary from roughly basaltic (- 10.4 wt% MgO) to picritic (- 16 wt% MgO) with increasing temperature (Fig. 4f). Over the interval 1270- 1300°C. silica decreases from 50.8 to 50.0 wt%, then increases to --SO.9 wt% as temperature increases to 1390°C (Fig. 4a). Note that based on our two-stage run, (#55T) the drop in silica between I270 and 1300°C may be less pronounced. A&O3 decreases nearly linearly from - 18 wt% (1270°C)to - I2 wt% (139O”C), while FeO* increases from -5.9 to 7.4 wt4: over this temperature interval (Fig. 4c,e). Most of the Fe in the partial melts is probably present as FeO, since the graphite capsule forces f0, to be at or below the GCO buffer, and since our peridotite starting material has a molar Fei ‘/Fe’ of -0.08. The trend of CaO with temperature (Fig. 4g) is concave-downward with a maximum at - 135O”C, coincident with the experimental brackets on the cpx-out temperature. The abundances of the minor oxides TiOz and NazO (Fig. 4b,h) drop rapidly over the interval 1270-1300°C and then fall more slowly with increasing temperature, while CI abundance increases monotonically between 1270 and 1390°C (Fig. 4d). Mg# in the melt increases with increasing temperature (Fig. 5) from 75.7(7) at 1270°C (#20) to 79.2(6) at 1390°C (#26). These high Mg#s in the liquids reflect the relatively low Fe/Mg ratio of our bulk peridotite mix. Using best-lit melt fractions (F) in of our long duration single-stage experiments and in the two-stage run at 1270°C (Table 2 and discussed below), we have calculated bulk partition coefficients (6) for selected oxides in the liquids using the batch melting equation, C, = C&F + (1 - F@) (SFIAR. 1970), where CO is the bulk concentration (Table I) and C~) is the liquid composition (Table 3). Between 1270 and 139O”C, & decreases sharply, from between 7(2) to i.3(2) ( 1o uncertainties calculated from errors on t;, C,, and C’,; all D values for #20 and #55T overlap at the In level). Bulk partition coefficients for A1203 and CaO also decrease

1260 1280 1300 1320 1340 1360 1380 1400 Temperature (“0 FIG. 5. Variation in mean molar Mg/(Mg + Fe*) in interstitial glass from the diamond layers of the longest duration single-stage experiments (circles), the reversal runs (horizontally pointing triangles). and the two-stage experiment (diamond). Uncertainties on the ratios are f I o and are calculated by propagating uncertainties in the MgO and Fe0 values. Uncertainties in temperature are all taken as +9T.

2819

Experimental study of partial melting in the mantle smoothly with increasing temperature, ranging from 0.16(1)0.08(2) and 0.26( l)-0.05(2), respectively. These variations in DA, and dca reflect the decreasing abundances of clinopyroxene and spine1 in the residues. The dca trend also appears to flatten with the disappearance of clinopyroxene. Calculated Ti and Na bulk partition coefficients are approximately independent of F, but show more scatter. The ranges for &, and &, are 0.1-0.05 and 0.05-0.01, respectively. Figure 6 compares our experimental liquid compositions with the results ofthe 10 kbar peridotite melting experiments of JAQUES and GREEN (1980) and HIROSE and KUSHIRO (1993). The latter study also used the diamond melt extraction technique. Although the absolute oxide concentrations in the liquids differ, reflecting differences between the bulk compositions (HPy f HK-66 # KLB-1 = MM3; see Table l), the temperature-liquid compositional trends are quite similar. For all four bulk compositions, the partial melts show a negative correlation between A&O3 content and temperature. The correlation between FeO* content and temperature is positive in all cases. CaO contents of the partial melts define concave downward trends when plotted against temperature for all the bulk compositions. The maximum CaO contents generally occur within the brackets on TcpX_Out (Fig. 6d). Below - 1300°C. all the trends differ by less than - I wt% CaO. Above 13OO”C, the HPy, HK-66, and KLB-I trends all drop rapidly to CaO contents that are -2 wt% lower than those of the MM3 trend at the same temperature. This divergence could reflect real differences in the melting behavior of HPy, HK-66, and MM3. However, KLB-1 and MM3 are compositionally similar, and produce similar amounts of melt at the same temperature (Fig. 9b). Given the Ca-poor nature of the residuum at temperatures above Tcpx_out, the slightly lower bulk CaO content of KLB-I compared to MM3 (Table I)

Temperature

(“C)

can only account for -0.5 wt% of the difference in liquid CaO content at these temperatures. The remaining 1.5 wt% discrepancy at > 1320°C is surprising. given the similar bulk compositions and experimental techniques, and we currently have no explanation for it. Both our single-stage data and the 10 kbar partial melts reported by HIROSE and KUSHIRO (1993) display a minimum in SiOz at - I3OO”C, while the silica contents in the calculated partial melts of JAQUES and GREEN ( 1980) increase monotonically with temperature (Fig. 6a). If we include the 1270°C two-stage melt, our SiOz trend is much closer to that generated by JAQUES and GREEN ( 1980). Further work is needed to clarify the slope of the temperature-SiOz trend at low melt fractions. The solid phase compositions also vary systematically with increasing temperature. The residual olivine in our experiments becomes more magnesian; Mg# values are 90.7(3) at 1270°C and 92.1(2) at 1390°C. The mean &‘$$& for our dataset is 0.33( 1). This value overlaps the 10 kbar average exchange coefficient calculated for 48 olivine-liquid pairs from the literature (0.32(I); TAKAHASHI and KUSHIRO, 1983; F~JJII and SCARFE, 1985: FALLOON and GREEN. 1987; FALLOON et al., 1988; ULMER. 1989). Mean Fe*-Mg &values for opx/ liq and cpx/liq from our experiments based on analyses of the rims are 0.3 l(I) and 0.32(l), respectively, and are consistent with pyroxene/liquid exchange coefficients calculated from other 10 kbar data (e.g., TAKAHASHI and KLJSHIRO, 1983: FALLOON and GREEN. 1987). The Wo-content of orthopyroxene rims coexisting with clinopyroxene increases only slightly from 1270- 133O”C, 4. I(5) in #55T and 3( 1) in #20 to 4.7(4) in #24, while the Wo-content in the chnopyroxene rims decreases from 36.8(6) in #55T and 38.0(7) in #20 to 32.6(8) in #24 (Fig. 7). At temperatures above cpxout. the Wo-content of the orthopyroxene begins to drop.

Temperature

(“C)

FIG. 6. Concentrations of selected oxides in partial melts of different peridotite compositions vs. temperature at IO kbar from our data (circles = single-stage, two-stage, and reversal runs) and from the literature. Squares = HPy (JAQUES and GREEN, 1980); triangles = KLB-I (HIROSE and KUSHIRO, 1993); diamonds = HK-66 (HIROSE and KUSHIRO. 1993). The symbols represent either analyzed glass compositions (this study: HIROSEand KUSHIRO, 1993) or calculated melt compositions (JAQUES and GREEN, 1980). The solid, dashed, and dot-dashed lines are either cubic-spline fits to the literature data (a, c, and d), or straight-line least-squares fits (b). The filled symbols denote an oliv + opx + cpx k sp residue: open symbols denote an oliv + opx + sp residue. The lowest temperature HK-66 and KLB-1 experiments contain plagioclase in addition to the olivine. orthopyroxene, clinopyroxene, and spinel. For data from this work, oxide error bars are + 1m. Uncertainties in temperature are all taken as *9”C.

M. B. Baker and E. M. Stolper

2820

g

1350 -

B

+

+ ~1300- + G 4 5

1250 0.0

‘k1 0. I

* . O.?

.

’

I 0.3

04

Ca/(Ca+Mg+Fe*) FIG. 7. Mean molar &aO/(CaO + MgO + FeO*) in pyroxene rims from the longest runs versus run tem~rature: circles = single-stage experiments, diamond symbol = two-stage experiment. Compositional uncertainties are +I6 and are based on propagating uncertainties in the three oxide concentrations; in some cases these errors are smaller than the symbols. Uncertainties in temperature are all taken as 5~9°C. Filled and open symbols are pyroxene compositions in a residual mineralogy of oliv + opx + cpx + sp and oliv t opx + sp, respectively.

Weight-based

pyroxene/liquid

partition coefficients calculated

for Ti02, A1203, and NazO appear to be nearly independent of temperature or liquid composition. Clinopyroxene values are 0.28-0.33, 0.29-0.34, and 0.17-0.20, respectively, while m/lx/fiq, Il)od;x/Jiq, and p&tiriq are 0.10-0.19, 0.21-0.28, and 0.03-0.05. Chromium pa~itioning between pyroxene and iiquid in our experiments shows a strong inverse correlation with temperature; for ciinopyroxene, the value decreases from I 1 (127O’C) to 5 (133O”C), while for orthopyroxene the value decreases from 7 (1270°C) to 2 (I 390°C). The negative temperature dependence of Dvfliq can also be seen in the onebar data of BARNES (1986) and in the 10 kbar data of STOLPER (1980), TAKAHASHIand KUSHIRO(1983), and FALL~ONand GREEN ( 1987). Bulk partition coefficients for TiO?, Al103, CrZ03, and NazO in each ofthe experiments calculated using the mineral/melt coefficients and the residual phase proportions (see Table 2) overlap with the D values calculated using the batch melting equation (see above). Figure 8 shows that the composition of spine1 in our experiments varies dramatically as a function of run temperature. The plotted points are mean compositions of grains close to the diamond layer in each capsule, and the uncertainties reflect Ijrain-to-grain variability. Cr# increases from 28.0(8) to .58.2(5) between 1270 and 1390°C, while Mg# decreases from 79.3(3) to 76. I(3) over the same temperature interval. This drop in Mg#, even as Mg/Fe* increases in the liquid and in the other solid phases, reflects the effect of (Cr/ Al)“” on the partitioning of Fe*+-Mg between spine1 and all other phases (e.g., IRVINE, 1965, EVANS and FROST. 1975; ENGI, 1983). Relationships Between Temperature and Phase Proportions The proportions of glass and crystalline phases in each long duration run (252 h) were calculated by fitting the bulk composition of the starting material to a linear combination of the analyzed phases. The algorithm incorporates errors in the phase and bulk compositions (ALBARBDE and PROVOST, 1977). The phase proportions reported in Table 2 for each run represent averages of multiple fits using different com-

binations of pyroxene core and rim compositions and spine{ compositions from either the top or bottom of the capsules. Best-fit melt fractions (F, by weight) for the single-stage experiments vary from 0.074( 10) at 1270°C to 0.27(l) at 1390°C. Note, that the calculated liquid fractions in both 1270°C runs (#20 and #55T) overlap at the I D level. For the partial melts coexisting with spine! Iherzolite, the data define a concave-downward trend in Fig. 9a. The curved 7”(‘C)-F trend indicates that melt production decreases with increasing temperature. A polynomial fit to the results of the singlestage experiments yields values of 0.2 1 wt%/“C at 1270°C to 0.11 wt%/*C at i 330°C. Extrapolating this quadratic fit to 0% melt suggests a 10 kbar solidus temperature of - 1240 t lO”C, similar to the experimental solidus bracket for KJBI (1250-127YC; TAKAHASHI, 1986; - 126O”C, Fig. 7, TAKAHASHI et al., 1993). Melt fractions calculated for the three high temperature experiments with only oliv + opx -t sp in the residue are consistent with a cusp in the ‘Z’(“C)-F’trend at cpx-out. In high-pressure melting experiments on both simple systems and natural peridotites. KUSHIRO (1969), MYSEN and KUSHIRO (1977) WENDLAND-~ and MYSEX ( 1980), and HARRISON ( 198 1) also observed discontinuities in T(“C)-Ftrends that were coincident with the disappearance of a residual phase. Figure 9b compares the fit to our data with calculated and ex~rimentally determined T( “C)-Fsystematics at - 10 kbar from the literature. Table 1 lists the bulk peridotite compositions associated with these literature trends. With the exception of the curves from MCKENZIE and BICKLE (1988) calculated from their third-order polynomial, and the cubicspline fit to the calculated points from KINZLER and GROVE (1993), the remaining trends from the literature shown in Fig. 9b are linear fits to the experimental data. Between 1260 and 129O’C (where the bulk of the literature trends apply to melts coexisting with spine1 lherzoiite assemblages), most of the trends define a relativefy narrow band in temperaturemelt fraction space (Fig. 9b). This overlap reflects the fact that similar bulk compositions produce similar melt fractions at the same temperature. The HK-66 line plots above the other data because this bulk composition is more fertile (i.e., has much hif&er Fe/Mg and alkali abundances and a lower solidus) relative to the other peridotites (see Table 1). With

0.74

0.7h

0 78

(1 x0

II U!

Mg/(Mg+Fe*) FIG. 8. Mean composition (molar Cr/(Cr + Al) -. Mg/(Mg -+Fe*)) of spinets closest to the diamond layer in each capsule; error bars arc +-lo and include both inter- and intra-grain variability. Open vs. filled symbols refer to residual mineralogy as described in Fig 3. Circles represent spinet compositions from single-stage experiments: the diamond symbol represents the spine1 composition from the twostage run.

Experimental study of partial melting in the mantle

2821

earlier estimates from high-pressure melting ex~~ments (e.g., KUSHIRO, 1969; PRESNALLet al., 1978; PRESNALLet al., 1979; KINZLERand GROVE, 1992af and with conventions wisdom, i.e., that normative clinopyroxene is a major constituent of basaltic liquids. In the following section we explore the stoichiometry of high-pressure peridotite melting in more detail. i 240

1280

1320

1360 I

‘i225

1275 I325 Temperature

1375 (“C)

1400

’

,

1425

FIG. 9. (a) Weight fraction of melt vs. run temperature from this study: circles = single-stage experiments. diamond symbol = twostage experiment. Phase proportions of melt and residual crystals in each experiment were determined by a feast squares mass balance procedure in which the analyzed com~sit~ons of the phases in the run product were required to sum to the MM3 bulk composition. ~Jnce~ainties on the melt fractions were estimated from multiple mass balance calculations using pyroxenc core and rim compositions, and the range of spine1 compositions observed in each charge. Uncertainties on temperature were ail taken to be rtP”C. Closed and open symbols indicate the presence or absence of clinopyroxene in the residual mineralogy. The melt fractions from the single-stage, spine1 Iherzolite-bearing experiments were fit with a quadratic [F = - 17.3 + O.OZS(Q-9.0 X lO-6(p)] shown as a solid curve over the tem~rat~re range of the ex~riments and as a dashed curve extrapolated to the solidus. ~nce~ainti~s on the coefficients are 6.3,O.OtO, and 3.7 x IO-*, respectively. (b) Comparison of our T-Ftrend with both IO kbar experimental (JAQUES and GREEN,1980; HIROSEand KUSWRO,1993) and theoretical (MCKENZIEand BICKLE, 1988: KINZLERand GROVE, 1993) results from the literature. Symbols are described in the caption to Fig. 6. Abbreviations are: J&G, HPy = JAQUESand GREEN(1980); M&B = MCKENZIE and BICKLE (1988): K&G = KINZLERand GROVE(1993); H&K, HK-66 = HiROSEand K
the exception of the calculation of KINZLER and GROVE (I 993), which produces a concave-upward trend and an increasingly higher rate of melt production at temperatures above - 1275”C, the sources compiled in Fig. 9b give rates of melt production that are comparable to our results. Figure 10 shows residual phase propo~jons in our experiments (determined from the mass balance fits; Table 2) plotted against temperature. The modal abundance of olivine is approximately constant, while the proportion of orthopyroxene decreases slightly with increasing temperature (Fig. lOa). Clinopyroxene and spine1 modal abundances both decrease with increasing temperature and, in the case of clinopyroxene, the variation is substantial (Fig. lob). These trends in the modal data indicate that clinopyrox~ne comprises the majority (>50%) ofthe mantle melting assemblage at IO kbar, a conclusion that is consistent with the results of

Ten Mar Melting Reactions Although the ~oichiomet~es of peridotite melting reactions exert fundamental controls on the compositions of liquids produced by mantle melting, a wide range of reaction coedficients for each phase exists in the literature (see summary in KosToPouLOs, 1991). For example, trace element geochemists have assumed melting proportions of clinopyroxene that vary from 0.40-0.73 (KOSTOPOULOS, 199 1). There is also no consensus on whether olivine is consumed or generated during melting (i.e., whether the melting reaction is “even” or “odd” in the terminology of MORSE, 1980). Based on Iiquidus relations in binary and ternary joins in the system Ca0-FeO-MgO-Alz03-SiOz 1- Cr.& KOSTOPOULOS (1991) concluded that fertile spine1 peridotite initially melts via the reaction oliv + opx + cpx t sp -+ liq (i.e., that the reaction is “even”). In contrast, using 1O- 12 kbar multiply saturated natural melts and both graphical and numerical analysis, KINZLEK and GROVE (1992a) concluded that the melting reaction is opx + cpx -t- sp + oliv + liq. They also noted that this “odd” reaction is consistent with melt production in the simple system CaO-MgO-A1203-Si02 at 9-25 kbar ( KUSHIRO, 1965; PRESNALL,1976: PRESNALLet al., 1979).

1260 1280 1300 13201340 136013RO14(X) Temperature (“C) FIG. 10. Variations in the proportions of residual crystalline phases as a function of run temperature. Best fit determination of phase proportions and errors as described in Fig. 9 and Table 2. (a) olivine and orthopyroxene: (b) clinopyroxene and spinel. Filled and open symbols indicate the presence or absence of ~linopyroxene in the residual mineralogy, while circles represent the longest duration singlestage experiments and the diamond symbol represents the two-stage experiment.

2822

M. B. Baker and E. M. Stolper

Figure 11 shows the liquid and pyroxene compositions from our four lowest temperature single-stage experiments projected from Sp ((Mg,Fe)(AI,Cr),O,) onto the Oliv-CpxQtz pseudoternary (coexisting olivines plot at the Qliv apex). Also plotted are MM3 and the field defined by primitive MORB glasses (PRESNALLand HOOVER. 1987). Figure I 1 illustrates that the tie lines between olivine and liquid pairs pierce the planes defined by coexisting clinopyroxene. orthopyroxene, and spine1 (note that the spinels in the experiments plot at the Sp apex of the Oliv-Cpx-Qtz-Sp pseudoquaternary). This chemographic relationship indicates that in our experiments the spine1 lherzolite melting reaction is opx + cpx + sp -+ oliv + liq. Note that the four 10 kbar experimental liquids ptot within or on the border of the primitive MORB held in pseudoternary composition space. While this overlap is suggestive that some primitive MORB glasses may be 10 kbar primary mantle melts, we will show below that the FeO*-MgO systematics of IO kbar mantle melts are substantially different from primitive MORBs, and thus that the overlap is an artifact of the projection scheme. Our data on mineral and melt compositions and proportions near the peridotite solidus are also ideal for providing quantitative constraints on the melting reaction. Figure 12

CPX

Ffc;. 12.Best-fit weight fractions of residuaf crystalline phases vs melt fraction in experiments from the singie-stage experiments (this study). For each experiment. the sum of the weight Fractions of the residual solids and the melt fraction equals I.O{see Table 2). The

solid phase uncertainties are generally smaller than the size of the symbols. The solid lines are weighted least-squares fits to the phase proportions in each experiment and the phase proportions in thr starting material (i.e.. at F = 0). The weighting forced the lines to go through the F = 0 points. Because the starting pyroxene and spine1 compositions will vary with P and r, the actual high-pressure subsolidus mode of MM3 may be slightly different from the proportions of phases mixed to produce MM3. Thus, in detail. the trends for each phase may not actually be linear at very low melt fractions. The equations of the fits are oliv = 0.508 + 0.22(6)f:; opx 7. 0.299 0.38(5)p, cpx = 0.167 - O.?l(5)I+‘;sp = 0.026 ().13(I)I-: The phase proportions in run #I5 (F ==0.092) were not included in the fmat fits because the olivine and o~hopyroxene proportions deviate from the simple linear trends defined by the starting material and the other runs.

shows proportions

Oliv

Qtz

FIG;.1 I, Projection of our four IO khar liquid and coexisting orthopyroxene and clinopyroxene compositions from Sp onto the OlivCpx-Qtz ternary; coexisting olivine compositions project onto the Oliv apex, and have not been plotted. Also shown is the field defined by the projected compositions of primitive (Mg# 2 67) MORB glasses (PRESNALLand HOOVER, 1987). The star = MM3, the bulk composition of our starting material: the dashed lines illustrate that the oliv-liq tie-line pierces the opx-cpx-sp plane indicating a melting reaction of the type opx + cpx + sp ==oliv + liq. The equations for transforming oxide weight percent compositions into the oxygennormalized mineral components Oliv. Cpx, Qtz. and Sp are listed below: Ofiv = 2 ([MgO + FeO*f - [A&O,] ~ [Cr,O,] --. (CaO] i [NaZO]). Cpx = 6 [CaO], Jadeite = 12 INa,Ol. Qtz = 2 ([SK&] < 0:5 ([Al,031 + [Cr20j] - [MgO] -. [FeO*]) .-- I.5 [CaO] - 4.5 [Na20]). Sp = 4 ([Alz03] - [NarO]). and Chromite = 4 [CrzO,], where bracketed oxides represent mole fractions. After normalizing all of the components to 1.0, Sp and Chromite are combined and the resulting Sp. Oliv, Cpx, and Qtz values are renormalized to 1.O. Projecting from Sp onto the Oliv-Cpx-Qtz plane simply involves normalizing the latter three values to I.O.

of residual olivine, orthopyroxene, clinopyroxene, and spine], calculated using the results of the single-stage experiments (Table 2) plotted as a function of F. 4t any value of F, the slope of the trend for each phase indicates how the proportion of that phase changes in the residue with increasing melt fraction. Multiplying the slopes by ~ I gives the melting reaction coefficients. The proportions of crystals in MM3 and the modes from the experiments (all in weight fractions) plotted in Fig. 12 were fit with straight lines using weighted least-squares and the resulting melting equation is (X38(5) opx + 0.71(5) cpx + 0.13(l) sp = 0.22(h) oliv i I Jig. The large value of the clinopyroxene coefficient confirms the results of the earlier studies cited above that indicate that clinopyroxene dominates the melting assemblage of dry. fertile peridotite at 10 kbar. Olivine occurs on the right-hand side of the equation, indicating that the melting reaction is “odd” under these conditions, and that otivine is the phase in reaction relation with the liquid. This result contrasts with that of KOSTOP~ULOS(1991), but agrees with the conclusions of KUSHIR~ (196.5) PRESNALI. (19761, PRESN’ALIet al. ( 1979), and KINZLER and GROVE (1992a) that at IO kbar. spine1 lherzolites melt by the reaction, opx + cpx -+-sp -+ oliv + liq. Our calculated coefficients are similar to those derived by KINZLER and GROVE (1992a); their values for orthopyroxene, clinopyroxene, spinel, and olivine are: 0.40(15), 0.82(15), 0.08(l), and -0.30(10). Although the reaction coefficients can be treated as constant over the spine1 lhenolite melting interval, the values for olivine, orthopyroxene, and spine1 must change dramatically

2823

Experimental study of partial melting in the mantle with the disappearance of clinopyroxene. The trend of the oliv-opx-sp-saturated liquids in pseudoternary composition space sugest that at temperatures above cpx-out, the melting reaction remains “odd”, i.e., opx + sp --+ oliv + liq. Linear, weighted least-squares fits to the modes of the three oliv + opx f sp bearing experiments yield the following melting reaction:

0 n X 0 +

Primitive MORB glasses: glass Mg#>67 Partial melts, this study: peridotite Mg#-90.5 Partial melts, H&K: peridotite Mg#-85.3 Partial melts, H&K: peridotite Mg#-89.6 Melts, “sandwich” experiments: peridotite+ basalt MgK88.1

MO& + oliv

1.06(36) opx + 0.04(3) sp = 0.1(3) oliv + 1 liq. The calculated coeacients are consistent with our inferences from chemographic relationships, but the uncertainties on the mineral coefficients are all large. These uncertainties stem from the limited spread in F-values, and in the case of olivine, from the small change in residual olivine abundance among the three experiments. Given the error associated with the olivine value, the exact nature of the melting reaction is still in question. Future higher temperature runs should help clarify the nature of the 10 kbar spine1 harzburgite melting reaction.

Comparison of ~x~rime~tal Primitive MORB Glasses

Liquid Com~sitions

and

Historically, the debate concerning the petrogenesis of MORBs centered around the composition of parental MORB liquids and the depth and extent of melting that produced these liquids. The two extreme positions were that parental melts were either (1) picritic liquids produced at pressures between 20 and 30 kbar by high extents of melting (20-30%) or (2) tholeiitic liquids (- 10% MgO) representing moderate degrees of melting (IO-20%) at - 10 kbar. This debate has been reviewed by BVSP ( 198 I), THOMPSON(I987), and ELTHON (1989), among others. While mantle melting models have evolved to include dynamic processes such as pooling low-degree melts produced over a wide range of pressures by adiabatic decompression (e.g., KLEIN and LANGMUIR, 1987: MCKENZIE and BICKLE, 1988), vestiges of simple batch melting models continue to appear in the literature (e.g., HIROSE and KUSHIRO, 1993). Figure 13 compares the FeO* and MgO contents of a suite of primitive MORB glasses (PRESNALLand HOOVER, 1987), with the compositions of oliv + opx + cpx + sp-saturated melts produced at 10 kbar in peridotite-diamond experiments (this study; HIROSE and KUSHIRO, 1993) and in peridotite-basalt “sandwich” experiments (TAKAHASHIand KUSHIRO, 1983; FUJII and SCARFE, 1985; FALLOON and GREEN, 1987; FALLOONet al., 1988). In order to illustrate the effects of low-pressure olivine fractionation, we have also calculated how the compositions of MORB liquids would change if they coexisted with Fog0 olivine (the dashed field labeled MORBs + oliv in Fig. 13). Some of the experimental glasses have compositions that are very similar to those displayed by the primitive MORB glasses, and these similarities have been used to support claims that at least some MORB glasses are near-primary melts produced at - 10 kbar (TAKAHASHIand KUSHIRO, 1983; FUJII and SCARFE,1985; FUJII, 1989; HIROSEand KUSHIRO, 1993). When the bulk Mg# of the experimental charge is limited to 289 (i.e., to values that are reasonable for the upper mantle), 10 kbar melt compositions (solid symbols in Fig. 13) do not

8 n

I

I

6

7 wt%

1

I

8

9

10

FeO*

FIG.13. Weight percent of FeO* and MgO in primitive MORB glasses (Mg# B 67: PRESNALL and HOOVER,I987) and in 10 kbar melts saturated with oliv + opx + cpx t sp S plagiuclase from both

~~dotite~iamond (this study: H&K = HIROSEand KLJSHIRO. 19Y31 and peridotite-basalt sandwich experiments (TAKAHASHIand Ku: SHIRO, 1983; FUJI! and SCARFE, 1985: FALLOONand GREEN, 1987; FALLOONet al., 1988). The Mg#s in the legend associated with the experimental liquids refer to the bulk Mg#s of the peridotite and peridotite + basalt used in the experiments. The arrows show how the MORB glass compositions move in FeO*-MgO space as a result of adding equilibrium olivine in 0.5 wt’% increments: the MORBs + oliv field encloses the calculated liquid compositions that would coexist with Fog0 olivine. MgO uncertainties are smaller than. while the errors in FeO* are slightly larger than. the size of the symbols.

overlap those of primitive MORBs. At the same MgO content, the experimental liquids from these bulk compositions all have much lower FeO* values than the primitive MORB glasses. Correcting the MORB glass compositions for the effects oflow-pressure olivine fractionation does not alter these conclusions, since none of the filled circles or squares plot within the MORBs + oliv field. Consequently, as has been stated in a number of recent papers (KLEIN and LANGMUIR, 1987; MCKENZIE and BICKLE, 1988; NIU and BATIZA, I99 1; KINZLER and

GROVE, 1992a,b; LANGMUIRet al., 1992). and in contrast to the conclusion of HIROSEand K~EHIRO(1993), Fig. 13 indicates that primitive MORB glasses cannot be primary or near primary melts generated from primitive mantle at pressures of 10 kbar or less. KL.EINand LANGMUIR(1987) noted that the FeO* content of peridotite partial melts is positively correlated with the pressure of melting. Thus, if MORB magmas are formed by melting mantle with an Mg# 2 89, the parental magmas must involve at least a component of FeO-rich melt generated at pressures significantly higher than 10 kbar in order to account for the magma compositions shown in Fig. 13. The inference that MORBs contain a component of melt from 20-30 kbar, i.e., from within the garnet lherzolite field, is consistent with the presence of a garnet-signature in the LuHf isotopic systematics and in the (230Th/238U)activity ratios in MORBs (SALTERS and HART, 1989: BEATTIE. 1993). The importance to IMORB petrogenesis of mixing melts from a

2824

M. B. Baker and E. M. Stolper

wide range of pressures is explicit in the polybaric fractional melting calculations of KLEIN and LANGMUIR (19871, MCKENZIE and BICKLE (1988), NIU and BATIZA (1991). KINZLER and GROVE (1992a,b, 1993), and LANC;MUIRet al. ( 1992). Note that higher pressure melts are even more olivinenormative than those generated in our experiments at 10 kbar, and thus, even the most primitive known MORB magmas must have evolved by significant olivine fractionation from more primitive liquids. Comparison of Ex~~~ntaI and Calculated Trends in Liquid Composition During Isobaric Melting of Peridotite As discussed briefly above, recent models of basalt petrogenesis involve pooling small degree partial melts produced by adiabatic melting of ascending mantle (KLEIN and LANGMUIR, 1987; MCKENZIE and BICKLE, 1988; NIWand BATIZA. 199 1; KINZLERand GROVE, 1992a,b; LANC;MIJIRet al., 1992). A critical component in all of these models is a quantitative understanding of the relationship between liquid composition and melt fraction as a function of pressure and temperature. KLEIN and LANGMUIR(1987) made the first effort to systematize the high-pressure experimental data and to generalize how major elements in the liquid varied during isobaric batch melting of a generic spine1 lherzolite. Subsequent workers (MCKENZIE and BICKLE, 1988; Ntu and BATIZA, 1991: WATSON and MCKENZIE, 199 1; KINZLER and GROVE, 1992a,b, 1993; LANGMUIR et al., 1992) have also tried to quantify the relationship between melt chemistry and melt fraction using expanded datasets and parameterizations. However, not all of these parameterizations can be correct, since at the same pressure and for similar bulk peridotite compositions, they predict different trends in liquid composition.

To evaluate these paramete~zations, we have compared calculated 10 kbar ~uilib~um (i.e., batch) melts from several recent studies (MCKENZIE and BICKLE, 1988: NHJ and BAI‘IZA, 1991; WATSON and MCKENZIE, 1991; LANC;MUIR et al., 1992; KINZLERand GROVE, 1993) with our experimental liquid compositions (Fig. 14). Although the peridotite compositions used by MCKENZIE and BICKI.E ( 1988). WA’I%J? and MCKENZIE (I 99 I), and NIU and BA,I?Z~X ( 199 I ) are dif: ferent from MM3, they are broadly similar (Table 1) and all are meant to represent primitive oceanic upper mantle. The KINZLER and GROVE (1993) and LANGMUIKet al. ( 1992) compositional trends were calculated using their parameterizations and the MM3 bulk com~sition. Figure 14a shows that for 5-30s melting, the parameterizations all predict a decrease in .A1203with increasing melt fraction, but the slopes of the predicted trends vary considerably. Similarly, all but the MCKENZIE and BICKLE (1988) parameterization predict an increase in FeO* with increasing melt fraction, although once again. there are considerable variations in slope among the predicted trends (Fig. IJbI. The trend calculated using the LANC;MIJIR et al. (1992) parameterization is an excellent fit to our experimenta data on F vs. FeO*. Figure 14~ shows the relationship between melt fraction and calculated and measured CaO contents in It) kbar batch melts. As discussed earlier, our data define a concave downward curve with a crest that is approximately coincident with clinopyroxene disappearance. The trends calculated by NIU and BATIZA (199 I ) and KINZLSR and GROW ( 1993) are consistent with the experimental data; the latter trend does not turn over because KINZLXR and GRovt: ( 1992b. 1993) stopped their calculations at cpx-out. Howcvcr. the MCKENZIE and BICKLL (1988) and WATSON and MCKENZIE ( 199 1) trends. which decrease monotonically with increasing I;, do not agree with either our experimental data (Fig. 14~) or the data of JAQUES and GREEN ( i 9801 or HIROSI-

filkd=olwJpx+cpx+sp open=oliv+opx+sp

/.

9

‘.‘I,.&

0

0 I

_d 0 2

0 4

Melt Fraction FIG. 14. The composition of batch melts of peridotite at 10 kbar as a function of weight fraction of melt in our experiments (open and closed data points, corresponding to lhenolitic and harzburgitic residues) compared with calculated trends (shown as curves) at IO kbar based on parameterizations in the literature. (a) A1,03, (b) FeO*. (c) CaO, and (d) Na10 all in w-t%,plotted as a function of melt fraction ( I Guncertainties). Literature trends are labeled as follows: M&B = MCKENZIEand BICKLEf 1988);N&B = NIU and BATIZAf 1991); W&M = WATSONand MCKENZIE(1991 f; K&G = KINZLERand GROVE(1993); L et al. = LANOMUIRet al. (1992) with modifications from T. Plank (pers. commun.).

2825

Experimental study of partial melting in the mantle and KUSHIRO (1993; Fig. 6d); this has undoubtedly led to significant inaccuracies in their calculations of pooled melt compositions. The parameterization of LANGMUIR et al. ( 1992) does not include CaO. Finally, Fig. 14d illustrates that the trends in liquid NazO values calculated by NIU and BATEA (1991), KINZLER and GROVE (1993), and LANGMUIR et al. (1992) agree with our measured glass compositions; results of all three studies are consistent with a bulk Na partition coeffcient in the range 0.01-0.03. In contrast, MCKENZIE and BICKLE (1988) and WATSONand MCKENZIE (199 1) assumed &, values of 0.17 and 0.2, and thus, their parameterizations produce much shallower and unrealistic F-Na20 trends. With the exception of the parameterizations of MCKENZIE and BICKLE ( 1988) and to a lesser extent, WATSON and MCKENZIE (1991), the mantle melting calculations are all grossly successful in reproducing the isobaric batch melting

trends plotted in Fig. 14. Although all these melting models represent significant advances over earlier more qualitative efforts, some of the parameterizations are more successful than others at repr~u~ng the 10 kbar experimental data. The paramete~zation of LAN~MUIR et al. (1992) is the most successful at reproducing both our F-FeO* and F-Na20 data, but their model does not include SiO2, Al~0~, Cr203, and CaO. We expect that given the expanding experimental database, these models will continue to improve and provide more accurate understanding of the details of compositional trends in natural magmas, (e.g., the contrasting global and local trends of FeO* vs. NazO in primitive MORBs).

peridotite starting composition is - 1240°C. Our experimental data provide estimates of the proportions of minerals that enter the liquid during the melting of spine1 lherzolite. At temperatures below the cpx-out curve, our bulk composition melts at 10 kbar by the “odd” reaction: 0.38 opx + 0.7 1 cpx + 0.13 sp + 0.22 oliv + 1 liq. The compositions of our 10 kbar partial melts and those of HIROSE and KUSHIRO (1993) show that when the bulk Mg# of the experimental system is limited to ratios 2 89 (the generally accepted lower limit for primitive mantle), the liquids have FeO* contents that are substantially lower than those measured in primitive MORB glasses. This observation is yet further evidence that primitive MORBs are not simple low-pressure partial melts, but involve a component of higher pressure partial melting, and have evolved by significant olivine fractionation from more primitive liquids. Recent efforts to parameterize the experimental database on peridotite melting and to calculate melt compositions as a function of P, T, and F are partially successful in reproducing the compositional trends determined in this study, and thus represent significant improvements over previous more qualitative approaches to modeling the products of mantle melting. Nevertheless, we urge caution in using these parameterizations at this time to interpret the details of compositional variations among natural basaltic magmas, because most are only partially successful at reproducing the mantle melting trends presented in this study. However, as they continue to be refined, they should provide powerful tools for inferring the details of melt generation under upper mantle conditions.

CONCLUSIONS In this paper, we have described a new experimental technique for circumventing the quenching problems that have plagued high-pressure pe~dotite melting studies. A thin layer of -50 pm diamonds is placed above a layer of peridotite powder. Partial melt extracted from the peridotite layer collects in the pore spaces between the diamonds and equilibrates diffusively with the residual peridotite mineralogy. Isolated from the crystalline residue on the timescale of the quench, the melt quenches to a glass that accurately represents the composition of the liquid coexisting with the residual crystalline phases under the conditions of the experiment. We have used this technique to investigate partial melting of a fertile mantle composition at 10 kbar. Oxide concentrations in the equilibrium liquids vary systematically with increasing temperature; TiOz, A1203, and NazO contents decrease monotonically, while FeO*, MgO, and Cr*O, contents increase steadily. CaO content shows more complicated behavior, first increasing and then decreasing, with the crest in the temperature-CaO trend approximately coinciding with the disappearance of clinopyroxene from the residue. Overall variation in silica content with temperature is small, although there appears to be a minimum at about 12% melting. The results of reversal and two-stage experiments suggest that the liquid compositions that define these trends represent close approaches to equilibrium. Calculated melt fractions also vary systematically with temperature. The slope of the T(“C)-F curve is not constant, but decreases as temperature increases from 1270 to 1330°C. Extra~lating the curve back to zero melt suggests that the anhydrous solidus temperature for our

A~knoM,~~~~m~~~ts--This research was supported by National Science Foundation grant EAR89-16707. We thank D. Bell for supplying the Kilbourne Hole nodule, and J. R. Beck&t and S. Newman for assistance in preparing minerai separates. After learning of our work, Professor Kushiro graciously provided preprints of his group’s experimental results using similar techniques, and we thank him for hisgenerosity. T. L. Grove, R. 3. Kinzler, and M. Rutherford provided thorough reviews of the manuscript. Caltech Division of Geological and Planetary Sciences Contribution 5248. Edilorial

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