Journal of Petrology

Journal of Petrology Advance Access published online on August 24, 2007
Journal of Petrology, doi:10.1093/petrology/egm039

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Phase Relations and Melting of Anhydrous K-bearing Eclogite from 1200 to 1600°C and 3 to 5 GPa

Carl Spandler^*, Greg Yaxley, David H. Green and Anja Rosenthal

Research School of Earth Sciences, Australian National University, Canberra, ACT 0200, Australia

Received November 23, 2006; Revised typescript accepted July 2, 2007

	ABSTRACT

TOP ABSTRACT INTRODUCTION EXPERIMENTAL SETUP RESULTS DISCUSSION CONCLUSIONS REFERENCES

 
To investigate eclogite melting under mantle conditions, we have performed a series of piston-cylinder experiments using a homogeneous synthetic starting material (GA2) that is representative of altered mid-ocean ridge basalt. Experiments were conducted at pressures of 3·0, 4·0 and 5·0 GPa and over a temperature range of 1200–1600°C. The subsolidus mineralogy of GA2 consists of garnet and clinopyroxene with minor quartz–coesite, rutile and feldspar. Solidus temperatures are located at 1230°C at 3·0 GPa and 1300°C at 5·0 GPa, giving a steep solidus slope of 30–40°C/GPa. Melting intervals are in excess of 200°C and increase with pressure up to 5·0 GPa. At 3·0 GPa feldspar, rutile and quartz are residual phases up to 40°C above the solidus, whereas at higher pressures feldspar and rutile are rapidly melted out above the solidus. Garnet and clinopyroxene are the only residual phases once melt fractions exceed 20% and garnet is the sole liquidus phase over the investigated pressure range. With increasing melt fraction garnet and clinopyroxene become progressively more Mg-rich, whereas coexisting melts vary from K-rich dacites at low degrees of melting to basaltic andesites at high melt fractions. Increasing pressure tends to increase the jadeite and Ca-eskolaite components in clinopyroxene and enhance the modal proportion of garnet at low melt fractions, which effects a marked reduction in the Al₂O₃ and Na₂O content of the melt with pressure. In contrast, the TiO₂ and K₂O contents of the low-degree melts increase with increasing pressure; thus Na₂O and K₂O behave in a contrasted manner as a function of pressure. Altered oceanic basalt is an important component of crust returned to the mantle via plate subduction, so GA2 may be representative of one of many different mafic lithologies present in the upper mantle. During upwelling of heterogeneous mantle domains, these mafic rock-types may undergo extensive melting at great depths, because of their low solidus temperatures compared with mantle peridotite. Melt batches may be highly variable in composition depending on the composition and degree of melting of the source, the depth of melting, and the degree of magma mixing. Some of the eclogite-derived melts may also react with and refertilize surrounding peridotite, which itself may partially melt with further upwelling. Such complex magma-genesis conditions may partly explain the wide spectrum of primitive magma compositions found within oceanic basalt suites.

KEY WORDS: eclogite; experimental petrology; mafic magmatism; mantle melting; oceanic basalts

	INTRODUCTION

TOP ABSTRACT INTRODUCTION EXPERIMENTAL SETUP RESULTS DISCUSSION CONCLUSIONS REFERENCES

 
Mafic magmatism on Earth is primarily derived from melting of mantle peridotite, yet there is a growing consensus that multiple mantle components contribute to the production of many intraplate and some mid-ocean ridge magmas (Hofmann, 2003, and references therein). In particular, most ocean island basalts (OIB) have isotopic and geochemical characteristics that are interpreted to require at least partial derivation from melting of high-pressure mafic rock types (e.g. Hofmann, 1997, 2003; Lassiter et al., 2000; Sobolev et al., 2005). Mafic rocks may be returned to the mantle via plate subduction or through delamination of the lower crust and are expected to form eclogite at pressures above 2·0 GPa (Ringwood & Green, 1966; Yasuda et al., 1994; Kogiso et al., 2004). Some eclogites have lower solidus temperatures than peridotite and hence have the potential to undergo extensive melting when present in the upper mantle (Yoder & Tilley, 1962; Green & Ringwood, 1968; Hirschmann & Stolper, 1996; Pertermann & Hirschmann, 2003a; Ito & Mahoney, 2005). Furthermore, relatively small variations in bulk composition can translate into substantial differences in solidus temperatures (Kogiso et al., 2004; Kogiso & Hirschmann, 2006), which, for a given mantle adiabat, directly controls the depth of eclogite melting in the mantle.

Xenolith suites or exhumed mantle sections can provide important information on the compositions, melting behaviour and partial melt compositions of mantle eclogite (Hirschmann & Stolper, 1996; Bodinier & Godard, 2003; Pearson et al., 2003), but undoubtedly the greatest insights into these processes are obtained from high-pressure, high-temperature experiments. Nonetheless, despite numerous experimental studies (e.g. Pertermann & Hirschmann, 2003b; Kogiso et al., 2004, and references therein) there remain fundamental gaps in our knowledge of the phase relations and melting conditions of eclogitic rocks as a function of pressure, temperature and bulk composition. Although a range of mafic compositions has been investigated, the effect that minor elements such as K, P and Ti have on phase relations and melt compositions is poorly understood. Technical limitations of available experimental apparatus have also hampered efforts to experimentally characterize mafic materials as a function of pressure and temperature. Experiments using the piston-cylinder apparatus allow precise control of pressure and temperature and have allowed detailed investigation of phase relations and melting of anhydrous eclogite or pyroxenite from 2·0 to 3·5 GPa (Yaxley & Green, 1998; Klemme et al., 2002; Pertermann & Hirschmann, 2003b; Keshav et al., 2004). However, most piston-cylinder systems are restricted to conditions of 3·5 GPa or lower. Multi-anvil devices have been employed for most experiments at higher pressure (> 4·0 GPa; Yasuda et al., 1994; Tsuruta & Takahashi, 1998; Tuff et al., 2005), although temperature uncertainties for these experiments are commonly around 50°C. Therefore, details of phase changes over relatively small temperature intervals cannot be accurately resolved by multi-anvil experiments. For this reason the effects of temperature and pressure on phase relations and melt compositions of eclogite have been little studied.

In this paper, we present results of piston-cylinder experiments conducted from 3·0 to 5·0 GPa and from 1200 to 1600°C on a composition representative of altered mid-ocean ridge basalt (MORB). This material is expected to be a significant component of subducting oceanic crust and hence may be present in appreciable quantities in the mantle. Melting of eclogite under these pressure–temperature conditions may contribute to the genesis of many oceanic magma suites (Kogiso et al., 2003; Pertermann & Hirschmann, 2003a; Ito & Mahoney, 2005; Sobolev et al., 2005; Yaxley & Sobolev, 2007). We define phase relations, solidus and liquidus temperatures, and melt compositions over the entire pressure range. These data provide a basis for evaluating the effects of pressure and temperature on mineral and melt compositions of anhydrous eclogite. These results, when considered with other experimental data, can be used to advance our understanding of melting of mantle heterogeneities.

	EXPERIMENTAL SETUP

TOP ABSTRACT INTRODUCTION EXPERIMENTAL SETUP RESULTS DISCUSSION CONCLUSIONS REFERENCES

 
Starting material synthesis
For all experiments we have used a synthetic starting material that is representative of sea-floor-altered MORB (Table 1; see also Yaxley & Green, 1994). The composition, labelled GA2, was designed to be similar to the GA1 composition of Yaxley & Green (1994, 1998), except with slightly higher TiO₂ contents to promote rutile saturation. Twenty grams of the starting material were prepared using a ‘sol–gel’ method (Luth & Ingamells, 1965) to produce a homogeneous material and to eliminate problems of sluggish reaction of refractory minerals during experiments. The composition also contains 23 trace elements doped to levels of between 20 and 1000 ppm, with the total trace-element content summing to c. 0·5 wt%. The trace-element results from the experiments will be presented elsewhere. Aluminium, Fe and all trace elements were combined as nitrate solutions dissolved in ethanol. Magnesium, Ca, Na, and K were added as carbonates that were pre-dissolved in nitric acid. The entire solution was combined with tetraethyl orthosilicate [Si(C₂H₅O)₄] and then slowly evaporated on a hot plate until dried into a gel (c. 15 h). The gel was then heated in a Pt crucible to 1000°C to decompose the nitrates. The residual material was combined with appropriate amounts of TiO₂ and MnO powders and milled for 30 min to a fine powder. The powder was then fused at 1350°C under controlled fO₂ conditions to prevent oxidation of Fe, and quenched to form a glass. Multiple analyses of the glass for major elements by energy dispersive spectrometry (EDS) and trace elements by laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) indicated that the glass is homogeneous to the level of analytical precision. The major-element composition of GA2 is presented in Table 1.

View this table: [in this window] [in a new window]   Table 1: Composition of mafic starting materials used in this and other experimental studies

 
Experimental techniques
All experiments were conducted in end-loaded Boyd–England type piston-cylinder apparatus at the Research School of Earth Sciences, Australian National University (ANU). Experiments at 3·0 and 4·0 GPa were conducted in standard 200 T presses, whereas the 5·0 GPa experiments employed a 500 T, ultrahigh-pressure press. For all experiments a small amount (10–20 mg) of GA2 sample was placed with or without vitreous carbon spheres (see below) within graphite capsules (inner diameter c. 1 mm), which were subsequently placed in 3·5 mm Pt capsules. Prior to final sealing of the Pt, the open capsules were heated to 200°C in an oven for a minimum of 2 h (mostly >10 h) to ensure completely anhydrous conditions in the capsules. After removal from the oven the capsules were immediately welded closed. All experiments used standard 1·27 mm salt–Pyrex assemblies with straight graphite heaters and MgO spacers above and below the capsules. Care was taken to position the capsules within the hotspot of the assembly and within 1 mm of the thermocouple tip. For this assembly, temperature gradients across the sample are expected to be less than 10°C based on calibration experiments using several thermocouples. Temperature was controlled using type-B thermocouples (Pt₉₄Rh₆/Pt₇₀Rh₃₀) and a Eurotherm temperature controller and is precise to within 3°C. Pressure is converted from load and is accurate to 0·1 GPa.

Experiments were initially heated to 600°C with minimal confining pressure ( 0·2 GPa) to soften the Pyrex before temperature and pressure were increased simultaneously until desired run conditions were reached. Because of the high run temperature and low-friction assemblies, all friction in the assemblies is expected to have been lost within the first 3–4 h of the experiments. As all experiment durations were 10 h (Table 2) no friction corrections were applied to the pressure measurements. Experiments were terminated by turning off power to the apparatus, resulting in quenching to below 200°C in <10 s. Recovered capsules were mounted in epoxy and polished to expose the sample. Capsules containing carbon spheres were reimpregnated with epoxy once exposed to prevent plucking from the capsule during polishing.

View this table: [in this window] [in a new window]   Table 2: Experimental conditions and run results of experiments using the GA2 mafic starting material

 
Use of vitreous carbon spheres
Processes of quench modification and insufficient melt pooling in experimental charges have long plagued experimentalists attempting to accurately determine the compositions of low-degree partial melts (e.g. Hirose & Kushiro, 1993; Baker & Stolper, 1994; Pickering-Witter & Johnston, 2000). To avoid these problems some workers included in their experimental charges a layer of diamond that acted as a trap for extracting melt from the residual minerals, thereby providing larger melt pools for analysis and avoiding modification of the melt during quenching (Johnson & Kushiro, 1992; Hirose & Kushiro, 1993; Baker & Stolper, 1994). However, the use of natural mineral starting materials, coupled with potential pressure gradients in experiments using the diamond trap technique, led others to question the reliability of the experimental results (Falloon et al., 1996, 1999). Indeed, the validity of partial melting or hydrothermal experiments using diamond traps remains an unresolved issue (see Spandler et al., 2007). More recently, vitreous carbon spheres have been employed by various workers as traps for experimental melts (Pickering-Witter & Johnston, 2000; Schwab & Johnston, 2001; Pertermann & Hirschmann, 2003b; Wasylenki et al., 2003). These workers have demonstrated communication between interstitial liquid in the peridotite and within the vitreous carbon sphere layer, and that in two-stage experiments, run times and starting materials can be optimized to obtain reliable results for melt composition determinations (Wasylenki et al., 2003). Starting materials should be homogeneous glasses or sintered oxide mixes rather than crushed mineral mixes, for which persistence of unreacted relict cores remains an issue. Unlike diamond, the vitreous carbon spheres are amenable to deformation and hence pressure gradients in the sample charge are likely to be significantly reduced. Instead, melt migration around the carbon spheres is thought to be driven by low surface tension between the melt and spheres, rather than pressure gradients (Pickering-Witter & Johnston, 2000; Pertermann & Hirschmann, 2003b).

In this study, experiments conducted close to solidus temperatures were loaded with a layer of vitreous carbon spheres (50–120 µm in diameter) at the base of the graphite capsule, prior to loading of the GA2 powder. The vitreous carbon sphere layer did not exceed 30% of the inner volume of the graphite capsule in any experiments, but in most cases the layer was at least four spheres in thickness (Fig. 1a). The carbon sphere layer serves as an effective melt trap, as in all experiments above solidus conditions melt isolated from the solid residue could be identified as thin films around the carbon spheres or filling pores within the spheres. Spheres close to the crystallized starting material often contained tiny mineral phases in addition to melt, so only melt rims deep within the carbon sphere layer were targeted for analysis (Fig. 1b). As all experiments were run in graphite capsules, the addition of the carbon spheres does not add any additional chemical components to the experimental samples. Furthermore, potential problems relating to disequilibrium melting of mineral mixes (e.g. Falloon et al., 1996) are not relevant here as we used a homogeneous synthetic starting material.

View larger version (145K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 1. (a) Backscattered electron images of an experimental charge containing vitreous carbon spheres as a melt trap. (b) Close-up image of the melt trap with thin films of melt preserved around the rims of the carbon spheres. It should be noted that only the melt rims furthest from the recrystallized starting material were analysed, as the spheres close to recrystallized starting material contained fine mineral grains together with melt.

 
Analysis of mineral and melt compositions
Sections through the experimental samples were examined using a JEOL 6400 scanning electron microscope (SEM), housed at the Electron Microscopy Unit of the ANU. Mineral grains and quenched melt pools from the experiments were analysed for major elements using an EDS detector mounted to the SEM. Some additional mineral analyses were performed by wavelength-dispersive spectrometry (WDS) using a Cameca SX 100 electron microprobe at the Research School of Earth Sciences, ANU, and the results obtained using the two methods were indistinguishable within analytical precision. For the EDS analysis, accelerating voltage, beam current and counting time were set to 15 kV, 1 nA and 100 s, respectively. Element concentrations were standardized against known mineral standards produced by Astimex Scientific Limited.

A focused beam was used for all mineral analyses and care was taken to avoid overlap with mineral inclusions or quench overgrowths on grain rims. For the melt analyses a defocused beam of >25 µm² was used where sufficiently large quenched glass pools were formed; typically in experiments with >25% melting. In these cases, quench modification of the glass pools adjacent to mineral grains could be avoided during analysis. Analyses conducted using this technique resulted in melt compositions with major-element totals of between 97 and 101%. For experiments containing low melt fractions, quenched glass preserved as thin (1–5 µm) rims around carbon spheres or in the pores of the carbon spheres were analysed using a focused beam. The low beam current used (1 nA) and anhydrous nature of the experiments reduce the likelihood for Na loss during analyses. Nonetheless, Na concentrations were continually monitored during acquisition to evaluate any Na loss. Sodium contents remained relatively constant in all cases, indicating that Na loss was negligible. In many cases the melt films around the carbon spheres were too thin to avoid overlap of the beam with the carbon spheres, which resulted in low major-element totals. Similar results were obtained from analysis of similar experimental run products by Pertermann & Hirschmann (2003b) and Wasylenki et al. (2003). Although carbon could not be quantified, increased height of the carbon peak in the EDS spectra was observed for many of the analyses. Nonetheless, for individual experiments, analyses of melt with major-element totals over 90% produced consistent element ratios and hence similar compositions when normalized to 100%. As an example, data for run MH25 are presented in Table 3. Therefore, only analyses that had totals of over 90% prior to normalization were used for calculating average melt compositions. For comparative purposes, melt compositions from all experiments were normalized to 100% major-element oxides. Similar procedures were used by Pertermann & Hirschmann (2003b) and Wasylenki et al. (2003).

View this table: [in this window] [in a new window]   Table 3: Analyses of melt from within the carbon sphere layer of experiment MH25 (5·0 GPa, 1400°C)

 
In all but two experimental runs the graphite capsules appeared to have remained intact during the experiments, which effectively prevented contact of the sample with the outer Pt capsule. However, in experiments MH15 and MH22 the graphite capsules had undergone some shearing and veins of melt could be observed extending through the graphite towards the Pt capsules. Melt compositions from these experiments contained noticeably lower Fe contents compared with other experiments conducted at similar pressure–temperature conditions and reasonable mass balancing of Fe could not be made. Melt from these experiments are deemed to have lost Fe as a result of contact and Fe alloying with the Pt capsule. There is a progressive variation in the Fe content of the quenched melt veins cutting the graphite capsules, with the most Fe-depleted melt located close to the contact with the Pt capsule. As Fe diffusion in high-temperature melts is very rapid, the preservation of this Fe zonation indicates that the shearing of the graphite capsules and Fe loss from the melt occurred at a late stage of the experiments, possibly during run quenching. In this case, Fe loss from the melt is unlikely to have significantly affected the residual mineral compositions. Melt compositions have been determined across the entire range of investigated conditions and these melt compositions vary systematically with pressure and temperature (see below). Therefore, by interpolating between melt compositions from experiments conducted under similar conditions (interpolation between runs MH17 and MH14 for run MH15, and interpolation between run MH23 and the GA2 bulk for run MH22) we calculate that the melt from experiments MH15 and MH22 lost 2·7 and 1·5 wt% FeO, respectively. Recalculation of the melt compositions from these two experiments to higher Fe contents not only is consistent with concentrations expected from the degree of partial melting, but also allows for reasonable calculations of mass balance.

Phase proportions of all experiments were calculated based on major-element mass balance of all phases against the GA2 starting composition using least-squares regressions. Microsoft Excel® and Analyse-it® were employed to calculate phase proportions in experiments containing four or fewer phases. Both programs gave very similar results. As K₂O is highly incompatible and is below detection limits in garnet and clinopyroxene (see below), weighting of K₂O was increased to allow for more accurate calculation of melt proportions. In contrast, for the low-temperature experiments that contain more than four phases these programs were unable to calculate accurate phase proportions. For these experiments, phase proportions were first estimated based on petrographic observations and extrapolations from phase proportions of the higher temperature experiments. These estimated proportions were then adjusted iteratively until the best fit (minimum residuals to GA2) was obtained. As a cross check, incompatible trace-element (e.g. Cs, Rb, U) concentrations of the melts were also scaled against the composition of GA2 to provide an estimate of melt proportion (Table 2). The trace-element data will be presented elsewhere, but in all cases there is excellent agreement between the calculated melt proportions and constraints from the trace-element data. Details of the experimental run conditions, phase relations and phase proportions are presented in Table 2.

Attainment of equilibrium
Many aspects of the experimental protocol were specifically designed to optimize equilibrium at run conditions, including using a homogeneous glass starting material, careful assembly of experiments, conducting long duration experiments under stable pressure and temperature conditions, and ensuring rapid quenching at run termination. Nonetheless, all experimental products were carefully scrutinized to evaluate attainment of equilibrium. Within error, no systematic variations in mineral compositions were found across the width or length of the experimental charges, although a few experiments conducted close to the solidus contained some garnets grains with cores slightly higher in Ti than the rims. These garnet cores are interpreted to represent garnet growth before complete equilibrium was reached, and so we consider only the garnet rims as part of the equilibrium mineral assemblages. Mineral zoning (exclusive of quenching effects) was not found in any other minerals. Temperatures calculated from coexisting garnet–clinopyroxene pairs using the geothermometer of Ellis & Green (1979) are within 50°C of the nominal run temperature for all experiments, and within 25°C of the nominal run temperature for all but four experiments (Table 2). These values are within the minimum uncertainty of ±50°C for the geothermometer (Ellis & Green, 1979). Melt compositions were also homogeneous throughout the capsules of individual experiments. Melt compositions around the carbon spheres were found to be similar to melt pools present in the mineral matrix, although the latter melt pools were more variable in composition as a result of quench modification. In these cases only analyses of melt from the carbon spheres were used to calculate representative melt compositions.

	RESULTS

TOP ABSTRACT INTRODUCTION EXPERIMENTAL SETUP RESULTS DISCUSSION CONCLUSIONS REFERENCES

 
Textural observations and phase proportions
In experiments conducted below or close to the solidus, the GA2 starting material recrystallized to form fine-grained (1–10 µm) eclogite dominated by clinopyroxene and garnet, with minor quartz–coesite, rutile and feldspar. In these experiments quartz–coesite is commonly observed in the matrix assemblage and as inclusions in garnet (Fig. 2a). With increasing temperature and, to a lesser extent, pressure the grain size of the experimental charges increases (Fig. 2), with individual clinopyroxene and garnet grains reaching hundreds of micrometres in size in some experiments containing high melt proportions (> 40%). In experiments with melt proportions in excess of 10–15%, melt appears to wet all grain boundaries and is completely interconnected throughout the charge (Fig. 2c). Phase proportions for all experiments are given in Table 2. From the quartz–coesite transition of Bose & Ganguly (1995), we expect quartz to be stable at 3·0 GPa and coesite to be stable in experiments at 4·0 and 5·0 GPa.

View larger version (80K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 2. Backscattered electron images of experimental runs at 3·0 GPa. (a) Subsolidus run MH1(1200°C) containing garnet (gt), clinopyroxene (cpx), quartz (qtz), rutile (rt) and feldspar (kspr). (b) Run MH18 (1260°C) containing gt, cpx, rt, kspr, qtz and a low-degree melt. (c) Run MH5 (1320°C) containing only gt, cpx and melt. The increase in grain size with increasing temperature and melt fraction should be noted.

 
The solidus for GA2 is located at close to 1230°C at 3·0 GPa, around 1260°C at 4·0 GPa and around 1300°C at 5·0 GPa, giving a solidus slope of 30–40°C/GPa (Fig. 3). At 3·0 GPa garnet, clinopyroxene, quartz, rutile and feldspar are all present up to 1260°C, but quartz, rutile and feldspar are completely melted out by 1280°C. At higher pressure, feldspar is rapidly melted out above the solidus, and the width of the temperature interval in which rutile is stable above the solidus appears to diminish with pressure (Fig. 3). Coesite is stable up to 70°C above the solidus. The liquidus surface is less well constrained, although projections of melting percentages with temperature (Fig. 4) indicate that the liquidus at 3·0 and 4·0 GPa are located at approximately 1460 and 1520°C, respectively. GA2 is completely liquid at 5·0 GPa and 1600°C. These results constrain the slope of the liquidus to around 70°C/GPa. Melting intervals range from 230°C at 3·0 GPa to around 300°C at 5·0 GPa.

View larger version (23K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 3. Experimentally determined phase relations for GA2. •, experiments containing melt; {cir}, subsolidus experiments; filled grey dots, experimental results on GA1 starting material from Yaxley & Green (1998). The numbers below the dots refer to the percentage melting of the experiments. (See text for details.) The location of the quartz–coesite transition is from Bose & Ganguly (1995).

 

View larger version (24K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 4. Calculated phase proportions of GA2 at (a) 3·0 GPa, (b) 4·0 GPa and (c) 5·0 GPa. The sharp inflections in melt productivity once minor phases are melted out at 3·0 GPa and at around 25–30% melting at all pressures should be noted.

 
Phase proportions calculated from mass balance (Table 2, Fig. 4), reveal distinct inflections in melt productivity (that is, the rate of increase of melt fraction) as a function of temperature. At 3·0 GPa melting increases rapidly above the solidus until quartz, feldspar and rutile are exhausted at 1280°C. Melt productivity then decreases markedly until the melt proportion reaches around 30%, at which point rapid melting again resumes until the liquidus is reached (Fig. 4a). At 4·0 GPa initial melting close to the solidus eliminates both feldspar and rutile. Again, melt productivity with garnet and clinopyroxene (± coesite) remains low until c. 25% melting is reached, after which melt productivity increases up to the liquidus (Fig. 4b). At 5·0 GPa, a similar trend in melting systematics can be discerned, although it is noteworthy that the increase in melt productivity at around 25–30% melt fraction is less pronounced. Proportions of both garnet and clinopyroxene decrease progressively with increasing melt fraction and show little variation with pressure (Fig. 5). The only notable exception is in experiments that contain less than 20% melt, where a distinct increase in garnet mode as a function of pressure is apparent (Fig. 5b). Near-liquidus experiments (MH20 and MH22) and projection of phase proportions (Fig. 4) indicate that garnet is the liquidus phase at all investigated pressures.

View larger version (12K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 5. (a) Clinopyroxene proportions, (b) garnet proportions, and (c) Al2O3 content of clinopyroxene vs percentage of melt in all experiments. (Note the increase in garnet proportion with pressure at melt fractions below 20%, and the increase in Al2O3 in clinopyroxene with pressure at high melt fractions.)

 
Mineral compositions
Garnet
Garnet is present in all subliquidus experiments and in most cases forms porphyroblastic grains of homogeneous composition. Representative analyses are presented in Table 4. Garnets from all experiments are almandine (20–40 mol%) and pyrope (40–60 mol%) rich, with around 20% grossular component. Compositional changes with pressure are minimal, with only a slight decrease in almandine content observed with increasing pressure. In contrast, large compositional changes in garnet accompany increasing temperature or melt production. Trends of increasing garnet Mg-number with temperature (Fig. 6a) compare closely with the trends of melt productivity (Fig. 4) at the respective pressure conditions. Most notably, the kinks in the garnet Mg-number trends at Mg-number 56–58 (Fig. 6a) directly correspond to the prominent kinks in the melt proportion trends at 25–30% melting (Fig. 4). Therefore, there is a direct linear relationship between garnet Mg-number and percentage melting (Fig. 6b). Reduction of grossular contents with increasing melt proportion is relatively minor. All analysed garnets also have relatively high TiO₂ contents (0·5–1·5 wt%) and detectable Na₂O. In contrast, K₂O was below detection limits (0·2 wt%) in all experiments.

View larger version (16K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 6. Experimental garnet compositions as a function of pressure, temperature and melt fraction.

 

View this table: [in this window] [in a new window]   Table 4: Representative compositions of experimental garnets

 
Clinopyroxene
Clinopyroxene is the dominant mineral phase to crystallize from GA2 in all experiments containing <40% melt. At higher degrees of melting garnet is the most abundant residual phase (Fig. 4). In all cases, clinopyroxenes are Al₂O₃ rich (13·5–15·7 wt%) and omphacitic (Table 5), and are comparable in composition with clinopyroxenes reported from other experimental studies of dry eclogite melting (e.g. Yaxley & Green, 1998; Klemme et al., 2002; Pertermann & Hirschmann, 2003b). As with garnet, clinopyroxene Mg-number trends tend to correlate directly with temperature or melt fraction. Jadeite contents decrease with increasing melt fraction, but increase significantly with increasing pressure, ranging from Jad_20–33 at 3·0 GPa to Jad_30–40 at 5·0 GPa (Fig. 7a). In contrast, Ca-Tschermaks and enstatite–ferrosilite components increase with temperature, but decrease with pressure (Fig. 7b and c). Diopside–hedenbergite components are around 30 mol% and show little change with temperature and only a slight decrease with pressure (Fig. 7d). Al₂O₃ contents of clinopyroxene decrease with increasing melting at 3·0 GPa, whereas at higher pressures the Al₂O₃ contents tend to increase with melt fraction (Fig. 5c). In contrast, the TiO₂ in clinopyroxene is seemingly unaffected by pressure, but is influenced by the presence of rutile and temperature. At all pressure conditions TiO₂ increases (up to 2·0 wt%) until rutile is melted out, then progressively decreases to around 0·5 wt% with increasing temperature (Fig. 7f). As with garnet, K₂O was below detection limits (0·2 wt%) in clinopyroxene from all experiments.

View larger version (26K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 7. Experimental clinopyroxene compositions as a function of pressure and temperature: (a) jadeite component; (b) Ca-Tschermaks component; (c) enstatite–ferrosilite component; (d) diopside–hedenbergite component; (e) Ca-eskolaite component; (f) TiO2 content.

 

View this table: [in this window] [in a new window]   Table 5: Representative compositions of experimental clinopyroxenes

 
Calculated molecular formulae for clinopyroxene always give low cation totals (Table 5), which are not an analytical artefact but are attributed to the presence of a Ca-eskolaite (Ca-Es) component (Wood & Henderson, 1978; Gasparik, 1986; Pertermann & Hirschmann, 2002). Ca-Es components were calculated according to Pertermann & Hirschmann (2002) and, like jadeite contents, increase with increasing pressure from 5–12 mol% at 3·0 GPa to 15 mol% at 5·0 GPa (Fig. 7e). At low melt fractions the Ca-Es component tends to be highest but then decreases markedly once melt fractions exceed 20–30%. This transition corresponds to melting out of quartz–coesite from the residue (Fig. 4), which is consistent with the high-pressure reaction

as proposed by Gasparik (1986) and Pertermann & Hirschmann (2002). This substitution is also consistent with the inverse relationship between Ca-Tschermaks and Ca-Es components observed in the experimental clinopyroxenes (Fig. 7b and e).

Feldspar
Fine-grained feldspar was identified in experiments up to 1260°C at 3·0 GPa and in subsolidus experiment MH32 at 5·0 GPa, but could only be quantitatively analysed in the 3·0 GPa experiments (Table 6). At 3·0 GPa, subsolidus feldspar (MH1) is sodian sanidine with minor Ca contents (San₄₁Ab₅₀An₉). Subsolidus feldspars of similar composition were reported for experiments on anhydrous granite at high pressure and temperature (Green & Lambert, 1965). In our experiments feldspar compositions change dramatically above the solidus, becoming progressively depleted in K and enriched in Na and (to a lesser extent) Ca with increasing melt fraction (Fig. 8). Comparison of the phase relations (Fig. 3) and the projected feldspar compositional trends with temperature indicates that the feldspar would be essentially K-free at the point of complete feldspar exhaustion, as shown in Fig. 8.

View larger version (21K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 8. Experimental feldspar compositions at 3·0 GPa. It should be noted that the extrapolated feldspar composition is K-free at the point of feldspar melting out.

 

View this table: [in this window] [in a new window]   Table 6: Representative compositions of experimental feldspars

 
Melt compositions
Average major-element compositions of melt from all melting experiments are presented in Table 7 and Fig. 9. Melt compositions vary significantly as a function of melt fraction and pressure. Compositions range from dacitic to andesitic with FeO, MgO, CaO and Al₂O₃ behaving compatibly and K₂O, TiO₂ and P₂O₅ behaving incompatibly. At melt fractions below 40% there is considerable variation in melt composition as a function of pressure. Trends in SiO₂ and MgO do not vary significantly with pressure, whereas FeO and CaO more readily enter the melt at higher pressure. There is a pronounced decrease in melt Al₂O₃ and Na₂O contents with increasing pressure. Most notably, Na₂O is incompatible at 3·0 GPa, with Na₂O contents increasing until feldspar is melted out, whereas Na₂O is compatible in the solid residue at 5·0 GPa (Fig. 9). The TiO₂ content of the dacitic melts also varies significantly with pressure, ranging from around 2·5 wt% TiO₂ at 3·0 GPa to 5·5 wt% TiO₂ at 5·0 GPa. Peak TiO₂ contents tend to be close to melting out of rutile. K₂O is highly incompatible at all pressure conditions, but is slightly less incompatible at 3·0 GPa where feldspar is still present. P₂O₅ is incompatible at all pressures.

View larger version (22K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 9. Experimental melt compositions as a function of pressure and melt fraction. Data on melting of the GA1 starting material from Yaxley & Green (1998) are also included.

 

View this table: [in this window] [in a new window]   Table 7: Average major-element (wt% oxides) compositions of experimental melts

 

	DISCUSSION

TOP ABSTRACT INTRODUCTION EXPERIMENTAL SETUP RESULTS DISCUSSION CONCLUSIONS REFERENCES

 
Phase relations and element partitioning
In this study we have conducted experiments over a large pressure and temperature range using the piston-cylinder apparatus. Therefore, our experimental data allow detailed evaluation of phase relations and compositions of anhydrous eclogite as a function of pressure and temperature. As with all high-pressure (>2·0 GPa) studies of mantle rocks of mafic composition, the dominant mineral phases of GA2 are garnet and clinopyroxene. Both of these phases become progressively Mg-rich and Ti- and Fe-depleted with increasing melt fraction. At very high melt fractions clinopyroxene is lost from the residue, leaving garnet as the sole liquidus phase of GA2 at all investigated pressure conditions.

The most striking feature of the clinopyroxene compositions is the increase in jadeite and Ca-eskolaite components and decrease in Ca-Tschermaks and enstatite–hedenbergite components with increasing pressure (Fig. 7). A direct outcome of these compositional changes is a shift to more aluminous clinopyroxene compositions at higher pressure and high melt fractions (Fig. 5c). The increase in jadeite stability effects a strong shift towards partitioning of Na into clinopyroxene rather than melt with increasing pressure (Fig. 10). In contrast, clinopyroxene–melt partitioning of Ti appears to decrease only slightly with increasing pressure (Fig. 10), despite significantly higher TiO₂ content in the melt at higher pressure (Fig. 9). At subsolidus to near-solidus conditions increased stability of jadeite with pressure is also likely to affect coexisting feldspar compositions. At the investigated conditions pure albite will break down to jadeite plus quartz–coesite at around 3·5 GPa (Holland, 1980). In our experiments at 3·0 GPa feldspar contains significant Na₂O, and becomes progressively K₂O depleted with increasing melting (Fig. 8), indicating that the albite component may be crucial for feldspar stability above the solidus at P < 3·5 GPa. Above 3·5 GPa the albite component is expected to react to form jadeite (Holland, 1980), so the subsolidus feldspars are likely to be K-rich. As K-feldspar appears to be rapidly melted out, this compositional change with pressure results in closing of the feldspar field against the solidus with increasing pressure (Fig. 3). The correlation of melt TiO₂ content with pressure at low melt fractions (Fig. 9) also indicates that rutile may be more readily melted out with increasing pressure, which is consistent with the determined phase relations (Fig. 3) and the assertions of Klemme et al. (2002). In contrast, pressure does not seem to have an effect on quartz–coesite stability with respect to the solidus, at least under the conditions investigated.

View larger version (15K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 10. Partitioning of (a) Na and (b) Ti between clinopyroxene and melt as a function of pressure and temperature. Y&G, Yaxley & Green (1998); P&H, Pertermann & Hirschmann (2003b). (Note the pronounced increase in DNacpx/melt with increasing pressure, whereas DTicpx/melt decreases with increasing pressure.)

 
Comparison with previous studies
Previous experiments on dry melting of eclogite or garnet pyroxenite included a large range of starting compositions and hence have collectively resulted in a large variation of potential melting conditions (e.g. Kogiso et al., 2004). A number of these previous studies used starting materials that are comparable in composition with GA2 used in this study (Table 1). The GA2 composition was synthesized to be similar to the GA1 composition of Yaxley & Green (1998) so our new experimental results are directly comparable with those of Yaxley & Green (1998). In particular, their phase relations and mineral and melt compositions for experiments at 3·5 GPa are broadly consistent with interpolations of our results from experimental runs at 3·0 GPa and 4·0 GPa (Fig. 3). A notable difference is the absence of rutile and feldspar reported for the near-solidus experiment (C338) of Yaxley & Green (1998). In this experiment rutile saturation may not have been achieved because of the lower TiO₂ content of GA1 (Table 1), and feldspar may have been misidentified as melt because of the similarity in composition of melt and feldspar expected at these conditions (e.g. Tables 6 and 7). Takahashi et al. (1998) conducted experiments on a K-bearing basalt (CRB 72-31) of similar composition to GA2 and located the liquidus at between 1450 and 1475°C at 3·0 GPa, in good agreement with our results. In contrast, the solidus for CRB 72-31 at 3·0 GPa was not bracketed, but assumed to lie close to 1400°C, despite the presence of only garnet and clinopyroxene as residual phases at 1400°C (Takahashi et al., 1998). In our experiments at 3·0 GPa, rutile, quartz and feldspar are present as residual phases to around 50°C above the solidus (Fig. 3), which indicates that the solidus for CRB 72-31 at 3·0 GPa should be located below 1350°C.

Yasuda et al. (1994), Pertermann & Hirschmann (2003a, 2003b) and Yaxley & Sobolev (2007) conducted experiments on MORB or oceanic gabbro starting materials (NAM-7, G2 and Gb108, respectively) that have lower K₂O contents than GA2 (Table 1). The solidus temperatures for 3·0–5·0 GPa quartz–coesite-bearing eclogite obtained from these studies are at least 50–150°C higher than for GA2. The difference in melting conditions is almost certainly due to the higher K₂O content of GA2, which stabilizes subsolidus sanidine feldspar and allows melting to begin at lower temperatures. This is also consistent with the very low solidus temperatures and phase assemblage of alkali basalt JB1 reported by Tsuruta & Takahashi (1998) and Wang & Takahashi (1999). Those researchers identified sanidine, rutile and quartz–coesite as subsolidus phases below 5·0 GPa, and determined solidus temperatures between 3·0 and 5·0 GPa as below 1200°C and 1300°C, respectively.

Most studies of eclogite melting have found that the melting interval is relatively small (< 200°C) and tends to decrease with increasing pressure (e.g. Yasuda et al., 1994; Pertermann & Hirschmann, 2003a). A notable feature of the phase relations determined for GA2 is that the melting interval increases with pressure from 230°C at 3·0 GPa to 300°C at 5·0 GPa, largely as a result of the low dT/dP of the solidus (Fig. 3). We attribute this difference in melting behaviour to depression of the solidus by the stability of accessory K-feldspar and rutile. With increasing pressure the stability fields of these phases shrink with respect to the solidus as Ti is increasingly partitioned into the melt (Fig. 9) and Na and K solubility in clinopyroxene and/or garnet increases (Fig. 10; Tsuruta & Takahashi, 1998; Wang & Takahashi, 1999; Harlow & Davies, 2004). Tsuruta & Takahashi (1998) also found a widening melting interval for alkali basalt JB1 with increasing pressure up to 5·0 GPa, which was also attributed to the presence of K-feldspar. However, between 6·0 and 7·0 GPa those workers observed a rapid increase in the solidus temperature, which correlates with a dramatic increase in K solubility in clinopyroxene and the loss of feldspar as a subsolidus phase (see also Wang & Takahashi, 1999). At pressures above 7·0 GPa phase relations and liquidus and solidus temperatures for alkali basalt (Tsuruta & Takahashi, 1998) and MORB (Yasuda et al., 1994) (both forming coesite eclogite) are remarkably similar. Given the similarity in phase relations of alkali basalt JB1 and GA2 between 3·0 and 5·0 GPa, we expect that phase relations for JB1 and GA2 at P above 5·0 GPa are also comparable.

Several other experimental studies have focused on melting of quartz–coesite-absent garnet pyroxenites that are considered to be comparable with many mantle xenoliths (Ito & Kennedy, 1974; Hirschmann et al., 2003; Kogiso et al., 2003; Keshav et al., 2004). These compositions have appreciably lower contents of K₂O, Na₂O, TiO₂ and SiO₂, are substantially richer in MgO and undergo melting at significantly higher temperatures (>200°C) than eclogites of basaltic protolith. Kogiso & Hirschmann (2006) investigated phase relations and melting of bimineralic eclogite with a bulk composition of SiO₂-depleted MORB at 3·0 and 5·0 GPa. The solidus temperature for this composition was also found to be significantly higher (around 100°C) than that of other MORB-like compositions such as G2 (Pertermann & Hirschmann, 2003a) or NAM-7 (Yasuda et al., 1994). Of note, a feature of the Mg-rich and bimineralic garnet pyroxenites is a very narrow melting interval of <150°C (Ito & Kennedy, 1974; Kogiso et al., 2003; Keshav et al., 2004; Kogiso & Hirschmann, 2006), which indicates that melt is produced very rapidly at temperatures above the solidus. These results are consistent with phase petrology theory, which predicts higher solidus temperatures and smaller melting intervals as the bulk composition approaches simple two-phase mixtures of pyrope–almandine and diopside–hedenbergite (Kogiso & Hirschmann, 2006).

Collective evaluation of the melting conditions of eclogite or garnet pyroxenite allows us to make some generalizations on high-pressure melting of mafic materials as a function of composition. Kogiso et al. (2004) and Kogiso & Hirschmann (2006) have already demonstrated that increasing alkalis (K₂O and Na₂O) will tend to reduce the solidus temperatures of eclogite or pyroxenite. However, such compositional changes seem to have little influence on liquidus temperatures. K-rich quartz–coesite eclogite will begin to rapidly melt at relatively low temperatures until minor phases (e.g. feldspar, rutile) are exhausted. After melting out of these minor phases, melt productivity significantly decreases until at a certain melt fraction the melting rate dramatically increases again (see Fig. 4). For GA2 this rapid increase in melting occurs at around 25–30% melting and corresponds to a clear change in garnet compositional trends (Fig. 6) and a reduction of variation in melt compositions as a function of pressure and temperature (Fig. 9). Similar abrupt changes from low melt productivity to high melt productivity have been reported from melting experiments on other alkali-rich mafic starting compositions (Tsuruta & Takahashi, 1998; Tuff et al., 2005). In particular, Tuff et al. (2005) documented a rapid increase in melting of alkali ferropicrite (97SB68) at melt fractions of 20–30% from 3·0 GPa to 7·0 GPa. Such inflections seem to be absent or unresolvable from melting curves of alkali-poor eclogite or bimineralic pyroxenites (e.g. Ito & Kennedy, 1974; Kogiso et al., 2003; Pertermann & Hirschmann, 2003a; Keshav et al., 2004; Kogiso & Hirschmann, 2006). In these cases, melt productivity is high and is exclusively or almost exclusively controlled by garnet and clinopyroxene. For our GA2 composition and the 97SB68 composition of Tuff et al. (2005), rapid melting conditions at melt fractions of >30% also correspond to melting of a garnet–clinopyroxene residue. The change in melt productivity at these melt fractions corresponds to a change from melting conditions that are influenced by minor phases such as feldspar, rutile and quartz–coesite, to melting conditions that are controlled exclusively by garnet and clinopyroxene. Chemically, these minor phases reflect the importance of K₂O, Na₂O, TiO₂ and SiO₂ in controlling solidus temperatures. For GA2 the position of the melting inflection is parallel to the liquidus (Fig. 3) and is roughly consistent with the solidus temperature expected of a bimineralic garnet clinopyroxenite (Kogiso et al., 2004; Kogiso & Hirschmann, 2006). Based on these observations we label this inflection surface the garnet–clinopyroxene residue melting surface (Fig. 3).

In sum, there are two conditions of rapid melting of eclogitic rocks with increasing temperature. The first is near-solidus melting that is dominated by minor phases (e.g. rutile, feldspar, apatite, quartz–coesite). The second condition corresponds to melting of the garnet–clinopyroxene residue, which approaches the solidus and melting behaviour of two-phase garnet pyroxenite. These two melting intervals of relatively high melt productivity are separated by an interval of low melt productivity in the pressure–temperature range where garnet and clinopyroxene are dissolving in the low-degree melt component (e.g. Fig. 4a).

Melt compositions as a function of temperature and pressure
At all pressure conditions investigated, low-degree partial melts of GA2 are dacitic with high K₂O contents and K/Na ratios. Similar melt compositions were reported for low-degree melting of GA1 (Yaxley & Green, 1998) and G2K (Pertermann & Hirschmann, 2003b); anhydrous starting materials that are of similar compositions to GA2. For compositions with even higher alkali element contents and K/Na ratios, low-degree melts tend to approach rhyolitic compositions (Tsuruta & Takahashi, 1998). At similar melt fractions, partial melts of low-alkali quartz eclogite tend to be andesitic with higher TiO₂ contents (Pertermann & Hirschmann, 2003b). Pertermann & Hirschmann (2003b) attributed these differences to the presence of alkali elements in the melt, which serve to increase SiO₂ and decrease TiO₂ contents of the melt while saturated in quartz–coesite and rutile. For garnet pyroxenites that lie on, or to the forsterite side of, the garnet–pyroxene thermal divide in CMAS or high-pressure projections [see Yaxley & Green (1998) and Kogiso et al. (2003) for details] melt compositions will be silica-undersaturated, Mg-rich picrites near the solidus (or at the ol–opx + cpx + ga ‘eutectic’), becoming basaltic with increased melting of residual garnet and clinopyroxene (Yaxley & Green, 1998; Kogiso et al., 2003; Kogiso & Hirschmann, 2006).

Based on high-pressure (3·0–7·5 GPa) multi-anvil experiments, Kogiso et al. (2003) and Kogiso & Hirschmann (2006) documented a general increase in the CaO/Al₂O₃ and decrease in the Na₂O contents of partial melts of garnet pyroxenite with increasing pressure. These compositional changes were respectively attributed to the increase in garnet stability relative to clinopyroxene (Kogiso et al., 2003) and increasing Na₂O partitioning in clinopyroxene (Kogiso & Hirschmann, 2006) with increasing pressure. Other than these two studies, the effect of pressure on eclogite partial melt compositions has not been investigated. From our experimental results we observe some striking compositional variations in low-degree (< 30%) partial melts of GA2 as a function of pressure (Fig. 9). With increasing pressure, FeO and CaO more readily enter the melt, whereas Na₂O and Al₂O₃ become increasing compatible in the solid residue. These effects are a direct consequence of increasing stability of garnet and jadeitic clinopyroxene with increasing pressure, which is in accordance with the finding of Kogiso et al. (2003) and Kogiso & Hirschmann (2006). The mode of garnet increases with pressure at melt fractions below 20% (Fig. 5b), which tends to produce melt with significantly lower Al₂O₃ contents. Furthermore, the slight shift to higher Mg-number in garnet with pressure (Fig. 6) effects the increase of melt FeO content with pressure. The increase in jadeite component of clinopyroxene with pressure is accompanied by a reduction in the Ca-Tschermaks and diopside components (Fig. 7), and increased CaO partitioning into the melt. At high melt fractions, garnet modes do not vary with pressure (Fig. 5b). However, the enhanced stability of jadeite with pressure produces pyroxenes with higher Al₂O₃ contents (Fig. 5c), and hence coexisting melts remain Al₂O₃ poor in these conditions (Fig. 9). Progressive increase in jadeite in pyroxene at high pressure also significantly affects clinopyroxene–melt partitioning of Na₂O (Fig. 10). At relatively low pressure (2·0–3·0 GPa) Na₂O behaves incompatibly during partial melting (Fig. 10; Pertermann & Hirschmann, 2003b), whereas at 5·0 GPa and low melt fractions, Na₂O is more compatible in the solid residue. A similar pressure effect on Na₂O concentration was also described for low-degree melts of lherzolite (Longhi, 2002).

Melt SiO₂ and MgO contents do not vary significantly with pressure. As expected, K₂O is highly incompatible, although slightly less so at 3·0 GPa, whereas feldspar remains stable in the residue. TiO₂ contents of low-degree melts increase dramatically with pressure; this is partly due to decreasing relative stability of rutile with increasing pressure (Fig. 3), but may also be due to decreasing TiO₂ partitioning between clinopyroxene (and garnet) and melt with pressure. Results from our experiments and those of Yaxley & Green (1998) indicate that Ti partitioning between clinopyroxene and melt decreases with increasing pressure and temperature, whereas data from Pertermann & Hirschmann (2003b) at 2·0–3·0 GPa do not indicate any pressure dependence (Fig. 10). Other than slight differences in bulk starting compositions we are unable to explain the inconsistency between these results. However, our results are again consistent with data from lherzolite melting experiments, which show a slight increase in TiO₂ contents of melts with increasing pressure (Longhi, 2002). Longhi (2002) also observed a decrease in Ti partitioning between clinopyroxene and melt with increasing melt fraction, which was largely attributed to changes in melt compositions.

At high degrees of melting of GA2 (> 40%) melt compositions become remarkably similar, regardless of the pressure (Fig. 9). The notable exception is the Al₂O₃, which remains relatively low in the 4–5 GPa melts as a result of stabilization of more aluminous clinopyroxene, as discussed above. At these high melt fraction conditions—above the so-called ‘garnet–clinopyroxene residue melting surface’—melting is completely controlled by garnet and clinopyroxene and the initial SiO₂–TiO₂–K₂O-rich liquid fractions become increasing diluted. A similar behaviour at high melt fractions is expected for a large range of high-pressure mafic rock-types, as under these conditions garnet and clinopyroxene are the principal, if not only, residual phases.

Expressed in phase equilibria terms, solidus temperatures and near-solidus melts in eclogitic rocks are strongly dependent on the minor phases and degree of silica undersaturation relative to the garnet + clinopyroxene compositional plane. The degree of silica undersaturation in garnet pyroxenite or eclogite is reflected in the presence of quartz–coesite or olivine (oversaturated and undersaturated, respectively) and more subtly in the presence of rutile, rutile + ilmenite, or ilmenite, and in the eskolaite content of clinopyroxene (silica oversaturated). Coupled with the silica-saturation control on partial melting, the presence of minor phases (ilmenite–rutile, corundum–kyanite, K-feldspar–phlogopite, apatite, carbonate–graphite) have the major control on solidus temperature, melt fraction and melt composition at temperatures below the garnet + clinopyroxene solidus. Once the latter solidus is reached, liquids approach the garnet + clinopyroxene plane in their compositions (i.e. approach basaltic compositions) from either the silica-undersaturated or -oversaturated (andesitic) side. It should be noted that this emphasis on silica saturation at high pressure uses a different measure [i.e. a high-pressure normative mineralogy (O’Hara, 1968; Yaxley & Green, 1998)] from that appropriate at low pressures (CIPW norm and basalt tetrahedron; Yoder & Tilley, 1962).

The contrasting behaviour of potassium and sodium at high P
A basic premise underpinning igneous geochemistry is that alkali elements Na and K are highly incompatible for a large range of igneous processes and hence in most cases these elements are considered to behave in a very similar manner [e.g. the total alkalis–silica (TAS) diagrams of Le Bas et al. (1986)]. Large variations in K or Na are often observed in volatile-rich alkaline rock suites (e.g. Foley, 1992a), but these rocks are relatively rare. In contrast, little attention has been paid to the behaviour of K and Na during volatile-deficient magma generation at high pressure, either directly from experimental data or indirectly through the geochemistry of oceanic basalts. In general, the K₂O and Na₂O contents of most basaltic alkaline rocks do directly correlate, but there are often significant variations in Na/K ratios, even within single rock suites (e.g. Stolper et al., 2004). The explanation for these variations has remained elusive.

An interesting aspect of our melt chemistry results concerns the behaviour of Na₂O and K₂O at high pressure during partial melting. At 3·0 GPa both K₂O and Na₂O behave incompatibly during partial melting, but with increasing pressure Na₂O becomes progressively retained in the solid residue, whereas K₂O becomes increasingly incompatible as the stability of supersolidus feldspar diminishes (Fig. 9). Therefore, Na/K ratios in the melts vary substantially over a range of pressure conditions. As Na₂O is also hosted by the jadeite component of clinopyroxene in lherzolite, we might expect increasing pressure to have a similar effect on Na/K ratios of melts of mantle lherzolite. However, the stability of super-solidus phlogopite in lherzolite to high temperatures and low water fugacities (Foley et al., 1986; Conceição & Green, 2004) means that garnet lherzolite melting involves relative partitioning of K and Na between melt, clinopyroxene and phlogopite. Our results with eclogite demonstrate a mechanism by which deep partial melting and melt extraction from K₂O- and Na₂O-bearing material can effectively fractionate K₂O from Na₂O. As melting of such rocks is likely to contribute to oceanic magmatism to some degree (e.g. Hirschmann & Stolper, 1996; Hofmann, 2003), this process may explain some of the Na and K variations observed in oceanic basalts. Some specific examples are discussed below. However, we do not suggest that the Na/K ratio of oceanic basalts could be used as an exclusive monitor of the depth of melting, as there are many other potential factors that may influence the geochemistry of these rocks, such as geochemical variation of source rocks, degree of partial melting and melt–peridotite reaction processes. These are also discussed in more detail below.

Eclogite in the mantle
Subduction of oceanic crust and delamination of dense mafic cumulates from the base of continental crust are probably the most important mechanisms for recycling mafic rocks into the mantle (Hofmann & White, 1982; Hirschmann & Stolper, 1996; Hofmann, 1997; Anderson, 2006). Oceanic crust may contain significant quantities of silica- and alkali-rich rock-types such as sea-floor-altered MORB and sedimentary rocks. However, there is a widely held assumption among the geochemical community that oceanic crust undergoes extensive loss of Si and alkali elements as a result of dehydration or partial melting during subduction, and so mafic rocks returned to the mantle would tend to be Si and alkali deficient (e.g. Hirschmann et al., 2003; Kogiso et al., 2003; Niu & O’Hara, 2003; Stracke et al., 2003). In contrast, examination of eclogites from palaeo subduction zones indicates that geochemical changes associated with subduction-zone dehydration are minimal (Hermann, 2002; Chalot-Prat et al., 2003; Spandler et al., 2004). The quantitative extraction of elements from the slab will be very different according to whether a fluid (H₂O-rich) or hydrous melt is liberated from the eclogite and many temperature–pressure paths predicted for subducted slabs do not enter the water-saturated or dehydration melting regimes for basaltic crust. Significant leaching of alkalis and silica by fluids is likely to be restricted to fluid channelways in the slab (Scambelluri & Philippot, 2001; Breeding et al., 2004; Hermann et al., 2006), so large sections of oceanic crust may be deeply subducted into the mantle without undergoing significant geochemical modification (Spandler et al., 2004; Hermann et al., 2006).

If significant amounts of alkali-rich, silica-saturated mafic rock are present in the mantle, then these rocks may be involved in the production of oceanic magmas. The composition of the GA2 starting material used in this study was designed to be representative of sea-floor-altered MORB (see also Yaxley & Green, 1994, 1998), which is an important component of subducting oceanic crust (Staudigel et al., 1996). Therefore, the data presented in this paper may be used for understanding melting of heterogeneous mantle domains. In the next section we speculate on possible melting processes that may lead to the formation of intraplate oceanic magmatism.

Implications for the generation of oceanic basalts
There is currently intense debate regarding the physical nature and dynamics of the upper mantle (e.g. Hofmann, 2003; Niu & O’Hara, 2003; Anderson, 2006; Nielsen et al., 2006), yet most contributors to this debate accept that heterogeneous mantle sources are required to form many intraplate basalt and MORB suites. In the paragraphs above we have compared our results with other experimental studies on diverse eclogitic to pyroxenitic compositions to illustrate the range in sub-solidus mineralogies, solidus temperatures and melt productivity among these compositions. We have not discussed the roles of volatile elements (particularly C, H, O, S), but the presence of these elements in intraplate basalts shows that they cannot be neglected in any realistic models. Nevertheless, it is useful to consider the processes that may operate during buoyant upwelling of heterogeneous asthenospheric mantle that consists of variably depleted peridotite (fertile lherzolite to harzburgite) entraining metre- to kilometre-sized eclogite bodies, as described in the ‘marble-cake’ or ‘plum-pudding’ models for the mantle (Allègre & Turcotte, 1986; Phipps Morgan, 2001). However, here we also consider that the eclogitic rocks range in composition from alkali-rich, quartz–coesite-bearing eclogite (such as GA2) to olivine-normative garnet pyroxenite (such as MIX1G; Kogiso et al., 2003). Such domains may be representative of a ‘mantle plume’ (Hofmann & White, 1982; Hofmann, 1997); buoyantly upwelling relicts of ancient subducted slabs (Presnall & Helsey, 1982) or possibly the entire upper mantle (Anderson, 2006). We employ a mantle potential temperature of 1430°C, based on the calculated thermal conditions for mantle melting in suboceanic regimes by Green et al. (2001). Using a cooler mantle potential temperature will not alter any basic outcome of the modelling other than to shift melting of various lithologies to lower pressures by around 1 GPa per 100°C.

At depths greater than 200 km (> 6·5 GPa) in the upper mantle all mafic rocks will consist almost entirely of garnet and clinopyroxene, and hence their solidus will be higher than the mantle adiabat, as shown in Fig. 11. During upwelling the more alkali-rich rocks will exsolve feldspar at around 6·5 GPa (Wang & Takahashi, 1999), which will cause a large decrease in the solidus temperature (Tsuruta & Takahashi, 1998). As the upwelling mantle adiabat crosses the solidus for alkali mafic rocks at a high angle, localized high melt productivity will result at 6·5–6·0 GPa (Fig. 11, stage 1) to produce dacitic–rhyodacitic magma with relatively high K₂O, TiO₂, CaO/Al₂O₃ and K₂O/Na₂O and residual garnet and clinopyroxene, the last with high jadeite and Ca-eskolaite components. Although acknowledging the speculative nature of eclogite permeability and melt mobility relations, textures from our experiments (Fig. 2) and the results of the ‘sandwich’ experiments of Yaxley & Green (1998) indicate that at melt fractions of around 10–15% eclogite melt permeability will be sufficient to allow melt extraction and infiltration into neighbouring peridotite. These melts will react rapidly with the peridotite to form refertilized orthopyroxene-rich rocks (Yaxley & Green, 1998; Yaxley, 2000; Sobolev et al., 2005, 2007). In contrast, melt extraction from the eclogite will leave a more refractory eclogite residue.

View larger version (41K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 11. Model for melting of an upwelling heterogeneous suboceanic mantle domain. The mantle adiabat is calculated using a mantle potential temperature of 1430°C from Green et al. (2001). GA2 and gt–cpx residue melting surface locations are from this study and extrapolations based on Tsuruta & Takahashi (1998). MORB solidus is from Yasuda et al. (1994) and Pertermann & Hirschmann (2003a). Mg-rich garnet pyroxenite (MIX1G) and mantle peridotite solidi are from Kogiso et al. (2003) and Hirschmann (2000), respectively. (See text for details.)

 
With further upwelling, the mantle adiabat will cross the solidus of MORB-like eclogite at close to 6 GPa, locally producing andesitic melts (Pertermann & Hirschmann, 2003a, 2003b). Extraction of this melt component into surrounding peridotite may also results in fertilization of the peridotite. At pressures of between 4 and 5 GPa the adiabat will cross the solidi for some refertilized peridotite (Yaxley, 2000), eclogitic gabbro (Yaxley & Sobolev, 2007), bimineralic eclogite (Kogiso & Hirschmann, 2006) and residual garnet pyroxenite that had undergone prior melt extraction (Fig. 11, stage 2). An effective melt continuum connecting the diverse lithologies may form, with the direction of melting and crystallization reactions, mixing and diffusion leading towards homogenization of phase compositions but not of phase proportions. Melts may be eliminated in an enlarged ‘refertilized mantle peridotite’ volume or may persist as picritic (in olivine-bearing domains) or basaltic (in eclogite-bearing, olivine-absent domains) compositions.

At around 3·5 GPa, relics of refractory high-MgO pyroxenites will begin to melt (Fig. 11, stage 3). High degrees of melting of these rocks will produce Si-undersaturated high-MgO alkali basalts (Kogiso et al., 2003). Finally, at 2·5–3·0 GPa the solidus of (volatile-absent) mantle lherzolite is crossed, causing partial melting and production of tholeiitic picrite with increasing melt fraction (Fig. 11, stage 4). By this stage all, or almost all, of the mafic lithologies would have partially melted to produce distinct magma types that may have reacted with neighbouring peridotite to varying extents. Furthermore, melting of these refertilized peridotites may also produce distinct magma compositions. If processes exist that allow extraction of diverse melts without reaction with channel walls or permeable wall rocks, then such melts may undergo mixing at varying stages and to varying extents.

Considering the scenario described above, the composition of oceanic magmas erupted at the surface (disregarding fractionation or crustal assimilation processes) may be the end products of a complex series of processes including multiple stages of partial melting and melt extraction, multiple stages and varying extents of reaction between these partial melts and peridotite, and multiple stages and variable extents of magma mixing. If melting of carbonate-rich lithologies or metasediments (Irifune et al., 1994; Yaxley & Brey, 2004; Dasgupta et al., 2006) and lithospheric refertilization and veining (Foley, 1992b; Pilet et al., 2005) are also considered the potential complexity is further increased. Because of the primary compositional variation of the source rocks, these processes may operate over a huge depth interval. Given this, it is perhaps not surprising that many previous workers have had difficulty in precisely explaining the composition of OIB suites by using a single melt composition or by simple end-member mixing models (e.g. see discussion by Hirschmann et al., 2003; Pertermann & Hirschmann, 2003b; Keshav et al., 2004).

Despite the potential complexity involved in the process of oceanic basalt genesis some constraints on the formation of these magma suites can be made based on the available experimental data. Hirschmann et al. (2003), Kogiso et al. (2003) and Kogiso & Hirschmann (2006) have shown that the compositions of silica-undersaturated alkali OIB have some similarities to melt compositions produced by deep melting of high-MgO or bimineralic garnet pyroxenite. Low-degree melting of peridotite that had previously been refertilized by eclogite-derived dacitic melt may be a viable source for voluminous flood basalt provinces (Yaxley, 2000). Sobolev et al. (2005, 2007) were able to precisely model the major- and trace-element geochemistry of a range of oceanic basalt suites and their crystals by mixing melts from two distinct mantle sources: (1) mantle peridotite; (2) pyroxenite formed by prior infiltration of peridotite by eclogite-derived silicious melts. In a similar way, Ito & Mahoney (2005) were able to model many of the isotopic and geochemical characteristics of OIB and MORB by considering melting of upwelling heterogeneous mantle domains. The models outlined by Ito & Mahoney (2005) and Sobolev et al. (2005, 2007) are comparable with our model outlined above. However, we emphasize the complexity of primary compositional variations in the recycled materials and variations in melt compositions as a function of pressure. These aspects may further explain some compositional features of Hawaiian lavas. For example, many of the distinctive chemical characteristics of the tholeiitic Koolau suite, such as low CaO/Al₂O₃ and K₂O/Na₂O and high SiO₂ (Frey et al., 1994), may derive from melting of peridotite that reacted with eclogite-derived melt at relatively low pressure (2–3 GPa). In contrast, the unusual K-rich, high CaO/Al₂O₃, high K₂O/Na₂O glasses reported from several Hawaiian volcanoes (Stolper et al., 2004) would be consistent with containing a component of eclogite melt generated at high pressure (Fig. 9).

	CONCLUSIONS

TOP ABSTRACT INTRODUCTION EXPERIMENTAL SETUP RESULTS DISCUSSION CONCLUSIONS REFERENCES

 
We have experimentally determined phase relations, melting conditions and mineral and melt compositions of an anhydrous, alkali-rich basaltic starting material (GA2) at pressure and temperature conditions from 3·0 to 5·0 GPa and from 1200 to 1600°C, respectively. Mineral assemblages across the pressure–temperature range are dominated by garnet and clinopyroxene, although quartz–coesite, rutile and feldspar are important minor phases at subsolidus to low-degree melting (< 20%) conditions. Because of the relatively high silica and alkali content, GA2 has lower solidus temperatures and wider melting intervals than many other experimentally investigated eclogites. Melt compositions progressively evolve from high-K dacitic compositions at low melt fractions to basaltic andesite at high degrees of melting. Both the garnet modal proportion and jadeite component of clinopyroxene increase with increasing pressure, which causes Al₂O₃ and Na₂O to be retained in the residue rather than partitioned into the melt with increasing pressure. This property leads to contrasting behaviour of Na and K at high pressure, which may help to explain variations in Na and K found in some oceanic magma suites.

Considering our results together with results of other experimental studies of high-temperature eclogite melting, we expect that melting of upwelling heterogeneous mantle domains may occur over a large pressure and temperature range. Melts from mafic lithologies are highly variable in composition and may react with mantle peridotite or mix with other melts during migration towards the surface. In this case, oceanic basalts may be the end products of a range of complex melting–refertilization–mixing processes in the mantle, which implies that simple mantle melting models may be inadequate to accurately model these magma suites.

	ACKNOWLEDGEMENTS

 
This work was supported by the ANU and the Australian Research Council. The Electron Microscope Unit, ANU, is acknowledged for access to the SEM, and we thank Dean Scott for technical advice and preparing and running many of the experiments. The manuscripts was significantly improved after critical reviews by Tetsu Kogiso, Maik Pertermann and Shantanu Keshav.

---

*Corresponding author. Present address: Institute of Geological Sciences, University of Bern, Baltzerstrasse 3, CH-3012 Bern, Switzerland. Email: Spandler{at}geo.unibe.ch

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