| ||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||
Journal of Petrology | Volume 44 | Number 7 | Pages 1163-1191 | 2003
© Oxford University Press 2003
Near-solidus Melting of the Shallow Upper Mantle: Partial Melting Experiments on Depleted Peridotite
DIVISION OF GEOLOGICAL AND PLANETARY SCIENCES, CALIFORNIA INSTITUTE OF TECHNOLOGY, MAIL CODE 170-25, PASADENA, CA 91125, USA
* Corresponding author. Present address: Department of Geological Sciences, 4044 Derring Hall, Virginia Polytechnic Institute and State University, Blacksburg, VA 24061, USA. Telephone: 540-231-2403. Fax: 540-231-3386. E-mail: lew{at}vt.edu
RECEIVED SEPTEMBER 18, 2001; ACCEPTED FEBRUARY 5, 2003
 |
ABSTRACT |
|---|
 TOP ABSTRACT INTRODUCTIONEXPERIMENTAL AND ANALYTICAL...RESULTSDISCUSSIONCONCLUSIONSREFERENCES |
|---|
We present the results of melting experiments on a moderately depleted peridotite composition (DMM1) at 10 kbar and 1250–1390°C. Specially designed experiments demonstrate that liquids extracted into aggregates of vitreous carbon spheres maintained chemical contact with the bulk charge down to melt fractions of
0·02–0·04 and approached equilibrium closely. With increasing melt fraction, SiO2, FeO*, and MgO contents of the partial melts increase, Al2O3 and Na2O contents decrease, and CaO contents first increase up to clinopyroxene-out at a melt fraction of 0·09–0·10, then decrease with further melting. A linear fit to melt fraction vs temperature data for lherzolite-bearing experiments yields a solidus of 1272 ± 11°C. The melting reaction is 0·56 orthopyroxene + 0·72 clinopyroxene + 0·04 spinel = 0·34 olivine + 1 liquid. Above clinopyroxene-out, the reaction is 1·24 orthopyroxene = 0·24 olivine + 1 liquid. Near the solidus, DMM1 glass compositions have lower SiO2, TiO2, Na2O, and K2O contents, higher FeO*, MgO, and CaO contents, and higher CaO/Al2O3 ratios compared with glasses from low-degree melting of fertile peridotite compositions. Recent computational models predict partial melting trends generally parallel to our experimental results. We present a parameterization of 10 kbar peridotite solidus temperatures suggesting that K2O and P2O5 have greater effects on solidus depression than Na2O, consistent with theoretical expectations. Our parameterization also suggests that abyssal peridotites have 10 kbar solidi of 1278–1295°C. KEY WORDS: depleted; experimental petrology; mantle melting; near-solidus; peridotite
|
INTRODUCTION |
|---|
|
TOPABSTRACTINTRODUCTIONEXPERIMENTAL AND ANALYTICAL...RESULTSDISCUSSIONCONCLUSIONSREFERENCES |
|---|
It is widely accepted that parental magmas of mid-ocean ridge basalts (MORBs) represent a mixture of liquids produced by pressure-release melting over a range of depths in the Earth's upper mantle (e.g. McKenzie, 1984; McKenzie & Bickle, 1988; Langmuir et al., 1992). Experimental and theoretical studies indicate that small amounts of melt (<2%) are interconnected in olivine-dominated, partially molten systems (e.g. Waff & Bulau, 1979; von Bargen & Waff, 1986; Daines & Richter, 1988) and that such low degrees of melt can move relative to the residual solid phases (McKenzie, 1985, 1989; Stevenson & Scott, 1991). The extremely depleted rare earth element patterns observed in residual clinopyroxenes from dredged abyssal peridotites have been used as evidence of this mobility and that melt production beneath mid-ocean ridges can approach the limit of fractional melting (Johnson et al., 1990; Johnson & Dick, 1992; Salters & Dick, 2002). Correlations between the major element compositions of spinels and the concentrations of moderately incompatible trace elements in clinopyroxenes from abyssal peridotites also suggest that these peridotites experienced near-fractional melting (Hellebrand et al., 2000). Finally, melt inclusions with variably depleted trace element concentrations relative to average MORB (e.g. Sobolev & Shimizu, 1993; Sobolev, 1996; Shimizu, 1998) suggest that erupted MORB magmas contain a component of liquid produced by melting of depleted peridotite. Taken together, these lines of evidence suggest that partial melts can be extracted from the mantle beneath mid-ocean ridges at low melt fractions and that at least some components of MORBs were produced by melting of depleted sources with low concentrations of incompatible major, minor, and trace elements relative to the fertile MORB source.
Most high-pressure experimental studies of peridotite melting have focused on fertile mantle compositions, and these studies have provided information on melt compositions over a wide range of pressures and temperatures (e.g. Jaques & Green, 1980; Takahashi, 1986; Hirose & Kushiro, 1993; Baker & Stolper, 1994; Baker et al., 1995; Kushiro, 1996; Robinson et al., 1998; Walter, 1998). In addition, Pickering-Witter & Johnston (2000) and Schwab & Johnston (2001) explored the effects of pyroxene and spinel abundances and pyroxene compositions on liquid compositions, melt productivities, and melting reactions at 10 kbar, and, by varying the orthopyroxene/clinopyroxene ratios in their starting materials, they were able to work on bulk compositions spanning wide ranges of MgO, Al2O3, CaO, and Na2O contents.
In a synthesis of available experimental data, Hirschmann et al. (1998a) showed the important role that alkalis play in controlling the compositions of olivine + orthopyroxene ± clinopyroxene-saturated melts of peridotite. Likewise, the broadly inverse correlation between total alkali contents of peridotites and their solidus temperatures (Herzberg et al., 2000; Hirschmann, 2000) suggests that, at a given pressure, depleted peridotites will begin to melt at higher temperatures than fertile peridotites. Because peridotites in the shallowest part of the melting zone beneath mid-ocean ridges are expected to have been depleted in incompatible major and minor elements (and especially alkalis) by prior melting, the expectation is that such peridotites would melt to a lower degree and produce different liquids compared with melting of fertile peridotite under similar conditions.
In this paper we present the results of melting experiments on a moderately depleted peridotite composition (DMM1) at 10 kbar and 1250–1390°C; these results complement the large amount of available data on more fertile peridotite compositions. We used a variant of the diamond aggregate melt-extraction technique developed by Johnson & Kushiro (1992), Hirose & Kushiro (1993), and Baker & Stolper (1994) to study near-solidus experimental glasses unmodified by quench crystal growth. In the experiments described here, the aggregate comprised vitreous carbon spheres or fragments (Wasylenki et al., 1996; Pickering-Witter & Johnston, 2000; Schwab & Johnston, 2001). A potential problem associated with the diamond aggregate technique in low melt fraction experiments is that pressure is initially low within the void spaces of the aggregate until the layer is completely filled with melt (Baker et al., 1996), possibly leading to the segregation of liquid that is far from equilibrium with the peridotite at the actual pressure of the experiment. The motivation for replacing the porous diamond aggregate used in our previous studies (Baker & Stolper, 1994; Baker et al., 1995) with vitreous carbon spheres is that, although vitreous carbon is strong enough to support open pore space at the start of an experiment, it is much less stiff than diamond (Noda et al., 1969; Sawa & Tanaka, 2002), so this open space collapses more quickly. Another advantage is that vitreous carbon is readily polished at the end of an experiment, allowing small pools of melt to be imaged and analyzed in situ with ease (e.g. Pickering-Witter & Johnston, 2000; Schwab & Johnston, 2001). In addition to modifying the melt-extraction technique in this way, we have addressed the controversy surrounding this technique (see Baker et al., 1996; Falloon et al., 1996, 1997, 1999; Wasylenki et al., 1996; Pickering-Witter & Johnston, 2000) with special experiments that demonstrate the reliability and close approach to equilibrium of melt-extraction experiments.
Following the presentation of the experimental results, we consider the effects of bulk composition on 10 kbar melting reaction coefficients and on melt productivity, and we compare our experimental glass compositions for the depleted peridotite DMM1 with those produced by melting more fertile peridotite compositions at the same pressure. We also compare our experimental glass compositions with liquid compositions calculated for the DMM1 composition using four peridotite melting models. Finally, by combining our experimentally determined 10 kbar solidus temperature with solidus determinations on other peridotite compositions, we develop a simple expression for predicting the solidus temperature at 10 kbar as a function of bulk composition.
|
EXPERIMENTAL AND ANALYTICAL TECHNIQUES |
|---|
|
TOPABSTRACTINTRODUCTIONEXPERIMENTAL AND ANALYTICAL...RESULTSDISCUSSIONCONCLUSIONSREFERENCES |
|---|
Starting materials
Starting materials for this study were prepared by mixing mineral separates (olivine, ol; orthopyroxene, opx; clinopyroxene, cpx; spinel, sp) from a Kilbourne Hole spinel lherzolite nodule (KBH), olivine from a Hawaiian dunite nodule (H1801i), and synthetic diopside (CaMgSi2O6). Electron microprobe analyses of the natural minerals are reported in Table 1. The crystallinity of the synthetic diopside was verified by X-ray diffraction, and the composition reported in Table 1 is that of stoichiometric diopside. The natural minerals were hand-picked from disaggregated nodules, ground, sieved to 16–28 mesh, cleaned in warm 2·4 N hydrochloric acid, and rinsed in deionized water. They were then crushed further, sieved to 200–325 mesh, and again washed in 2·4 N HCl and deionized water. Finally, the grains were ground and sieved to
13 µm. The synthetic diopside was ground by hand until most of the grains were <15 µm, but the powder still contained a few grains as large as 40 µm in size. The minerals were mixed in the weight proportions 0·432 KBH olivine, 0·202 H1801i olivine, 0·274 KBH orthopyroxene, 0·014 KBH clinopyroxene, 0·060 diopside, and 0·018 KBH spinel to generate the depleted peridotite starting material (DMM1). The mixture was ground by hand for 1 h under ethanol to ensure homogeneity.
|
The bulk composition of DMM1 in terms of major oxides is reported in Table 1. With the exceptions of K2O and P2O5, these values were computed from the mineral compositions and their proportions in the mix (see preceding paragraph). K2O and P2O5 contents for DMM1 were calculated using the K2O and P2O5 contents of glasses in selected experiments, mineral–melt partition coefficients from the literature, and melt fractions and residual mineral proportions for these experiments derived from mass balance constraints using the other oxides (see footnotes to Tables 1 and 3).
|
The starting DMM1 composition was chosen to represent a moderately depleted mantle peridotite. When compared with the residues in the 10 kbar melting experiments of Baker & Stolper (1994) and Baker et al. (1995), DMM1 is roughly equivalent to the residue formed by 12–13% batch melting of MM3. Figure 1 compares the concentrations of MgO, Al2O3, CaO, and Na2O in DMM1 with other peridotites studied experimentally at 9–10 kbar, with estimates of the primitive mantle, and with reconstructed abyssal peridotites. DMM1 has lower Al2O3, CaO, and Na2O contents than all estimates of primitive mantle, and, unlike other peridotite compositions that have been studied experimentally, it falls on the well-defined MgO–oxide trends for abyssal peridotites (Fig. 1). Figure 1a and b shows that DMM1 is lower in Al2O3 than all other experimentally studied peridotite compositions and lower in CaO than the others, except those investigated by Sen (1982; 5 in Fig. 1b) and by Pickering-Witter & Johnston (2000; 9 in Fig. 1b; note that this composition is far removed from the MgO–CaO trend defined by abyssal peridotites and primitive mantle estimates). Figure 1c shows that DMM1 also has lower Na2O than most peridotite compositions previously studied.
|
: 
: Experimental methods
All experiments were run in a 1·27 cm piston cylinder apparatus using CaF2 cells, straight-walled graphite furnaces, and inner pieces of crushable MgO dried at 1000°C for at least 8 h. Pressure was applied using the hot-piston-in technique with no friction correction. Experiments at 1300°C with identical assemblies bracketed the Ca-Tschermak breakdown reaction to lie at 11–14 kbar. This pressure range encompasses the reaction boundary (1300°C, 13 kbar) determined by Hays (1966). W97Re3/W75Re25 thermocouples were used to monitor and control temperature to within
1°C of the set point. No pressure correction was applied to the nominal e.m.f.–temperature relation. Based on past experiments with double thermocouples, temperatures are estimated to be accurate to within ±15°C. Run durations and experimental conditions are reported in Table 2. Except in experiments 1, 2, and 3, N2 gas was bled into the slot in the thermocouple plate during each experiment to minimize oxidation of the thermocouple wires within and just below the steel base plug. At the end of each experiment, the thermocouple wires just below the base plug (i.e. within the run assembly) were inspected for signs of oxidation; no evidence of significant oxidation was observed on any of the wires, including those from experiments 1–3. After the power was turned off, samples cooled to below 1000°C within several seconds. As each run assembly was taken apart, the position of the capsule relative to the center of the furnace and that of the thermocouple junction relative to the top of the capsule were measured to ensure that the capsule was properly positioned within the furnace and that the thermocouple tip was 0·5–1 mm from the top of the capsule. Each capsule was sliced vertically with a diamond wafering blade, mounted in epoxy, and polished for electron microprobe analysis.
|
This study includes three types of melting experiments. The first type of experiment consisted of two stages (see Baker & Stolper, 1994). For the first stage, 3–6 mg of peridotite powder that had been dried in a vacuum oven at 110°C for
2 h was loaded into a graphite inner capsule. The inner capsule was then placed in a 0·15 inch o.d. Pt capsule. The composite capsule assembly was then dried (see Table 2); and finally the crimped end of the Pt capsule was welded. After drying and welding, the capsule was run at temperature and pressure for 52–136 h. For the second stage, the silicate charge was removed from the first-stage capsule and loaded into a new graphite capsule along with 80–100 µm diameter vitreous carbon spheres that had been dried for at least 1 h in a 110°C vacuum oven. The mass of vitreous carbon (Table 2) was 3–9% of the mass of the silicate sample. The loaded second-stage graphite capsule was then placed in a Pt capsule, after which the composite assembly was dried and welded shut. This second-stage capsule was then run for 18–135 h at the same temperature and pressure as its corresponding first-stage run. At the end of each second-stage experiment, the charge contained quenched glass associated with the vitreous carbon. The glass formed rinds generally 3–15 µm thick around individual vitreous carbon spheres. Varying amounts of glass and crushed carbon spheres filled the spaces between the intact spheres (Fig. 2a). The idea behind this type of experiment was to avoid the possibility of formation and segregation of disequilibrium liquids at the start of the experiment that would not subsequently equilibrate with the main mass of the sample. In two-stage experiments, liquids maintain intimate contact with the peridotite and approach equilibrium during the first stage and are segregated only in the second stage.
|
The second type of experiment was run as a single stage, but the temperature was changed during the experiment to evaluate whether liquid that had segregated into the interstices between the vitreous carbon spheres could respond chemically to changes in temperature; these experiments are indicated by a T next to the run number in Table 2. Graphite capsules were loaded with vitreous carbon and peridotite powder, placed in Pt outer capsules, and then dried and welded shut as for the previously described charges. Each sample was run at 10 kbar and an initial temperature as listed in Table 2 for
26–68 h. Following this initial period, the temperature was raised either 15 or 40°C, and the charge was held at this higher temperature for an additional 88–117 h. The potential concern addressed by these experiments is whether liquid segregated into the vitreous carbon at the start of the experiment is in sufficient chemical contact with the main mass of the experimental charge that its composition can change by diffusive interaction with the peridotite and its interstitial liquid.
We refer to the third type of experiment as ‘compositional-convergence’ runs; these experiments are marked with a C next to the run number in Table 2. For these seven experiments, glass-bearing vitreous carbon aggregates were prepared as follows. An oxide mix of a synthetic ferrobasalt or a powdered rhyolitic glass from Glass Buttes, Oregon (both compositions are reported in Table 1), were mixed with conchoidally fractured vitreous carbon shards (80–147 µm in size). Approximately 85% (by weight) of the material in each capsule was vitreous carbon. Each of these two mixtures was loaded into a separate Pt capsule and then held at 10 kbar and 1350°C for a few hours. At the end of these two preparatory runs, the silicate liquids had impregnated most of the pore spaces between the vitreous carbon fragments. Fragments (10–41 µg; see Table 2) of the glass-impregnated vitreous carbon aggregates were loaded into graphite–Pt capsules with DMM1 powder (3–6 mg; see Table 2) as described above for the other two types of experiments. The samples were run for 127–188 h. In one experiment (20C), a layer of vitreous carbon spheres was also loaded into the capsule at the opposite end of the charge from the ferrobasalt-impregnated vitreous carbon aggregate. The purpose of these experiments was to examine the degree to which the ferrobasaltic and rhyolitic glasses in the interstices of the vitreous carbon aggregate were able to converge in composition toward each other and toward the compositions of glasses within vitreous carbon aggregates that were not pre-filled with glass (i.e. glass within the layer of vitreous carbon in 20C or that in charges from the previously described experiments). Because the ferrobasaltic and rhyolitic glasses are extremely distant in composition from the equilibrium partial melts of the DMM1 peridotite, the extent to which they can shift compositionally during an experiment can be used to demonstrate that liquids segregated into vitreous carbon aggregates can maintain chemical contact with the main mass of the peridotite during an experiment. This means that even if a disequilibrium liquid migrates into the vitreous carbon aggregate at the start of an experiment, it can evolve substantially via exchange with the bulk charge over the course of the experiment. Moreover, if the initial ferrobasaltic and rhyolitic compositions converge to a single composition similar to glasses from our two-stage and temperature-change experiments, we can infer with confidence that an equilibrium liquid composition was approached closely.
Oxygen fugacity was not controlled in our experiments, but the presence of graphite inner capsules constrains fO2 to below the graphite–CO vapor buffer (GCO). Following the approach discussed by Bertka & Holloway (1988) and Gudmundsson et al. (1988), we placed 13 µg of platinum wire in the middle of the peridotite powder in experiment 25 (1325°C). During the experiment the Pt wire absorbed Fe. Based on the compositions of the resulting Fe–Pt alloy and coexisting orthopyroxene and olivine, and using the equations and solution models of Jamieson et al. (1992), the log10 fO2 in run 25 was –9·2, or 2·4 log units below the quartz–fayalite–magnetite buffer (QFM) at 10 kbar (Huebner, 1971). The slightly different activity–composition models of Gudmundsson & Holloway (1993) yielded a log10 fO2 of -8·5, or 1·7 log units below QFM. The inferred fO2 values of our experiments are thus consistent with estimates for the MORB source region (Christie et al., 1986; Green et al., 1987; O'Neill & Wall, 1987; Mattioli & Wood, 1988; Mattioli et al., 1989).
Analytical techniques
Experimental charges were analyzed at Caltech with a five spectrometer JEOL 733 electron microprobe using an accelerating voltage of 15 keV. Crystalline phases in all experiments (including the Pt alloy in run 25) were analyzed with a beam current of 30 nA and a rastered beam at magnifications greater than 40 000x (the resulting spot size is less than 2 µm x 2 µm). Glasses were analyzed with a 2·5, 5, or 10 nA beam current and a rastered area as large as possible for the glass pools being analyzed (generally 5 µm x 5 µm). Repeated analyses on the same region of glass within several experiments indicated little or no Na loss. All data were processed using CITZAF (Armstrong, 1988). Two basaltic glasses, VG-2 and BGIO (Jarosewich et al., 1979), were analyzed as secondary standards during each microprobe session to estimate accuracy and precision of the instrument on a long-term basis; when crystalline phases were analyzed, Johnstown hypersthene and Natural Bridge diopside (Jarosewich et al., 1979) were also analyzed to provide additional estimates of accuracy and precision.
Raw oxide sums for the experimental glasses were typically 95–97%, but were occasionally lower, and in one case down to 89%. We consistently observed higher totals for glass pools that were more than 5 µm in the shortest dimension and lower totals for smaller glass pools. The low totals probably reflect the inclusion of vitreous carbon in the electron beam analysis volume. This inference is supported by the observation that when all of the glass analyses for a given sample are normalized to 100%, there are no systematic differences in the normalized oxide concentrations as functions of the raw oxide sums. Within the vitreous carbon aggregate of a given experiment, normalized glass compositions display no systematic variation as a function of distance from the DMM1 peridotite. Mean compositions for normalized glasses, crystalline phases, and secondary standards are reported in Table 3.
Water contents of the glasses in two charges (3 and 22) were measured by secondary ion mass spectrometry with the modified Cameca IMS-3f ion microprobe at Lawrence Livermore National Laboratories using the techniques described by Kent et al. (1999). Measured H+/30Si+ ratios in our experimental glasses were converted to water contents with a calibration curve constructed by analyzing a set of Marianas back-arc basin basaltic glasses (Stolper & Newman, 1994) and synthetic MORB glasses (Dixon et al., 1995) whose water and silica contents were determined by independent techniques. Errors on the water contents of the experimental glasses are conservatively estimated at ±15% relative. Charges 3 and 22 contain 1·2 ± 0·2 and 0·3 ± 0·1 wt % water, respectively. These have not been corrected for the possible effects of vitreous carbon in the analysis volume. As vitreous carbon can dissolve hydrogen in the presence of water-bearing melt, but contains no Si (L. Wasylenki, unpublished data, 1995), such a correction would probably lead to lower estimates of the water contents of our experimental glasses.
|
RESULTS |
|---|
|
TOPABSTRACTINTRODUCTIONEXPERIMENTAL AND ANALYTICAL...RESULTSDISCUSSIONCONCLUSIONSREFERENCES |
|---|
Evaluation of the approach to equilibrium
Although the melt-extraction technique used in this study provides a simple way to avoid the quench modification problems encountered in conventional peridotite melting experiments close to the solidus, the technique must be applied with care. In particular, the issue of whether the glasses within the vitreous carbon aggregates (this study; Pickering-Witter & Johnston, 2000; Schwab & Johnston, 2001) or diamond aggregates (e.g. Johnson & Kushiro, 1992; Hirose & Kushiro, 1993; Baker & Stolper, 1994; Baker et al., 1995; Kushiro, 1996) actually represent equilibrium melt compositions has sparked much debate (Baker et al., 1996; Falloon et al., 1996, 1997, 1999; Wasylenki et al., 1996). For example, as explained above, one possible problem with the one-stage experiments of Baker & Stolper (1994) is that at the onset of an experiment, low pore pressure within the diamond layer could affect the composition of liquid initially filling these pores, i.e. the initial liquid may not be in equilibrium with the bulk peridotite at the nominal pressure of the experiment. Similarly, the kinetics of melting could lead to generation of a disequilibrium liquid at the start of an experiment. In either of these cases, liquid migrating rapidly into the pores in the diamond or vitreous carbon layer might not be in equilibrium with the bulk of the material in the capsule, and, if this segregated liquid were subsequently cut off from exchange with the adjacent peridotite, results would be erroneous. Liquids segregated into diamond aggregates in short-duration, one-stage experiments do indeed differ from those produced in long-duration runs at the same temperature and pressure (Johnson & Kushiro, 1992; Baker & Stolper, 1994), so this general class of potential problems cannot be disregarded. Our two-stage experiments, however, were designed to minimize these problems in that the liquid moving into a vitreous carbon layer at the beginning of a second-stage run is remelted glass from the first-stage run and thus should represent a near-equilibrium melt.
Temperature-change experiments
Although our two-stage experiments were explicitly designed to minimize problems of melt re-equilibration in the vitreous carbon layers, we nevertheless made further efforts to demonstrate directly that liquids in the vitreous carbon layers in our experiments can change composition over the course of an experiment. The temperature-change experiments described above provide evidence of continued interaction between segregated melt and the peridotite over the course of an experiment.
We conducted three temperature-change experiments to demonstrate that segregated liquids can change composition in response to changes in experimental conditions. In each such experiment, after initial run durations of 26–68 h at 1285 or 1310°C, temperatures were raised either 15 or 40°C and the runs continued for an additional 88–117 h. We assume that during the initial stages of these experiments, liquids in the vitreous carbon layers approached the glass compositional trends defined by the two-stage experiments at the same initial temperatures. In two of the three temperature-change experiments (17T, 1285–1300°C; 25T, 1285–1325°C), the final liquid compositions are similar to those in two-stage experiments run at the same final temperatures and different from liquid compositions expected at 1285°C (see Fig. 4). (The third experiment, 23T, appears to have had a significant temperature gradient, as evidenced by modal variations from top to bottom of the charge, and to have had a final temperature of 1310°C rather than 1325°C; see the section ‘Phase relations and liquid compositions' below.) Based on mass balance calculations using the quenched glass and mineral compositions and the bulk DMM1 composition, the melt fractions of 17T and 25T during their respective high-temperature steps were 0·03 and 0·11 (Table 2). Thus, over the course of 93–117 h at these melt fractions, the segregated liquids in the vitreous carbon layers were able to change in response to the increase in temperature via diffusive exchange with liquid retained in the peridotite. We also attempted temperature-change experiments in which the temperature was lowered in the second phase of the experiment, but, as one would expect, this resulted in extensive crystallization within the vitreous carbon aggregates.
|
) from run 20C are marked with small black circles. It should be noted that these two compositions overlap at the 1Compositional-convergence experiments
Successful ‘compositional-convergence’ experiments demonstrate not only that liquids in vitreous carbon aggregates can change significantly, but also by convergence that an equilibrium result has been approached from more than one direction in multi-component composition space. As described above, these experiments involved placing
10–41 µg of a vitreous carbon aggregate that was pre-impregnated with either a synthetic ferrobasalt or a rhyolite in a capsule with DMM1 powder (Fig. 2b). Based on the results of our two-stage experiments, both the ferrobasalt and the rhyolite glass compositions were initially far from equilibrium with DMM1 at 10 kbar and any temperature. With the exceptions of K2O and, to a lesser extent, Na2O for the two rhyolite-bearing experiments, the few micrograms of glass added to the system have little effect on the bulk composition. The potassium budget in the two rhyolite-bearing runs (26C and 34C) is dominated by the rhyolite glass; for example, we estimate that the bulk K2O content in 34C increased from 0·006 wt % (the estimated bulk composition of DMM1) to 0·02 wt % as a result of the addition of the rhyolite-bearing vitreous carbon aggregate. The change in the sodium content in run 34C is less extreme; the Na2O concentration increased from 0·06 to 0·07 wt %.
Figure 3 illustrates the large shifts in composition that the ferrobasaltic and rhyolitic glasses underwent as a result of diffusive equilibration with 2–12 wt % melt in the peridotite and shows that the quenched liquids in all of these experiments define a single compositional trend approached from opposite directions in most compositional dimensions. As shown in Fig. 4, this compositional trend is consistent with the trend defined by glass compositions from the two-stage and temperature-change experiments. It is important to emphasize that the compositions of the liquids in the vitreous carbon chips in these experiments have converged to their final compositions from opposite directions in terms of several oxides and that the compositional changes achieved in these experiments are in some cases very large (e.g. SiO2, CaO, MgO, and FeO* for the initially rhyolitic liquid). It is also striking that three of the convergence experiments are within <5°C of our estimated solidus for the DMM1 composition and have only 2–5% melt (Table 2). These results should thus dispel controversy about the validity of melt-segregation experiments into diamond or vitreous carbon aggregates, as they demonstrate large compositional shifts, convergence from melts initially in the aggregates that are extremely far in composition from the final compositions, and that the results of two-stage, temperature-change, and ‘compositional-convergence’ experiments form coherent compositional trends as a function of melt fraction near the solidus.
|
Homogeneity of crystalline phases
We have also examined the compositions of the residual peridotite minerals in all of our experiments (Table 3), as true achievement of equilibrium would require homogeneous crystals and mineral compositions that are independent of distance from the glass-bearing aggregates.
In all experiments olivine grains (15–40 µm in size) were homogeneous within analytical error in each charge, both from core to rim of individual analyzed grains and from top to bottom of the residual peridotite.
Orthopyroxene grains with cross-sectional diameters larger than 10 µm often displayed small, incompletely reacted cores, visible with high-gain, backscattered electron imaging. We estimate that these cores occupy <35% of the volume of opx in a given charge. The cores have compositions between the initial KBH opx composition and analyzed rim compositions, indicating some compositional change toward equilibrium. We did not observe any significant differences in opx core and rim compositions as a function of distance from the vitreous carbon aggregates in any of our experiments.
Clinopyroxene grains are typically <15 µm in size. Unreacted cpx cores were not clearly visible in backscattered electron images, but microprobe analyses indicate that cpx grains in the run products are chemically heterogeneous. In particular, CaO and Al2O3 contents vary within and among the cpx grains in each charge. In our study, more than 20 cpx analyses were collected from each charge, and those that clustered toward the low-Al and low-Ca end of the trend in Al2O3–CaO–MgO space defined by all analyses within a charge were averaged for Table 3. None of the cpx analyses from any of the run products were close in composition to the starting KBH cpx composition or pure diopside, suggesting that all of the cpx in our experiments at least partially re-equilibrated. We did not observe any systematic variations in cpx compositions as a function of distance from the vitreous carbon aggregates, but in some experiments modal cpx abundance increases with increasing distance from the top of the capsule (i.e. the end closest to the thermocouple junction). The experiments in which this modal variation occurs probably experienced larger thermal gradients than those experiments that show a more uniform cpx distribution (Lesher & Walker, 1988).
Spinel is a minor phase in our experiments and, as run temperatures fall, shows more compositional heterogeneity within runs than olivine or pyroxenes. Spinel grains in the run products are small (<10 µm in size), and generally only 3–6 grains are visible on the polished surface of each charge. Thus we were not able to assess the degree of compositional variation within individual sp grains or as a function of the position of sp grains in the charge relative to the melt-bearing vitreous carbon aggregates. We did observe that in a given charge, grains of sp in direct contact with melt tended to be more compositionally similar than those sp grains that were not obviously in contact with melt. The compositions reported in Table 3 are averages of the sp grains that had similar compositions within each charge. As we discuss below, these mean compositions are consistent with the compositional trends of sp from other 9–12 kbar peridotite melting experiments.
Overall, although the pyroxenes and spinels in the experimental charges are not homogeneous within analytical error, our run durations were long compared with most previous peridotite melting experiments, and the observed heterogeneity is attributable to incompletely reacted cores that would be difficult to react fully without much finer starting materials or much longer experiments. As these cores do not make up a major fraction of the residual phases, it is unlikely that they have a significant influence on the results. More importantly, the compositional heterogeneities observed in pyroxenes and spinel within each charge do not appear to correlate with distance from the vitreous carbon aggregates.
Phase relations and liquid compositions
Our experiments span the 1250–1390°C temperature range. Glass was present in all experiments run at 1270°C; no glass was detected in high-magnification, backscattered electron images of the 1250°C experiment (36), and although such visual determinations can be misleading, we tentatively place the solidus for DMM1 between 1250 and 1270°C. Melt coexists with ol, opx, cpx, and sp at 1270–1300°C. Experiments at 1310–1390°C contain liquid, ol, and opx, with the exception of experiment 23T (nominally 1325°C), which contains 1 wt % of cpx at the end of the capsule far from the vitreous carbon aggregate and the thermocouple junction. We interpret these results to indicate that cpx-out (the temperature at which cpx disappears from the solid residue) and sp-out are at 1300–1310°C. We attribute the cpx in run 23T to a deviation in nominal temperature at the limit of our stated ±15°C uncertainty, coupled with a temperature gradient along the length of the capsule. Comparison of the composition of glass in 23T with glasses from our other experiments (Fig. 4) shows that it is consistent with being just at cpx-out, and examination of compositional trends in 23T and the other experimental glasses with temperature (not shown) suggests that the temperature of run 23T was 1310°C.
Figure 4 shows the concentrations of selected oxides in glasses from our experiments as functions of MgO content. At constant pressure, the MgO content of peridotite partial melts is positively correlated with temperature (e.g. Hirose & Kushiro, 1993; Baker & Stolper, 1994; Walter, 1998; Falloon et al., 1999), but our trends in glass composition are smoother when plotted against MgO content rather than temperature, presumably reflecting uncertainties in temperature in our experiments. Glass compositions from 1270 to 1390°C vary from basaltic (11–12 wt % MgO) at the low end of the temperature range to picritic (17 wt % MgO) at the high end. Over this range of MgO contents, SiO2, Cr2O3, and FeO* concentrations increase with increasing MgO, whereas TiO2, Al2O3, and Na2O abundances all decrease with increasing MgO. CaO contents display a more complex pattern, first increasing with increasing MgO content until cpx-out, after which the CaO content of melt decreases with increasing MgO content (and temperature). Except for SiO2 and TiO2, the trends displayed by the glasses are qualitatively similar to those observed in near-solidus melting experiments on fertile mantle compositions at 10 kbar (e.g. Hirose & Kushiro, 1993; Baker & Stolper, 1994; Baker et al., 1995; Kushiro, 1996; Hirschmann et al., 1998a). We will return to a more detailed comparison of 10 kbar depleted and fertile peridotite melting trends in the discussion section below.
Solid phase compositions
Olivine
The mg-number [defined as 100 MgO/(MgO + FeO*), molar] of ol in our experiments increases from 89·8 at 1270–1275°C to 90·8 at 1390°C. 
values [
, where liq indicates liquid] are between 0·28 and 0·36, with most values between 0·30 and 0·34. For comparison with our experimental results, we have compiled mineral and glass compositions from 142 experiments conducted at 9–12 kbar and 1200–1400°C that contain the residual solid phases ol + opx ± cpx ± sp ± plagioclase (Takahashi & Kushiro, 1983; Falloon & Green, 1987; Kinzler & Grove, 1992a; Hirose & Kushiro, 1993; Baker & Stolper, 1994; Kushiro, 1996; Gaetani & Grove, 1998; Hirschmann et al., 1998a; Wagner & Grove, 1998; Falloon et al., 1999, 2001; Pickering-Witter & Johnston, 2000; Schwab & Johnston, 2001; M. B. Baker, unpublished data, 1996). In Fig. 5 we plot our 
values and those from our literature dataset as a function of the alkali content of the coexisting glasses [this figure also includes the 10 kbar ol–liq KD,Fe*-Mg values of Ulmer (1989) and Draper & Green (1999)]. Our dataset lies within and to the low-alkali end of the field defined by the literature values (Fig. 5). The decrease in with increasing alkalis in the melt in our data is also manifest in the main trend of the literature values (Fig. 5). The correlation between and alkalis is also seen in 1 atm experiments (e.g. Baker et al., 1996) and reflects, in part, the effect of alkalis on increasing the Fe3+/Fe2+ ratio in silicate melts at constant temperature, pressure, and fO2 (e.g. Sack et al., 1980), although the Fe2+–Mg olivine–liquid exchange reaction also appears to be sensitive to bulk composition (e.g. Gee & Sack, 1988; Kushiro & Mysen, 2002).
|
Orthopyroxene
With rising temperature, opx mg-numbers increase from 90·0–90·5 near the solidus (1270–1275°C) to
91·5 at 1370–1390°C. Over this temperature range, 
values increase with temperature from 0·29 to 0·32 and are also inversely correlated with Na2O + K2O in the coexisting glasses. Figure 6 shows that our vs glass alkali trend is consistent with a large body of experimental data. 
values (CaOopx/CaOliq, both in wt %) from our experiments are consistent with values from other high-pressure experiments containing opx + ol ± cpx-saturated liquids. from our cpx-free runs show a slight decrease with rising temperature, but the mean values from the lherzolite- and harzburgite-saturated experiments are identical at 0·16 ± 0·02 (1
). These means also overlap with the mean values for cpx-bearing and cpx-free experiments from the literature (0·20 ± 0·04 and 0·16 ± 0·05, 1, respectively).
|
The opx–liq partition coefficient for Al2O3,
, shows a weak negative correlation with temperature in our experiments, decreasing from 0·22–0·28 at 1270–1275°C to 0·15–0·23 at 1370–1390°C; the mean values from the cpx-bearing and cpx-free experiments overlap at the 1 level (0·24 ± 0·06 and 0·20 ± 0·03). values from the literature show considerable scatter, with ranges of 0·14–0·51 (cpx-bearing) and 0·10–0·56 (cpx-free). The mean values for the two sets of literature data (0·30 ± 0·09 and 0·25 ± 0·07, 1) overlap with our lherzolite- and harzburgite-saturated experimental values at the 1 level.
The opx–liq partition coefficients for Na2O, 
, from our experiments are positively correlated with temperature: ranges of values are 0·02–0·04 (0·028 ± 0·008, 1) for cpx-bearing assemblages and 0·03–0·07 (0·041 ± 0·012, 1) for cpx-free assemblages. However, the uncertainty on individual determinations is large (25–50%) because of the low concentration of Na2O in the orthopyroxenes (0·02–0·05 wt %; Table 3). The mean value for all 9–12 kbar experiments compiled from the literature is 0·04 ± 0·01 (1), which overlaps with the mean values from our experiments (the literature mean includes all experiments containing glasses in the same range of Na2O contents as our experiments, regardless of whether they contain cpx). The literature data also suggest a positive correlation with temperature, although again uncertainties on individual determination are large because of the low Na2O contents of the opx.
values in our experiments are 0·03–0·20 (0·11 ± 0·05, 1) for cpx-bearing experiments and 0·07–0·18 (0·12 ± 0·04, 1) for cpx-free experiments and do not correlate with temperature. Again, individual values have large uncertainties (25–100%), as a result of the low concentrations of TiO2 in the orthopyroxenes (0·01–0·08 wt %). Mean literature values from lherzolite- and harzburgite-saturated experiments are 0·20 ± 0·07 and 0·15 ± 0·08 (both 1) and overlap with our results.
Clinopyroxene
values in our experiments are 0·29–0·36 with a mean of 0·32 ± 0·02 (1). Figure 7a shows that these values are consistent with those from the literature and both sets of data display a weak inverse correlation with alkalis in the coexisting glass. Figure 7b compares 
values from our experiments and from cpx–liq pairs saturated with opx from the literature with the weight percent sum of Na2O and K2O in the coexisting glasses. values from the DMM1 cpx–liq pairs are consistent with the literature values; the positive correlation between and Na2O + K2O reflects the effect of alkalis on the activity coefficient of calcium complexes in the melt (e.g. CaSiO3; Hirschmann et al., 1999a).
|
values from our experiments increase from 0·28 at 1270°C to 0·34 at 1300°C and fall in the field of literature values. Mean calculated for our data and for the compilation of literature data overlap at the 1 level (0·32 ± 0·03 for our DMM1 results vs 0·36 ± 0·08 for cpx–liq pairs from other peridotite melting experiments). 
values from our experiments and for the compilation of literature data have identical means (0·16 ± 0·04, 1). values do not correlate with temperature in either dataset. 
values for the DMM1 experiments have large uncertainties (25–75%), reflecting the low concentrations of titanium in the clinopyroxenes. The mean value for our data is 0·22 ± 0·05 (1). This is lower than the mean of the literature values (0·38 ± 0·13, 1), although there is considerable overlap between the two datasets. As discussed by Hirschmann et al. (1999a), at constant pressure is positively correlated with the SiO2 content of melt and negatively correlated with alkali content. The offset in between the two sets of data thus plausibly reflects the generally lower SiO2 and Na2O + K2O contents of the DMM1 glasses relative to the glasses in the 9–12 kbar literature compilation (47·8–48·9 vs 46·9–55·5 wt % SiO2 and 0·8–2·1 vs 0·8–7·8 wt % Na2O + K2O, respectively).
Spinel
Although the spinel mg-number is a complex function of a number of variables, the most important are the mg-numbers of the coexisting phases, the partitioning of Al and Cr between sp and these phases, and the fO2. Figure 8a shows that the mean sp compositions in our experiments overlap the inverse correlation between mg-number and cr-number [100Cr/(Cr + Al), molar] displayed by spinels from other 9–12 kbar peridotite melting studies. In addition, the Al2O3 contents of sp and coexisting liquid in our experiments (Fig. 8b) are consistent with the trend of data from other 9–12 kbar experiments.
|
|
DISCUSSION |
|---|
|
TOPABSTRACTINTRODUCTIONEXPERIMENTAL AND ANALYTICAL...RESULTSDISCUSSIONCONCLUSIONSREFERENCES |
|---|
Melting reactions and melt productivity
Proportions of glass and crystalline phases in each experiment were calculated by mass balance using an algorithm that incorporates uncertainties in both the bulk composition and the compositions of the coexisting phases (Albarède & Provost, 1977). Liquid fractions increase from
0·02–0·04 at 1270–1275°C to 0·17 at 1390°C (Fig. 9). A weighted linear fit of melt fraction (F) vs temperature (T), using errors on both F (Table 2) and T (±15°C; Reed, 1992), of the spinel lherzolite-saturated glasses yields a solidus temperature of 1272 ± 11°C (1), within our experimental bracket of 1250–1270°C (because of its uncertain temperature, run 23T was not included in the fit). Weight fractions of ol and opx change from 0·64 and 0·27 at 1270°C to 0·69 and 0·19 at 1390°C. Weight fractions of cpx and sp decrease from 0·06 and 0·005 at the solidus to 0 at 1300–1310°C.
|
Figure 10 shows the fractions of residual solid phases plotted against melt fraction. The lines are weighted least-squares fits to the ol + opx + cpx + sp-saturated and ol + opx-saturated melts, respectively. As discussed by Baker & Stolper (1994), multiplying the slopes of these lines by -1 yields the melting reaction coefficients. In the spinel lherzolite field, the resulting equation is
 |
|
|
At temperatures above cpx + sp-out, the reaction for melting of DMM1 is based on the best-fit lines in Fig. 10 and is
 |
Isobaric melt productivity, (
F/T)P, is an important term in the equation describing melt production during adiabatic mantle upwelling (e.g. Asimow et al., 1997). By analogy with simple systems and based on MELTS (Ghiorso & Sack, 1995) calculations on peridotite compositions (including DMM1 and MM3), we expect a steady increase in (F/T)P up to cpx-out (Hirschmann et al., 1999b). At cpx-out, a discontinuous drop in productivity is expected, and the T vs F trend is expected to remain concave-up with further rises in temperature (Asimow et al., 1997; Hirschmann et al., 1999b). To determine average values for (F/T)P for ol + opx + cpx + sp-saturated and ol + opx-saturated melting of the DMM1 composition, we fitted our T(°C) vs F data (Table 2 and Fig. 9; excluding run 23T) to straight lines below and above cpx-out. The best-fit average melt productivities are 0·3 ± 0·2%/°C (1) for cpx-saturated melting and 0·08 ± 0·01%/°C (1) above cpx-out. For comparison, the fertile MM3 peridotite composition has best-fit productivities of 0·21 ± 0·02%/°C (1) for cpx-saturated melting and 0·12 ± 0·01%/°C (1) above cpx-out (Baker & Stolper, 1994; Hirschmann et al., 1998a; M. B. Baker, unpublished data, 1996).
Both the DMM1 and MM3 compositions show the expected discontinuous drop in productivity associated with the exhaustion of cpx (Asimow et al., 1997). However, given the large uncertainties associated with the isobaric melt productivity of DMM1, we cannot evaluate whether (F/T)P for DMM1 near the solidus is larger than that of MM3 (for identical low values of F where both compositions have the same residual phase assemblage), as expected based on the calculations of Hirschmann et al. (1999b). Moreover, the T–F data for DMM1 in the spinel lherzolite field (Fig. 9) are not adequate to constrain whether (F/T)P is constant or dependent on F, as expected from analysis of simple systems and MELTS calculations (Asimow et al., 1997; Hirschmann et al., 1999b). However, for MM3 the T–F data in the spinel lherzolite field cannot be fitted with a concave-upward curve (Fig. 9 inset) and the 15 kbar peridotite melting experiments of Robinson et al. (1998) also yield nearly constant melt productivities in the spinel lherzolite field. In contrast, 10 kbar experiments on the FER-E and INT-E compositions (Pickering-Witter & Johnston, 2000; Schwab & Johnston, 2001) suggest that melt productivities do indeed increase with rising temperature in the spinel lherzolite field. We do not have an explanation for the diversity of isobaric productivity functions that have been observed experimentally for lherzolite melting.
The effects of peridotite composition on 10 kbar near-solidus partial melts
Partial melts generated from depleted peridotite differ in several important ways from partial melts of fertile peridotite. Figure 11 illustrates some of these differences by comparing liquid compositions from this study as functions of melt fraction with trends of 10 kbar partial melt compositions from experiments on fertile peridotites. We have included the glass data of MM3 (Baker & Stolper, 1994; Baker et al., 1995; Hirschmann et al., 1998a; M. B. Baker, unpublished data, 1996), KLB-1 (Hirose & Kushiro, 1993), PHN1611 (Kushiro, 1996), and FER-E (Pickering-Witter & Johnston, 2000).
|
At melt fractions of 0·1–0·2, SiO2 contents of the DMM1 experimental glasses overlap those of experimental glasses from the fertile peridotite compositions (Fig. 11a). At lower melt fractions, however, partial melts of all the fertile peridotites display increases in their silica contents, whereas the silica contents of the DMM1 partial melts continue to decrease with decreasing F. This difference in the behavior of SiO2 in near-solidus melts of fertile and depleted peridotite reflects the effects of alkalis on the liquid silica activity coefficient [at 1 atm, these effects have been discussed by Kushiro (1975) and Ryerson (1985)]. As emphasized by Hirschmann et al. (1998a), the silica activity of opx + ol-saturated liquids varies little with melt fraction, but the activity coefficient of silica in the liquids decreases significantly with increasing alkali content. Consequently, as Na2O + K2O concentrations of liquids increase to high levels with decreasing F in melting experiments on fertile lherzolites, the SiO2 concentrations likewise increase. However, the alkali contents of glasses in the DMM1 experiments are low relative to partial melts of more fertile peridotite, and the effects of increasing alkali contents on the SiO2 contents of partial melts would only be expected to be expressed at very low melt fractions. For example, Fig. 11e shows that at F
0·02, the Na2O concentration in the DMM1 melt is only 2 wt %, much lower than partial melts of more fertile peridotites such as MM3, where at F 0·02 the Na2O contents can reach concentrations >7 wt %. It should be noted also that the SiO2 contents of the partial melts of the fertile peridotites begin to increase only when Na2O contents reach 3 wt %. Thus the lack of increasing SiO2 contents in the DMM1 glasses at melt fractions down to 0·02 (where Na2O content is only 2 wt %) is consistent with the results from fertile peridotites. MELTS calculations predict that for the DMM1 composition alkali contents build up sufficiently to cause increases in silica contents only at melt fractions lower than 1 wt % (Hirschmann et al., 1998b).
Although we have emphasized the role of Na2O, which is the most abundant alkaline component, concentrations of K2O in the melts of fertile peridotite are also important: at F 0·02, the 10 kbar partial melt of MM3 contains 0·43 wt % K2O (Hirschmann et al., 1998a); at F 0·04, the 10 kbar partial melt of PHN1611 contains 3·2 wt % K2O (Kushiro, 1996). Overall, such high values of Na2O + K2O can lead to significant changes in melt chemistry in addition to increasing SiO2 contents (Fig. 11a), such as less incompatible behavior of Ti and other highly charged incompatible trace elements (e.g. U4+, Th4+, Ta5+, Hf4+) in cpx and decreasing ol–liq, opx–liq, and cpx–liq values (e.g. Baker et al., 1995; Hirschmann et al., 1998a, 1999a; see Figs 5, 6, and 7). The deviation in the F vs FeO* trend of DMM1 partial melts relative to the trends of melts of fertile peridotites at low F (Fig. 11c) also reflects the low alkali contents in the near-solidus melts of DMM1 relative to those of fertile peridotites. At constant melt fraction a decrease in the alkali content in an ol + opx-saturated liquid decreases the stability of ol (the most Fe-rich silicate phase) relative to opx (Kushiro, 1975), which in turn increases the FeO* content in the melt.
Alumina contents of partial melts from the five peridotite compositions plotted in Fig. 11b decrease with increasing melt fraction, but partial melts of DMM1 contain 2–4 wt % less Al2O3 than melts from the more fertile bulk compositions, with the deviation increasing slightly with increasing F. The lower Al2O3 contents in the partial melts of DMM1 compared with those of the fertile peridotites at all melt fractions probably reflect several different factors, including the 18–80% lower alumina content of DMM1 relative to the fertile peridotites and the effect of alkali content on the activity coefficient of Al2O3 in peridotite melts (Hirschmann et al., 1999a).
For DMM1, CaO contents in cpx-saturated liquids increase from 13 to nearly 14 wt % with increasing melt fraction and then decrease with increasing F beyond the exhaustion of cpx (Fig. 11d). In contrast, the CaO contents of cpx-saturated partial melts of the fertile peridotites increase by up to a factor of two with increasing melt fraction from F 0·02 up to the point of cpx exhaustion, after which the CaO contents of the partial melts decrease rapidly with further melting. For DMM1 and the fertile peridotites, the maximum CaO contents occur at different melt fractions (0·09–0·10 for DMM1, 0·10–0·30 for the fertile peridotites); this is because the maximum CaO content occurs at the exhaustion of cpx from the solid residue, which occurs at different values of F for peridotites with different initial cpx contents. An important result is that the maximum CaO contents in the DMM1 partial melts are 1–2·5 wt % higher than the maximum values of melts from the more fertile compositions even though the DMM1 bulk composition has 1–3 wt % less CaO than the more fertile mantle compositions. The elevated CaO contents of melts of the depleted DMM1 composition relative to partial melts of the fertile compositions are even more dramatic at low F; for example, at F 0·02, partial melts of DMM1 contain 13 wt % CaO compared with 6·5 wt % for partial melts of MM3. The explanation for this significant difference again lies in the much higher alkali contents of near-solidus melts of fertile peridotite relative to melts of DMM1 (Hirschmann et al., 1999a). The alkali-rich partial melts of the fertile peridotites have higher CaO activity coefficients relative to alkali-poor melts of depleted lherzolite. The higher CaO activity coefficients in turn result in lower concentrations of CaO in alkali-rich, cpx-saturated liquids compared with alkali-poor, cpx-saturated liquids, as the activity of CaO is roughly fixed in lherzolite-saturated melts (Hirschmann et al., 1999a). It should be noted that the difference in CaO contents of cpx-saturated melts of fertile and depleted peridotites, first discussed quantitatively by Hirschmann et al. (1999a), is somewhat unintuitive. The fertile peridotites, although richer in CaO, generate partial melts that are poorer in CaO, and experimental confirmation of this phenomenon is an important result of this study.
CaO/Al2O3 ratios in partial melts of DMM1 and four fertile peridotite compositions are plotted in Fig. 11f. These ratios lie in a narrow band for the melts of fertile peridotite, increasing from 0·4 to 0·7 as F increases from 0·02 to 0·15. This simple trend has led to the use of CaO/Al2O3 in primitive MORBs as a proxy for the extent of melting (e.g. Niu & Batiza, 1991; Niu & Hékinian, 1997). However, for melts of depleted peridotites, the CaO/Al2O3 ratios of low-degree partial melts of DMM1 are as high as those of high-degree melts of fertile peridotite and increase slowly with increasing F (Fig. 11f; Hirschmann et al., 1999a). The high CaO/Al2O3 ratios and their small range in the DMM1 melts reflect primarily the low degree of variation in CaO contents before cpx exhaustion compared with CaO in partial melts of fertile peridotite; above cpx-out, CaO and Al2O3 trends generally have similar slopes and thus less variation in CaO/Al2O3 ratios (Fig. 11b and d). The key point of Fig. 11f, as anticipated by Hirschmann et al. (1999a), is that variations of CaO/Al2O3 with F in partial melts depend strongly on source peridotite composition; thus CaO/Al2O3 must be used cautiously as an indicator of degree of partial melting in a series of basalts.
Comparison of experiments and mantle melting models
A wide range of models have been developed for calculating the compositions of partial melts of peridotite under upper-mantle conditions (e.g. McKenzie & Bickle, 1988; Niu & Batiza, 1991; Kinzler & Grove, 1992a, 1992b; Langmuir et al., 1992; Ghiorso & Sack, 1995; Kinzler, 1997; Ghiorso et al., 2002; Longhi, 2002). Most high-pressure experiments that have been available to calibrate these models have come from melting of fertile peridotites and pyroxene-enriched compositions, so the degree to which they can model the melting behavior of more depleted compositions is uncertain. Yet, as emphasized in the Introduction, current thinking about the dynamics of partial melting processes suggests that melt extraction occurs throughout the melting process, producing variously depleted peridotite compositions in much of the MORB source region. Therefore accurate modeling of MORB generation must handle well the melting behavior of depleted peridotitic compositions. Our experimental results provide an opportunity to evaluate the performance of these models on a depleted peridotite composition far removed from the data used to calibrate them.
Figure 12 shows melt fractions vs major element concentrations for the experimental glasses presented in this paper and compares them with the results of calculations on the DMM1 bulk composition using four of the available models of peridotite melting. Calculations are shown for the thermodynamically based MELTS (Ghiorso & Sack, 1995) and pMELTS (Ghiorso et al., 2002) algorithms. pMELTS is a better calibrated version of the MELTS model, based on a much larger high-pressure experimental dataset for calculating crystal–liquid equilibria in silicate systems by thermodynamic potential minimization. pMELTS also incorporates a revised thermodynamic model and equation of state for the silicate melt. Also shown in Fig. 12 are liquids calculated using BATCH (Longhi, 2002), based on parameterizations of multiply saturated phase boundaries and mineral–melt distribution coefficients, and the expressions of Langmuir et al. (1992), based on selected mineral–melt distribution coefficients parameterized using T, pressure (P), and liquid compositions. It should be noted that the expressions of Langmuir et al. (1992) cannot calculate liquid SiO2, Al2O3, or CaO contents.
|
The four models shown in Fig. 12 were chosen because they could handle the DMM1 composition with either little [the expressions of Langmuir et al. (1992)] or no modification (MELTS, pMELTS, BATCH). Although we show the MELTS trends in Fig. 12 for completeness, we concentrate on the comparison between the calculated pMELTS liquids and DMM1 glass compositions because the MELTS–DMM1 comparison was discussed by Hirschmann et al. (1998b). The calculations for the Langmuir et al. model were performed using a spreadsheet provided by C. H. Langmuir (personal communication, 1998). In using this spreadsheet, we adjusted the variable that controls the relationship between temperature and melt fraction so as to match the experimental solidus temperature of DMM1. We also entered the modal abundance of cpx for DMM1 at its solidus (from the intercept at F = 0 of the linear fit to cpx abundance shown in Fig. 10) and changed the cpx melting reaction coefficient in the spreadsheet so as to reproduce the trend of F vs cpx abundance shown by our results in Fig. 10.
For SiO2 (Fig. 12a) above cpx-out, the liquids calculated with pMELTS match the glasses produced by melting DMM1. Although those liquids calculated with BATCH have the correct slope, they are on average 1 wt % too low. Below cpx-out, both calculated trends have approximately constant liquid SiO2 contents for F 0·02, but the pMELTs-calculated liquids are on average 0·8 wt % too high in SiO2, and the BATCH-calculated liquids are on average 0·3 wt % too low. The pMELTS trend shows a slight upturn in liquid SiO2 content at the lowest calculated melt fractions; as described above, this effect reflects the elevation in alkalis at very low melt fractions. It should be noted that the pMELTS-calculated occurrence of cpx-out at F 0·09 corresponds well to the occurrence of cpx-out in our experiments (F 0·09–0·10), indicating that pMELTS is successful at reproducing a key aspect of the melting behavior. For BATCH, cpx-out occurs at F 0·08. MELTS-calculated liquids are displaced to significantly lower silica contents than the experimental data; this inaccuracy in MELTS is well known and reflects the overstability of opx and understability of ol for peridotite compositions (e.g. Hirschmann et al., 1998b; Ghiorso et al., 2002).
Al2O3 contents of partial melts of DMM1 calculated using MELTS, pMELTS, and BATCH are all displaced downwards, but are parallel to the experimental data (Fig. 12b). Mean absolute deviations between the experimental glass compositions and the liquid compositions calculated using pMELTS and BATCH are 1·37 and 1·16 wt % Al2O3, respectively. For pMELTS, the main cause of this systematic deviation is the 3·5–4·5 times higher sp contents in the calculated near-solidus residues compared with the experimental modes, and the fact that above cpx-out, sp is still a residual phase in the pMELTS calculations over the range of F values plotted in Fig. 12 (recall that sp- and cpx-out occur over the same melting interval in our experiments). The systematic deviations in liquid Al2O3 contents in the BATCH calculations reflect (at the same melt fractions) the higher Al2O3 contents in the calculated cpx and opx compared with the Al2O3 contents in the our experimental pyroxenes. Although F vs liquid Al2O3 contents from both models show systematic deviations from the experimental data, both models predict the approximately linear decrease in liquid Al2O3 content with increasing melt fraction (Fig. 12b).
At the same melt fraction, FeO* and MgO contents of DMM1 partial melts calculated using pMELTS and BATCH are similar, but their F vs FeO* and MgO trends plot above, but parallel to, the experimental glass compositions (Fig. 12c and d). The mean absolute deviations are 0·67 wt % FeO* and 1·65 wt % MgO for the pMELTS-calculated trends and 0·76 wt % FeO* and 1·55 wt % MgO for the BATCH-calculated trends. Both models reproduce the slope changes in F vs FeO* and MgO at cpx-out. The Langmuir et al. model predicts liquids with lower FeO* and similar MgO contents to the experimental glasses. The mean absolute deviations in FeO* and MgO are 0·52 and 0·35 wt %, respectively. In contrast to the pMELTS and BATCH calculations, the F vs FeO* and MgO trends of these Langmuir et al. calculations show only a subtle change in slope near cpx-out; the near linearity of the Langmuir et al. FeO* and MgO trends reflects the model constraint that the Mg/Fe ratio of residual olivine and the Mg/Fe ratio of the bulk residue are equal. The offsets between the MELTS trends and the experimental data in the F–MgO and F–FeO* space are, as for SiO2, the result of overstabilization of opx and understabilization of ol in the MELTS model (Hirschmann et al., 1998b).
Figure 12e shows that the liquid compositions calculated with pMELTS, BATCH, and MELTS all parallel the F–CaO trend of the experimental DMM1 glasses, with BATCH matching most closely (as noted above, the Langmuir et al. model does not provide liquid CaO contents). Mean absolute deviations for the pMELTS and BATCH trends are 1·05 and 0·58 wt % CaO, respectively. With increasing melt fraction from the solidus up to cpx-out, the calculated liquid CaO contents increase gradually, as do the CaO concentrations in the experimental glasses. Above cpx-out, the CaO contents of the experimental glasses and all the calculated liquid compositions decrease rapidly with increasing F. Although BATCH reproduces the F vs CaO partial melting trend of DMM1, it is less successful in reproducing the low CaO contents in experimental glasses close to the fertile peridotite solidus; for example, using BATCH with MM3 as the starting composition, the calculated 10 kbar CaO content of the liquid at F = 0·03 is 2·3 wt % higher than the experimental glass (Baker & Stolper, 1994; Hirschmann et al., 1998a; M. B. Baker, unpublished data, 1996). In contrast, for the same conditions, the pMELTS-calculated liquid is only 0·5 wt % low in CaO relative to the experimental results.
Although DMM1 was not included in the calibrations of either pMELTS or BATCH, both of these programs reproduce the essential characteristics of the melt fraction vs major element trends observed in the experimental glasses. The successes of these calculations include their generally close matches to the absolute values of the major oxide concentrations, their successful predictions of the overall shapes of the observed trends, their predictions of the changes in slope in the F vs SiO2, FeO*, MgO, and CaO trends that occur at or near cpx-out, and their predictions of cpx-out at values of F either within (pMELTS) or close to (BATCH) the experimental bracket. Nevertheless, liquid compositions calculated by all of the models display systematic offsets relative to the experimental glass compositions. Finally, although the Langmuir et al. model provides the best matches to the F vs FeO* and MgO glass trends for DMM1, the absence of expressions for SiO2, Al2O3, or CaO and the necessity of adjusting the subsolidus mode and melting reaction coefficient for cpx to optimize the calculated liquid compositions limit the usefulness of this approach.
Variations in solidus temperature as a function of peridotite bulk composition
McKenzie & Bickle (1988), Langmuir et al. (1992), Iwamori et al. (1995), Herzberg et al. (2000), and Hirschmann (2000) have compiled and parameterized available data on the solidi of fertile peridotite compositions in pressure–temperature space. The parameterizations of these studies are broadly consistent, and the rise in solidus temperature (Ts) with increasing pressure from 1 bar to 100 kbar is reasonably well constrained. The relationship between Ts and bulk composition at any given pressure, however, is less well quantified. Qualitatively, solidus temperature at constant pressure decreases in peridotites with increasing concentrations of incompatible elements (of which Na2O, K2O, and H2O are likely to be the most important) and decreasing mg-number (Herzberg et al., 2000; Hirschmann, 2000). Deconvolving the influence of these two compositional effects is generally difficult, as those peridotites with high alkali contents also tend to have relatively low mg-numbers. Linear fits of bulk peridotite Na2O + K2O vs Ts and peridotite mg-number vs Ts by Herzberg et al. (2000) suggest that at 15 kbar, increasing the alkali content from 0·2 to 0·4 or decreasing the mg-number from 90 to 80 will produce a 30–35°C drop in Ts. However, mg-numbers and bulk alkali contents in the peridotite compositions fit by Herzberg et al. (2000) show an inverse correlation (correlation coefficient = 0·68), making it difficult to evaluate independently these two parameters.
Using an idealized ternary system as a framework for considering the effects of composition on solidus temperature, Hirschmann (2000) was able to model the negative correlation between solidus temperatures and incompatible element content (e.g. Na2O + K2O). His analysis, based on the cryoscopic expression (e.g. Carmichael et al., 1974, pp. 170–173), showed that the bulk partition coefficients and concentrations of incompatible elements at the peridotite solidus are important parameters influencing the solidus temperature. In particular, the more incompatible an element, the larger its effect on the solidus (assuming melting is a continuous reaction). He further showed that the conclusions drawn from the idealized ternary system could be applied to understanding variations in the solidus temperatures for MM3 (Baker & Stolper, 1994; Hirschmann et al., 1998a; M. B. Baker, unpublished data, 1996), FER-B, FER-D, and FER-E (Pickering-Witter & Johnston, 2000) at 10 kbar, and MPy and Tinaquillo lherzolite (Robinson et al., 1998) at 15 kbar.
We have extended the analysis of Hirschmann (2000) by considering eight peridotites for which the 10 kbar solidus has been determined and by including the effect of P2O5 in addition to Na2O and K2O. Because P2O5 contents in the bulk compositions of Pickering-Witter & Johnston (2000) and Schwab & Johnston (2001) are unknown, we have not included the solidus temperatures for these bulk compositions in our parameterization. The complete dataset includes DMM1 (this study), MM3 (Baker & Stolper, 1994; Hirschmann et al., 1998a; M. B. Baker, unpublished data, 1996), syn-MM3 (Falloon et al., 1999), KLB-1 and HK-66 (Takahashi & Kushiro, 1983; Hirose & Kushiro, 1993), HPy-40 (Jaques & Green, 1980), PHN1611 (Kushiro, 1996), and DW (Bertka & Holloway, 1994; this is a model Martian mantle composition). Bulk compositions are listed in Tables 1 and 5. These peridotites have solidus temperatures ranging from 1193 to 1272°C. Their concentrations of Na2O, K2O, and P2O5 are highly variable, ranging from 0·06 to 0·95, 0 to 0·22, and 0 to 0·16, and their mg-numbers range from 90·5 to 75·0. Experimentally determined solidus assemblages at 10 kbar include both spinel lherzolite and plagioclase + spinel lherzolite.
|
Most of these peridotites lack the necessary experimental data to estimate near-solidus liquid compositions. Thus, we have used pMELTS (Ghiorso et al., 2002) to calculate 10 kbar near-solidus liquid compositions for these eight peridotites. The Na2O, K2O, and P2O5 contents of the pMELTS-calculated liquids at 10 kbar, QFM, and F = 0·005 range from 2·8 to 11·6, 0 to 11·3, and 0 to 8·6 wt %, respectively, and all of the calculated melts coexist with ol, opx, cpx, and sp. At F = 0·005, alkali feldspar is predicted to be a stable phase in the three peridotite compositions with the highest bulk alkali contents (PHN1611, HK-66, and HPy-40; these peridotites have bulk Na2O + K2O contents of 0·47–1·17 wt %). pMELTS also predicts that whitlockite is stable at F = 0·005 in the five bulk compositions that have >0·017 wt % P2O5 (PHN1611, KLB-1, HPy-40, HK-66, and DW). It should be noted that, in this respect, results from pMELTS calculations differ from the expectations of the idealized ternary system analysis of Hirschmann (2000), in which it was assumed that the incompatible components never become a major structural constituent of a crystallizing phase and that the concentrations in near-solidus melts are controlled entirely by partitioning into the major peridotite phases.
Figure 13a shows that the experimentally determined solidus temperatures are well correlated with the concentrations of alkalis plus phosphorus in the calculated near-solidus melts. In addition to calculating liquid compositions at F = 0·005, we continued each pMELTS calculation down temperature and determined the model solidus for each peridotite (Ts,pMELTS). Figure 13b shows that the calculated Ts,pMELTS values are correlated with the experimentally determined solidus temperatures, although the model temperatures are lower than the experimental determinations by 25–70°C (the deviations are crudely correlated with the sum of the bulk alkali and phosphorus contents). The differences in temperature between the experimental and model solidi may reflect in part the difficulties in observing minute amounts of melt in experimental charges (each experimentally determined solidus temperature may be systematically too high), but they also probably reflect inaccuracies in the thermodynamic models that underpin pMELTS. The important point is that experimentally determined solidus temperatures are correlated with solidus temperatures and compositions of low-degree melts predicted by pMELTS, allowing us to utilize the thermodynamically consistent results of pMELTS as guides to the major factors controlling solidus temperature.
|
We have fitted the data shown in Fig. 13a to an equation of the form Ts(°C) = a1 + a2(Na2O) + a3(K2O) + a4(P2O5). The choice of this linear function is based on the simple case of solidus depression of a single phase by an incompatible element at infinite dilution (e.g. Hirschmann, 2000). That concept represents an oversimplification, and the presence of minor phases in the pMELTS calculations also complicates the application of this equation. The best fit (a1 = 1295·8, a2 = –3·628, a3 = –7·218, a4 = –5·100) reproduces the experimental solidus temperatures with a mean deviation of 3°C (less than half of the temperature uncertainty associated with any of the Ts determinations; Fig. 14). The largest temperature mismatch is for HK-66 (12°C), but the calculated Ts value is still within the experimental solidus bracket for this bulk composition (Fig. 14). Given the compositional complexities of natural peridotites, the simple functional form of the fit, and the difficulties in determining the solidus experimentally, the quality of the fit is actually somewhat surprising.
|
Although the peridotite bulk compositions span a range of mg-numbers (Tables 1 and 5), inclusion of an mg-number term in the equation does not improve the fit (i.e. the mean temperature deviation remains
3°C). It should be noted, however, that the model Martian mantle composition, DW, is the peridotite composition with the lowest mg-number, but it also has the highest P2O5 content and one of the highest Na2O contents in our dataset. Thus, as stated above, mg-number and incompatible element concentrations tend to be correlated in peridotites, making it difficult to separate the effect of mg-number from those of alkalis. However, to evaluate the effect of mg-number on solidus temperature, we increased the mg-number of DMM1 (keeping the total number of moles of FeO and MgO constant) from 89·9 to 92 and used pMELTS to calculate the solidus temperature, which changed by only 2°C. This suggests that over the limited range of mg-numbers observed in the vast majority of spinel peridotite samples (88–92; Herzberg et al., 1988), mg-number will have no effect on Ts at 10 kbar.
In our parameterization, the coefficients for K2O and P2O5 have similar absolute values that are larger than that for Na2O. The larger effects on lowering of solidus temperature of the more incompatible components is precisely what was anticipated by Hirschmann (2000), suggesting that this simple fit has some validity as a predictive tool. Our expression for Ts suggests that the 10 kbar solidus temperature for an alkali- and phosphorus-free peridotite will be 1296°C. This is broadly consistent with the 1287–1289°C range for the 10 kbar solidus temperatures for the model peridotite compositions DF80 and JDS (with mg-numbers of 89·6 and 90·0, respectively) of Gudfinnsson & Presnall (2000) in CMASF based on mass balance constraints and their experimentally calibrated expressions for liquid and solid compositions in the plagioclase lherzolite field. Many abyssal peridotites have bulk compositions more strongly depleted in alkalis than DMM1 (Baker & Beckett, 1999; Fig. 1), and using our Ts expression and the abyssal peridotite compositions from Baker & Beckett (1999) that have the lowest and highest MgO contents, we predict that 10 kbar abyssal peridotite solidus temperatures would lie in the range of 1278–1295°C.
|
CONCLUSIONS |
|---|
|
TOPABSTRACTINTRODUCTIONEXPERIMENTAL AND ANALYTICAL...RESULTSDISCUSSIONCONCLUSIONSREFERENCES |
|---|
(1) We conducted melting experiments on a depleted peridotite composition (DMM1) at 10 kbar using a modification of the diamond-aggregate melt-extraction technique in which diamonds were replaced by vitreous carbon spheres. Several approaches were used to demonstrate that melts segregated into aggregates of vitreous carbon remained in chemical contact with the bulk charge over the course of the experiments down to melt fractions of
0·02–0·04 and closely approached the compositions of equilibrium partial melts of the source composition. (2) The most important differences between the compositions of partial melts of DMM1 and fertile peridotites at 10 kbar occur at low melt fractions. The differences in SiO2 and CaO contents between low-degree melts of depleted and fertile peridotite are particularly striking. In contrast to results for fertile peridotites, SiO2 contents of partial melts of DMM1 do not increase as melt fractions decrease from 0·1 to 0·02, resulting in differences of up to 6–8 wt % SiO2 between partial melts of fertile and depleted lherzolites at 2% melting. Likewise, CaO contents of partial melts of depleted lherzolites are higher at cpx-out than partial melts of fertile lherzolites, and they decrease much less with decreasing melt fraction, resulting in 6–8 wt % more CaO in the partial melts of depleted lherzolites at 2% melting. Additionally, FeO* and MgO contents do not decrease as dramatically in the partial melts of DMM1 as they do in partial melts of fertile peridotites over this range of melt fractions, and Na2O contents are substantially lower in melts of DMM1. We infer that the substantially lower alkali contents in the near-solidus partial melts of DMM1 relative to those of fertile peridotite account for the differences in their SiO2, FeO*, MgO, and CaO concentrations.
(3) The melting reaction for DMM1 at 10 kbar in the spinel lherzolite field is 0·57 opx + 0·72 cpx + 0·04 sp = 0·35 ol + 1 liq. The reaction coefficients are similar to those determined for ol + opx + cpx + sp-saturated melting of fertile peridotite and analog mantle compositions in the CaO + MgO + Al2O3 + SiO2 + FeO system. In particular, cpx dominates the melting reaction over a wide range of peridotite compositions, and the cpx reaction coefficient is insensitive to the modal abundance of cpx in the lherzolite. Above cpx- and sp-out, DMM1 melts via the reaction 1·24 opx = 0·24 ol + 1liq, which is also qualitatively similar to harzburgite melting reactions previously determined for fertile peridotites.
(4) pMELTS (Ghiorso et al., 2002), BATCH (Longhi, 2002), and the expressions presented by Langmuir et al. (1992) were used to calculate partial melts of the DMM1 composition at 10 kbar. Although melts of depleted peridotites were not included in the calibrations of either pMELTS or BATCH, both of these programs reproduce the essential characteristics of the melt fraction vs major element trends observed in experiments on DMM1. The successes of these calculations include their generally close matches to the absolute values of the major oxide concentrations, their successful predictions of the overall shapes of the observed trends, their predictions of the changes in slope of the F vs SiO2, FeO*, MgO, and CaO trends that occur at or near cpx-out, and their predictions of the melt fraction at which cpx-out occurs. Nevertheless, liquid compositions calculated by these models display systematic offsets relative to the experimental glass compositions. The Langmuir et al. expressions provide the best match to the F vs FeO* and MgO glass trends for DMM1, but cannot be used to model SiO2, CaO, or Al2O3 contents of partial melts.
(5) Combining our experimentally determined solidus temperature with solidus temperatures determined on seven other peridotites, we parameterized Ts at 10 kbar as a function of the Na2O, K2O, and P2O5 contents in near-solidus liquids at F = 0·005 calculated with pMELTS. The resulting equation, Ts(°C) = 1295·8 - (3·628)Na2O – (7·218)K2O – (5·100)P2O5 (oxide quantities in wt %), recovers the input temperatures with a mean deviation of 3°C. This expression predicts that the 10 kbar solidi of abyssal peridotites are 1278–1295°C and is consistent with the theoretical expectation that more incompatible elements have a larger influence on the solidus.
|
ACKNOWLEDGEMENTS |
|---|
The authors are grateful for helpful discussions with and/or laboratory support from Marc Hirschmann, Glenn Gaetani, John Beckett, and Paul Asimow. We thank Ian Hutcheon and Doug Phinney for assistance with the ion microprobe measurements. We also thank Charles Lesher, Keith Putirka, and Brandon Schwab for helpful reviews. This work was supported by National Science Foundation grant OCE93-14505. This research was also funded in part by stipend support for L.E.W. from the National Physical Science Consortium and NASA-Johnson Space Center.
|
FOOTNOTES |
|---|
Present address: Department of Geosciences, Wilkinson Hall, Oregon State University, Corvallis, OR 97331, USA 
|
REFERENCES |
|---|
|
TOPABSTRACTINTRODUCTIONEXPERIMENTAL AND ANALYTICAL...RESULTSDISCUSSIONCONCLUSIONSREFERENCES |
|---|
Albarède, F. & Provost, A. (1977). Petrological and geochemical mass-balance equations: an algorithm for least-square fitting and general error analysis. Computers and Geosciences 3, 309–326.[CrossRef]
Allègre, C. J., Poirer, J.-P., Humler, E. & Hofmann, A. W. (1995). The chemical composition of the Earth. Earth and Planetary Science Letters 134, 515–526.[CrossRef][Web of Science]
Armstrong, J. T. (1988). Quantitative analysis of silicate and oxide minerals, comparison of Monte Carlo, ZAF and 
(
z) procedures. In: Newbury, D. E. (ed.) Microbeam Analysis. San Francisco, CA: San Francisco Press, pp. 239–246.
Asimow, P. D., Hirschmann, M. M. & Stolper, E. M. (1997). An analysis of variations in isentropic melt productivity. Philosophical Transactions of the Royal Society of London, Series A 355, 255–281.[CrossRef]
Baker, M. B. & Beckett, J. R. (1999). The origin of abyssal peridotites, a reinterpretation of constraints based on primary bulk compositions. Earth and Planetary Science Letters 171, 49–61.[CrossRef][Web of Science]
Baker, M. B. & Stolper, E. M. (1994). Determining the composition of high-pressure mantle melts using diamond aggregates. Geochimica et Cosmochimica Acta 58, 2811–2827.[CrossRef][Web of Science]
Baker, M. B., Hirschmann, M. M., Ghiorso, M. S. & Stolper, E. M. (1995). Compositions of near-solidus peridotite melts from experiments and thermodynamic calculations. Nature 375, 308–311.[CrossRef]
Baker, M. B., Hirschmann, M. M., Wasylenki, L. E., Stolper, E. M. & Ghiorso, M. S. (1996). Reply to, Quest for low degree mantle melts. Nature 381, 286.[CrossRef]
Bertka, C. M. & Holloway, J. R. (1988). Martian mantle melts: an experimental study of iron-rich garnet lherzolite minimum melt composition. Proceedings of the 18th Lunar and Planetary Science Conference, 723–739.
Bertka, C. M. & Holloway, J. R. (1994). Anhydrous partial melting of an iron-rich mantle. 1. Subsolidus phase assemblages and partial melting phase relations at 10 to 30 kbar. Contributions to Mineralogy and Petrology 115, 313–322.[CrossRef][Web of Science]
Bishop, F. C., Smith, J. V. & Dawson, J. B. (1978). Na, K, P and Ti in garnet, pyroxene and olivine from peridotite and eclogite xenoliths from African kimberlites. Lithos 11, 155–173.[CrossRef][Web of Science]
Brunet, F. & Chazot, G. (2001). Partitioning of phosphorus between olivine, clinopyroxene and silicate glass in a spinel lherzolite xenolith from Yemen. Chemical Geology 176, 51–72.[CrossRef][Web of Science]
Cameron, M. & Papike, J. J. (1981). Structural and chemical variations in pyroxenes. American Mineralogist 66, 1–50.[Abstract]
Carmichael, I. S. E., Turner, F. J. & Verhoogen, J. (1974). Igneous Petrology. New York: McGraw–Hill.
Christie, D. M., Carmichael, I. S. E. & Langmuir, C. H. (1986). Oxidation states of mid-ocean ridge basalt glasses. Earth and Planetary Science Letters 79, 397–411.[CrossRef][Web of Science]
Daines, M. J. & Richter, F. M. (1988). An experimental method for directly determining the interconnectivity of melt in a partially molten system. Geophysical Research Letters 15, 1459–1462.[Web of Science]
Dixon, J. E., Stolper, E. M. & Holloway, J. R. (1995). An experimental study of water and carbon-dioxide solubilities in mid-ocean ridge basaltic liquids. 1. Calibration and solubility models. Journal of Petrology 36, 1607–1631.
Dobson, P. F., Epstein, S. & Stolper, E. M. (1989). Hydrogen isotopic fractionation between coexisting vapor and silicate glasses and melts at low pressure. Geochimica et Cosmochimica Acta 53, 2723–2730.[CrossRef][Web of Science]
Draper, D. S. & Green, T. H. (1999). P–T phase relations of silicic, alkaline, aluminous liquids: new results and applications to mantle melting and metasomatism. Earth and Planetary Science Letters 170, 255–268.[CrossRef][Web of Science]
Falloon, T. J. & Green, D. H. (1987). Anhydrous partial melting of MORB pyrolite and other peridotite compositions at 10 kbar and implications for the origin of primitive MORB glasses. Mineralogy and Petrology 37, 181–219.[CrossRef][Web of Science]
Falloon, T. J., Green, D. H., O'Neill, H. S. C. & Ballhaus, C. G. (1996). Quest for low degree mantle melts. Nature 381, 285.[CrossRef]
Falloon, T. J., Green, D. H., O'Neill, H. S. C. & Hibberson, W. O. (1997). Experimental tests of low degree peridotite partial melt compositions, implications for the nature of anhydrous near-solidus peridotite melts at 1 GPa. Earth and Planetary Science Letters 152, 149–162.[CrossRef][Web of Science]
Falloon, T. J., Green, D. H., Danyushevsky, L. V. & Faul, U. H. (1999). Peridotite melting at 1·0 and 1·5 GPa, an experimental evaluation of techniques using diamond aggregates and mineral mixes for determination of near-solidus melts. Journal of Petrology 40, 1343–1375.[CrossRef][Web of Science]
Falloon, T. J., Danyushevsky, L. V. & Green, D. H. (2001). Peridotite melting at 1 GPa: reversal experiments on partial melt compositions produced by peridotite–basalt sandwich experiments. Journal of Petrology 42, 2363–2390.
Gaetani, G. A. & Grove, T. L. (1998). The influence of water on melting of mantle peridotite. Contributions to Mineralogy and Petrology 131, 323–346.[CrossRef][Web of Science]
Gee, L. L. & Sack, R. O. (1988). Experimental petrology of melilite nephelinites. Journal of Petrology 24, 256–265.
Ghiorso, M. S. & Sack, R. O. (1995). Chemical mass-transfer in magmatic processes 4. A revised and internally consistent thermodynamic model for the interpolation and extrapolation of liquid–solid equilibria in magmatic systems at elevated temperatures and pressures. Contributions to Mineralogy and Petrology 119, 197–212.[Web of Science]
Ghiorso, M. S., Hirschmann, M. M., Reiners, P. W. & Kress, V. C. (2002). The pMELTS: a revision of MELTS for improved calculation of phase relations and major element partitioning related to partial melting of the mantle to 3 GPa. Geochemistry, Geophysics, Geosystems 3 (Paper 2001GC000217).
Green, D. H., Falloon, T. J. & Taylor, W. R. (1987). Mantle derived magmas, role of variable source peridotite and variable C–H–O fluid compositions. In: Mysen, B. O. (ed.) Magmatic Processes, Physicochemical Principles. University Park, PA: United States Geochemical Society, pp. 139–154.
Gudfinnsson, G. & Holloway, J. R. (1993). Activity–composition relationships in the system Fe–Pt at 1300 and 1400°C and at 1 atm and 20 kbar. American Mineralogist 78, 178–186.[Abstract]
Gudfinnsson, G. & Presnall, D. C. (2000). Melting behavior of model lherzolite in the system CaO–MgO–Al2O3–FeO at 0·7–2·8 GPa. Journal of Petrology 41, 1241–1269.
Gudfinnsson, G., Holloway, J. R. & Carmichael, I. S. E. (1988). Pressure effect on the ferric/ferrous ratio in basaltic liquids. EOS Transactions, American Geophysical Union 69, 1511.
Halliday, A. N., Lee, D.-C., Tommasini, S., Davies, G. R., Paslick, C. R., Fitton, J. G. & James, D. E. (1995). Incompatible trace elements in OIB and MORB and source enrichment in the sub-oceanic mantle. Earth and Planetary Science Letters 133, 379–395.[CrossRef][Web of Science]
Hart, S. R. & Zindler, A. (1986). In search of a bulk-earth composition. Chemical Geology 57, 247–267.[CrossRef][Web of Science]
Hays, J. F. (1966). Lime–alumina–silica. Carnegie Institute of Washington Yearbook 65, 234–239.
Hellebrand, E., Snow, J. E., Dick, H. J. B. & Hofmann, A. W. (2000). Coupled major and trace elements as indicators of the extent of melting in mid-ocean-ridge peridotites. Nature 410, 677–681.
Hervig, R. L. & Smith, J. V. (1982). Temperature-dependent distribution of Cr between olivine and pyroxenes in lherzolite xenoliths. Contributions to Mineralogy and Petrology 81, 184–189.[CrossRef][Web of Science]
Herzberg, C. T., Feigenson, M., Skuba, C. & Ohtani, E. (1988). Majorite fractionation recorded in the geochemistry of peridotites from South Africa. Nature 332, 823–826.[CrossRef]
Herzberg, C. T., Raterron, P. & Zhang, J. (2000). New experimental observations on the anhydrous solidus for peridotite KLB-1. Geochemistry, Geophysics, Geosystems 1 (Paper 2000GC000089).
Hirose, K. & Kushiro, I. (1993). Partial melting of dry peridotites at high pressures: determination of compositions of melts segregated from peridotite using aggregates of diamond. Earth and Planetary Science Letters 114, 477–489.[CrossRef][Web of Science]
Hirschmann, M. M. (2000). Mantle solidus, experimental constraints and the effects of peridotite composition. Geochemistry, Geophysics, Geosystems 1 (Paper 2000GC000070).
Hirschmann, M. M., Baker, M. B. & Stolper, E. M. (1998a). The effect of alkalis on the silica content of mantle-derived melts. Geochimica et Cosmochimica Acta 62, 883–902.[CrossRef][Web of Science]
Hirschmann, M. M., Ghiorso, M. S., Wasylenki, L. E., Asimow, P. D. & Stolper, E. M. (1998b). Calculation of peridotite partial melting from thermodynamic models of minerals and melts. I. Review of methods and comparison with experiments. Journal of Petrology 39, 1091–1115.[CrossRef][Web of Science]
Hirschmann, M. M., Ghiorso, M. S. & Stolper, E. M. (1999a). Calculation of peridotite partial melting from thermodynamic models of minerals and melts. II. Isobaric variations in melts near the solidus and owing to variable source composition. Journal of Petrology 40, 297–313.[CrossRef][Web of Science]
Hirschmann, M. M., Asimow, P. D., Ghiorso, M. S. & Stolper, E. M. (1999b). Calculation of peridotite partial melting from thermodynamic models of minerals and melts. III. Controls on isobaric melt production and the effect of water on melt production. Journal of Petrology 40, 831–851.[CrossRef][Web of Science]
Huebner, J. S. (1971). Buffering techniques for hydrostatic systems at elevated pressures. In: Ulmer, G. C. (ed.) Research Techniques for High Pressure and High Temperature. New York: Springer, pp. 123–177.
Ihinger, P. D., Zhang, Y. & Stolper, E. M. (1999). The speciation of dissolved water in rhyolitic melt. Geochimica et Cosmochimica Acta 63, 3567–3578.[CrossRef][Web of Science]
Iwamori, H., McKenzie, D. & Takahashi, E. (1995). Melt generation by isentropic mantle upwelling. Earth and Planetary Science Letters 134, 253–266.[CrossRef][Web of Science]
Jamieson, H. E., Roeder, P. L. & Grant, A. H. (1992). Olivine–pyroxene–PtFe alloy as an oxygen geobarometer. Journal of Geology 100, 138–145.[Web of Science]
Jaques, A. L. & Green, D. H. (1980). Anhydrous melting of peridotite at 0–15 kb pressure and the genesis of tholeiitic basalts. Contributions to Mineralogy and Petrology 73, 287–310.[CrossRef][Web of Science]
Jarosewich, E., Nelen, J. A. & Norberg, J. A. (1979). Electron microprobe reference samples for mineral analysis. Smithsonian Contributions to Earth Sciences 22, 68–72.
Johnson, K. T. M. & Dick, H. J. B. (1992). Open system melting and temporal and spatial variation of peridotite and basalt in the Atlantis Fracture Zone. Journal of Geophysical Research 97, 9219–9241.
Johnson, K. T. M. & Kushiro, I. (1992). Segregation of high-pressure partial melts from peridotite using aggregates of diamonds: a new experimental approach. Geophysical Research Letters 19, 1703–1706.[Web of Science]
Johnson, K. T. M., Dick, H. J. B. & Shimizu, N. (1990). Melting in the oceanic upper mantle: an ion microprobe study of diopsides in abyssal peridotites. Journal of Geophysical Research 95, 2661–2678.
Kent, A. J. R., Norman, M. D., Hutcheon, I. D. & Stolper, E. M. (1999). Seawater assimilation in an oceanic volcano: evidence from matrix glasses and glass inclusions from Loihi seamount, Hawaii. Chemical Geology 156, 299–319.[CrossRef][Web of Science]
Kinzler, R. J. (1997). Melting of mantle peridotite at pressures approaching the spinel to garnet transition: application to mid-ocean ridge basalt petrogenesis. Journal of Geophysical Research 102, 853–874.[CrossRef]
Kinzler, R. J. & Grove, T. L. (1992a). Primary magmas of mid-ocean ridge basalts 1. Experiments and methods. Journal of Geophysical Research 97, 6885–6906.
Kinzler, R. J. & Grove, T. L. (1992b). Primary magmas of mid-ocean ridge basalts 2. Applications. Journal of Geophysical Research 97, 6907–6926.
Kushiro, I. (1975). On the nature of silicate melt and its significance in magma genesis, regularities in the shift of the liquidus boundaries involving olivine, pyroxene, and silica minerals. American Journal of Science 275, 411–431.
Kushiro, I. (1996). Partial melting of a fertile mantle peridotite at high pressures, an experimental study using aggregates of diamond. In: Basu, A. & Hart, S. (eds) Earth Processes, Reading the Isotopic Code. American Geophysical Union Monograph 95, 109–122.
Kushiro, I. & Mysen, B. O. (2002). A possible effect of melt structure on the Mg–Fe2+ partitioning between olivine and melt. Geochimica et Cosmochimica Acta 66, 2267–2272.[CrossRef][Web of Science]
Langmuir, C. H., Klein, E. M. & Plank, T. (1992). Petrological systematics of mid-ocean ridge basalts, constraints on melt generation beneath ocean ridges. In: Morgan, J. P., Blackman, D. & Sinton, J. (eds) Mantle Flow and Melt Generation at Mid-Ocean Ridges. American Geophysical Union Monograph 71, 183–280.
Lesher, C. E. & Walker, D. (1988). Cumulate maturation and melt migration in a temperature gradient. Journal of Geophysical Research 93, 10295–10311.
Libourel, G., Deloule, E. & Toplis, M. J. (1994). Phosphorus partitioning in basalts: an experimental and ion probe study. Mineralogical Magazine 58A, 527–528.
Longhi, J. (2002). Some phase equilibria systematics of lherzolite melting. Geochemistry, Geophysics, Geosystems 3, U1–U33.
Mattioli, G. S. & Wood, B. J. (1988). Magnetite activities across the Mg2AlO4–Fe3O4 spinel join, with application to thermobarometric estimates of upper mantle oxygen fugacity. Contributions to Mineralogy and Petrology 98, 148–162.[CrossRef][Web of Science]
Mattioli, G. S., Baker, M. B., Rutter, M. J. & Stolper, E. M. (1989). Upper mantle oxygen fugacity and its relationship to metasomatism. Journal of Geology 97, 521–536.[Web of Science]
McDonough, W. F. & Sun, S.-S. (1995). The composition of the Earth. Chemical Geology 120, 223–253.[CrossRef][Web of Science]
McKenzie, D. (1984). The generation and compaction of partially molten rock. Journal of Petrology 25, 713–765.
McKenzie, D. (1985). The extraction of magma from the crust and mantle. Earth and Planetary Science Letters 74, 81–91.[CrossRef][Web of Science]
McKenzie, D. (1989). Some remarks on the movement of small melt fractions in the mantle. Earth and Planetary Science Letters 95, 53–72.[CrossRef][Web of Science]
McKenzie, D. & Bickle, M. J. (1988). The volume and composition of melt generated by extension of the lithosphere. Journal of Petrology 29, 625–679.
Morgan, J. W., Wandless, G. A. & Petrie, R. K. (1980). Earth's upper mantle: volatile element distribution and origin of siderophile element content. Lunar and Planetary Science XI, 740–742.
Niu, Y. & Batiza, R. (1991). An empirical method for calculating melt compositions produced beneath mid-ocean ridges: application to axis and off-axis (seamounts) melting. Journal of Geophysical Research 96, 21753–21777.
Niu, Y. & Hékinian, R. (1997). Spreading rate dependence of the extent of mantle melting beneath ocean ridges. Nature 385, 326–329.[CrossRef]
Noda, T., Inagaki, M. & Yamada, S. (1969). Glass-like carbons. Journal of Non-Crystalline Solids 1, 285–302.[CrossRef]
O'Neill, H. S. C. & Wall, V. J. (1987). The olivine–orthopyroxene–spinel oxygen geobarometer, the nickel precipitation curve, and the oxygen fugacity of the Earth's upper mantle. Journal of Petrology 28, 1169–1191.
Pickering-Witter, J. & Johnston, A. D. (2000). The effects of variable bulk composition on the melting systematics of fertile peridotitic assemblages. Contributions to Mineralogy and Petrology 140, 190–211.[CrossRef][Web of Science]
Reed, B. C. (1992). Linear least-squares fits with errors in both coordinates. II: Comments on parameter variances. American Journal of Physics 60, 59–62.[CrossRef]
Ringwood, A. E. (1979). Origin of the Earth and Moon. New York: Springer.
Robinson, J. A. C., Wood, B. J. & Blundy, J. D. (1998). The beginning of melting of fertile and depleted peridotite at 1·5 GPa. Earth and Planetary Science Letters 155, 97–111.[CrossRef][Web of Science]
Ryerson, F. J. (1985). Oxide solution mechanisms in silicate melts—systematic variations in the activity-coefficient of SiO2. Geochimica et Cosmochimica Acta 49, 637–649.[CrossRef][Web of Science]
Sack, R. O., Carmichael, I. S. E., Rivers, M. L. & Ghiorso, M. S. (1980). Ferric–ferrous equilibria in natural silicate liquids at 1 bar. Contributions to Mineralogy and Petrology 75, 369–376.[CrossRef]
Salters, V. J. M. & Dick, H. J. B. (2002). Mineralogy of the mid-ocean-ridge basalt source from neodymium isotopic composition of abyssal peridotites. Nature 418, 68–72.[CrossRef][Medline]
Sawa, T. & Tanaka, K. (2002). Nanoindentation of natural diamond. Philosophical Magazine A—Physics of Condensed Matter: Structure, Defects, and Mechanical Properties 82, 1851–1856.
Schwab, B. E. & Johnston, A. D. (2001). Melting systematics of modally variable, compositionally intermediate peridotites and the effect of mineral fertility. Journal of Petrology 42, 1789–1811.
Sen, G. (1982). Composition of basaltic liquids generated from a partially depleted lherzolite at 9 kbar pressure. Nature 299, 336–338.[CrossRef]
Shimizu, N. (1998). The geochemistry of olivine-hosted melt inclusions in a FAMOUS basalt ALV19-4-1. Physics of the Earth and Planetary Interiors 107, 183–201.[CrossRef][Web of Science]
Smith, D., Griffin, W. L. & Ryan, C. G. (1993). Compositional evolution of high-temperature sheared lherzolite PHN 1611. Geochimica et Cosmochimica Acta 57, 605–613.[CrossRef][Web of Science]
Sobolev, A. V. (1996). Melt inclusions in minerals as a source of principle petrological information. Petrology 4, 209–220.
Sobolev, A. V. & Shimizu, N. (1993). Ultra-depleted primary melt included in an olivine from the Mid-Atlantic Ridge. Nature 363, 151–154.[CrossRef]
Stevenson, D. J. & Scott, D. R. (1991). Mechanics of fluid–rock systems. Annual Review of Fluid Mechanics 23, 305–339.[CrossRef]
Stineman, R. W. (1980). A consistently well-behaved method of interpolation. Creative Computing 6, 54–57.
Stolper, E. M. & Newman, S. (1994). The role of water in the petrogenesis of Mariana trough magmas. Earth and Planetary Science Letters 121, 293–325.[CrossRef][Web of Science]
Stosch, H.-G. (1981). Sc, Cr, Co and Ni partitioning between minerals from spinel peridotite xenoliths. Contributions to Mineralogy and Petrology 78, 166–174.[CrossRef][Web of Science]
Sun, S.-S. (1982). Chemical composition and origin of the earth's primitive mantle. Geochimica et Cosmochimica Acta 46, 179–192.[CrossRef][Web of Science]
Takahashi, E. (1986). Melting of a dry peridotite KLB-1 up to 14 GPa, implications on the origin of peridotitic upper mantle. Journal of Geophysical Research 91(B8), 9367–9380.
Takahashi, E. & Kushiro, I. (1983). Melting of a dry peridotite at high pressures and basalt magma genesis. American Mineralogist 68, 859–879.[Abstract]
Ulmer, P. (1989). The dependence of the Fe2+–Mg cation partitioning between olivine and basaltic liquid on pressure, temperature and composition, an experimental study to 30 kbars. Contributions to Mineralogy and Petrology 101, 261–273.[CrossRef][Web of Science]
von Bargen, N. & Waff, H. S. (1986). Permeabilities, interfacial areas and curvatures of partially molten systems, results of numerical computations of equilibrium microstructures. Journal of Geophysical Research 91, 9261–9276.
Waff, H. S. & Bulau, J. R. (1979). Equilibrium fluid distribution in an ultramafic partial melt under hydrostatic stress conditions. Journal of Geophysical Research 84, 6109–6114.
Wagner, T. P. & Grove, T. L. (1998). Melt/harzburgite reaction in the petrogenesis of tholeiitic magma from Kilauea volcano, Hawaii. Contributions to Mineralogy and Petrology 131, 1–12.[CrossRef][Web of Science]
Walter, M. J. (1998). Melting of garnet peridotite and the origin of komatiite and depleted lithosphere. Journal of Petrology 39, 29–60.[CrossRef][Web of Science]
Wänke, H., Dreibus, G. & Jagoutz, E. (1984). Mantle chemistry and accretion history of the earth. In: Kröner, A., Hanson, G. & Goodwin, A. (eds) Archean Geochemistry. Berlin: Springer, pp. 1–24.
Wasylenki, L. E., Baker, M. B., Hirschmann, M. M. & Stolper, E. M. (1996). The effect of source depletion on equilibrium mantle melting. EOS Transactions, American Geophysical Union 77(46), F847.
CiteULike 
Connotea 
Del.icio.us What's this?
This article has been cited by other articles:
|
|
 |
|
  E. Medard, C. A. McCammon, J. A. Barr, and T. L. Grove Oxygen fugacity, temperature reproducibility, and H2O contents of nominally anhydrous piston-cylinder experiments using graphite capsules American Mineralogist, November 1, 2008; 93(11-12): 1838 - 1844. [Abstract] [Full Text] [PDF]
|
|
|
|
 |
|
 T. Kuritani, T. Yokoyama, and E. Nakamura Generation of Rear-arc Magmas Induced by Influx of Slab-derived Supercritical Liquids: Implications from Alkali Basalt Lavas from Rishiri Volcano, Kurile Arc J. Petrology, July 1, 2008; 49(7): 1319 - 1342. [Abstract] [Full Text] [PDF]
|
|
|
|
 |
|
 S. Pilet, M. B. Baker, and E. M. Stolper Metasomatized Lithosphere and the Origin of Alkaline Lavas Science, May 16, 2008; 320(5878): 916 - 919. [Abstract] [Full Text] [PDF]
|
|
|
|
|
|
I. Baziotis, E. Mposkos, and P. D. Asimow Petrogenesis of Ultramafic Rocks from the Ultrahigh-pressure Metamorphic Kimi Complex in Eastern Rhodope (NE Greece) J. Petrology, May 1, 2008; 49(5): 885 - 909. [Abstract] [Full Text] [PDF]
|
|
|
|
|
|
C. Spandler, G. Yaxley, D. H. Green, and A. Rosenthal Phase Relations and Melting of Anhydrous K-bearing Eclogite from 1200 to 1600{degrees}C and 3 to 5 GPa J. Petrology, April 1, 2008; 49(4): 771 - 795. [Abstract] [Full Text] [PDF]
|
|
|
|
|
|
G. P. Brey, V. K. Bulatov, A. V. Girnis, and Y. Lahaye Experimental Melting of Carbonated Peridotite at 6-10 GPa J. Petrology, April 1, 2008; 49(4): 797 - 821. [Abstract] [Full Text] [PDF]
|
|
|
|
|
|
T. J. Falloon, D. H. Green, L. V. Danyushevsky, and A. W. McNeill The Composition of Near-solidus Partial Melts of Fertile Peridotite at 1 and 1{middle dot}5 GPa: Implications for the Petrogenesis of MORB J. Petrology, April 1, 2008; 49(4): 591 - 613. [Abstract] [Full Text] [PDF]
|
|
|
|
 |
|
 K. D. Putirka Thermometers and Barometers for Volcanic Systems Reviews in Mineralogy and Geochemistry, January 1, 2008; 69(1): 61 - 120. [Full Text] [PDF]
|
|
|
|
|
|
Z. S. De Souza, H. Martin, J.-J. Peucat, E. F. Jardim De Sa, and M. H. D. F. Macedo Calc-Alkaline Magmatism at the Archean Proterozoic Transition: the Caico Complex Basement (NE Brazil) J. Petrology, November 1, 2007; 48(11): 2149 - 2185. [Abstract] [Full Text] [PDF]
|
|
|
|
 |
|
 X. Liu and H. St.C. O'Neill EFFECTS OF P2O5 AND TiO2 ON THE PARTIAL MELTING OF SPINEL LHERZOLITE IN THE SYSTEM CaO-MgO-Al2O3-SiO2 AT 1.1 GPa Can Mineral, June 1, 2007; 45(3): 649 - 655. [Abstract] [Full Text] [PDF]
|
|
|
|
|
|
N. S. C. Simon, R. W. Carlson, D. G. Pearson, and G. R. Davies The Origin and Evolution of the Kaapvaal Cratonic Lithospheric Mantle J. Petrology, March 1, 2007; 48(3): 589 - 625. [Abstract] [Full Text] [PDF]
|
|
|
|
|
|
E. MEDARD, M. W. SCHMIDT, P. SCHIANO, and L. OTTOLINI Melting of Amphibole-bearing Wehrlites: an Experimental Study on the Origin of Ultra-calcic Nepheline-normative Melts J. Petrology, March 1, 2006; 47(3): 481 - 504. [Abstract] [Full Text] [PDF]
|
|
|
|
|
|
J. L. Mosenfelder, N. I. Deligne, P. D. Asimow, and G. R. Rossman Hydrogen incorporation in olivine from 2-12 GPa American Mineralogist, February 1, 2006; 91(2-3): 285 - 294. [Abstract] [Full Text] [PDF]
|
|
|
|
|
|
X. LIU, H. ST. C. O'NEILL, and A. J. BERRY The Effects of Small Amounts of H2O, CO2 and Na2O on the Partial Melting of Spinel Lherzolite in the System CaO-MgO-Al2O3-SiO2 {+/-} H2O {+/-} CO2 {+/-} Na2O at 1{middle dot}1 GPa J. Petrology, February 1, 2006; 47(2): 409 - 434. [Abstract] [Full Text] [PDF]
|
|
|
|
 |
|
 T. Furman, J. Bryce, T. Rooney, B. Hanan, G. Yirgu, and D. Ayalew Heads and tails: 30 million years of the Afar plume Geological Society, London, Special Publications, January 1, 2006; 259(1): 95 - 119. [Abstract] [PDF]
|
|
|
|
|
|
T. FURMAN, J. G. BRYCE, J. KARSON, and A. IOTTI East African Rift System (EARS) Plume Structure: Insights from Quaternary Mafic Lavas of Turkana, Kenya J. Petrology, May 1, 2004; 45(5): 1069 - 1088. [Abstract] [Full Text] [PDF]
|
|
|
|
|
|
M. PERTERMANN and M. M. HIRSCHMANN Anhydrous Partial Melting Experiments on MORB-like Eclogite: Phase Relations, Phase Compositions and Mineral-Melt Partitioning of Major Elements at 2-3 GPa J. Petrology, December 1, 2003; 44(12): 2173 - 2201. [Abstract] [Full Text] [PDF]
|
|
| ||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||