High-pressure experimental trace-element partitioning between clinopyroxene and basaltic melts

CHEMICAL GEOLOGY ISOTOPE (;q,.'OS('ll~.'\('lq

ELSEVIER

Chemical Geology 117 ( 1994 ) 127-147

High-pressure experimental trace-element partitioning between clinopyroxene and basaltic melts T h o m a s S k u l s k i " , W i l l i a m M i n a r i k ~, E. B r u c e W a t s o n Rensselaer Polytechnic Institute, Department of Earth and Environmental Sciences, Troy N Y 12180-3590, USA

Received 3 December 1993; revision accepted 28 March 1994

Abstract Clinopyroxenes play an important role in determining the distribution of trace elements in magmatic systems. In order to evaluate the extent and source of variation of clinopyroxene-silicate melt partition coefficients (D), experiments were conducted on natural basalts (picrite, alkaline basalt, nephelinite, high-alumina basalt: 12351300°C: 1-2.8 GPa) doped with a multi-element spike (Ti, Cr, V, Sr, Y, Zr, Nb, Ta, Hf, La, Ce, Nd, Sm, Dy and Yb) and water (2%). Trace-metal concentrations were determined by synchrotron XRF (SXRF) microprobe and a subset of samples were analyzed by SXRF superconducting wiggler and ion microprobe (IMP). Mid-range Dvalues (by IMP) for clinopyroxene in a picritic starting composition at 1300°C and 1 GPa are: Dye, 0.273; Dv, 2.77; Dcr, 7.32; Dsr, 0.079; Dy, 0.376; Dzr, 0.089; Dnb, 0.003; Dnf, 0.179: Dta, 0.010; DLa, 0.03; Dee, 0.054; Dn0, 0.112; Dsm, 0.201: DDy, 0.280; and Dvb , 0.255. D-values for Ta and Zr are lower than adjacent D-values for REE (La and Nd, respectively) on a spider diagram; however, there is no decoupling observed between Dr,, DHf and adjacent DREE (Dv.vb). Many of the experiments resulted in sector-zoned augite, and these constrain the nature of crystal chemical controls on partitioning at fixed T,P,X. Sector-zoned augite in an alkaline basalt at 1250°C and 1 GPa shows enrichment in Ti, AI, Ca, Cr, trace HFSE and REE and depletion in St, Mg and Fe in the last-growing (100) sector relative to the slower growing (010) sector. Reverse-zoning in incompatible elements in the ( 100 ) sector adjacent to normally-zoned (010) sectors confirms that sector zoning in augite arises from differences in surface kinetic processes (adsorption-desorption) during crystal growth. A positive correlation between IVA1and HFSE concentrations emphasizes the importance of coupled substitution involving highly charged cations in M sites and AI in T sites. I. Introduction Clinopyroxene is an i m p o r t a n t repository for trace elements in b o t h the m a n t l e source regions Present addresses:

Geological Survey of Canada, Continental Geoscience Division, 601 Booth Street, Ottawa, Ont. K1A 0E8, Canada. Lawrence Livermore National Laboratory, Earth Sciences Department, P.O. Box 808, L202, Livermore CA 94550, USA.

o f basalts and a m o n g s t mafic cumulates in crustal m a g m a chambers. As such, it has been the focus o f n u m e r o u s detailed trace-element and isotopic studies a i m e d at unravelling the nature o f m a n t l e melting, or the crystallization history of derivative mafic melts. Modelling constraints on the nature of h i g h - t e m p e r a t u r e processes require partition coefficients as input, and the usefulness of these calculations is contingent in part, on the trace-element partition coefficients used

0009-2541/94/$07.00 © 1994 Elsevier Science B,V. All rights reserved SSD1 0009-2541 ( 94 ) 00072-G

128

7". Skulski et al. / Chemical Geology 117 (1994) 12 7-14 7

beam techniques, such as the ion microprobe and proton-induced X-ray emission method (PIXE), it has become possible to attain fine-scale spatial resolution, detect elements at low concentrations, and obtain simultaneous multi-element analyses. As a result it is now possible to extract detailed trace-element variations from individual natural crystals (e.g., Shimizu, 1990; O'Reilly et al., 1991 ), or determine multi-element partition coefficients from a single experimental charge (e.g., Hart and Dunn, 1993). These technological advances provide a means of address-

and the assumptions made on their constancy. To date, most mantle melting models, whether forward (e.g., Johnson et al., 1990), or inverse (e.g., McKenzie and O'Nions, 1991 ), have assumed constant partition coefficients (D; weight concentration ratio in mineral/melt) over a large pressure and temperature range. This approach may be too simple, particularly since contemporary models emphasize the importance of small melt fractions, and the pooling of melts derived from a large pressure range. With the advent of highly sensitive microTable 1 Starting compositions for basalt experiments

l PD-27 PIC SiO2 TiO2 A1203 Fe203 Cr203 MgO FeO MnO CaO NazO K20 P205 H20

Total Sc Cr Ni Sr Y Zr V Nb Ta Hf La Ce Nd Sm Gd Dy Yb

2 HF-13 NEP

3 HF-24 AOB

4 AT-112 HAB

5 Cr-Di +HFSE

6 5+ REE

7 Garnet glass

8 HF-24+ 5

9 PD-27+ 6

10 HF-13+ 6

I1 HF-24+ 6

I2 AT-II2 +6

13 HF-13 + 7

46.94 1.03 13.37

40.07 2.37 11.43 4.88

46.94 2.08 13.78 2.88

49.0 1.02 19.0

37.96 0.32 4.15

37.23 0.31 4.07

36.34 0.16 19.23

45.04 0.94 12.19

10.91 8.91 0.23 10.69 4.95 1.86 1.42

10.19 9.20 0.17 9.31 3.08 1.04 0.37

6.18 9.47 0.21 11.40 2.56 0.65 0.11

0.57 11.51 1.93 0.04 15.36 0.87

0.56 11.29 1.89 0.04 15.06 0.85

2.06 19.61 4.18 0.15 4.42

0.05 14.04 9.74 0.16 11.67 1.32 0.02 0.08 2.0

38.99 2.12 10.48 4.30 0.05 10.73 8,04 0,21 10.90 4.45 1.64 1.25 2.0

45.05 1.86 12.55 2.54 0.05 10.09 8.30 0.15 9.69 2.80 0.92 0.32 2.0

46.87 0.93 17.16

14.67 10.83 0.18 11.55 1.40 0.02 0.09

45.13 1.86 12.56 2.54 0.06 10.11 8.30 0.15 9.72 3.57 0.92 0.32 2.0

0.05 6.56 8.53 0.19 11.53 2.34 0.58 0.10 2.0

38.91 2.11 11.97 4.30 0.21 11.54 8.27 0.22 9.86 4.37 1.64 1.25 2.0

100.38

97.72

99.04

99.60

72.71

71.30

86.15

97.24

97.25

95.16

96.32

96.84

96.65

922 443 106 18 54 249 9

330 240 1,498 24 304 221 94

300 217 525 16 133 197 28

123 646 191 464 163 8,985 173 2,497 2,474 6,685

121 1,190 391 93 162 8,745 220 2,432 2,426 6,555 888 293 142 74 74 79 78

121 664 212 1,321 168 8,965 195 2,508 2,426 6,555 888 293 142 74 74 79 78

121 638 191 464 160 8,814 173 2,449 2,426 6,555 888 293 142 74 74 79 78

121 373

1,255 3,888

1,231 3,813

14,097

1,515 90,481

1,486 88,735

297 1,690

25,225 25,242 68,207

24,738 24,755 66,891 9,064 2,988 1.449 750 746 810 799

7,001 16,747 1,663 72,696 8,453 3,448

146 8,697 2,425 2,426 6,555 888 293 142 74 74 79 78

1,673 212 1,321 50 434 195 769 16,412 163 7,125 828 338

I = Baffin Island picrite with 25% olivine (ol) phenocrysts; 2 = Hirshfield nephelinite with 10% ol phenocrysts and ~ 2% ol xenocrysts; 3 = Hirshfield alkaline olivine basalt with trace ol; 4=Atka high-alumina basalt with 7.4% ol, 42.9% plagioclase and trace magnetite phenocrysts (mode from B. Marsh ); 5 = doped glass ( HFSE ) based on Cr-diopside from a lherzolite xenolith at Alligator Lake (~; 10047; Francis, 1987 ); 6 = Cr-diopsidic glass from 5 with REE; 7=doped glass based on experimentally synthesized garnet (run 43; Takahashi, 1986); 8= weighted composition of alkaline basalt with 10% of 5 and 2% water; 9-12 = various rock powders with 10% of 6 and 2% water; 13 = nephelinite with 10% of 7 and 2% water.

T. Skulski et al. / Chemical Geology 117 (1994) 1 2 7 - 1 4 7

ing an important problem in the study of trace element partitioning. Partition coefficients are known to vary with temperature (e.g., Irving, 1978), pressure (e.g., Green and Pearson, 1983), melt composition (e.g., Nielsen, 1985) and crystal composition (e.g., Blundy and Wood, 1991 ). Isolating the individual contributions of these variables is not straightforward, since they are interrelated in multicomponent systems. Clinopyroxene crystals commonly show chemical zoning in elements such as A1, Ti and Cr in both natural samples (e.g., Shimizu and Le Roex, 1986) and experimentally synthesized pyroxenes (Hart and Dunn, 1993 ). In principle, crystal chemical controls on trace-element partitioning can be isolated from changes in intensive variables and melt composition, by growing chemically zoned minerals. In this study, we report the results of a series of high- temperature ( 1235-1300 ° C ) and -pressure (1-2.8 GPa) experiments designed to determine clinopyroxene-melt trace-element partition coefficients in a series of natural basalts (nephelinite, alkaline basalt, high-alumina basalt and picrite; Ti, Cr, V, Sr, Y, Zr, Nb, Ta, Hf, La, Ce, Nd, Sm, Dy and Yb). Trace-element concentrations were determined for all of the experiments by synchrotron X-ray fluorescence microprobe (SXRF; Ti, Y, Zr, Nb, Ta, Hf), and for a subset of samples by SXRF superconducting wiggler (La, Ce and Nd) and ion microprobe

129

(Ti, Cr, V, Sr, Y, Zr, Nb, Ta, Hf, La, Ce, Nd, Sm, Dy and Yb). Some of the samples formed sectorzoned augite crystals. One particularly wellformed crystal was studied in detail by electron microprobe and ion microprobe, and the results show the important role played by coupled substitution of highly charged cations and A1 between silicate melt and augite.

2. Experimental method

Powdered rock samples ( < 300 mesh; nephelinite, alkaline basalt, picrite and high-alumina basalt) were doped with a glass powder containing a multi-element spike (Table 1 ). Water was added (liquid form) to ensure rapid growth of large crystals and promote faster diffusion of trace cations in the silicate melts. The diopsidic glass spike (Table 1) is based on natural Crdiopside from a lherzolite xenolith containing FO89.8olivine (sample 1004 7; Francis, 1987 ) and the garnet glass on a experimentally synthesized garnet (run 43; Takahashi, 1986). Trace-metal oxides were added to these glasses, and the resulting mixtures were multiply fused. The starting mixtures consisted of rock powder (91.2%), doped glass (7% diopside or garnet glass + 2.8% trace metals) and water (2%; Table 1 ). The experimental assemblies consisted of powdered starting mixtures sealed in graphite-lined, welded Pt capsules (3-ram ID), packed in

Table 2 Experimental r u n c o n d i t i o n s Run

Bulk composition

T (~C)

P (GPa)

Time (hr)

% liq

% cpx

% ol

TSI02 TS137 TS139 TSI41 TS166 TS167 TSI81

HF-24+CP P D - 2 7 + CP HF-24+CP AT-112+CP HF-24+CP HF-13+CP HF-13+GA

1,250 1,300 1,250 1,235 1,270 1,270 1,270

1.0 1.0 1.0 1.5 2.5 2.5 2.8

24 24 24 24 20 20 24

85 90 89 85 64 73 91

10 1 6 14 30 18

5 9 5 6 9 9

% zr

lr. tr.

Sectorzoned cpx

x x x x

% Fc loss

Xr 2

% diff. ZrsxRv m.b.

16 27 6 14 8 10 9

0.48 0.20 0.21 0.05 0.01 0.38 0.70

+20 + 1 + 18 -21 -21 +1 +38

C h e m i c a l m o d e s a n d Fe loss calculated by least squares. A b b r e v i a t i o n s used are l i q = l i q u i d (glass): cpx = c l i n o p y r o x e n e ; o1= olivine; zr = zircon; L'r 2 = s u m o f residuals s q u a r e d ( bulk-calculated c o m p o s i t i o n ); % diff. ZrsxRv m.b. = difference ( % ) between calculated Zr ( S X R F d a t a - m o d a l l y weighted) a n d e s t i m a t e d bulk c o m p o s i t i o n (Table 1 ).

130

72 Skulski et al. / Chemical Geology 117 (1994) 12 7-14 7

Table 3

Average glass compositions in basalt experiments TSI02

TS137

TS139

TSI41

HF-24

2a

PD-2 7

2a

HF-24

20

AT- 112

2~

SiO2 TiO? A1703 MgO FeO MnO CaO Na20 K20

47.48 2.06 14.60 7.74 9.43 0.19 9.36 3.66 1.15

0.71 0.28 0.36 0.27 0.20 0.08 0.26 0.32 0.05

47.17 1.08 13.65 11.06 6.95 0.13 12.74 1.69 0.03

0.36 0.15 0.20 0.03 0.23 0.03 0.06 0.12 0.02

46.53 1.85 14.05 7.92 10.30 0.15 9.78 3.51 1.07

0.12 0.19 0.26 0.13 0. l l 0.04 0.08 0.14 0.02

48.41 0.79 19.01 5.60 7.85 0.16 10.63 2.87 0.64

0.31 0.27 0.13 0.11 0.17 0.03 0.11 0.07 0.04

Total

95.67

94.50

95.16

95.96

0.68

0.35

0.73

0.79

Fe/Mg

Ti V Cr Sr Y Zr Nb Hf Ta La Ce Nd Sm Dv Yb

EMP/

EMP/

SXRF

SXRF

12,349

6,475

784 290 12,981 3,709 5,240 1,920

163 231 9,781 2,736 4,330 1,540

IMP

EMP/

IMP

2a

EMP/

( n = 3)

SXRF

12,669 178 404 591 144 12,983 3.741

993 7 25 37 8 1,80l 484

4,736

1,049 403 202 89 123 86

97 35 17 5 5 5

SXRF 6,409 297 1,156 127 149 12,343 3,517

1,154 385 192 87 126 83

11,091

667 240 11,947 3,354 4,800 1,750 1,110 364 142

620 201 7,798 3,025 3,237 1,677 1,040 345 155

IMP

4,766 204 116 533 123 7,949 3,485

1,000 349 161 71 101 62

VG2 TS166

TS167

TS181

HF-24

2a

HF-13

2a

HF-13

2a

accepted value

this study

(n=9) 2a

SiO2 TiO2 AI~O3 MgO FeO MnO CaO Na20 KzO

44.72 2.60 16.26 5.12 10.58 0.60 7.13 4.05 1.41

0.16 0.26 0.32 0.11 0.18 0.00 0.08 0.20 0.08

39.40 2.81 12.76 5.98 12.11 0.23 9.99 5.99 2.31

0.46 0.21 0.25 0.15 0.29 0.04 0.15 0.47 0.04

40.73 2.64 13.29 8.81 10.94 0.25 10.50 5.34 1.94

0.60 0.48 0.14 0.36 0.33 0.10 0.15 0.32 0.04

50.80 1.85 13.90 6.80 I 1.90 0.22 11.10 2.70 0.21

50.60 1.85 13.73 6.82 11.66 0.20 11.06 2.82 0.20

0.29 0.17 0.26 0.08 0.17 0.08 0.18 0.16 0.02

Total

92.47

91.58

94.44

99.48

98.94

0.93

1.14

0.70

Fe/Mg

T. Skulski et al. / Chemical Geology 117 (I 994) 127-14 7

131

Table 3 (continued)

Ti V Cr Sr Y Zr Nb Hf Ta La Cc Nd

TS166 HF-24

TS167 HF-13

EMP/ SXRF

EMP/ SXRF

IMP

EMP

IMP

2c~ (n= 3 )

15,587

16,846

15,826

784 244 10,404 3,739 4,550 2,155 1,150 291 151

2,288 242 11,491 3,496 4,640 1,935 645 509 276

21,350 122 86 3,045 231 18,827 5,822

18,588 212 940 2,246 75 771 1,464

454 21 91 80 2 26 91

Sm Dy Yb

TSI81 HF-13

1,955 892 394 144 177 114

E M P / S X R F = electron microprobe (Ti)/synchrotron XRF; IMP = ion probe.

graphite cylinders. This assembly was contained in oven-fired brucite cups with lids, that were packed along with crushable alumina and Pyrex ® spacers in a cylindrical graphite furnace tube. The furnaces were inserted into a Pyrex ® sleeve contained within a cylindrical pressure medium of compressed salt (19-ram diam.). All experiments were conducted using modified Boyd and England (1960) -type piston cylinders in the cold piston in configuration. Pressure and temperature were stepped sequentially to run conditions and held constant for typical run durations of 24 hr (Table 2). Ulmer ( 1989 ) found that D-values for high field strength elements (HFSE) were stable for run times between 12 and 51 hr, and based on a number of our own test experiments, we found stable D-values for run times in excess of 24 hr. Temperature was controlled using W R e 3 % - W R e 2 5 o / o thermocouple wire calibrated against manufacturers specifications, and without emf or pressure correction.

3. Analytical techniques The run products were mounted in epoxy, sectioned and polished. The samples were analyzed

and photographed [back-scattered electron (BSE) images] at Rensselaer using a 5-spectrometer, JEOL ® 733 electron microprobe. Analytical conditions for glass analysis were 15-keV accelerating voltage, 10-nA beam current, ~ 10/tm beam diameter for glass analysis ( 1 /tm for crystals) with 40-s counting time. The accuracy and precision of the analyses was determined by the replicate analysis of basalt and rhyolite glasses (Table 3), and augite secondary standards (Table 4). Na loss in hydrous glasses was minimized by using a defocussed beam, and by analyzing Na at the start of the electron microprobe (EMP) analysis. For SXRF analyses, thin wafers of the microprobe mounts were cut, doubly polished (0.3/tm A1203 grit) so as to produce a single-phase, thin layer of the run products, and mounted with Crystal Bond ® on Spectrosil ® slides. The density, thickness and chemical compositions of the phases analyzed were determined for reduction of the SXRF data. The thickness of the sample wafers was measured optically and corrected for refractive index. The density of the samples was calculated from their measured chemical corn-

132

72 Skulski et al. / Chemical Geology 117 (I 994) 127-147

Table 4 Clinopyroxene compositions in basalt experiments TSI02 HF-24 core SiO2 TiO2 AI20~ Cr203 MgO FeO MnO CaO Na20 Total Fe/Mg

Ti V Cr Sr Y Zr Nb Hf Ta La Ce Nd Sm Dy Yb

51.60 0.76 5.23 0.59 16.86 5.23 0.15 18.98 0.69

TSI37 PD-27

TS139 HF-24

TSI41 AT-I12

TSI41 ATI12 BSE d

2~

core

20

(100)

(010)

BSE b

0.44 0.14 0.66 0.17 0.13 0.18 0.03 0.23 0.03

51.45 0.34 5.09 1.38 18.08 3.45 0.15 19.61 0.38

0.13 0.04 0.11 0.12 0.20 0.13 0.04 0.26 0.01

51.52 0.63 4.81 0.61 17.24 5.45 0.13 17.80 0.70

52.96 0.44 3.38 0.40 18.86 5.92 0.20 17.35 0.60

46.53 0.70 13.29 0.36 12.67 5.56 0.22 19.94 0.68

50.14 0.41 7.99 0.39 15.40 6.12 0.22 19.23 0.55

100.01

99.96

98.89

100.12

99.99

0.17

0.11

0.18

0.18

0.25

0.22

100.5

SXRF

SXRF

IMP

SXRF

IMP

SXRF

IMP

SXRF

SXRF

IMP

4,556

2,038

3,777

2,327 544 3,033 25 48 737 2.5

4,197

2,458

46 76 972 53 917 23

4,540 794 4,481 43 74 1,838 15

2,638

73 141 1,368 64 1,220 55

1,748 824 8,466 10 56 1,098 10

81 210 4,529 149 3,100 185 149 66 97

67 185 3,648 93 2,480 128 123 73 80

3,364 1,260 3,122 36 133 4,453 79

35 21 22 18 35 21 TS166

TS167

HF-24

2a

SiO2 TiO2 A1203 Cr203 MgO FeO MnO CaO Na20

49.90 1.09 7.92 0.13 15.10 6.85 0.08 17.49 0.92

1.46 0.37 1.66 0.13 1.20 0.65 0.04 0.24 0.12

Total

99.48

101.25

0.26

0.19

Fe/Mg

TS139 HF-24

HF-13

49.29 1.85 7.32 0.57 13.71 4.67 0.23 22.15 1.39

91 109 1,900 53 1,600 42 79 40 40

77 49 53 38 70 39

74 74 919 42 717 34 37 15 19

Kakanui augite 2o

accepted value

this study

2a

1.18 0.42 1.63 0.20 0.89 0.17 0.02 0.21 0.14

50.73 0.74 8.73 0.14 16.65 5.37 0.22 15.82 1.27

50.20 0.73 8.24 0.16 16.10 6.21 0.12 15.99 1.35

0.41 0.26 0.14

99.53

99.08

0.15 0.19 0.11 0.16 0.18

32 23 27 19 41 24

121 74 80 57 113 65

T. Skulski et al. / Chemical Geology 117 (1994) 12 7-14 7 Table 4 (continued)

Ti V Cr Sr Y Zr Nb Hf Ta La Ce Nd Sm Dy Yb

TS166 HF-24

TS167 HF-13

SXRF

SXRF

IMP

6,535

11,091

12,297 79 253

132 188 2.234 99 1,670 57

200 103 2,736 50 2,760 50 103 65 64

123

154 113 100 51 74 46

BSE-b = bright backscattered electron sector; BSE-d = dark backscattered electron sector.

positions using the method of Niu and Batiza (1991). The SXRF analyses were conducted at the National Synchrotron Light Source, Brookhaven National Laboratory. Most analyses were conducted with five bunches of accelerated electrons to insure maximum beam current and duration. On beamline X 2 6 A (white light), spectra were accumulated for 5 min (live time) using a Canberra ® energy-dispersive spectrometer (EDS). The photon beam was collimated to 8 /tm, and depending on run conditions, spectral intensities were corrected for Be, C and air filtration of the incident beam. Deadtime was kept below 65% by using a 170-/zm A1 foil filter. The beam centring was monitored and adjusted to give optimal Sr peak/background count values on an anorthite standard. The superconducting wiggler, line X17B, was used to obtain La, Ce and Nd concentrations. The photon source was filtered with Ta, and collimated to 25 #m using Pb windows. Fifteen-minute count times (live) were used. The principles of SXRF data reduction are discussed in Sutton et al. ( 1987 ). An interactive,

133

graphical peak fitting program was used to calculate peak areas. This involved choosing an energy region of interest between 3 and 17 keV, and fitting the background with a 3 ° polynomial. Individual peaks were fit by Gaussian approximation with peak energy and full width at half maximum (FWHM) as seed values. Elemental concentrations were calculated from the peak areas of K-Ks, Ca-Ks, Fe-K~, Ti-K~, Cr-K~, SrK~I, Y-K~I, Nb-K~, Zr-K~, Ta-L,~ and Hf-L~. Iron (as determined by EMP) was the internal reference element used. For any element z, the concentration C t : ~ was computed from (Sutton et al., 1987): C< :) = S t: ) ( I t : ) / I t Ref) ) Ct Ref)

where It:)=integrated count from element z; ItRer)=integrated count from Fe; C
134

71 Skulski et al. / Chemical Geology 117 (1994) 12 7-14 7 10000

1000

13..

~

E 100 o...

E

I000

o. ._~ IO0

0

i

10

0

100

10 10

1000

100

1000

10000

La (pprn) SXRF

Ce (ppm) SXRF 100000

10000

. . . . . . . .

,

. . . . . . . .

,

. . . . . . . .

,

Q.

E r~

E 100 c~

~

....

I

1000 13-0000

•

.0 Z

1ooo

I0

1O0 700

1000

10000

100000

Zr (ppm) SXRF

1

10

1 O0

1000

10000

Nb (pprn) SXRF

Fig. 1. Comparison between synchrotron XRF microprobe (SXRF) and ion microprobe analyses (IMP) of basaltic glasses and clinopyroxene crystals. The error bars shown in this and subsequent figures reflect +_15% analytical uncertainty in elements determined by SXRF, and _+5% for trace elements determined by IMP.

Cam6ca ® IMS 3-fion microprobe at Woods Hole Oceanographic Institution. Operating conditions of the ion microprobe are the same as described in Shimizu and Le Roex ( 1986 ) and Sisson (1991). Energy filtering was used to eliminate molecular ion interference in the mass spectrum. The secondary accelerating voltage (4500 eV) was reduced by 90 eV to measure 28Si, 47Ti ' 5~V, S2Cr' 88Sr' 89y and 9°Zr, and by 35 eV to measure 3°Si, 139La, 14°Ce, 146Nd, 147Sm, 163Dy, 167Er, 174Yb, 18°Hf and 181Ta. The primary O beam was adjusted to yield 3-105-5 • l05 cps of 28Si on an amphibole reference standard. A focused primary beam of ~ 15-#m diameter was used for rare-earth element (REE) analysis, and

8 #m for other trace elements. Background and secondary ion intensities were ratioed to the intensity of 163Dyfor REE, and 28Si for other trace elements. The intensity ratios were converted to concentrations through the use of correction factors derived from working curves (Shimizu and Le Roex, 1986). In the case of H f and Ta, concentrations were not calculated, but partition coefficients were derived from the intensity ratios ~8°Hf/28Si and ~SJTa/28Si of the respective phases. The uncertainties in major-element concentration determined by ion microprobe is believed to be less than + 1%, to approximately _+5% for the trace elements. A number of samples was analyzed by both

72 Skulski et a[. / Chemical Geology 117 (I 994) 12 7- I4 7

SXRF and ion microprobe (IMP) (Fig. 1 ). The IMP and SXRF data for La, Zr and Ce data correlate well; however, at the lower Nb concentrations found in clinopyroxene, the concentrations determined by SXRF are higher by a factor of 6. Unlike the IMP which is essentially a near-surface technique, the SXRF probe induces X-ray fluorescence over the entire sample thickness for the thin samples analyzed here. However, though some of the high Nb SXRF values in pyroxene crystals may reflect inadvertent analysis of subsurface glass inclusions, the systematic shift in NbsxRF concentrations relative to the IMP data, and the greater precision and lower detection limit of IMP vs. SXRF analyses, suggest an analytical bias in the SXRF data set at low Nb abundances.

A

135

4. Experimental results The results of 7 experiments are presented in Tables 2-6. Augite crystals ( 1-30% ) were found in experiments conducted over a pressure range from 1 to 2.8 GPa, temperatures from 1235 ° to 1300 ° C, and bulk compositions including alkaline basalt, nephelinite, picrite and high-alumina basalt. All of the runs, with the exception of TS141 high-alumina basalt, crystallized olivine and olivine are the liquidus phase in a nephelinite run at 2.8 GPa (TSI81-HF-13; Table 2). Trace amounts of zircon were found in runs TS141 and TS166 (alkaline basalt). Sectorzoned augite was observed in runs TS137 (picrite), TS 139, TS 166 (alkaline basalt; Fig. 2 ) and TS 141 (high-alumina basalt).

B

Fig. 2. A. Back-scattered electron image (72 × ) of alkaline basalt + doped diopsidic glass (TS 139HF-24:1350 ~C; 1 GPa ). The black region near the base of the photo is the graphite liner. Large black crystals on middle right and small crystals at base of charge are olivine (Fos4). The charge is 1.43 mm long. B. Back-scattered electron image of larger sector-zoned augite from TS139-HF-24. The crystal is 220 by 135 ~tm. The chemical composition of this crystal is shown in Figs. 3-5 and in Table 4. Points marked a-a' are the coordinates of a line scan ( Fig. 4 ) through the (100)sector and b-b' is through the (010) sector.

136

A

C

E

T. Skulski et al. / Chemical Geology 117 (1994) I27-147

B

13

F

Fig. 3. X-ray maps of a sector-zoned augite (TSI39-HF-24; Figs. 2, 4 and 5 ): (A) Ti X-ray map showing (100) and (010) sectors of augite; (B) A1 X-ray map; (C) Cr X-ray map; ( D I Mg X-ray map; (E) Fe X-ray map); and (F) Ca X-ray map. The maps measure 289 by 188/tm and are composed of focused spot analyses spaced at 2-/~m intervals in each direction. The sample stage was rastered under the stationary electron beam ( 15-keV accelerating voltage, 100-nA cup current ), dwelling 2.8 s on each point. The scan took a total of 14 hr to complete. The elements analyzed by wavelength-dispersive spectroscopy (WDS) (Ti, A1, Cr, Mg and Fe ) were first standardized using pyroxene, and the intensities reported are in weight percent concentrations uncorrected for matrix effects (believed to be minor). These wt% concentrations are indicated on the WDS maps for the core and rim of each sector. Ca was analyzed by energy-dospersive spectroscopy (EDS) and is reported as total X-ray counts. Dark pits in the crystal were formed during ion microprobe analysis.

T. Skulski et al. / Chemical Geology 117 (1994) 12 7-14 7

137 A

(100

i

5.5

\

4.5

8t~

4

,,~

r¿

II

3.5 I

i

I'

I I

I

2.5

Ir~

j! I

I II

fr I

2

1

I

I

r"

J

2.2 "~ diffusional pileup

b. j~ I

~ l~r f -b II '

J

k

1.8

glas.~

glass

1.6

o~ 1.4

B

f

/ ion hole / /

.o 1.2 /

core rim

El' 0.8 I(1001

rim

I c°relIJ ~ ~J" ,,~" y

0.6

0.4 -150

I

i

i

:

I

-100

-50

0

50

100

150

microns

Fig. 4. Line scans from X-ray maps of sector-zoned augite (TSI39-HF-24) shown in Figs. 2 and 3. The line scans are located in Fig. 2B, where a-a' is a scan through the (100) sectors and b-b' is a scan through the (010) sectors. The scans are composed of swaths that represent the average composition of a 3-pixel-wide line (6/~m). A. A1203 line scan through a sector-zoned augite shown as coordinates a-a' in Fig. 2B. The (100) scan starts in glass on the far right, and ends at the rim of an adjacent crystal. The spikes in AI concentration on the left are where the scan passes through a hole formed during ion microprobe analysis. The shorter scan on the bottom, passes through the (010) sectors (coordinates bb' in Fig. 2B). The spike in the middle of this scan is where the scan passes through a (100) domain in the centre of the crystal. B. TiO2 scan through sector-zoned augite. A concentration gradient adjacent to the rim of the crystal is thought to be a result of rapid quench growth and diffusional pile-up.

138

T. Skulski et al. / Chemical Geology 117 (1994) 127-147 100 Augite / AOB 1250 °C, 1 GPa

10 1

D eff..1

~/~ "'"'-....

.01 .001 .0001

i

,

Nb

Ta

i

i

La

i

Sr

Ce

h

i

i

Sm

Nd

i

1010) h

Zr

Ti

,

,

Y

Hf

~

i

V

Yb

Cr

Fig. 5. Clinopyroxene-melt effective partition coefficients determined by ion microprobe for the large crystal (run TSI39-HF-24) shown in Figs. 2-4. The dark region is the (010) sector, and the white area corresponds to the (100) sector.

Single clinopyroxene crystals have A1203 c o n tents that show a significant variability (e.g., A1203 in TS166 _+2a 21%). The greatest variation is shown by sector-zoned augite (see below). Sample TSI41 shows up to 40% difference in A1203 content between adjacent sectors. Two length scales of compositional zoning are measured in sector-zoned augite from run TS139HF-24 (Figs. 3 and 4; alkaline basalt; 1250°C, 1 GPa; Table 2). The (100) sector is enriched in Ti, A1, Cr and Ca relative to the adjacent (010) sector which has higher Si, Mg and Fe contents (Fig. 3 ). Zoning is also observed within individual sectors; (100) sectors have AI-, Ti- and Crrich cores, and Mg- and Fe-rich rims (Fig. 4). However, zoning in the (010) sector is just the reverse; the rims are rich in A1, Ti and Ca, relative to cores that are rich in Mg, Fe and Cr. The trace-element contents of this augite crystal are similarly zoned (Table 4). The (100) sector is enriched in REE, HFSE, Sr, Y and V relative to the (010) sector (Fig. 5; Table 4). The cores of the (100) sector are enriched in REE, HFSE, Sr, Y and V, relative to the rims, whereas the reverse is true in the (010) sectors (Table 4). The (100) augite sector in run TS 139-HF-24 represents the preferred direction of growth [relative to (010) ], and zoning within this sector may have been established early, as a result of relatively rapid crystal growth. The chemical zoning within the

(010) augite sectors are directly in stride with the chemical evolution of the silicate melt during the run. Within the glass, an 11% enrichment in Nb occurs between the base (4017 ppm) and top of the charge (3563 ppm ), and this corresponds to the total amount of crystallization calculated by least-squares mixing (Table 2 ). The experiments produced glass compositions (Table 3) whose major- and trace-element chemical variability were within, or close to analytical uncertainties. One exception is run TS139-HF-24, in which the glass shows chemical zoning which reflects failure of the silicate melt to equilibrate with the crystals of augite and olivine at the base of the charge. The results of a mass-balance calculation on the measured (SXRF) Zr content of the run products (modally weighted) show that the calculated bulk Zr concentration is almost within analytical uncertainty ( + 15%) of the estimated bulk compositions (Table 1 ). Larger discrepancies in Zr mass balance (e.g., TS181-HF-13) arise from uncertainties in the estimated bulk composition (Table 1 ). The estimated bulk compositions reflect pre-fusion reagent weights, and ignore possible sample loss during glass preparation (i.e. wetting of Pt crucibles) and potential inhomogeneity in the spike + rock powder starting mixtures. All of the samples appear to have lost some Fe to the Pt capsules (as constrained by least-squares mixing calculations; Table 2). Although this problem is partly mitigated by using graphite-lined Pt capsules, the presence of water in the run must enhance the diffusivity of Fe in the silicate melt. Iron-magnesium exchange coefficients (Kd-values) for olivine and clinopyroxene in run TSI37-PD-27 exceed accepted values (clinopyroxene-melt Kacve/Mg)= 0.29, Thompson, 1974; and olivine-melt Kd~w/Mg~=0.30, Roeder and Emslie, 1970), and this can be shown quantitatively to reflect Fe loss during the course of the run. Runs TSI02, -166, -167 and -181 are characterized by Kd(v~/Mg)-valuesfor clinopyroxenemelt and olivine-melt that are lower than accepted values, and this probably reflects accumulation of ferric iron in the charge. Although some ferric iron is produced during Fe loss to Pt encapsulating materials, our calculations show

77. Skulski et al. / Chemical Geology I 17 (1994) 12 7-14 7

that enrichment of original ferric iron in the starting materials during crystallization of the charge, adequately explains the low Kd(ve/Mg)" values. Despite the fact many of the experiments in this study failed to attain major-element equilibrium, either because of Fe loss, or due to growth of sector-zoned augite crystals, useful constraints on the nature of crystal chemical controls on trace-element partitioning can still be extracted from the data set.

1.0

0.8

%=

141~

D~

1.32D-0.18

r2--0.9~/~

0.6

0.4

0.2

0.0 0.6

5. D-values: high field strength elements (HFSE: Ti, Zr, Nb, Hf, Ta) The concentrations of HFSE measured in clinopyroxene crystals in this study (Table 4) exceed common natural abundance levels (e.g., Fujimaki and Tatsumoto, 1984). In order to demonstrate that Henry's law applies at the concentrations encountered in this study, we compare the D-values of the geochemically similar element pairs, Hf-Zr and Ta-Nb, and test for

139

0.5

139d?D • I t

' 102 i

t

i

A

T

o:079014r=08 / I ]

0.4

g £3

0.3 0.2 0.1 0.0

0.0

0.2

0.4

0.6

0.8

1.0

1.0 2

Dzr 0.62DH(0.04(r = 0.96) Fig. 7. A. DHf vs. DTi diagram for clinopyroxene-haplobasalt, and -basaltic melt experiments. B. Dz~ vs. Dxi for clinopyroxene-basaltic melt experiments (see Fig. 6 for details).

0.8

0.6

a 0.4

0.2

'~167

13~9b 0.0 0.0

0.2

0.4

'

0.6

'

0.8

1.0

D Hf Fig. 6. Dzr vs. DHf diagram for clinopyroxene and basaltic melt experiments. Tie lines join adjacent sectors in sectorzoned augite (TS 141 and TS139; Table 5 ). Sectors with high mean atomic number, Z, are denoted b for back-scattered bright, and sectors denoted d are back-scattered dark sectors. Zr and Hfdata in runs abbreviated 102, 141b, 166 and 167 were determined by SXRF and unadorned error bars reflect +_21% propagated uncertainties. All other basalt data were analyzed by IMP, and error bars (with small bar) reflect + 7% uncertainties.

constancy. The rationale of this test is that primitive mafic lavas have approximately constant Zr/Hf- and Nb/Ta-values ( Z r / H f ~ 35-40, Nb/ Ta.~ 17; Jochum et al., 1989; with the exception of some intraplate basalts; Dupuy et al., 1992), and, therefore, their mantle source regions must have approximately constant values of the bulk partition coefficients Dzr/Di-lfand DNb/DTao v e r the range of natural abundance levels for these elements (Watson, 1985). In so far as clinopyroxene is likely to be present in the mantle source region of a wide compositional spectrum of basaltic magmas, ratios of experimentally determined D-values for Zr-Hf and Nb-Ta, should be constant in the Henry's law interval. Linear regression of the D H f and Dzr data in 1-GPa ex-

140

T. Skulski et al. / Chemical Geology I 17 (1994) 127-147

Table 5 Clinopyroxene partition coefficients TSI02 HF-24

Ti V Cr Sr Y Zr Nb Hf Ta La Ce Nd Sm Dy Yb

TS137 ~

PD-27 BSE-b

TS139 ~ HF-24

TS139 a HF-24

BSE-b

BSE-d

TS141"

AT- 112 BSE-b

EMP/ SXRF

EMP/ SXRF

IMP

EMP/ SXRF

IMP

EMP/ SXRF

IMP

EMP/ SXRF

0.369

0.315

0.340

0.193 3.07 7.10 0.052 0.344 0.059 0.001 0.122 0.004 0.033 0.061 0.141 0.232 0.337 0.288

0.886

0.282 0.329 0.099 0.019 0.212 0.015

0.359 4.61 10.99 0.076 0.517 0.143 0.004 0.321 0.016 0.072 0.122 0.260 0.421 0.571 0.494

0.238

0.093 0.486 0.105 0.017 0.233 0.029

0.273 2.77 7.32 0.079 0.376 0.089 0.003 0.179 0.010 0.030 0.054 0.112 0.20 ! 0.280 0.255

TS141 a AT-112

TS166 a HF-24

TS167 HF-13

0.076 0.442 0.145 0.016 0.153 0.022 0.071 0.1 l0 0.282

0.124 0.290 0.069 0.009 0.136 0.013 0.033 0.041 0.134

0.131 1.71 0.581 0.049 0.958 0.116 0.143 0.191 0.626

BSE-d

Ti V Cr Sr Y Zr Nb Hf Ta La Ce Nd Sm Dy Yb

EMP/ SXRF

IMP

EMP/ SXRF

EMP/ SXRF

IMP

0.519

0.706 6.18 26.91 0.068 1.09 0.560 0.023 0.925 0.098 0.121 0.210 0.494 0.804 1.124 1.047

0.419

0.658

0.662 0.648 2.94

0.168 0.770 0.215 0.026 0.367 0.026 0.144 0.117 0.397

0.087 0.426 0.238 0.014 0.595 0.026 0.160 0.128 0.232

0.108 0.925 0.468 0.031 0.766 0.080 0.118 0.212 0.516

0.532

0.006 0.007 0.079 0.127 0.255 0.359 0.418 0.405

BSE-b = bright backscattered electron sector; BSE-d = dark backscattered electron sector; EMP / SXRF = analysis by electron microprobe (Ti) and synchrotron XRF probe; IMP =analysis by ion microprobe. a Effective partition coefficients (for sector-zoned clinopyroxene) are shown.

periments (Fig. 6) results in a correlation coefficient (r 2) of 0.96. The excellent fit of this expression supports our belief that the DHr and

Dzr-values were measured within the Henry's law interval. A somewhat poorer fit is found for DNb and DTadata.

T. Skulski et al. / Chemical Geology 117 (1994) 127- I47

141

Table 6 Olivine compositions in natural basalt experiments HF-24 102

SiO2 Cr203 MgO FeO MnO NiO CaO

Total Fe/Mg Kcl(Fe.Mg )

39.90 0.09 45.06 15.45 0.19

PD-27 2or 0.34 0.04 0.49 0.57 0,08

137 40.25 0.11 47.29 12.28 0.21 0.00 0.25

HF-24 2t~ 0,18 0,01 0,39 0,74 0.05 0,00 0.02

139 40.06 0.12 45.08 15.59 0.18 0.00 0.21

HF-24 2a

166

0.09 0.05 0.12 0.08 0.01 0.00 0.01

39.59 0.00 42.10 18.63 0.11 0.00 0.14

HF-13 2o 0.15 0.00 0.45 0.40 0.01 0,00 0,06

167 39.92 0.08 43.87 16.36 0.25 0.14 0.44

HF-13 20 0.19 0.03 0.86 0.93 0.01 0.08 0.12

181 39.92

1.54

43.05 17.09 0.34

2.10 1.77 0.17

0.40

0.13

100.69

100.40

101.29

100.60

101.14

100.20

0.19 0.28

0.15 0.43

0.19 0.26

0.26 0.28

0.21 0.18

0.22 0.31

Ti V Cr Sr Y Zr Nb

20

IMP

D

32 23 541 0.8 0.2 1.4 1.7

0,002 0.189 6.291 0.00036 0.0027 0,0018 0.0012

IMP = analysis by ion microprobe.

The D-values for Ti, Zr, Nb, Hf and Ta in alkaline basalt (TSIO2-HF-24; Table 5 ) are 0.37, 0.105, 0.017, 0.233 and 0.029, respectively, at 1250°C and 1 GPa. Considerably higher Dzrvalues (0.27) are found at lower temperatures and for more silicic melt compositions, such as the 950-1000°C data (0.9-1 GPa) reported by Watson and Ryerson (1986). In higher-temperature experiments on picrite, D-values for Ti, Zr, Nb, Hfand Ta are 0.27, 0.089, 0.003, 0.179 and 0.01, respectively (TS137-PD-27; Table 5). U1mer (1989) reported clinopyroxene D-values from experiments at ~ 1340°C (2.8 GPa) on a picrite that are higher than values reported here for Nb (0.02) and Hf (0.22), and lower for Zr (0.03) and Ti (0.18 ). Slightly higher DnvsE-Values are reported in the higher-pressure experiments of Johnson (1993; 2-3 GPa; 13001470°C) relative to our data on picrite. Forsythe et al. (1991 ) reported linear correlations between DTi- and DHFsE-values for clinopyrox-

ene in natural basalts synthesized at 1 atm, where Dzr = 0.55 × DTi and DHf= 1.5 ~
142

72 Skulski et al. / Chemical Geology I 17 (1994) 127-147

6. D-values: Sr, V and Cr D-values for Sr in unzoned clinopyroxene ranges from 0.079 in picrite (TSI37-PD-27; Table 5) to 0.087 in nephelinite (TS167-HF-13) and falls within the range of published values reported by Green (1994 in this issue). Sectorzoned augites have Dscvalues that are up to a factor of 2 higher than those in unzoned pyroxene (Table 5 ). D-values for V vary from 2.77 in picrite to 0.653 in nephelinite (Table 5 ). Ulmer (1989) reported a Dr-value of 1.57 for picrite at ~ 1340°C. Values of Dcr are also higher in the picrite starting composition (7.32) relative to the nephelinite (2.93) and slightly higher than published values of experimentally determined Dc~ (4-5; Longhi and Pan, 1989). D-values for olivine-glass range from 0.00036 for Sr, to 6.29 for Cr, and the latter is a factor of 10 higher than that measured by Kennedy et al. ( 1993 ).

7. D-values: rare-earth elements (REE: La, Ce, Nd, Sm, Dy, Yb) and Y Smooth chondrite-normalized REE profiles in volcanic rocks demonstrate that ratios of partition coefficients for adjacent REE are nearly constant in the minerals (including clinopyroxene) making up the mantle source regions of basaltic magmas, over the range of individual REE concentrations in mantle minerals (Watson, 1985 ). Therefore, nearly constant values Of DLa/ Dc~ (r2=0.94), DNd/DSm (r2=0.99) and Dyb/ DDy (r2=0.99) in clinopyroxene (Table 5) suggest that the REE doping levels used are within the limits of Henry's law. Furthermore, the geochemical similarity between Yb and Y requires that the ratio Dyb/D v in the minerals comprising the mantle source regions ofbasalts is insensitive to variations in Yb and Y concentrations. A nearly constant Dyb/Dy-value in the experiments conducted here (r 2= 0.98) indicates that the Y doping levels are also within the limits of Henry's law. The D-values for REE were measured over a temperature interval from 1300°C (TS137-PD27picrite) to 1270°C (TS167-HF-13 nephelin-

100 10 1 a

0.1 0.01 ~,"/~

0.001

-Z

•

AOB ( BSE dark ) 1250 ~'C, 1 GPa

O

NEP 1270 °C, 25GPa

'l G'Pa

.

.

.

.

0.0001

Ta Ce Nd Hf Ti Yb Cr Nb La Sr Zr Sm Y V Fig. 8. Multi-element clinopyroxene partition coefficient data for natural basalt samples. Shown are data from runs TS 137PD-27 (picrite); TS139-HF-24 (alkaline basalt, AOB; 2 sectors); TSI41-AT-112 (high-alumina basalt, ttAB); and TS167-HF-13 (nephelinite, NEP).

ite). The values of DLa , Dce, DNd, Dsm, DDy and Dyb vary considerably over this temperature interval (e.g., DLa 0.03-0.079). The DREE-Values for unzoned pyroxenes reported in this study (TS167, Table 5), fall within the range of published values (Green, 1994 in this issue); however, the DREE-Values of sector-zoned augites in TS141 exceed published values. Apparent D-values measured in sector-zoned augite are almost a factor of 2 higher in the high-alumina basalt TS141-AT-112 (DLa 0.143, Table 5). Interestingly the overall pattern of D-values for clinopyroxene vary systematically, regardless of whether the pyroxene was sector-zoned or not (Fig. 8). The pattern of D variations when plotted on an extended trace-element incompatibility diagram (Fig. 8 ) shows lower DNb and DTa relative to DLa , and for a few runs, slightly lower Dzr relative to DNd and Dsm. Johnson ( 1993 ) also reports lower DHFSErelative to DREEfor clinopyroxene-melt on incompatibility diagrams. Our results also agree with those of Hart and Dunn ( 1993; alkaline basalt 1380°C and 3 GPa) who report lower Dzr relative to Dsm, and no decoupling between Dxi, DHf and adjacent DREEon spider diagrams.

T. Skulski et al. / Chemical Geology 117 (1994) 127-147

8. Origin of sector zoning in augite Studies of naturally occurring sector-zoned augites have shown that the prismatic sectors of the form (100), (010) and (110) are enriched in Ti relative to basal sectors of the form (111) (e.g., Strong, 1969; Hollister and Gancarz, 1971; Wass, 1973; Downes, 1974; Leung, 1974). Dowry (1976) expanded the protosite concept of Nakamura (1973) and proposed a model in which each crystal face has an adsorption layer that is superimposed on a surface of minimum potential bonding energy. Depending on the crystal configuration the surfaces of least bonding will expose different proportions of protosites into which cations will be adsorbed. Cations with high charge and small radius will be bound most firmly to these partly completed sites. Dowry showed that the (100) surface would contain 4 half-filled M~ protosites, two Ml 4/6 sites on both (110) and (010), and one MI 4/6 site on (111). The experimental study of sector-zoned clinopyroxerie (CaMgSi206-CaTiA1206) by Kouchi et al. (1983) showed that interface kinetics (adsorption-desorption of cations) plays a central role in the formation of sector zoning, when the layer growth mechanism takes place. Within individual sectors, the TiO2 and A1203 contents increased with increase in growth rate and degree ofundercooling (AT). But at constant AT, different sectors in a single crystal show a negative correlation between crystal growth rate and TiO2 and A1203 contents. The order of TiO2 enrichment between sectors is (100) > (110) >/ (010) >/ ( i 11 ) at delta AT= 13 ° and 18 °C (Kouchi et al., 1983). The Dowty (1976) model of sector zoning is supported by trace-element data on sectorzoned augite (Shimizu, 1981) in which the slower growing prism sector ( 100 ) is found to be enriched in both compatible and incompatible elements relative to the faster growing (711) sector, and this enrichment is proportional to the ratio of charge to radius squared. These observations cannot be explained by coupling between crystal growth rate and trace-metal diffusion in the adjacent melt (e.g., Albarrde and Bottinga, 1972), but reflect surface kinetic processes as proposed by Dowty (1976).

143

The relative enrichment in non-quadrilateral constituents in the (100) vs. (010) augite sectors in run TSI39-HF-24 are consistent with the relative enrichment data of Kouchi et al. ( 1983 ) and of other studies of natural augites. The relative elongation of the crystal in the a-direction is interpreted to indicate more rapid growth perpendicular to the (100) sector as opposed to the adjacent (010) sector (Fig. 2B). Assuming constant growth (see below), an average rate of growth of 7.10-s cm s-~ is calculated for the (010 ) sector and 1- 10- 7 cm s- ~for ( 100 ). Both higher rate of growth and effective partition coefficients for incompatible elements such as Ti in the core of ( 100 ) (e.g., 0.36 ), relative to ( 010 ) (e.g., 0.19) are consistent with the data of Kouchi et al. (1983) at low AT (13°C). At first glance, the fact that the fast growing sector (100) is enriched in incompatible elements relative to the slower growing (010) sector could be interpreted to indicate that diffusion in the melt phase controlled development of sector zoning. If growth of the (010) sector was controlled only by diffusion in the liquid and interface equilibrium was attained, an unzoned crystal would be expected (Tsuchiyama, 1985). Furthermore, the reverse zoning observed in the (100) sector (i.e. decrease in A1 from core to rim; Fig. 4) clearly indicates that enrichment of incompatible elements in the melt phase adjacent to the growing crystal was not responsible for sector zoning. The isothermal experiments of Kouchi et al. ( 1983 ) produced reverse zoning in all sectors. We conclude that interface kinetics was responsible for generation of sector zoning in TS139-HF-24, and the greater proportion of partly coordinated M and M2 protosites, as predicted by Nakamura (1973) and Dowry (1976), can explain the enrichment in incompatible elements between the (100) and (010) sectors. The reverse zoning in non-quadrilateral constituents in the (100) sector of TS139 can be explained by the results of experiments in the simple system CaMgSi2Or-CaA12Si208 (Tsuchiyama, 1985). Tsuchiyama showed that isothermal crystallization resulted in diopside crystals showing reverse zoning in AI. The A1 content of the rim compositions decreased initially, and

144

T. Skulski et aL / Chemical Geology 117 (1994) 12 7-14 7

then became a constant value corresponding to the equilibrium value. Both Tsuchiyama and Kouchi et al. (1983) noted that the effective partition coefficient increases as AT or the rate of crystal growth increase. The reverse zoning process is attributed by Tsuchiyama to rapid crystal growth during initial large supercooling, followed by decreasing growth rate with time as equilibrium with the residual liquid is approached. Interface equilibrium is not maintained at the beginning of growth of the crystals, and interface kinetics by which excess amounts of solute (A1) are incorporated into the crystal during growth result in reverse zoning. Although this mechanism can explain the reverse zoning in the (100) sector of pyroxene in TS139-HF24, it cannot explain the concurrent normal zoning in the (010) sector. Development of a normal zoning profile in the (010 ) sector of the TS139-HF-24 augite is consistent with slower growth of this face relative to (100). As a result the (010) sector grew from a melt that became progressively depleted in compatible constituents such as Si, Mg, Fe (in part because of Fe loss) and Cr, and enriched in incompatible constituents. The theoretical treatment of Shimizu ( 1983, his equation 8) predicts that the greater the effect of surface kinetics (reaction or attachment/detachment kinetics), the less pronounced will be the core to rim zoning for incompatible elements and the more pronounced will be zoning in compatible elements. The concentration differences between core and rim in the (010) face of TS139-HF-24 are greater for incompatible elements (e.g., IAAII =1.3) and smaller for compatible elements (e.g., IACrl = 0.13 ) relative to the zoning observed in (100) (IdAll = 1.0 and IdCrl =0.36; Figs. 3 and 4A). These results are interpreted to indicate that surface kinetic effects were less important in the case of the (010 ) face relative to ( 100 ). We envisage initial fast growth parallel to the (100) and (001) axes, forming a skeletal crystal with a bowtie form, followed by continued growth in the (010) direction to fill out the form of the now euhedral crystal.

9. Crystal chemical controls on trace-element partitioning One of the interesting implications of the sector-zoned augite crystals is that it is possible to isolate the influence of crystal chemical controls on effective partition coefficients from system variables such as temperature and pressure, and extensive variables such as melt composition. The negative correlation between Mg (or Fe) and Ti, A1 (Fig. 3) and trace HFSE in the sectorzoned augite in run TS139-HF-24 (Table 4), is consistent with ionic-radius (Whittaker and Muntus, 1970) considerations that indicate that 5000

'

141d

141b

4000 "Zr (cpx) = 15181x ( ~VAIc p x ) + 7 2 ~ (r = 0.67) 3000

r ,67 T ,66

v J.-'O

2ooo

N

102

I

,o00

,,37 J

139d A

0

!

200 141b

150

2 ~v Nb (cpx) = 674 x ( AI cpx) - 54 (r =0.8)

j .

100

JJ

166~ _ . ~

j~-

141d ~ . . ~

/

Z

139d 0

0.0

139bli" ~P Q 137 lilt" I

0.1

167

B I

0.2

0.3

IVAI(cpx) Fig. 9. A. Tetrahedrally coordinated A1 in clinopyroxene vs. Zr content. Tie-lines j oin adjacent sectors in sector-zoned augite, where b stands for back-scattered bright (high mean atomic number, Z) and d stands for back-scattered dark (low mean atomic number, Z). B. Tetrahedrally coordinated AI in clinopyroxene vs. Nb content (see Fig. 6 for details).

T. Skulski el a l. / Chemical Geology l 17 (I 994)

HFSE substitute for Mg and Fe in the M~ site. A positive correlation between tetrahedrally coordinated A1 content (X~Av ) and Zr, as shown in Fig. 9A, in both zoned and unzoned augite underscores the importance of coupled substitution of cations such as Ti, Zr, H f into the M ~site of pyroxene, which charge balance al in the T site. The exchange reaction may be written as: CaMgSi206 +ZrO2~L~ +A1203~L) CaZrA12 O6 + 2SiO2~n) +MgOin I

( 1)

Nb also correlates with Xkv and this may reflect a coupled substitution where Nb enters the M~ site and Na enters the M2 site to charge balance A1 in the T site (Fig. 9B ). The exchange reaction can be written: CaMgSi206 +Nb2Os~L. ) +A1203~LI +Na203~cl ~2NaNbAI206 +CaO

(2) Positive correlations between Nb, Zr and A1 content are evident in the natural basalt experiments. One means of testing whether reactions such as ( 1 ) or (2) represent a realistic representation of the exchange equilibria, is to see if the compositional dependence of D vs. temperature is reduced by recasting in terms of an exchange equilibrium constant. However, there are two limitations that hinder our progress in this regard. First of all, the majority of our natural basalt runs are clearly out of equilibrium, in that the crystals are both sector zoned, and in the case of TS139-HF-24, zoned on the scale of individual sectors. Furthermore, the uncertainties associated with trace-element measurements result in propagated errors that are too large to permit meaningful evaluation of the calculation (Ray et al., 1983).

10. Conclusions Multi-element partition coefficients for clinopyroxene and melt (HFSE, REE, Sr, Cr and V) are reported over a temperature range of ( 12351300°C) and pressure range (1-2.5 GPa) in

127-147

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natural basalts including nephelinite, alkaline basalt, picrite and high-alumina basalt. Equilibrium partition coefficients were measured for an alkaline basalt at 1 GPa and 1250°C, a picrite at 1 GPa and 1300°C, and a nephelinite at 2.5 GPa and 1270°C. The importance of crystal chemical controls on trace-element partitioning in clinopyroxene are demonstrated by non-equilibrium experiments which produced sector-zoned augite, and permit the separation of crystal chemical controls independent of temperature, pressure and bulk composition. D-values for the HFSE, REE, Cr and V show a dependence on the content of tetrahedrally coordinated A1 in clinopyroxene. This reflects the importance of coupled substitution of highly charged cations into M sites and A1 into T sites of clinopyroxene. These results indicate that parameterized mantle melting models must consider the influence of crystal chemical controls on the partitioning of trace elements between silicate melts and clinopyroxene.

Acknowledgements T.S. was supported by a Natural Sciences and Engineering Research Council Postdoctoral Fellowship. He is grateful to Don Francis and Don Baker for providing lava samples, John Ayers and Janet Manchester for their help in the experimental lab, and to David Wark for helping with the EMP analyses. Funding for the use of the SXRF was provided by the National Synchrotron Light Source and Mark Rivers and Steve Sutton are thanked for all of their help. We thank Nobu Shimizu for his help with the ion microprobe and for stimulating discussions on traceelement partitioning. Research at Rennselaer Polytechnic Institute was funded by an NSF grant (EAR-91-05055) to E.B.W. This paper has benefitted from thoughtful reviews by Trevor Green, Simon Kohn and Stephen Foley. The first author thanks the Geological Survey of Canada for providing the necessary resources to complete this project.

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