Journal of Petrology Advance Access published online on November 6, 2007
Journal of Petrology, doi:10.1093/petrology/egm067
| ||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||
Geobarometry for Peridotites: Experiments in Simple and Natural Systems from 6 to 10 GPa
1Institut Für Mineralogie, J. W. Goethe-Universität, Altenhöferallee 1, D-60438 Frankfurt Am Main, Germany
2Vernadsky Institute Of Geochemistry and Analytical Chemistry, Russian Academy of Sciences, UL. Kosygina 19, Moscow, 119991 Russia
3Institute For Geology Of Ore Deposits, Petrography, Mineralogy and Geochemistry, Russian Academy of Sciences, Staromonetny 35, Moscow, 119017 Russia
Received October 20, 2006; Revised typescript accepted October 9, 2007
 |
ABSTRACT |
|---|
 TOP ABSTRACT INTRODUCTIONEXPERIMENTAL METHODSSIMPLE SYSTEMSNATURAL SYSTEMAPPLICATION TO MINERAL BAROMETRYCONCLUDING REMARKSREFERENCES |
|---|
Experiments with peridotite minerals in simple (MgO–Al2O3–SiO2, CaO–MgO–SiO2 and CaO–MgO–Al2O3–SiO2) and natural systems were conducted at 1300–1500°C and 6–10 GPa using a multi-anvil apparatus. The experiments in simple systems demonstrated consistency with previous lower pressure experiments in belt and piston–cylinder set-ups. The analysis of spatial variations in pyroxene compositions within experimental samples was used to demonstrate that pressure and temperature variations within the samples were less than 0·4 GPa and 50°C. Olivine capsules were used in natural-system experiments with two mineral mixtures: SC1 (olivine + high-Al orthopyroxene + high-Al clinopyroxene + spinel) and J4 (olivine + low-Al orthopyroxene + low-Al clinopyroxene + garnet). The experiments produced olivine + orthopyroxene + garnet ± clinopyroxene assemblages, occasionally with magnesite and carbonate-rich melt. Equilibrium compositions were derived by the analysis of grain rims and evaluation of mineral zoning. They were compared with our previous experiments with the same starting mixtures at 2·8–6·0 GPa and the results from simple systems. The compositions of minerals from experiments with natural mixtures show smooth pressure and temperature dependences up to a pressure of 8 GPa. The experiments at 9 and 10 GPa produced andradite-rich garnets and pyroxene compositions deviating from the trends defined by the lower pressure experiments (e.g. higher Al in orthopyroxene and Ca in clinopyroxene). This discrepancy is attributed to a higher degree of oxidation in the high-pressure experiments and an orthopyroxene–high-P clinopyroxene phase transition at
9 GPa. Based on new and previous results in simple and natural systems, a new version of the Al-in-orthopyroxene barometer is presented. The new barometer adequately reproduces experimental pressures up to 8 GPa. KEY WORDS: garnet; mineral equilibrium; multi-anvil apparatus; orthopyroxene; geobarometry
|
INTRODUCTION |
|---|
|
TOPABSTRACTINTRODUCTIONEXPERIMENTAL METHODSSIMPLE SYSTEMSNATURAL SYSTEMAPPLICATION TO MINERAL BAROMETRYCONCLUDING REMARKSREFERENCES |
|---|
Equilibrium relationships between pyroxenes and garnets are the basis for thermobarometry in peridotitic system; for example, for the two-pyroxene thermometer and the orthopyroxene–garnet barometer. The former is based on the mutual solubility of enstatite and diopside components in coexisting clinopyroxene and orthopyroxene; that is, the reactions
and 
. Orthopyroxene–garnet barometry makes use of the maximum Al solubility in orthopyroxene, which can be described by the reaction (
. These equilibria have been studied experimentally in simple and complex systems since the mid-1970s. Brey & Koehler (1990
) derived widely used formulations of the two-pyroxene thermometer and the Al-in-orthopyroxene barometer on the basis of experiments in simple (Perkins et al., 1981; Gasparik & Newton, 1984; Nickel & Brey, 1984) and complex peridotitic systems (Brey et al., 1990). Most of these experiments are restricted to pressures up to 6 GPa. There is growing evidence that some peridotite xenoliths and inclusions in diamonds were equilibrated at higher pressures (e.g. Menzies et al., 2004; Phillips et al., 2004). The present work was aimed at extending the experimental database to higher pressures and reducing uncertainties associated with extrapolation of empirical barometric expressions beyond the experimentally studied range. The experiments were made feasible through the increased performance and better calibration of multi-anvil apparatuses. Considerable progress has also been made in electron microprobe techniques, which allow the detailed characterization of experimental products using back-scattered electron images, element mapping and exact positioning of the electron beam. The analysis of touching mineral grains and outer zones of newly formed crystals allows better evaluation of the attainment of equilibrium in the experiments.
This study takes advantage of the new techniques and extends experimental geobarometry to pressures higher than 6 GPa. Equilibria amongst mantle minerals were investigated at 6–10 GPa and 1300–1500°C in the simple systems MgO–Al2O3–SiO2 (MAS), CaO–MgO–SiO2 (CMS), and CaO–MgO–Al2O3–SiO2 (CMAS) with the multi-anvil apparatus (Walker-type) of the Institut für Geowissenschaften, Frankfurt University. The simple system experiments were used (1) to check the accuracy of experimental parameters and possible pressure–temperature gradients in the experimental assembly and (2) to obtain high-pressure experimental constraints on the mineral reactions that are employed in peridotite geothermobarometry. Subsequently, experiments were carried out in the same P, T range with natural peridotite mineral mixes available from the study of Brey et al. (1990) with the aim of developing an improved version of the Al-in-orthopyroxene barometer.
|
EXPERIMENTAL METHODS |
|---|
|
TOPABSTRACTINTRODUCTIONEXPERIMENTAL METHODSSIMPLE SYSTEMSNATURAL SYSTEMAPPLICATION TO MINERAL BAROMETRYCONCLUDING REMARKSREFERENCES |
|---|
MAS, CMS and CMAS starting mixtures
Two types of starting mixtures were used: (1) previously synthesized end-members forsterite, enstatite, pyrope, and diopside mixed together in equimolar proportions; (2) glasses of similar bulk compositions prepared from gels. The preposition is that the crystalline phases adjust their composition during the experiment and approach equilibrium from the pure end-member side (i.e. enstatite gains Ca in CMS and Al in MAS experiments), and that devitrification of the glass at the beginning of the experiment yields incipient crystalline phases of intermediate compositions (i.e. high-Ca and high-Al orthopyroxenes in CMS and MAS experiments, respectively). The newly formed metastable compositions react during the experiment and approach equilibrium from the opposite side. Thus, experiments with these two types of starting mixtures should bracket equilibrium pyroxene and garnet compositions.
Starting mixtures with natural minerals
The natural mineral mixtures were those used by Brey et al. (1990) (Table 1). Mixture SC1 consists of olivine [mg-number = Mg/(Mg + Fe) = 0·89], Al-rich orthopyroxene, Al-rich clinopyroxene and spinel from a spinel peridotite. Mixture J4 consists of a similar olivine, Al-poor orthopyroxene, Al-poor clinopyroxene, and garnet from a garnet peridotite. Thus, the starting pyroxenes differed mainly in Al2O3 and the high-Al pyroxenes were oversaturated in Al2O3 at all experimental conditions. The pyroxenes from the two starting mixtures approach equilibrium from different sides, mostly in adjacent charges, which should help to define equilibrium compositions.
|
Experimental procedure
Experiments were carried out in a Walker-type multi-anvil apparatus with WC cubes with 8 mm truncation edges. The pressure assembly consisted of a 95% MgO + 5% Cr2O3 octahedron, a zirconia sleeve, a rhenium foil heater and MgO inserts (Fig. 1). Temperature was measured with a W–Re thermocouple inserted through an Al2O3 ceramic sleeve. The free space was filled with Ciatronics cement. Pressure was calibrated at room temperature (Bi phase transitions) and at high temperatures by S. Buhre and H. Steinberg (two unpublished PhD theses, 2005) using the
Mg2SiO4–ßMg2SiO4 transition at 1200°C and 13·6 GPa (Morishima et al., 1994), the coesite–stishovite transition at 1200°C and 9·2 GPa (Yagi & Akimoto, 1976) and the CaGeO3 garnet–perovskite transition at 1100°C and 6·1 GPa (Susaki et al., 1985). The results of high-temperature calibration experiments and the calibration curve are shown in Fig. 2.
|
|
Starting materials for the MAS, CMS and CMAS experiments were loaded into Pt capsules with 1·6 mm outer diameter and 0·1 mm wall thickness. The capsules were 0·7–0·8 mm in height and contained about 1·0 mg of the starting materials. They were sealed by electric welding and pressed to cylinders to minimize the deformation of the cell assembly during experiments. Usually, two capsules were placed symmetrically in the centre of the furnace, and the two ‘reversal’ starting mixtures for each composition were used simultaneously in one run to ensure similar P–T conditions and a reliable bracketing of equilibrium composition. In a few experiments the two starting mixtures were placed in a single capsule separated by a Pt disc (Fig. 3). Oxygen and water activities were not controlled in the experiments.
|
In the experiments with natural starting minerals, the starting mixtures were loaded into containers (0·7–0·8 mm high, 1·4 mm outer diameter and 0·5 mm inner diameter) machined from single-crystal San Carlos olivine. Two such capsules filled with the different starting materials (0·2–0·3 mg each) were placed in a platinum tube with 0·1 mm wall thickness, which was sealed by electric welding. After pressing the capsule was about 2 mm high and 1·6 mm wide. The capsule was placed in the centre of the furnace. Experimental products were about 0·5–0·7 mm in diameter and 0·2–0·4 mm high with 0·2–0·3 mm distance between the two starting mixes (Fig. 4). Oxygen fugacity was not controlled, and substantial oxidation was observed in some experimental products, especially at 9 and 10 GPa (see below). No attempt was made to eliminate trace amounts of H2O and CO2 from the starting mixes, because small amounts of fluid or melt accelerate recrystallization and equilibration without affecting the relations of anhydrous phases.
|
After the experiment the capsule was mounted in epoxy, cut, polished and analysed with a Jeol Superprobe 8900 electron microprobe equipped with five wavelength-dispersive spectrometers. An accelerating voltage of 15 kV and a beam current of 20 nA were used; counting times varied between 20 and 40 s on peak and background with a focused electron beam. Standards were natural silicates and pure oxides and metals; the ZAF algorithm was used for matrix correction.
|
SIMPLE SYSTEMS |
|---|
|
TOPABSTRACTINTRODUCTIONEXPERIMENTAL METHODSSIMPLE SYSTEMSNATURAL SYSTEMAPPLICATION TO MINERAL BAROMETRYCONCLUDING REMARKSREFERENCES |
|---|
Experimental results
Experimental conditions and composition of run products are given in Table 2. Experiments with crystalline starting materials yielded aggregates of subhedral to euhedral crystals, commonly 5–10 µm and up to 30 µm in size (Fig. 5a). Rim zones of large crystals were preferentially analysed, but there was generally no compositional difference between core and rim of garnet and both pyroxenes. Experiments with glassy starting materials produced much smaller crystals, measuring less than 10 µm (Fig. 5b).
|
|
Because the most reacted mineral zones are those that bracket equilibrium composition, Table 2 shows only high-Ca orthopyroxene and garnet compositions and low-Ca clinopyroxene compositions for experiments with crystalline starting materials, and low-Ca orthopyroxene and garnet and high-Ca clinopyroxene compositions for experiments with glassy starting materials.
Pyrope and enstatite were the run products in MAS experiments at 8 GPa and 1300 and 1400°C. The pyropes from the glassy starting material always yielded an appreciable excess of silica (up to 3·1 Si in the structural formula), whereas almost stoichiometric garnets were obtained from the crystalline starting material. Because of grain-size variations, we assume that the excess silica reflects fluorescence from neighbouring orthopyroxenes rather than a majorite component in garnet. The Al2O3 contents of orthopyroxenes from the glassy and crystalline starting materials overlap, but the overlap intervals are small and comparable with the analytical uncertainty (Fig. 6). The Al2O3 content of orthopyroxene from the 1400°C experiment is about twice as high as that from the 1300°C experiment (Table 2; Fig. 6).
|
Two experiments were performed in CMS at 6 and 8 GPa and 1300°C. Both experiments produced stoichiometric pyroxenes. In terms of their Ca content, there is no path-looping in the 6 GPa experiment and a very minor overstepping at 8 GPa in the pyroxenes of the two types of starting material (Fig. 7). The decrease of diopside solubility in orthopyroxene and the increase of enstatite solubility in clinopyroxene with pressure are consistent with the results from lower pressure experiments (Fig. 7).
|
Garnet and two pyroxenes were present in all CMAS experiments (Table 2). The garnet compositions from the two starting materials always showed a small overstepping interval of less than 0·005 Ca/(Ca + Mg) (Fig. 8). Previous work at lower pressure (Brey et al., 1986
) has shown that this ratio for garnet in equilibrium with orthopyroxene and clinopyroxene is insensitive to temperature and decreases with increasing pressure. The decrease from about 0·11 at 6 GPa to about 0·09 at 8 GPa (Fig. 8) is entirely consistent with these observations.
|
Pyroxene compositions in CMAS can be fully described by the Ca fraction on the M2 site and the Al fraction on the M1 site (= Al/2). The compositions of experimental pyroxenes are rather variable, with distinct correlations between CaM2 and AlM1 but with very different slopes for experiments with glassy and crystalline starting material (Fig. 9). The compositional trends in experiments with the crystalline starting material emanate from pure diopside and enstatite and show a considerable range in Ca coupled with a moderate gain in Al. The pyroxenes from the glassy starting material show a very large range in Al and minor Ca variations (Fig. 9). Exceptions are several high-Ca and low-Al orthopyroxene compositions observed in the 1300°C experiment with the glassy starting material (Fig. 9). The origin of these orthopyroxenes is unclear, but their compositions are obviously far from equilibrium. The two trends converge towards common Al and Ca contents with clear temperature dependence of Ca and only a moderate temperature effect for Al. The pyroxenes from the crystalline and glassy starting materials show mostly overlap in CaM2 except for clinopyroxene at 1500°C and orthopyroxene at 1400°C (Fig. 9). Alumina always shows a gap with the end-points for the glassy starting materials exhibiting stronger P–T dependence than those for the crystalline starting material.
|
Consistency with results from lower pressures
The results from the simple systems are used to evaluate the consistency of the multi-anvil experiments with lower pressure experimental data from the literature. Such a comparison implies a considerable extrapolation in P and T of the previous data with the aid of existing thermodynamic models. We used the models of Gasparik (2000
) and Nickel & Brey (1984), which are in agreement with most of our, and independent, experimental data in MAS and CMS for pressures up to 6 GPa.
The orthopyroxenes in equilibrium with garnet at 8 GPa in MAS are compared in Fig. 6 with the model of Gasparik (2000). Our experimental bracket at 1400°C lies at higher AlM1 and that at 1300°C at lower AlM1 relative to the model. Taken alone, our experimental data suggest a much stronger temperature effect on the solubility of Al in orthopyroxene than the model. This is not consistent with the data of Perkins et al. (1981), and we interpret the differences as resulting from the combined effect of uncertainties in the experimental P–T conditions. The discrepancy with the Gasparik (2000) model corresponds to a pressure difference of about –0·4 GPa at 1400°C and +0·3 GPa at 1300°C (or a combined P–T effect). The earlier model of Gasparik & Newton (1984) strongly underestimates the Al content of orthopyroxene at high pressures (Fig. 6).
Figure 7 compares pyroxene compositions from our CMS experiments at 1300°C with the model of Nickel & Brey (1984). The new results at 6 GPa are in very good agreement with the Nickel & Brey (1984) experiments in a belt apparatus and the model. The result at 8 GPa with an overlap in composition suggests a decrease of Ca in clinopyroxene between 6 and 8 GPa (Fig. 7), which is opposite to the well-established pressure effect from lower pressures (Lindsley & Dixon, 1976). It may also mean that the actual temperature in the experiment was about 50°C higher than the target value (1300°C). However, this is in conflict with the result for orthopyroxene, which is in good agreement with the model. Most probably, the discrepancy is a combined effect of an error in temperature smaller than 50°C and an inadequacy of the Nickel & Brey (1984) model for accurate extrapolation to 8 GPa. As stated above, the garnet compositions are in good agreement with the empirical relationship obtained by Brey et al. (1986) from lower pressure experiments.
The analysis of our data in simple systems shows that there is reasonable agreement with previous work at pressures up to 6 GPa within 0·3 GPa and 50°C. Moreover, the agreement between the new high-pressure experiments and existing thermodynamic models suggests that the composition of phases changes smoothly at least up to 8 GPa. This allows us to use the composition of experimental pyroxenes to evaluate possible P–T gradients in our samples from clinopyroxene–orthopyroxene and orthopyroxene–garnet equilibria.
P–T gradients in the capsules
Temperature gradients in multi-anvil assemblages may be rather high (Watson et al., 2002; van Westrenen et al., 2003; Hernlund et al., 2006) and possible pressure gradients must also be taken into account. Assuming local equilibration, we can use the distribution of the composition of phases within the capsules from our experiments as temperature and pressure sensors.
Mapping of pyroxene compositions in runs 81 (MAS) and 82 (CMS) provides insight into possible pressure and temperature gradients within these Pt capsules (Fig. 2). The Al2O3 content of orthopyroxene shows a gap of 0·1 wt % on the left- and the right-hand side of the capsule in run 81, and the difference between the midpoints of the gaps is 0·06 wt %. Assuming that the half-widths of the gaps can be treated as standard deviations of the average Al2O3 contents in the left and right parts of the capsule, the confidence interval for the difference between them is 0·005–0·11 wt % Al2O3 at the 95% confidence level. The 0·06 wt % difference in the Al2O3 content of orthopyroxene corresponds to a pressure difference of about 0·4 GPa. The difference in Al2O3 is comparable with the accuracy of electron microprobe analysis, and the relations observed in Fig. 2 are also consistent with a zero pressure gradient across the sample. Clinopyroxene in run 82 (Fig. 3) shows a gap of 0·2 wt % CaO on the left- and right-hand side of the capsule, and the confidence interval for the difference between the midpoints is 0·03–0·27 wt % CaO (95% confidence level). This corresponds to a temperature difference of about 15°C and is again within the analytical uncertainty. There is an overstepping in CaO in orthopyroxene of about 0·1 wt % at a negligible difference between the midpoints of the overstepping intervals in the left and right parts of the sample (0·02 ± 0·05 wt % CaO). Taken together, there is no compelling evidence to suppose any significant temperature gradient within the capsule.
The capsules used in simple system experiments were too thin (up to 300 µm) to estimate gradients along the furnace axis. This information is especially important for the experiments in natural systems, when the two samples are separated by an olivine layer of about 0·4 mm thickness (Fig. 4). Mapping within each of the sample layers (about 0·2 mm high and 0·2 mm in diameter) shows no systematic compositional variations, suggesting negligible pressure and temperature gradients. To assess possible differences in P–T conditions between the positions of the two samples, two experiments were performed at nominally identical parameters (6 GPa and 1400°C) with a reversed sample arrangement (J4 above SC1 in run 77 and SC1 above J4 in run 74). The average Al contents in the M1 of orthopyroxene were 0·0140 for SC1 and 0·0135 for J4 in run 74 and 0·0160 for SC1 and 0·0163 for J4 in run 77. The reversal of sample arrangement changed the differences between the orthopyroxene compositions from the two materials by 0·0008, which is equivalent to about 0·1 GPa. However, the difference of Al in M1 between the orthopyroxenes from the same starting materials (SC1 or J4) in runs 74 and 77 is about three times greater. This difference suggests that there is an actual pressure difference of about 0·3 GPa between these two experiments. Therefore, the reproducibility of pressures between runs is a more serious problem than a pressure gradient within a capsule from one run. We estimate that the minimum errors in P and T for repeat experiments lie within 0·3 GPa and 50°C.
Equilibrium composition of phases in simple systems
Having established that P and T gradients are small within a capsule we used the midpoints of either the compositional gaps or the compositional overlaps between the two reversal starting materials as equilibrium compositions. Overlap or path-looping is common in reversal experiments (e.g. Lane & Ganguly, 1980; Perkins & Newton, 1980). If path-looping takes place, the equilibrium composition does not necessarily lie within the overlap region. However, when the overlap region is small, it is usually assumed that the midpoint is a valid approximation for the equilibrium composition (Perkins & Newton, 1980). The interpreted equilibrium values are given in Table 3.
|
|
NATURAL SYSTEM |
|---|
|
TOPABSTRACTINTRODUCTIONEXPERIMENTAL METHODSSIMPLE SYSTEMSNATURAL SYSTEMAPPLICATION TO MINERAL BAROMETRYCONCLUDING REMARKSREFERENCES |
|---|
Experimental results
Run conditions and products are shown in Table 4. The mineral mixtures recrystallized to massive aggregates of euhedral to subhedral grains, 10–30 µm in size (Fig. 10). The presence of small, but unknown amounts of carbonate–silicate melt in the charges promoted the recrystallization process. The melt was found always between the olivine and the Pt capsule (Fig. 4) but must have originated within the charges and migrated to the rims during the run. Its presence was unforeseen and the source of CO2 or carbonate is unclear. It was probably derived from small amounts of absorbed carbon species in starting materials and olivine capsules and from CO2 fluid inclusions. To elucidate this problem we carried out a time study with runs for 30 min, 2 h, 10 h and 23 h at 6 GPa and 1500°C with the same experimental setup but with a small amount of magnesite added to the charges. From the previous experiments we expected olivine, pyroxenes, garnet, magnesite and melt in all runs. However, in the two shorter runs melt did not appear and we obtained a subsolidus assemblage of olivine, pyroxenes, garnet and tiny graphite specs. Melt appeared in the 10 h run between the olivine and Pt capsule and along grain boundaries in the recrystallized olivine capsule. In the 23 h run melt was confined only to interstices between olivine and Pt. Our interpretation is as follows: the initial conditions within the charges were such that carbon species were reduced to graphite. With increasing run time Fe loss occurred from olivine to the Pt walls, which liberated oxygen. This changed conditions towards the EMOG buffer (Eggler & Baker, 1982
) and oxidation of graphite and redox melting occurred. The compositions of melts and their relationships with solid phases will be described in detail in a separate publication (Brey et al., in preparation).
|
|
Because of the variable amounts of melt in the runs, the presence or absence of clinopyroxene and magnesite does not strictly depend on experimental conditions (Table 4). These phases tended to survive at high pressures and low temperatures. Euhedral magnesite occurred both in the peridotitic part and in the space between the olivine and the Pt capsule.
The compositions of minerals were examined in great detail with emphasis on possible zoning in grains and spatial variations within the samples. Most analyses were made near grain rims, but we took care to avoid the influence of adjoining phases. Many of the samples appeared to consist of loose aggregates of grains, which resulted in the isolation of single crystals by thin epoxy films. This allowed us to obtain highly reliable analyses and to document subtle compositional variations. The main features may be summarized as follows.
- Olivines are very homogeneous both within single grains and throughout the sample volume. The range of mg-numbers [= Mg/(Mg + Fe)] of olivine from single samples was usually smaller than 0·004. The mg-number of olivine in the samples is generally similar to that of the adjacent olivine capsule. Because the olivine capsules vary in mg-number from 0·91 to 0·93 their mg-number probably buffers that of the experimental phases.
- Orthopyroxene completely recrystallized, and no relics of initial high-Al pyroxenes were found. There are, however, rather large variations in CaO in most experiments (Fig. 11), and the CaO contents are often considerably lower than in the starting orthopyroxenes, especially in clinopyroxene-free experiments. The appearance of melt is the explanation: once originated it eliminates clinopyroxene and the remaining orthopyroxene and garnet equilibrate at lower CaO contents. In several clinopyroxene-free experiments (runs 62, 59, 60), measurement of touching garnet and orthopyroxene grains revealed two discrete populations with higher and lower Ca contents in the minerals. This is probably indicative of local equilibrium in the respective charges, and the average compositions across both populations were considered further. The orthopyroxene alumina content is almost not affected by the appearance of melt because it is buffered by the coexistence of garnet.
- The amount of clinopyroxene was always very small and only a few analyses could usually be performed. Similar to orthopyroxene, clinopyroxene displays no regular zoning and no relics of starting compositions were found. Its composition may be strongly variable (e.g. run 66) probably as a consequence of incomplete equilibration with orthopyroxene and garnet.
- Garnets from J4 often show concentric zoning patterns with starting compositions surviving in the cores (Fig. 10). In SC1 garnet had to nucleate probably with considerable compositional scatter, because garnets with both Ca-rich and Ca-poor cores were found together with unzoned grains. The core-to-rim zoning often shows an increase in Ca where clinopyroxene coexisted, but more complex patterns were also found (Fig. 12). Reasoning again from the effect of melt appearance and the consequent disappearance of clinopyroxene, garnets in these runs equilibrate at lower CaO because the buffering effect of clinopyroxene vanishes.
|
|
Evaluation of equilibrium composition and P, T variations
‘Equilibrium’ compositions were selected according to the following criteria: (1) the analyses from rim zones or small crystals were preferably used; (2) if the analyses formed compact clusters with minor variations, the mean analysis of this cluster was used; (3) garnets with regular zoning patterns allowed us to track changes in compositions and the direction of approach to equilibrium; the outermost zones were selected; (4) clinopyroxenes from the two starting materials formed compositional fields separated by very narrow gaps (not shown); taking the direction of approach to equilibrium into account, the compositions near the convergence of the two arrays (brackets) were used in both samples. These criteria usually did not allow us to select a unique equilibrium composition, but rather the range in which the equilibrium composition probably lies (Figs 11 and 12). The average of analyses falling within this range was used as a proxy for the equilibrium composition (Table 5).
|
The mg-numbers of experimental phases are variable, which is probably due to the competing influence of the buffering through the olivine capsules and interaction with the carbonate-rich melt. Nevertheless, the ratios of the mg-numbers of olivines and coexisting orthopyroxenes are close to unity and their values are always higher than those for garnets and clinopyroxenes. A distinct correlation between temperature and Fe–Mg distribution is observed for olivine–garnet and orthopyroxene–garnet, consistent with the results of O’Neill & Wood (1979
) and Harley & Green (1982).
Where coexisting with clinopyroxene (runs at 1300°C and 7–10 GPa and at 1400°C and 8 and 10 GPa), the Ca content of orthopyroxene decreases with increasing pressure (Fig. 13). The results are adequately described by the Nickel & Brey (1984) model for CMS up to pressures of 8 GPa (Fig. 13). In contrast, the Ca contents of orthopyroxene from the 9 and 10 GPa experiments are significantly lower than predicted by the Nickel & Brey (1984) model. This difference could be a consequence of the orthopyroxene–high-P clinopyroxene transition at pressures of around 9 GPa (Pacolo & Gasparik, 1990; Angel et al., 1992; Woodland & Angel, 1997). The C2/c high-P clinopyroxene polymorph is unquenchable and transforms to the P21/c structure during decompression (Angel et al., 1992). The Ca concentration in clinopyroxene is much lower than for CMS (Fig. 13) but consistent with iron-bearing systems and with the results of Brey et al. (1990).
|
The concentrations of Al in orthopyroxene are plotted against pressure in Fig. 14 for experiments in natural systems (data from Brey et al., 1990
, and present study) and compared with the model of Gasparik (2000) for MAS. There is good consistency between the results from the natural systems at 6 GPa. Compared with MAS the Al content in natural orthopyroxene is significantly lower at pressures below 7 GPa and almost identical at higher pressures. In contrast to orthopyroxene, the Al content in clinopyroxene is practically independent of pressure (Fig. 14).
|
Figure 15 shows the variation of
= Ca/(Ca + Mg + Fe + Mn) for garnets coexisting with both orthopyroxene and clinopyroxene from the experiments of this work and of Brey et al. (1990) as a function of pressure. For comparison, the temperature-invariant pressure dependence of in CMAS is shown as obtained by Brey et al. (1986). shows a gentler slope compared with CMAS up to 8 GPa and negative temperature dependence in agreement with the equation of Brenker & Brey (1997). In contrast to the gradual decrease with pressure, increases dramatically above 8 GPa, which translates into CaO contents in garnet of more than 6 wt % at 1300°C and 9 and 10 GPa. Simultaneously, Al2O3 decreases in the 1300°C runs to about 17 and 14·5 wt %, respectively, FeO increases to 9 and 12 wt % and MgO slightly decreases to 20·5 and around 19·5 wt %. SiO2 remains largely unchanged or may even slightly decrease in the 9 and 10 GPa runs (Table 5). To maintain charge balance and compensate for the lower Al2O3 contents, a substantial amount of Fe2O3 must be present in the form of an andradite component. Where clinopyroxene is missing garnets become subcalcic and their compositions are consistent with those in the lower pressure experiments (i.e. Al2O3 is in the 20 wt % and FeO in the 7 wt % range). For the 1500°C experiments SiO2 in garnet increases from slightly more than 43 wt % at 7 GPa to about 44 and 44·5 wt % at 8 and 10 GPa, indicating an increasing majorite component.
|
|
APPLICATION TO MINERAL BAROMETRY |
|---|
|
TOPABSTRACTINTRODUCTIONEXPERIMENTAL METHODSSIMPLE SYSTEMSNATURAL SYSTEMAPPLICATION TO MINERAL BAROMETRYCONCLUDING REMARKSREFERENCES |
|---|
Figure 16 compares the experimental pressures of the present study for natural systems with values calculated by four Al-in-orthopyroxene barometers for mantle peridotites. The MacGregor (1974
) barometer overestimates pressure by more than 1 GPa at 6 GPa, and gives reasonable results at 7 and 8 GPa. The barometer of Brey & Koehler (1990) gives systematically higher pressures (by 0·5–1·0 GPa) at 6–8 GPa and reproduces experimental pressures at 9 and 10 GPa. The Nickel & Green (1985) barometer (designed for lherzolitic lithologies) is in agreement with the experimental data at 7 and 8 GPa, but the calculated pressures for the 9 and 10 GPa experiments are lower by more than 1 GPa than the experimental values. The Girnis et al. (1999) barometer underestimates pressure over the whole experimental range. For the following, it is important to bear in mind that an Al-in-orthopyroxene barometer must not reproduce pressures above 8 GPa, where low-Ca clinopyroxene coexists with garnet under the experimental conditions. The higher pressure experiments are also those with the unusually oxidized, high-Ca garnets, and the reason for the inconsistency may also lie there. We therefore take only experiments up to 8 GPa into consideration to redesign the Brey & Koehler (1990) Al-in-orthopyroxene barometer.
|
The Brey & Koehler (1990
) barometer was anchored on the Gasparik & Newton (1984) equation for MAS, which was obtained assuming ideal Mg–Al mixing on the M1 site of orthopyroxene, but taking thermal expansion and compressibility data into account. Their equation was supplemented by additional terms accounting for contributions from Fe, Ca, and Cr from more complex systems. The final adjustment came from the 3–6 GPa experiments of Brey et al. (1990) in natural systems. As can be seen from Fig. 6, the Gasparik & Newton (1984) MAS model failed to reproduce experimental orthopyroxene compositions at 8 GPa, which could be the reason why the Brey & Koehler (1990) barometer overestimates pressure at 6–8 GPa. In contrast, the MAS model of Gasparik (2000) adequately reproduces our experimental results for this system at 8 GPa (Fig. 6). Both the Gasparik & Newton (1984) and Gasparik (2000) models employ the same empirical equations describing the dependence of equilibrium constants of reactions (K) on pressure (P) and temperature (T), RTlnK = a1 + a2T + a3T1·2 + a4P + a5PT + a6P2, where a1 ... a6 are the adjusted parameters.
Given the adequacy of the Gasparik (2000) model for MAS, we adopted the same form of barometric expression as used by Brey & Koehler (1990) to take into account our new experimental data [equations (1)–(4)]. All the coefficients were adjusted to minimize the target function 
(Pcalc – Pexp)2, where Pcalc and Pexp are the calculated and experimental pressures, respectively.
For the fitting procedure we used the 6–8 GPa experiments from this study and the experiments in natural systems at 2·8–6·0 GPa by Brey et al. (1990). The database also included the experiments on MAS and CMAS from this study and from the literature (Perkins & Newton, 1980; Perkins et al., 1981; Nickel et al., 1985; Klemme & O’Neill, 2000), which were used as additional constraints. The model values of RTln K were constrained to lie between RTln Kexp + 2
and RTln Kexp – 2, where Kexp is the experimental value of the equilibrium constant obtained in MAS and CMAS experiments, and is the standard deviation of RTln Kexp propagated from the estimated uncertainties in compositions and experimental parameters. The problem was solved by constrained least-squares optimization. The modified version of the Brey & Koehler (1990) barometer is as follows (the new barometer is implemented in the PTEXL3 MS Excel spreadsheet, which can be downloaded from the Internet site of the Institut für Geowissenschaften, J. W. Goethe-Universitat at http://www.mineralogie.uni-frankfurt.de/index.html):
|
| (1) |
|
| (2) |
|
| (3) |
|
| (4) |
|
|
Site occupancy:
(a) Orthopyroxene
|
|
(b) Garnet
|
|
|
|
) experiments to ± 0·3 GPa (1) (Fig. 17). The uncertainty is slightly higher than the 0·22 GPa obtained by Brey & Koehler (1990), but the new equation reproduces the experiments from piston–cylinder, belt and multi-anvil apparatuses from 2·8 to 8·0 GPa and 900 to 1500°C. The barometer underestimates pressures by 0·5–1·5 GPa for the experiments at 10 and 9 GPa with Ca- and Fe-rich garnets. The discrepancy could hardly be due to the unusually high Fe3+ contents in garnets from these experiments, because the substitution of Fe3+ for Al in garnet must be accompanied by a decrease in Al content in coexisting orthopyroxene, and pressure overestimation by an Al-in-orthopyroxene barometer should be expected. In our opinion, a more reasonable explanation is the effect of the orthopyroxene–high-Ca clinopyroxene transition. This implies that the solubility of Al in high-P C2/c clinopyroxene is higher than in orthopyroxene. This also means that the derived barometer and other existing Al-in-orthopyroxene barometers are not suitable for pressures higher than 8 GPa.
|
To illustrate changes associated with the new calibration, pressures and temperatures were calculated for a suite of peridotite xenoliths from kimberlites containing diamond or graphite compiled by Grütter et al. (2006
) and inclusions in diamonds from the De Beers Pool reported by Phillips et al. (2004). The xenoliths are garnet and chromite–garnet lherzolites or harzburgites, mainly from southern Africa and Yakutia. According to the empirical barometric expression proposed by Grütter et al. (2006), they were equilibrated at pressures up to 6 GPa. Figure 18 shows the equilibrium pressures and temperatures for these xenoliths calculated iteratively using the orthopyroxene barometer of Brey & Koehler (1990) (PBKN) and its modified version (PBBG) in combination with the orthopyroxene–garnet (Harley, 1984) and orthopyroxene–clinopyroxene (Brey & Koehler, 1990) thermometers. Most of the xenoliths plot along a geotherm slightly hotter than 40 mW/m2. The P–T parameters calculated by equations (1)–(4) are systematically lower than those obtained using the Brey & Koehler (1990) expression. The difference is up to 1 GPa for the highest pressures (6–7 GPa) and less than 0·2 GPa at 3–5 GPa. Most of the diamond-bearing inclusions fall within the diamond stability field or near the diamond–graphite transition boundary, and all the graphite-bearing samples but one fall within the graphite field. However, there is a gap between the diamond- and graphite-bearing inclusions, and this gap lies slightly below the diamond–graphite boundary. The difference is about 0·3 GPa for the Brey & Koehler (1990) barometer and up to 0·5 GPa for the new calibration (Fig. 18). The discrepancy does not necessarily imply the incorrect calibration of Al solubility in orthopyroxene, because it also depends on the reliability of the thermometers. Another possible reason is the decoupling of Fe–Mg and Ca–Mg exchange equilibria from the Al-in-orthopyroxene solubility in equilibrium with garnet. This problem requires a combined analysis of thermometric and barometric expressions and is beyond the scope of the present study.
|
The diamond inclusion suite includes touching and non-touching garnet–orthopyroxene pairs (sometimes with olivine or chromite) coexisting within single diamonds. Equilibrium pressures and temperatures were estimated using the Brey & Koehler (1990
) barometer and equations (1)–(4) together with the Fe–Mg orthopyroxene–garnet (Harley, 1984) and garnet–olivine thermometers (O’Neill & Wood, 1979) (Fig. 18). Similar to the xenoliths, the new calibration yields lower pressures (by up to 1 GPa) compared with the Brey & Koehler (1990) barometer. The touching and non-touching inclusions form two separate fields (Fig. 18). The temperature difference was explained by Phillips et al. (2004) by re-equilibration of the touching inclusions to ambient temperatures on a conductive geotherm. It is to be noted that the P–T conditions of the non-touching pairs cluster around a conductive geotherm of around 40 mW/m2. |
CONCLUDING REMARKS |
|---|
|
TOPABSTRACTINTRODUCTIONEXPERIMENTAL METHODSSIMPLE SYSTEMSNATURAL SYSTEMAPPLICATION TO MINERAL BAROMETRYCONCLUDING REMARKSREFERENCES |
|---|
The experimental results presented here have two important applications for understanding upper mantle processes. One of them is the extension of orthopyroxene barometry to pressures up to 8 GPa. Our results show that extrapolation of many popular mineral barometers to high pressures may be associated with large systematic errors. In particular, the widely used garnet–orthopyroxene barometer of Brey & Koehler (1990
) adequately reproduces pressures up to 5 GPa but overestimates pressure within the 6–8 GPa range by up to 1 GPa. Therefore we caution against uncritical extrapolation of barometeric expressions far beyond the range of experimental data used for their calibration. The second result that should be emphasized is the importance of the orthopyroxene–high-P clinopyroxene transition at a pressure of c. 9 GPa. This transition is probably accompanied by a slight increase in Al content in low-Ca pyroxene, which propagates into considerable errors in pressure estimates obtained by Al-in-orthopyroxene barometers calibrated at low pressures. We noted also a slight decrease in the Ca content of low-Ca pyroxene in equilibrium with high-Ca clinopyroxene at 9 and 10 GPa. This effect can be important for two-pyroxene thermometry.
|
ACKNOWLEDGEMENTS |
|---|
The authors thank S. Klemme, H. St. C. O’Neill and H. Grütter for helpful comments and suggestions, and Gareth Davies for editorial assistance. This study was financially supported by the Deutsche Forschungsgemeinshaft (DFG) and the Russian Foundation for Basic Research.
*Corresponding author. Telephone: +49 (069) 798 40124. Fax: +49 (069) 798 40121. E-mail: brey{at}em.uni-frankfurt.de
|
REFERENCES |
|---|
|
TOPABSTRACTINTRODUCTIONEXPERIMENTAL METHODSSIMPLE SYSTEMSNATURAL SYSTEMAPPLICATION TO MINERAL BAROMETRYCONCLUDING REMARKSREFERENCES |
|---|
1 Angel RJ, Chopelas A, Ross NL. Stability of high-density clinoenstatite at upper-mantle pressures. Nature (1992) 358:322–324.[CrossRef]
2 Brenker FE, Brey GP. Reconstruction of the exhumation path of the Alpe Arami garnet-peridotite body from depths exceeding 160 km. Journal of Metamorphic Geology (1997) 15:581–592.[CrossRef][Web of Science]
3 Brey GP, Koehler T. Geothermobarometry in four-phase lherzolites II. New thermobarometers, and practical assessment of existing thermobarometers. Journal of Petrology (1990) 31:1353–1378.
4 Brey GP, Koehler T, Nickel KG. Geothermobarometry in four-phase lherzolites I. Experimental results from 10 to 60 kbar. Journal of Petrology (1990) 31:1313–1352.
5 Brey GP, Nickel KG, Kogarko LN. Garnet–pyroxene equilibria in the system CaO–MgO–Al2O3–SiO2 (CMAS): prospect for simplified (‘T-independent’) lherzolite barometry and an eclogite-barometer. Contributions to Mineralogy and Petrology (1986) 92:448–455.[CrossRef][Web of Science]
7 Eggler DH, Baker DR. Reduced volatiles in the system C–O–H: implications to mantle melting, fluid formation, and diamond genesis. In: High Pressure Research in Geophysics—Akimoto S, Manghnani MH, eds. (1982) Tokyo: Centre for Academic Publications. 237–250.
8 Gasparik T. An internally consistent thermodynamic model for the system CaO–MgO–Al2O3–SiO2 derived primarily from phase equilibrium data. Journal of Geology (2000) 108:103–119.[CrossRef][Web of Science][Medline]
9 Gasparik T, Newton RC. The reversal Al-contents of orthopyroxene in equilibrium with spinel and forsterite in the system MgO–Al2O3–SiO2. Contributions to Mineralogy and Petrology (1984) 85:186–196.[CrossRef][Web of Science]
10 Girnis AV, Stachel T, Brey GP, Harris JW, Phillips D. Internally consistent geothermobarometers for garnet harzburgites: model refinement and applications. In: The J. B. Dawson Volume, Proceedings of the VIIth International Kimberlite Conference—Gurney JJ, Gurney JL, Pascoe MD, Richardson SH, eds. (1999) Cape Town: Red Roof Design. 247–254.
11 Grütter H, Latti D, Menzies A. Cr-saturation arrays in concentrate garnet compositions from kimberlite and their use in mantle barometry. Journal of Petrology (2006) 47:810–820.
12 Harley SL. An experimental study of the partitioning of Fe and Mg between garnet and orthopyroxene. Contributions to Mineralogy and Petrology (1984) 86:353–373.
13 Harley SL, Green DH. Garnet–orthopyroxene barometry for granulites and garnet peridotites. Nature (1982) 300:697–700.[CrossRef]
14 Hernlund J, Leinenweber K, Locke D, Tyburczy JA. A numerical model for steady-state temperature distributions in solid-medium high-pressure cell assemblies. American Mineralogist (2006) 91:295–305.
15 Kennedy CS, Kennedy GC. The equilibrium boundary between graphite and diamond. Journal of Geophysical Research (1976) 81:2467–2470.
16 Klemme S, O’Neill H. StC. The near-solidus transition from garnet lherzolite to spinel lherzolite. Contributions to Mineralogy and Petrology (2000) 138:237–248.[CrossRef][Web of Science]
17 Koehler T, Brey GP. Calcium exchange between olivine and clinopyroxene calibrated as a geothermobarometer for natural peridotites from 2 to 60 kb with applications. Geochimica et Cosmochimica Acta (1990) 54:2375–2388.[CrossRef][Web of Science]
18 Lane DL, Ganguly J. Al2O3 solubility in orthopyroxene in the system MgO–Al2O3–SiO2: a re-evaluation and mantle geotherm. Journal of Geophysical Research (1980) 85:6963–6972.
19 Lindsley DH, Dixon SA. Diopside–enstatite equilibria at 850 to 1400°C, 5 to 35 kbars. American Journal of Science (1976) 276:1285–1301.
20 MacGregor ID. The system MgO–Al2O3–SiO2: solubility of Al2O3 in enstatite for spinel and garnet peridotite compositions. American Mineralogist (1974) 59:110–119.[Web of Science]
21 Menzies A, Westerlund K, Grütter H, Gurney J, Carlson J, Fung A, Nowicki T. Peridotitic mantle xenoliths from kimberlites on the Ekati Diamond Mine property, N.W.T. Canada: major element compositions and implications for the lithosphere beneath the central Slave craton. Lithos (2004) 77:395–412.[CrossRef][Web of Science]
22 Morishima H, Kato T, Suto M, Ohtani E, Urakawa S, Utsumi W, Shimomura O, Kikegawa T. The phase boundary between - and ß-Mg2SiO4 determined by in situ X-ray observation. Science (1994) 265:1202–1203.
23 Nickel KG, Brey GP. Subsolidus orthopyroxene–clinopyroxene systematics in the system CaO–MgO–SiO2 to 60 kb: a re-evaluation of the regular solution model. Contributions to Mineralogy and Petrology (1984) 87:35–42.[CrossRef][Web of Science]
24 Nickel KG, Green DH. Empirical geothermobarometry for garnet peridotites and implications for the nature of the lithosphere, kimberlites and diamonds. Earth and Planetary Science Letters (1985) 73:158–170.[CrossRef][Web of Science]
25 Nickel KG, Brey GP, Kogarko LN. Orthopyroxene–clinopyroxene equilibria in the system CaO–MgO–Al2O3–SiO2 (CMAS): new experimental results and implications for two-pyroxene thermometry. Contributions to Mineralogy and Petrology (1985) 81:44–53.
26 O’Neill H. StC, Wood BJ. An experimental study of Fe–Mg-partitioning between garnet and olivine and its calibration as a geothermometer. Contributions to Mineralogy and Petrology (1979) 70:59–70.[CrossRef][Web of Science]
27 Pacalo R. EG, Gasparik T. Reversals of the orthoenstatite–clinoenstatite transition at high pressures and temperatures. Journal of Geophysical Research (1990) 95:15853–15858.[CrossRef]
28 Perkins D, Newton RC. The compositions of coexisting pyroxenes and garnet in the system CaO–MgO–Al2O3–SiO2 at 900°–1100°C and high pressures. Contributions to Mineralogy and Petrology (1980) 75:291–300.[Web of Science]
29 Perkins D. III, Holland T. JB, Newton RC. The Al2O3 content of enstatite in equilibrium with garnet in the system MgO–Al2O3–SiO2 at 15–40 kbar and 900–1000°C. Contributions to Mineralogy and Petrology (1981) 78:99–109.[CrossRef][Web of Science]
30 Phillips D, Harris JW, Viljoen KS. Mineral chemistry and thermobarometry of inclusions from De Beers Pool diamonds, Kimberley, South Africa. Lithos (2004) 77:155–179.[CrossRef][Web of Science]
31 Pollack HN, Chapman DS. On the regional variation of heat flow, geotherms, and lithospheric thickness. Tectonophysics (1977) 38:279–296.[CrossRef][Web of Science]
32 Steinberg HK. Hochdruckexperimente an Calciumsilikatphasen zur Rekonstruktion der Aufstiegsgeschichte von Diamanten. (2005) Ph.D. dissertation. J. W. Goethe-Universität, Frankfurt am Main, 141 pp.
33 Susaki J, Akaogi M, Akimoto S, Shimomura O. Garnet–perovskite transformation in CaGeO3: in-situ X-ray measurements using synchrotron radiation. Geophysical Research Letters (1985) 12:729–732.[Web of Science]
34 Van Westrenen W, Van Orman JA, Watson H, Fei Y, Watson EB. Assessment of temperature gradients in multianvil assemblies using spinel layer growth kinetics. Geochemistry, Geophysics, Geosystems (2003) 4:1036. doi:10.1029/2002GC000474.[CrossRef]
35 Watson EB, Wark DA, Price JD, Van Orman JA. Mapping the thermal structure of solid-media pressure assemblies. Contributions to Mineralogy and Petrology (2002) 142:640–652.[Web of Science]
36 Woodland AB, Angel RJ. Reversal of the orthoferrosilite–high-P ferrosilite transition, a phase diagram for FeSiO3 and implications for the mineralogy of the Earth's upper mantle. European Journal of Mineralogy (1997) 9:245–254.
37 Yagi T, Akimoto S. Direct determination of coesite–stishovite transition by in-situ X-ray measurements. Tectonophysics (1976) 35:259–270.[CrossRef][Web of Science]
CiteULike 
Connotea 
Del.icio.us What's this?
This article has been cited by other articles:
|
|
 |
|
  S. A. Gibson, J. Malarkey, and J. A. Day Melt Depletion and Enrichment beneath the Western Kaapvaal Craton: Evidence from Finsch Peridotite Xenoliths J. Petrology, October 22, 2008; (2008) egn048v1. [Abstract] [Full Text] [PDF]
|
|
| ||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||