Oxygen fugacity of the diamond + CO fluid assemblage and CO2 fugacity at 8 GPa

EPSL ELSEVIER

Earth and Planetary Science Letters 128 (1994) 439-451

Oxygen fugacity of the diamond + C - O fluid assemblage and CO 2 fugacity at 8 GPa Tom LaTourrette a,1, John R. Holloway b b

a Department of Geology, Arizona State University, Tempe A Z 85287-1404 USA Department of Chemzstry, Artzona State University, Tempe A Z 85287-1404 USA •

•

Received 2 May 1994; accepted 13 October 1994

Abstract

We have bracketed the oxygen fugacity ( f O 2) of the diamond + C - O fluid buffer (CCO) relative to the wiistite-magnetite (WM) and nickel-nickel oxide (NNO) buffers at 8 GPa and 950-1550°C using a Walker-style multi-anvil press. The intersection of CCO with WM is between 1050° and 1150°C and thus the logfO 2 of CCO at 1100°C is constrained to be - 5 . 1 0 _+ 0.59. From 1100° to 1550°C CCO is between NNO and WM, and below 1100°C CCO is more oxidized than both NNO and WM. Although the intersection of CCO and NNO was not located, previous studies indicate that CCO has shallower slope than NNO and with this constraint the logfO 2 of CCO at 1550°C is - 1.76 _+ 0.95. The f O 2 of CCO at 8 GPa and 950-1550°C can be expressed as logfO 2 = 8.4 _+ 0.8 - 18570 + 7000/T(°K) and is consistent with other recent experiments. The results of this study define the f O 2 of experiments conducted in graphite capsules and regions of the mantle saturated with diamond and a C - O fluid. CCO is more oxidizing than diamond-carbonate buffers and thus the existence of CO2-rich fluids in natural samples at ~ 8 GPa probably requires an olivine-free local environment such as eclogite. CCO lies in the reduced half of the range of estimated mantle f O 2 values and thus diamond will be stable only in the more reduced regions. CO 2 fugacities estimated from these results are at the low end of the range predicted by equations of state at low temperature but show a greater thermal expansion for CO 2 and, hence, a greater increase in f C O 2 with temperature than the equations of state. This results in lower predicted decarbonation reaction temperatures near 8 GPa compared to existing equations of state.

1. Introduction T h e oxygen fugacity ( f O 2) in the E a r t h ' s m a n tle influences (or in some cases is i n f l u e n c e d by) several i m p o r t a n t p r o p e r t i e s a n d process including: extraction of the core, crust and a t m o s p h e r e ;

1 Present address: Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, CA 91125, USA.

the stability of d i a m o n d s , fluids, c a r b o n a t e s a n d hydrous minerals; the m e l t i n g t e m p e r a t u r e a n d melt structure a n d rheology; and, possibly, trace e l e m e n t mobility. I n particular, the speciation of c a r b o n c o m p o u n d s and, therefore, the stability of e l e m e n t a l c a r b o n (graphite or d i a m o n d ) or fluids d e p e n d s strongly o n the a m b i e n t f O 2. U n d e r r e d u c i n g conditions, graphite or d i a m o n d is stable, while u n d e r m o r e oxidizing c o n d i t i o n s carb o n will be p r e s e n t as fluid or c a r b o n a t e . T h e

0012-821X/94/$07.00 © 1994 Elsevier Science B.V. All rights reserved SSDI 0 0 1 2 - 8 2 1 X ( 9 4 ) 0 0 2 2 1 - 5

440

T. LaTourrette, J.R. Holloway / Earth and Planetary Science Letters 128 (1994) 439-451

most oxidizing conditions under which graphite or diamond is stable may be expressed as:

xC+Oz=(Z-x)CO2+(Zx-2)CO

(1)

At a fixed t e m p e r a t u r e and pressure, the coexistence of graphite or diamond with a C - O fluid therefore defines the f O 2 of the system. Reaction (1), combined with decarbonation reactions, also determines conditions of fluid stability. The relevance of this equilibrium is further illustrated by the recent report of CO 2 inclusions in diamond [1]. The f O 2 of this assemblage is constrained by reaction (1), demonstrating that the f O 2 of the diamond + C - O fluid assemblage is applicable to portions of the Earth's mantle. In addition, reaction (1) determines the f O 2 of C - O fluid saturated high p r e s s u r e - t e m p e r a t u r e experiments run in graphite capsules and thus has obvious interest in experimental petrology. The f O 2 defined by reaction (1) can also be used to estimate the fugacity of CO 2. Present equations of state for CO 2 are based largely on data for pressures below about 4 and above 10 GPa. These results provide new data at 8 GPa. Accurate knowledge of CO 2 fugacity is critical to predicting decarbonation reaction boundaries at high pressure. It is also necessary for quantifying CO 2 solubility m e a s u r e m e n t s in silicate melts. For example, to understand the effect of the partial pressure of CO 2 in the vapor on the solubility of CO 2, melt concentrations must be compared to CO 2 fugacity. Also, a good equation of state is required for estimating e n t r a p m e n t pressures of CO2-rich fluid inclusions from their measured densities. In order to help understand the oxidation-reduction behavior of the carbon-oxygen system and CO 2 fugacity under mantle conditions, we have experimentally bracketed the position of the f O 2 of the diamond + C - O fluid assemblage at 8 G P a from 950 ° to 1550°C.

2. Procedure

A technique for determining the f O 2 of graphite or diamond coexisting with a C - O fluid (the CCO equilibrium) previously developed by

Thompson and Kushiro [2], Ulmer and Luth [3] and Fei et al. [4] was adopted for this study. The basic procedure entails bracketing the position of the f O 2 of CCO relative to that of known oxygen buffers in t e m p e r a t u r e - f O 2 space. This is done by running samples that act as redox sensors in experiments that are buffered by CCO; the fate of the sensor then determines the position of CCO relative to that of the sensor. In this study Fe304 and NiO were chosen as sensors because their redox boundaries (the wiistite-magnetite (WM) and nickel-nickel oxide (NNO) oxygen buffers, respectively) are well known and lie near CCO at the conditions of our experiments. The CCO curve has a shallower slope than those for wiistite-magnetite and nickel-nickel oxide and thus crosses these redox reactions at some temperature. The goal of the experiments is therefore to run sensors at various temperatures, at a constant pressure, until the point at which CCO and the metal oxide buffers cross is bracketed by the occurrence of the reduced and oxidized phase of the sensors. The procedure can be expressed in terms of the intersection of reactions (1) and (2), which represent the CCO and W M buffers, respectively, which results in reaction (3):

xC+Oz=(Z-x)CO2+(Zx-2)CO

(1)

2Fe304 = 6FeO + 0 2

(2)

2Fe304 + x C = 6FeO + ( 2 - x ) C O

2 + (2x-

2)CO

(3)

Thus, if Fe304 is stable, the intersection reaction proceeds to the left and the f O 2 of CCO is therefore greater than that of WM. If F e O (FeO is used throughout this paper to represent wiistite, FexO, where x < 1) is stable, the intersection reaction proceeds to the right and CCO is more reduced than WM. Analogous reactions can be written for the N N O buffer. In order for this procedure to be useful, the positions of the wiistite-magnetite and nickelnickel oxide oxygen buffers must be known well. These two systems have recently been carefully calibrated at 0.1 MPa (1 atmosphere) pressure by O'Neill and co-workers [5,6]. These authors state an accuracy of better than _+0.015 in l o g f O 2 ,

T. LaTourrette, J.R. Holloway / Earth and Planetary Science Letters 128 (1994) 439-45l

which is much less than the other errors introduced in the present measurements. The change in molar volume of the metals and oxides with pressure was calculated with a B i r c h - M u r n a g h a n equation of state [7]. Molar volumes for 0.1 MPa pressure and 298°K are from Robie et al. [8] and thermal expansions, bulk moduli and pressure derivatives of the bulk moduli are from Fei et al.

[4]. While there is some discrepancy between the static and dynamic measurements of the bulk modulus of wiistite, the change in f O 2 of the wiistite-magnetite buffer resulting from the measured range of bulk moduti (150-180 G P a [9]) is less than 0.2 log units. The effect of changing wiistite stoichiometry on the f O 2 of this buffer is less obvious because there are no data on the composition of wiistite coexisting with magnetite at high pressure. It is unlikely that the wiistite compositions in our quenched charges are representative of those at run conditions, as wiistite tends to change composition during quenching [3,10,11]. Therefore, we have estimated the composition of wfistite in equilibrium with magnetite at 8 G P a from the internally consistent data base of Fei and Saxena [9]. We then used their nonideal solution model to calculate the activity of stoichiometric F e O in wiistite at each tempera-

441

ture for our experiments relative to those for the 0.1 MPa wiistite compositions of Darken and Gurry [10]. Assuming that the wiistite compositions in the calibration of O'Neill [5] are consistent with those of Darken and Gurry (both measured at 0.1 MPa), this factor will compensate for the change in the f O 2 of the wiistite-magnetite buffer (reaction 2) with pressure. This correction ranges from + 0.6 to + 1.0 l o g f O 2 units over the temperature range of this study and we have assigned errors on this correction equal to half of these values.

3. Experimental and analytical methods All experiments were run in a Walker-style multi-anvil press [12]. The sample assembly is compressed within the octahedral cavity formed by the truncated corners of 8 tungsten carbide cubes (either 8 or 12 m m edge length). The sample assembly is shown schematically in Fig. 1. The sample is mixed with graphite and placed in a graphite capsule. In an effort to minimize the effect of thermal gradients, the sample was kept less than 0.5 m m in length (Fig. 1). This capsule, in turn, is inserted into a platinum capsule which is sealed by welding. The capsule is centered in a

A

Pt

T

3mm

Graphite

........

.....

J

-3_0_~mm _

~ - - - 3.8 mm - - ' - ~

Fig. 1. Schematic cross section of sample assembly.Left: sample capsule. Note that sample height is kept thin to minimize thermal gradients. Right: complete assembly,with capsule in center. MgO octahedron transfers pressure from 8 tungsten carbide cubes. Re furnaces are thicker around capsule to minimize thermal gradients.

442

T. LaTourrette, J.R. Holloway / Earth and Planetary Seience Letters' 128 (1994) 439-451

hole drilled through an MgO octahedron, which acts as the pressure medium in the press. The areas above, below and around the capsule are filled with crushable alumina spacers. A P t Ptg0Rh m thermocouple is inserted longitudinally so that the bead rests on the top of the capsule. The furnaces used in this study were made of either rhenium or inconel metal. Rhenium furnaces are prone to large thermal gradients, so, in an attempt to minimize thermal gradients, stepped furnaces were made by doubling the thickness of the furnace wall around the capsule. Although hydrogen fugacity was not buffered in these experiments, great care was taken to keep H 2 contents as low as possible. Ceramic parts were fired at 1000°C, capsules were kept at 120°C for at least 12 h before welding and the assembled octahedron was kept at 120°C for at least 12 h before the experiment. In addition, reconnaissance experiments using the same technique with CoO + graphite at 2.5 GPa in a piston cylinder match we[1 with the results of Ulmer and Luth [3], who minimized f H 2 with a hematite-magnetite buffer. While the sample assembly for the piston cylinder differs from that for the multianvil, the graphite furnaces used in the former are likely to co?Itain more water and, hence, pose a greater potential for H 2 contamination than the metal multi-anvil furnaces. Pressure was calibrated at 25°C with the Bi I - I I and Bi I I 1 - V transitions, at 1000°C with the CaGeO~ garnet-perovskite transition, and at 1200°C with the coesite-stishovite transition. Pressure uncertainty is estimated to be + 5 % . Details of the pressure calibrations are given by Pawley [13]. The temperature distribution across the capsule region was measured by running an EnsoDis~ ~ oxide mix with the same capsule assembly materials (Re furnace) and dimensions as for the CCO experiments. Compositions of the pyroxenes from the run were measured with an electron microprobe and appear to be well equilibrated, with Kt~ ((1 - CaCVX)/(1 - Ca°P×)) varying smoothly (0.154-0.220) across the sample. The compositions were converted to temperature using the calibration of Nickel et al. [14] and the resulting temperature distribution, shown in Fig. 2, indicates a temperature gradient of 100°C

Then

± 1380-1415 ° . 0.5 mm

1345-1380° 1310-1345°.

3--

Fig. 2. T e m p e r a t u r e d i s t r i b u t i o n in a s a m p l e as r e c o r d e d by the c o m p o s i t i o n s of coexisting pyroxenes. A s s e m b l y m a t e r i a l s and g e o m e t r y are the s a m e as for C C O e x p e r i m e n t s . N o m i n a l t h e r m o c o u p l e t e m p e r a t u r e was 1300°C: p r e s s u r e c o r r e c t i o n is a p p r o x i m a t e l y 25°C. R e c o r d e d t e m p e r a t u r e s indicate an unc e r t a i n t y of _+50°C.

across the sample. Similar experiments using inconel furnaces give smaller gradients [Alison Paw[ey, pers. commun.], so the maximum temperature uncertainty in our experiments is +50°C. The mean sample temperature determined from the 2 pyroxene thermometer is 50°C greater than the measured thermocouple temperature. About 25 ° of this difference is due to the effect of pressure on the thermocouple emf [15]. The resuiting sample temperature is Tsa,npl e = ( T n..... + 50 o) + 50 o. Most runs were conducted with the oxidized phase (Fe~O 4 or a mixture of NiCO 3 + NiO) plus graphite as the starting material. The composition of the Fe304 + graphite starting mixture was confirmed with X-ray diffraction analysis. NiCO 3 was hydrothermally synthesized from NiCO 3 • xH20 at 400°C and 2700 bars for 47 h [16]. X-ray analysis showed that the resulting compound was a mixture of NiCO 3 and a lesser amount of NiO. In runs with the NiCO 3 + NiO + graphite mixture, the NiCO 3 breaks down to NiO + CO 2, and samples were saturated with fluid, while in the case of the Fe304 + graphite, fluid formed upon reduction of Fe304 to FeO (reaction 3). In addition, a reversal run was made with the nickelnickel oxide system: a mixture of Ni metal + graphite was place in a graphite capsule, then this capsule, along with some PdO, was sealed inside a platinum capsule. During heating, the Pd al-

T LaTourrette, J.R. Holloway/Earth and Planetary Science Letters 128 (1994) 439-451 Table 1 Experimental

Run

d e t a i l s (all r u n s at 8 G P a p r e s s u r e )

T (°C)

t (hr)

Slatting material(l)

Run productS(l)

FeOC1

1450

1

mt + gph

wii, mt, sid, gph, diam

FeOC3

1250

3

mt + gph

wti, mt, sid, gph, diam

FeOC4

1150

6

mt + gph

wii, mt, sid, gph, diam

FeOC7

1050

16

mt + gph

mt, gph, diam

FeOC6

950

17

mt + gph

mt, gph, diana

NiCOC10

1550

1

nc + gph

no, nc, gph, diam

NiCPdO 1

1250

3

n + gph+ PdO

n, no, nc, gph, diam

1 mt=

Fe304;

= diamond;

wii = FeO; sid=

FeCO3;

gph = graphite; diam

nc = NiCO3; no = NiO; n = Ni.

loyed with the Pt, releasing oxygen necessary for oxidation of the Ni. In addition, we attempted to use the R e - R e O 2 buffer, which has recently been calibrated by Pownceby and O'Neill [17]. This was unsuccessful, however, because of the formation of ReC. Samples were pressurized at room temperature over 4 - 6 h, then heated to run temperature in 10-15 min. Run times varied from 1 to 17 h, depending on the temperature. Runs were quenched by shutting off the power to the furnace and pressure was released over 6 - 8 h. Run products were examined optically as polished sections and grain mounts, and with X-ray diffraction. Experimental details are listed in Table 1.

4. Results

The experimental results are summarized in Table 1. In experiments where a phase change occurred, the reactions did not go to completion. The degree of reaction increases with increasing temperature and we interpret this as the result of early formed reaction products armoring the reactant phase and forming a kinetic barrier to further reaction. Thus, in forward experiments, the reduced phase is taken as the stable phase when it occurs. In the case of the F e - C - O system, this corresponds to the formation of F e O at temperatures above 1050°C. In the N i - C - O system, the temperature at which NiO is reduced to Ni is unknown; in the highest temperature experi-

443

ment (1650°C) the run failed from a cube failure (blow-out) within seconds of reaching temperature. The sample recovered was highly deformed and X-ray analysis shows peaks for graphite and platinum. It appears that the sample melted. While the melting may have occurred during the transient pressure and temperature changes during the blow-out, it is possible that it resulted from the eutectic melting of Ni + C (1326°C at 0.1 MPa [18]). In this case the NiO would have to have been reduced and thus Ni would be the stable phase. This interpretation is tenuous, however, and is not used to constrain the results. The reversal run in the N i - C - O system formed a small amount of NiO (experiment NiCPdOI), confirming that NiO is the stable phase at 1250°C. Significant solubility of C in Ni metal would increase the apparent f O 2 of the N N O buffer and, therefore, introduce an uncertainty in bracketing the position of the CCO buffer relative to NNO. At 1326°C and 0.1 MPa, the solubility of C in Ni metal is only 0.6 wt% [18], although this amount will increase with pressure. In any case, as the reversal run was at least 300°C and 2 l o g f O 2 units below the intersection of CCO and NNO, the solubility of C in Ni metal has no effect on our results. The experiments were all run in the diamond stability field and, while varying amounts of diamond were present in all runs (easily identified optically and during polishing as extremely hard regions embedded within the graphite capsule), metastable graphite was always much more abundant. This has an insignificant effect on the results, however, because the difference in f O 2 (and derived f C O 2) between graphite and diamond is < 0.2 log units. The samples were too small to produce a measurable amount of fluid ( < 0.0006 g would be formed upon complete reaction of Fe304 to FeO), so the presence of fluid must be assumed. As stated above, fluid was always present in the N i - C - O experiments, while in the F e - C - O experiments fluid formed upon reduction of Fe304 to FeO. Calculated mole fractions of CO in the fluid in these experiments are quite small. While the exact value depends upon the equation of state used, at 1500°C, where the relative fraction of CO is highest, estimates range

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T. La Tourrette, Z R. Holloway / Earth and Planetary Science Letters 128 (1994) 439-451

from 0.0001 to 0.034 for the equations of state of Holloway [19] and Saxena and Fei [20], respectively. Thus, the fluid is estimated to be nearly pure CO 2. Carbonates are present in several of the run products in both the F e - C - O and N i - C - O experiments (Table 1). If stable, carbonates would form at the expense of a C - O fluid and, hence, the experiments would not record the true intersections of CCO with WM or NNO. We believe that these carbonates are not stable but rather form during the quench. The reasons for this are: (1) In the F e - C - O experiments, the amount of F e C O 3 present increases with increasing temperature, which is the opposite of what is expected, since carbonates are favored by low temperature. If carbonate formation ceases due to slowing reaction rates at a fixed blocking temperature, experiments quenched from higher temperatures will be cooling more slowly near this t e m p e r a t u r e and, hence, spend more time in the carbonate stability field than experiments quenched from lower temperatures. (2) The ratio of N i O / N i C O 3 in the 1550°C run increased significantly from that of the starting materiaI~. (3) The carbonates are extremely fine grained ( < 1 /zm). The stabilities of F e C O 3 and NiCO 3 have been measured at low pressure ( P < 1 G P a [16,21,22]), and both phases dissociate below about 500°C in this region. Extrapolations to 8 G P a were unsuccessful, however, because of absent or poor thermodynamic data. A remaining possibility is that the carbonates were dissolved in the fluid or a melt phase. The solubility of Fe- or Ni-carbonate in CO 2 at 8 G P a is unknown and, hence, the possibility of dissolved carbonate cannot be ruled out. We found no evidence for melting (such as crystal-liquid segregation) in any of the runs. In addition, based on Ulmer and Luth's [3] falling sphere experiment, which precludes melting at 1180°C and 2 GPa, we infer that no melting took place in our experiments at similar temperatures and 8 GPa. The experimental results are plotted in terms of 104/T against l o g f O 2 in Fig. 3. The experi-

2

1600

1400 1300

i

q

T (°C) llO0

1200

i

t

1000

J

900

i

0

i

8 GPa

-2

-8

-10 -12

i

I

5.5

i

I

6

i

I

6.5

i

[

L

I

7 7.5 104/T(°K)

i

I

8

i

I

i

8.5

Fig. 3. Plot of l o g m f O 2 against 1 0 4 / T ( ° K ) showing experimental results. Each experiment is shown as a parallelogram on the wiistite-magnetite (WM) or nickel-nickel oxide (NNO) buffer curve. The length and width of the parallelogram represent the temperature and l o g f O 2 uncertainties, respectively. The point at 1400°C is from a similar study by Fei et al. [4]. Open parallelograms are for experiments in which the oxidized phase ( F e 3 0 4 or NiO) was stable and shaded parallelograms are for experiments in which the reduced phase (FeO or Ni) was stable. The N N O point at 1250°C is a reversal experiment (see text). The f O 2 of C C O must lie between the open and shaded points.

mental points are shown as parallelograms, where the length represents the t e m p e r a t u r e uncertainty in the experiment and the width represents the uncertainty in the position of the WM or N N O buffer, which is due to uncertainties in the experimental pressure and the wfistite composition. The N N O point at 1400°C is from the work of Fei et al. [4]. O p e n parallelograms indicate runs in which the oxidized phase was stable, and shaded parallelograms indicate runs in which the reduced phase was stable. The CCO buffer curve must therefore lie between the open and shaded parallelograms. As shown in Fig. 3, the intersection of CCO with WM is located between 1050 ° and 1150°C and thus the l o g f O 2 of CCO at ll00°C is constrained to be - 5 . 1 0 _+ 0.59. From 1100 ° to 1550°C CCO is constrained to lie between N N O and WM, and below 1100°C CCO is more oxidized than both N N O and WM. Although the intersection of CCO and N N O was not located (see explanation above), reasonable constraints can be placed on the f O 2 of

T. La Tourrette, ZR. Holloway / Earth and Planetary Science Letters 128 (1994) 439-451

445

tainty in l o g f O 2 for the two points discussed above. 0

8 GPa

-2

cf

5. Discussion -4

o

5.1. Oxygen fugacity

-6 -8

-10 900

~

Equation (4) Bracketed re

1000 1100 1200 1300 1400 1500 1600 T (°C)

Fig. 4. Plot of l o g f O 2 against temperature for the diamond + C - O fluid assemblage (CCO) determined in this study. Both the value calculated from Eq. (4), with errors as discussed in the text, and the bracketed region from Fig. 3 are shown.

CCO at temperatures up to 1550°C. The CCO buffer must lie between the 1450°C W M run and the 1550°C N N O run. However, this bracket can be reduced by considering the limitations on the slope of CCO. Because the CCO buffer curve has a shallower slope than that of WM, and also has a shallower slope than N N O at lower pressure [3], we can assume CCO has shallower slope than N N O at 8 GPa as well. In this case the maximum f O 2 of CCO at 1550°C is defined by a projection parallel to N N O from the intersection of CCO with W M at 1100°C. Combined with the experimental point at 1550°C, this constrains the l o g f O 2 of CCO at 1550°C to be - 1 . 7 6 _+ 0.95, which is closer to N N O than WM. This bias towards N N O is supported by the results of the failed 1650°C run which, as noted above, suggest that Ni metal may have formed, putting the C C O - N N O intersection between 1550 ° and 1650°C. Since f O 2 buffer curves are linear in 1 / T l o g f O 2 space, these results can be used to derive a linear expression for the f O 2 of CCO at 8 GPa. The resulting equation is: 18570 _+ 7000 l o g f O 2 = 8.4 +_ 0.8

T(°K)

(4)

This curve is plotted in Fig. 4, along with the f O 2 region bracketed in the experiments. Error bars at 1100 ° and 1550°C represent the uncer-

Our determination of the f O 2 of CCO is compared to other estimates in Fig. 5. Three different equations of state for CO 2 [23-25] were combined with thermodynamic data for CO2, 0 2 and diamond [26] to derive estimates of the f O 2 of the CCO buffer. These are shown by the dashed curves in Fig. 5. The curve determined in this work is consistent with the lowest calculated f O 2 values at low temperature, and has a noticeably steeper slope than the calculated curves. Our results are roughly consistent with those of Belonoshko and Saxena [23] above approximately 1100°C. Also shown in Fig. 5 is an extrapolation of the recent experimental measurements of Yasuda and Fujii ([27], extrapolated from data at 3, 5.3, and 7 GPa), who utilized a solid electrolyte oxygen sensor. The position of their CCO f O 2 buffer curve agrees reasonably well with the curve determined

,I1....°/ i ....i -2

-6 k ~ / / ' 7

o

[ ~ .15 / -7 ~ " / ~ ' ~ -8

i

900

,

i

YF93

I

-- -- -BSgl . . . . . ~I-B91 i

i

i

i

i

I I i

1000 1100 1200 1300 1400 1500 1600 T (°C)

Fig. 5. Comparison of recent estimates of the f O 2 of CCO. Dashed curves are calculated from equations of state for CO z (BS91 = Belonoshko and Saxena, [23]; MB91 = M~ider and Berman [24]; SPIP = Sterner and Pitzer, [25]). Diamonds show an extrapolation of recent experiments by Yasuda and Fujii [27]. Note the steeper slope of the experimental results relative to the calculations.

T. LaTourrette, J.R. Holloway / Earth and Planetary Science Letters 128 (1994) 439-451

446

in this work. Most notably, the slopes of the 2 curves are similar, resulting in consistent deviation from a given equation of state with changing temperature. Our curve is about 0.5 log units more oxidizing than Yasuda and Fujii's, which is within the experimental uncertainties. While they do not report their uncertainty, we have assigned an error equivalent to the error in the position of the W M buffer, since they report their f O 2 relative to WM. Graphite is a very popular capsule material for high pressure, high t e m p e r a t u r e experiments because it prevents iron loss to metal capsules and places an u p p e r limit on the f O 2. Our experimental results can be used to estimate the f O 2 of experiments conducted in graphite capsules. For C - O fluid-saturated conditions, the f O 2 is buffered by CCO and Eq. (4) describes the f O 2 at 8 GPa from 1000 ° to 1550°C. For experiments in which the fluid contains additional components or experiments in which fluid is absent, such as melting, the C - O fluid fugacity is lower and Eq. (4) represents an u p p e r limit in f O 2 [28]. As shown in Fig. 6, the f O 2 of CCO lies within the field of estimated mantle f O 2 values and, therefore, experiments conducted in graphite capsules at these pressures are applicable to igneous petrology. The f O 2 of CCO is relevant to studies of the

-2 -4 -6 -8

-10 -12

i

900

, 1 " "

i

I

~

I

h

I

L

I

I

i

1 0 0 0 1 1 0 0 1 2 0 0 1 3 0 0 1 4 0 0 1 5 0 0 1600 T (°C)

Fig. 6. Results for the f O 2 of CCO from this work compared to various f O 2 values relevant to the mantle. I W = ironwiistite, EMOD = enstatite + magnesite = olivine + diamond buffer (Eq. 5), and field labelled Mantle Peridotites is from measurements of Fe 3+ in natural mantle samples.

f O 2 of the mantle because regions containing diamond or graphite and C - O fluids in equilibrium are constrained to lie on the C C O buffer curve. CO2-rich fluid inclusions are quite common in mantle minerals (e.g., [29-31]), although the presence of graphite or diamond is less so. Recently, however, Schrauder and Navon [1] have reported CO 2 micro-inclusions in natural diamond. Based on shifts in infrared absorption peaks for C O 2, they estimate a trapping pressure of 7 to 8.5 GPa. If an equilibrium assemblage, this sample confirms the presence of a free fluid phase near 8 GPa and its coexistence with diamond puts the f O 2 of this sample on the C C O buffer curve. While in some cases the relationship between diamonds and their inclusion assemblages can be complex [32], these samples suggest that at least local portions of the mantle lie on the CCO. Schrauder and Navon [1] note that the existence of a fluid at 8 G P a is surprizing, because CO 2 should react with forsteritic olivine to form carbonates. Our results confirm that this is a problem. As shown in Fig. 6, our determination of C C O lies above the f O 2 defined by the E M O D reaction: MgSiO 3 + MgCO 3 = Mg2SiO 4 + C + 0 2

(5)

enstatite magnesite olivine diamond = E M O D as recently measured by Wei and Luth [33]. Fluid will be stable only at a t e m p e r a t u r e above the intersection of CCO and E M O D . Given the variation in f O 2 of this and related d i a m o n d carbonate buffers, due to solid solutions with Fe and Ca [34,35], the exact temperature at which these buffers can intersect CCO will vary, but it is always likely to be quite high. Along a continental shield geotherm, the temperature at 8 GPa ( ~ 250 km) is about 1375°C [36], which is well below the C C O - E M O D intersection inferred in Fig. 6. Thus, it is unlikely that this intersection is often reached, and carbonate should normally be stable. In addition, natural samples will probably begin melting at a temperature below the C C O E M O D intersection, in which case volatiles would dissolve in the melt and fluid would not be stable. Experiments by Canil and Scarfe [37] indicate

7] LaTourrette, J.R. Holloway / Earth and Planetary Science Letters 128 (1994) 439 451

that the peridotite + CO 2 solidus at 8 GPa is in fact below the C C O - E M O D intersection. Thus, our results confirm the suggestion of Schrauder and Navon [1] that the presence of a CO2-rich fluid at 8 GPa appears to require an olivine-absent local environment. An evaluation of the stability of diamond in the mantle can be made by comparing the f O 2 of CCO with estimates of the mantle f O 2. The most common indicator of the oxidation state of the mantle is the Fe 3+ content in pcridotite minerals and quenched M O R B glasses. Estimates based on such measurements have been reported by several workers (e.g., [38] and references within, [39-43]), and while there are systematic differences between different calibrations, in all cases the overall spread is quite large. From 1.5 to 9 GPa the estimated mantle f O 2 based on these samples spans a range from about - 3 to + 1.5 log units relative to the f a y a l i t e - m a g n e t i t e - q u a r t z (FMQ) oxygen buffer. This field is superimposed on Fig. 6 and it is apparent that the CCO buffer curve lies in the middle of this range. At f O 2 values above CCO diamond is not stable and this implies that diamond is stable only in the more reduced regions of the mantle. This is consistent with the limited occurrence of diamond-bearing mantle xenoliths. It has also been proposed (e.g., [41,44]) that parts of the mantle have an f O 2 defined by a d i a m o n d - c a r b o n a t e equilibrium such as the E M O D assemblage (reaction 5). This requires the presence of a carbonate mineral, which, while uncommon, may not be as rare as once thought, based on new reports of primary mantle carbonates [45,46]. In addition, experiments by Canil [47] show that carbonates may dissociate even during very rapid depressurization and thus would not be found at the Earth's surface. Diamondbearing, CO2-rich kimberlites lend support to the proposal that parts of the mantle may be buffered by an assemblage similar to E M O D . However, the higher f O 2 values in many mantle samples relative to E M O D and CCO (Fig. 6) and limited distribution of diamonds in mantle samples makes it unlikely that large portions of the mantle have a n f O 2 defined by a d i a m o n d - c a r b o n a t e assemblage.

447

5.2. CO 2 fugacity Assuming that the fluid in these experiments is nearly pure CO 2 (see above), we can use our results to calculate CO 2 fugacity at 8 GPa. Setting x = 1 in reaction (1) and combining with the expression AG ° = - R T l n k , we see that: lnfCO2 = lnfO2 -

( A G ° - fV~dP) RT

(6)

where V~ = the volume of carbon as graphite or diamond. The f O 2 of CCO from this work and thermodynamic data ( A H °, S °, Cp, V, coefficients of thermal expansion and compressibility) from the compilation of Holland and Powell [26] have been used to calculate f C O 2 from Eq. (6) and the results are listed in Table 2. These results are compared with various equations of state for CO 2 in Fig. 7. As was found for the comparison of oxygen fugacities, the CO 2 fugacity calculated from this work is generally consistent with equations of state giving the lowest f C O 2 at low temperature. Although the error on our f C O 2 estimate increases at high temperature, it appears that f C O 2 increases relative to any given equation of state with increasing temperature. The reason for this discrepancy is unclear, but the absence of volumetric or fugacity data for CO 2 in this pressure region (no data between 6 and 10 GPa) may be a

Table 2 C o m p a r i s o n of l o g f C O 2 f r o m this w o r k a n d v a r i o u s e q u a tions of state ~' T /°C)

This work H77

SF87

BS91

SPIP

HPg0

MB91

900

11.85

1000

11.60

14.82 13.03 13.06 12.41 12.02 13.96 14.06 12.38 12.47 11.81 11.62 I3.30

1100

11.39

13.39 11.82 11.93 11.27 11.26 12.72

1200

11.20

12.82

1300

11.04

12.32 10.90 11.10 10.41 10.65 11.79

1400

10.90

11.88 10.53 10.76 10.05 10.39 11.41

1500

10.76

11.49 10.19 10.46

9.73

1600

10.65

11.12

9.44

! 1.33

9.88

11.49 10.81 10.94 12.23

10.19

10.15 11.06 9.92

10.76

~ This w o r k f r o m Eq. (6) in text; H 7 7 = Holloway [19]; SF87 S a x e n a a n d Fei [20]; BS91 = B e l o n o s h k o a n d S a x e n a [23]; S P I P = S t e r n e r and Pitzer [25]; H P 9 0 = H o l l a n d a n d Powell [26]; MB91 = Milder a n d B e r m a n [24].

T. LaTourrette, J.R. Holloway / Earth and Planetary Science Letters 128 (1994) 439-451

448

contributing factor. Another potential cause for the difference is that the fluid in our experiments is not pure CO 2. If a significant amount of H 2 was present in the capsule, the measured f O 2 (and hence the calculated f C O 2) in our experiments would be lower than that for a pure C - O fluid. This effect will be negligible, however, as even the extreme case of all2 onuia= 0.5, the f O 2 and f C O 2 would only be lowered by 0.3 log units. Conversely, a significant amount of CO in the fluid would cause our calculated CO 2 fugacities to be too high. As the proportion of CO increases with increasing temperature, this could potentially explain the increase in f C O 2 relative to the equations of state with increasing temperature. However, the mole fraction of CO at 1500°C is estimated to be < 0.04, which will have a negligible effect. Finally, the presence of carbonates in our experiments, if stable, would have the effect of lowering the measured f O 2 , as there would be no fluid present. For the reasons discussed above, however, we feel that the carbonates are not stable and hence have no effect on the measured f O 2. The difference between our results and the equations of state thus appears to result from a difference between the measured and calculated fugacity of CO 2 at 8 GPa. Our results require that the compressibility of CO 2 is greater than

1015 •14 ~- x. ~. 1u

~

-

-.. 13 I•

] I --

8 GPa ~

I

x . ~. _

This work - SF87 --

- H I I

I ..... I .....

" ~ _

-

.

.0101 10 9 900

--

_

SPIP BS91

....

..... 1000 1100 1200 1300 1400 1500 1600 T (°C)

Fig. 7. C o m p a r i s o n o f e s t i m a t e o f C O 2 f u g a c i t y d e t e r m i n e d in this w o r k ( f r o m E q . 6) w i t h t h o s e o f v a r i o u s e q u a t i o n s o f s t a t e f o r C O 2 ( a b b r e v i a t i o n s as in T a b l e 2). N o t e t h a t o u r results a r e c o n s i s t e n t with t h e e q u a t i o n s giving l o w e r C O 2 f u g a c i t i e s at low t e m p e r a t u r e a n d c r o s s e s to h i g h e r v a l u e s w i t h i n c r e a s ing t e m p e r a t u r e .

Table 3 E s t i m a t e d e q u i l i b r i u m t e m p e r a t u r e f o r r e a c t i o n (7) at 8 G P a

This work HP900)

950°C 1025

SPIP

1075

SF87

1250

BS91

1350

MB91

1700

H77

1875

1 A b b r e v i a t i o n s as in T a b l e 2

predicted at low temperatures and suggest that its thermal expansion is greater than predicted at high pressures. One important implication of such a difference is its effect on calculated mineral equilibria because the temperatures of decarbonation reactions depend strongly on CO 2 fugacity. The reaction: CaMg(CO3) 2 + 2SiO 2 = CaMgSi206 + 2CO 2 dolomite coesite diopside fluid

(7) may be important in determining the stability of carbonates in eclogites [34]. Using the same set of thermodynamic data [48], the equilibrium temperature at 8 G P a for reaction (7) was estimated for the CO 2 fugacities compared in Fig. 7. The results are listed in Table 3. The estimated temperatures span over 900°C, illustrating both the sensitivity of decarbonation reactions to f C O 2 and the difficulty of extrapolating reactions to extreme P - T conditions. The results from this work give the lowest t e m p e r a t u r e for reaction (7), implying that decarbonation reaction temperatures at high pressure may be lower than calculations predict. Recently, however, Luth [pers. commun.] has bracketed this equilibrium between 1450 ° and 1500°C at 6 GPa, which implies that at 8 G P a the t e m p e r a t u r e would be among the highest values in Table 3. This discrepancy suggests that the temperatures in Table 3 are too low and that the thermodynamic data used to calculate them are in error. Note that only a 1% change in the enthalpy of

T. LaTourrette, J.R. Holloway / Earth and Planetary Science Letters 128 (1994) 439-451

formation or a 2% change in the first term of the heat capacity equation for the phases in reaction (7) will result in a 500-600°C increase in the calculated temperature.

449

cal assistance with the multi-anvil press, J. Tyburczy for helpful discussions about wiistite, and R.W. Luth, C. Ballhaus and an anonymous reviewer for thorough reviews. [CL]

6. Summary Our experiments bracketing the f O 2 of the CCO at 8 GPa and 950-1550°C agree within error with the recent experiments of Yasuda and Fujii [27] and show that CCO is fairly reduced relative to existing equations of state, especially at low temperature. C - O fluid-saturated experiments run in graphite capsules will have a n f O 2 given by Eq. (4), which lies within the field of estimated mantle f O 2 values. The results of this study describe the f O 2 for regions of the mantle saturated with diamond and a C - O fluid at 8 GPa, such as the sample recently reported by Schrauder and Navon [1]. The existence of a free, CO2-rich fluid phase at 8 GPa appears to preclude the presence of olivine and fluid is thus favored in eclogitic regions in the mantle. The f O 2 of CCO lies in the reduced half of the range of mantle f O 2 values determined from Fe 3+ in peridotites and M O R B glasses, limiting diamond stability to the more reducing regions. Portions of the mantle may lie along EMOD, although the limited distribution of diamonds suggests that this assemblage is not widespread. The CO 2 fugacity calculated from our experimental results is generally consistent with the equations of state giving the lowest CO 2 fugacities and requires that CO 2 is more compressible at high pressure than most equations predict.

Acknowledgements This study was supported by NSF grant EAR9205061 (JRH). The multi-anvil facility is funded by Arizona State University and NSF grant DMR-8406823 to the Materials Research Group. The X-ray equipment is funded by NSF grant DMR-8406823. We thank A.R. Pawley for techni-

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