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Garnet in the Earth’s Mantle Bernard J. Wood*, Ekaterina S. Kiseeva* and Andrew K. Matzen* 1811-5209/13/0009-421$2.50 DOI: 10.2113/gselements.9.6.421 luminous garnet, (Mg,Fe2+,Ca)3 (Al,Cr) 2Si3O12 , is an important constituent of mantle peridotite (~10%) and of the other abundant upper mantle rock, eclogite (~50%). Its unusual crystal chemistry means that it strongly prefers some trace elements and confers a “garnet signature” on mantle melts. As depth increases from 250 to 600 km, garnet increases in abundance in mantle rocks, dissolving large fractions of the other silicates and becoming Si rich (majoritic). These compositional changes are observed in some garnets found as inclusions in diamond. Garnet disappears from mantle assemblages at about 700 km depth, where it is replaced by an even denser silicate, perovskite. A roxene in addition to olivine and orthopyroxene. Harzburgites and lherzolites are generally thought to be related to one another, with harzburgite being the solid residuum after partial melting of more fertile “primitive” lherzolite. This relationship led Ringwood (1962) to propose that the primitive, or unmelted, mantle has a composition that is a mixture of residual peridotite (75%) and basalt (25%). In this mantle composition, which Ringwood KEYWORDS : phase transformations, garnet, mantle peridotite, eclogite, diamond named “pyrolite,” garnet becomes inclusions, majorite stable at the expense of spinel at pressures above 2.8 GPa (>85 km INTRODUCTION depth; FIG. 1) at the solidus. The reaction relationship Garnet is a common mineral in crystalline rocks of the between the two aluminous minerals (spinel and garnet) Earth’s upper mantle and lower crust and occasionally may be represented approximately as: occurs in volcanic liquids. Garnet is unlike most other crustal minerals in that Mg and Fe2+ are 8-coordinated by 0.4Ca(Mg,Fe2+)Si2O6 + 3.2(Mg,Fe2+)SiO3 + (Mg,Fe2+)Al2O4 = oxygen while Al occupies a site with 6-fold coordination clinopyroxene orthopyroxene spinel (1) (Mg and Fe2+ are usually 6-fold and Al 4-fold coordinated (Ca,Mg,Fe2+) 3Al2 Si3O12 + (Mg,Fe2+) 2 SiO4 in other minerals) – garnet crystal chemistry is discussed garnet olivine further in Geiger (2013 this issue). As a result of their high cation coordination numbers, garnets have relatively o high densities, low compressibilites and increasing stability At the solidus temperature of 1460 C, the transition from spinel lherzolite to garnet lherzolite takes place over a depth with increasing pressure. In consequence, garnet-bearing rocks have higher seismic velocities and densities than interval of about 5 km (Robinson and Wood 1998). The low-pressure, garnet-free rocks of similar composition, a transformation has a positive pressure–temperature slope of characteristic that has been used in attempts to constrain ~40 bar/degree but is strongly curved because, as temperature decreases, both the solubility of alumina in pyroxene the compositions of the lower crust, upper mantle and transition zone. The high coordination numbers also mean and the disorder of Mg and Al in the spinel decrease. These changes have the effect of flattening the slope of the transithat garnet residual to, or fractionating from, silicate liquids imparts a characteristic geochemical signature on the tion in pressure–temperature space (FIG. 1). The overall result is that garnet peridotite is stable only in relatively trace element pattern of the product igneous rock. These cool and deep parts of the mantle (conditions that are often properties make garnet a petrologically and geochemically important constituent of the upper mantle and transition observed beneath thick cratonic lithosphere); elsewhere, spinel peridotite is the stable form. zone. In this article, we explore garnet’s occurrence in the mantle and its influence on the physical properties of Compositionally, peridotitic garnets are composed of about mantle rocks and on the geochemical properties of mantle- 75% pyrope (Mg Al Si O ), 10% grossular (Ca Al Si O ) 3 2 3 12 3 2 3 12 derived silicate melts. and 15% almandine (Fe 2+ Al Si O ). They have lower 3 GARNET IN UPPER MANTLE PERIDOTITES Mantle rocks are observed at the surface either as tectonic fragments (kilometre scale) or as inclusions in explosive eruptive rocks (centimetre scale). The majority of samples are peridotites, either harzburgites (~85% olivine, 15% orthopyroxene) or lherzolites, which contain clinopy- * Department of Earth Sciences, University of Oxford South Parks Road, Oxford OX1 3AN, UK E-mail: [email protected] E LEMENTS , V OL . 9, PP. 421–426 2 3 12 magnesium numbers [Mg# = molar Mg/(Mg+Fe2+)] than coexisting olivines, and the partitioning of Mg and Fe2+ between garnet and olivine provides a useful geothermometer for peridotites (O’Neill and Wood 1979). At pressures above the spinel stability field, garnet incorporates the chromium and Fe 3+ that are released on spinel breakdown. These elements substitute for Al3+ in the octahedral site. However, in contrast to the calcium end-members, Ca3Cr2Si3O12 (uvarovite) and Ca3Fe3+2Si3O12 (andradite), which are stable at crustal pressures, the Mg– Cr and Mg–Fe3+ end-member garnets are very unstable. 421 D ECEMBER 2013 100 2.5 Garnet peridotite 75 2 Spinel peridotite 1.5 1 800 Solidus 50 of the transition zone (410–660 km depth). In this depth interval, fertile peridotitic mantle becomes a bimineralic rock consisting of ~65% wadsleyite or ringwoodite (both high-pressure forms of olivine) and ~35% garnet (Ringwood 1991). The increase in the modal proportion of garnet takes place gradually over a wide depth interval due to pyroxene breakdown and its dissolution into the garnet structure as Mg and Si are transferred from 6- and 4-coordination in pyroxene to mixtures of 8- and 6-coordination and 6- and 4-coordination, respectively, in garnet. The continuous reactions result in increasing densities and seismic velocities of the mantle and may be simplified as follows: 900 1000 1100 1200 1300 1400 1500 Temperature (oC) The stability fields of spinel peridotite and garnet peridotite in pressure–temperature space. The intersection with the solidus is from Robinson and Wood (1998). The extension of the field boundary to lower temperatures is from a number of experimental studies. FIGURE 1 olivine (2) orthopyroxene As mentioned earlier, garnets have higher coordination numbers and higher densities than olivines or pyroxenes; “skiagite” garnet, Fe 2+ 3Fe3+ 2 Si3O12 , is no different. This means that, as pressure increases, the reaction will favour the reactant, or “skiagite,” side. As a result, at a fi xed oxygen fugacity relative to a standard buffer, such as fayalitemagnetite-quartz (FMQ), garnet will become richer in Fe3+ with increasing pressure. Alternatively, with increasing depth, a garnet of fi xed composition (Fe3+ content) will be stabilised at progressively lower oxygen fugacity relative to FMQ. Woodland and Koch (2003) measured garnets from Kaapvaal craton peridotites and found modest increases in Fe3+/Fe 2+ with increasing pressure, due, in part, to transfer of Fe3+ from clinopyroxene to garnet with increasing temperature. Calculated oxygen fugacities, however, decline relative to the FMQ buffer in the same P–T interval (Woodland and Koch 2003) because of the stabilisation of skiagite with increasing pressure. Some of the interesting side effects of this change are that coexisting C–H–O fluid phases should change from CO2 rich to CH4 rich with increasing depth and that the mantle may have a low enough oxygen fugacity to precipitate Ni-rich metallic alloy at depths greater than 300 km. With increasing pressure, mantle peridotite undergoes several important phase transitions in which garnet is a key player (Ringwood 1991). FIGURE 2 shows the proportions of different phases in peridotitic mantle as a function of depth. The modal amount of garnet increases from about 10% at the spinel–garnet transition to 35% in the middle E LEMENTS 2Ca(Mg,Fe2+)Si2O6 = Ca 2 (Mg,Fe2+)(MgSi)VISi3O12 clinopyroxene majoritic garnet (4) The result of dissolution of “excess” Mg (and other divalent cations) and Si into garnet (the “majorite” substitution) is that its Al 2O3 content declines gradually from ~22 to 8 wt% while its CaO content increases from ~4 to 7 wt% as pressure increases from 3 to 19 GPa and pyroxenes are absorbed into its structure (Wood 2000). Beyond 19 GPa (550 km depth), CaSiO3 perovskite becomes stable in peridotite (FIG. 2) and increasing depth leads to progressive extraction of Ca from garnet; for example, at ~23 GPa, which corresponds to the appearance of (Mg,Fe2+)SiO3 perovskite, the CaO contents of garnet are lower than 4 wt% (Wood 2000). Perovskite initially forms at the expense of ringwoodite, but as pressure increases, the majoritic [(Mg,Fe2+) 3 (MgSi)VISi3O12 ] component of the garnet begins to dissolve in the (Mg,Fe)SiO3 perovskite, increasing the alumina content of the remaining garnet (Wood 2000). Eventually the (Mg,Fe 2+)SiO3 perovskite dissolves the remaining aluminous garnet as the components Al 2O3, Fe3+ AlO3 and (Mg,Fe 2+)(Al,Fe3+)O2.5 (Frost and Langenhorst 2002), and garnet disappears from the peridotite assemblage. 200 Depth km The 6-fold-coordinated sites in garnet can incorporate enough Fe3+ that their contents can be easily measured by Mössbauer spectroscopy and calibrated as an oxygen barometer for garnet peridotites (Gudmundsson and Wood 1995; Stagno et al. 2013). There are several possible ways to represent the equilibrium between the three iron-bearing phases, the simplest being: garnet (skiagite) (3) NaAlSi2O6 + Ca(Mg,Fe2+)Si2O6 = NaCa(Mg,Fe2+)(AlSi) (5) VI Si O 3 12 clinopyroxene majoritic garnet The substitution of Cr and Fe3+ for Al is facilitated by the addition of Ca to the garnet, an example of the “reciprocal solution” effect (Wood and Nicholls 1978). 2Fe2+ 3Fe3+2 Si3O12 = 4Fe2+2 SiO4 + 2Fe2+ SiO3 + O2 4(Mg,Fe2+)SiO3 = (Mg,Fe2+) 3 (MgSi)VISi3O12 orthopyroxene majoritic garnet 300 Olivine (Mg,Fe) SiO 2 Pyroxene 4 400 13 Wadsleyite Garnet 500 600 700 800 Ringwoodite 21 (Mg,Fe)SiO3 perovskite with up to 5% Al O 2 CaSiO3 perovskite 3 + (Mg,Fe)O 900 0.4 0.5 0.6 0.7 0.8 0.9 Volume fraction FIGURE 2 Mineral proportions in fertile peridotite (pyrolite) as a function of depth in the mantle (Ringwood 1991). Note the increasing modal amount of garnet with depth from 200 to 450 km due to the dissolution of pyroxene into garnet. Garnet becomes more majoritic in this depth interval (see text). 422 Pressure GPa 3 Depth km GPa 3.5 D ECEMBER 2013 29 1 GARNET AND SEISMIC VELOCITIES IN THE MANTLE The densities and elastic properties of garnet and the other silicates of the mantle may, in principle, be used to investigate radial and lateral variations in mantle bulk composition. For example, using predominantly low-pressure–low-temperature elastic data, Anderson and Bass (1986) suggested that the transition zone (410–660 km) is composed of a relatively garnet-rich composition, which they called “piclogite.” Piclogite is similar to pyrolite in that it is a mixture of depleted peridotite and basalt, but piclogite has much more basalt. The result is that, under upper mantle conditions, a piclogitic mantle would have abundant pyroxene and garnet and only ~30% olivine (compared to ~65% olivine for pyrolite). A transition zone of piclogite would move the mantle Mg/Si ratio closer to that of the Sun, something that is cosmochemically attractive. Piclogite’s lower olivine content generates smaller 410 and 660 km p-wave discontinuities (olivine–wadsleyite and ringwoodite–perovskite transitions; FIG. 2) than pyrolite, but a steeper velocity gradient between the 410 and 660 km discontinuities because of the greater extent of transformation of pyroxene to garnet and higher modal proportion of the latter (Anderson and Bass 1986). P-wave velocities of piclogite are reasonably consistent with those of the Preliminary Reference Earth Model (PREM; Dziewonski and Anderson 1981), which has discontinuities of 2.6% and 4.8% in p-wave velocity at 400 and 670 km, respectively. PREM is based largely on the free oscillations of the Earth, however, and a model based purely on p- and s-wave velocities, such as ak135 (Kennett et al. 1995), gives larger discontinuities of 4.0% at 410 km and 5.8% at 660 km. The difficulty of arriving at a defi nitive compositional model based solely on wave speeds is compounded by lateral variations in seismic velocities and uncertainties in the elastic properties of the minerals of interest at high pressures and temperatures. Most recent work suggests a mantle that is broadly peridotitic (pyrolitic) but with an increasing proportion of basaltic component in the transition zone and lower mantle (Cobden et al. 2009). ECLOGITES AND RELATED ROCKS IN THE MANTLE Although most upper mantle xenoliths are peridotites (FIG. 3B), eclogites are a second major lithology in kimberlite and dominate the xenolith lithologies in five welldocumented diamondiferous kimberlite pipes: Roberts Victor and Bobbejaan (South Africa), Zagadochnaya (Russia), Orapa (Botswana) and Koidu (Sierra Leone). Eclogites are rocks composed mostly (≥75%) of subequal amounts of garnet and omphacitic clinopyroxene (FIG. 3C). Unlike garnets from peridotite lithologies, garnets from A FIGURE 3 mantle eclogite xenoliths are Cr-poor, Ca-rich pyrope– grossular–almandine mixtures with wide variations in the grossular content (Jacob 2004). Arguably, the most important feature of eclogites is that their bulk compositions often closely resemble the compositions of basalts. The amount of eclogite in the mantle, its origins, and the part it plays in petrogenesis are subjects of considerable debate. Noting that basalts and eclogites are compositionally similar but that garnet-rich eclogites are the denser, highpressure form, it was suggested on a number of occasions in the early to mid 20th century that the Moho corresponds to the phase transformation of basalt to eclogite and that the upper mantle is eclogitic rather than peridotitic. This model, which would explain basalt genesis as near total melting of the upper mantle, was discredited by the experimental work of D. H. Green and A. E. Ringwood on the basalt–eclogite transformation (Green and Ringwood 1967). These authors found that the conditions of the transformation are such that eclogite is stable throughout most of the continental crust and hence could not be responsible for the Moho discontinuity under the continents. Furthermore, the transformation is spread out over a considerable pressure interval, unlike the seismically sharp Moho. With increasing pressure, the pyroxene–plagioclase assemblage of basalt is joined by garnet at about 1 GPa, which grows in modal amount at the expense of plagioclase and pyroxenes. Eventually plagioclase is eliminated when the pressure is sufficiently high for clinopyroxene to accommodate all the sodium in the rock as NaAlSi2O6 and NaFe3+ Si2O6 components (~2.5 GPa). Although not responsible for the Moho, the basalt–eclogite transformation must occur at the tops of subducting slabs where mid-ocean ridge basalt (MORB) is converted to the eclogite assemblage of garnet plus clinopyroxene. At still higher pressures, eclogite undergoes further phase transformations analogous to those in peridotite, principally, the conversion of pyroxene components into garnet. FIGURE 4 (Irifune et al. 1986; Ricolleau et al. 2010) illustrates the phase relations of eclogite as a function of depth in the Earth. The dissolution of pyroxene components into garnet leads to a rock that is ~95% garnet at depths of about 500 km, the remainder being small amounts of excess SiO2 which crystallises as stishovite at any pressure above 9 GPa. As pressure increases, as for peridotitic compositions, CaSiO3 perovskite becomes stable at about 18 GPa. As the modal proportion of CaSiO3 perovskite increases, the CaO content and modal proportion of garnet gradually decline until (Mg,Fe)SiO3 perovskite becomes stable at about 23 GPa (~660 km). Thereafter, the fraction of garnet rapidly declines such that beyond 750 km depth, (Mg,Fe)SiO3 perovskite, CaSiO3 perovskite, stishovite and a new aluminous (NAL) and/or a calcium ferrite phase make B C (A) Garnet inclusion in diamond. (B) Garnet peridotite xenolith. (C) Eclogite xenolith. IMAGES COURTESY OF Z. SPETSIUS (A) AND K. MACDONALD (B) E LEMENTS 423 D ECEMBER 2013 of other high-pressure minerals trapped as inclusions in diamonds (Harte 2010) reinforce the hypothesis that some of the diamonds have originated well below the cratons. 300 Pressure GPa Depth km 200 Pyroxene 400 GARNET AND GEOCHEMISTRY OF VOLCANIC ROCKS 13 Garnet 500 600 Considerable effort in both experimental and observational geochemistry has been expended in trying to understand the controls on the major and trace element compositions of melts that are erupted at the surface of the Earth. Garnet does not crystallise at low pressures in most igneous systems but, in a fertile peridotite, it is stable below ~85 km depth (2.8 GPa) at the mantle solidus (see FIG. 1). Thus, if we can identify a “garnet signature” in eruptive products we have good reason to believe that there was, at least, some melting taking place at depths greater than 85 km. Seismological evidence of melt at the lithosphere–asthenosphere boundary to at least 75 km depth (Qi et al. 1997) suggests that garnet is likely to be present in the source region of at least some basalts. In this section we review how the presence, or absence, of garnet in the source region affects the trace element geochemistry of melts that we observe at the surface. 21 700 (Mg,Fe)SiO3 perovskite 800 900 0 0.2 SiO2 NAL stish phase 0.4 0.6 Volume fraction CaSiO3 perovskite 0.8 29 1.0 Composite diagram (Irifune et al. 1986; Ricolleau et al. 2010) showing phase transformations and proportions for an eclogite composition as a function of depth in the Earth. Note that eclogite becomes essentially a garnetite in the deeper parts of the transition zone (410–660 km). Stish = stishovite. NAL = new aluminous phase, Na-rich FIGURE 4 up the modal mineralogy. Ringwood (1991) argued that subducted oceanic crust would be dominantly composed of garnet at 650 km depth and that this “garnetite” should be less dense than the surrounding peridotite and have the potential to form a perched basaltic layer in the mantle. More recent density data indicate, however, that eclogite is denser than ambient mantle throughout the lower mantle (Ricolleau et al. 2010) and that recycled basalt can subduct to the core–mantle boundary. INCLUSIONS IN DIAMONDS Although experiments provide constraints on the majorite substitution mechanisms (reactions 3–5) in mantle garnet, natural occurrences of majoritic garnet are rare because most xenoliths have apparent equilibration pressures of <7 GPa (<220 km). Under these conditions, as can be seen from FIGURE 2, the dissolution of pyroxene into garnet has hardly begun. Therefore, in order to fi nd natural examples we need to turn to mineral inclusions in diamond (FIG. 3A). As the most common silicate inclusion in diamonds, garnet has the potential to provide unique information about old continental cratons, whose keels extend to depths of more than 200 km. FIGURE 6 shows the chondrite-normalised rare earth element (REE) contents of clinopyroxene and garnet from peridotite xenoliths found in the Vitim volcanic field (Ionov et al. 1993). Although the xenoliths all contain olivine, orthopyroxene, clinopyroxene and either garnet or spinel, clinopyroxene and garnet are the only minerals that contain significant concentrations of REEs. In spinel peridotites clinopyroxene is the dominant reservoir of REEs and in garnet peridotite both clinopyroxene and garnet contain significant amounts of REEs. Garnet has a strong preference for the heavy REEs (HREEs), which means that clinopyroxene coexisting with garnet is depleted in these elements (FIG. 6). This leads to the question: if the residue of partial melting contains garnet, can its presence in the source region be inferred by examining the REE pattern of the eruptive products? Garnets trapped as inclusions in diamonds are generally attributed to either eclogitic or peridotitic paragenesis (E-type or P-type, respectively), although there are rare websteritic varieties. E-type garnet is Cr poor (<0.4 wt% Cr2O3) and Ca rich, and has lower and more variable Mg# (45–70). P-type garnet is Cr rich and Ca poor, and has higher and more consistent Mg# (~80–92) (Stachel and Harris 2009). FIGURE 5 shows the compositions of high-pressure garnets from diamond inclusions (Harte 2010) plotted so as to highlight the majorite substitutions. As can be seen, increasing majorite content leads to increasing Si4+, Ti4+ and Na + at the expense of Al3+, Cr3+ and divalent cations. Although most of the world’s diamonds are interpreted to have formed within the cratonic lithosphere at depths to ~250 km (Stachel and Harris 2009), comparison of majoritic garnet inclusions with those of garnet generated in highpressure experiments (Kiseeva et al. 2013) indicates that these particular inclusions come from deeper in the mantle at transition zone depths (410–660 km). Moreover, findings E LEMENTS Compilation of majoritic garnet compositions from inclusions in diamonds. Note that end-member majorite would plot at x = 4, y = 4. M = divalent cations; pfu = per formula unit. 424 FIGURE 5 D ECEMBER 2013 U4+ is closer to ro than Th4+. Therefore, in the presence of garnet, the higher U bulk partition coefficient leads to a melt enriched in Th compared to U. This means that the excesses of 230 Th relative to its parent, 238U, observed in many MORB samples are most readily explained by melting in the garnet peridotite stability field (Beattie 1993). Chondrite-normalized rare earth element concentrations in clinopyroxene (Cpx) and garnet (Grt) from spinel (Sp) peridotite and garnet peridotite of Vitim Plateau (Ionov et al. 1993) FIGURE 6 FIGURE 7 shows garnet–liquid and clinopyroxene–liquid partition coefficients determined in one experiment at 3.0 GPa (Pertermann et al. 2004). The partition coefficients are given by: [i] xtal (Di = ), where [i] xtal and [i] melt [i] melt refer to weight concentrations in the crystal and melt phases. As anticipated, the garnet–liquid partition coefficients for the heavy REEs, (those with the smallest ionic radii) are significantly higher than the clinopyroxene– liquid values. While the clinopyroxene–liquid partition coefficients for all REEs are below 1, the garnet–liquid partition coefficients for the smaller HREEs are close to 10, with the light REEs (LREEs) being progressively excluded as the ionic radius increases. The result means that melting in the presence of significant amounts of residual garnet will generate liquids exhibiting depletions in the HREEs relative to the LREEs. The roughly parabolic relationship between the partition coefficient and ionic radius was fi rst observed in the 1960s (Onuma et al. 1968) and quantified in terms of strain energy by Blundy and Wood (1994). The latter authors showed that the elastic strain energy introduced into the crystal lattice when an ion of radius r i is partitioned into a site of radius ro can be systematized and quantified. If the ion has radius ro then there is no strain energy of substitution and the partition coefficient, Do, is a maximum for ions of that charge (Blundy and Wood 1994). The strain energy of substitution increases in an approximately parabolic way with the difference between r i and ro , yielding a parabolic decrease of D i with radius difference, as observed (FIG. 7). The ro value of the garnet X-site is smaller than the ro value of the M2 site in clinopyroxene, and the garnet site is much “stiffer,” with a higher Young’s modulus than the clinopyroxene site (Van Westrenen et al. 2001). This means that smaller ions prefer the garnet site and that garnet discriminates much more strongly than does clinopyroxene against ions that are different from ro in radius. Thus, garnets preferentially retain the smaller HREEs (3 + ions) and are effective at separating, or fractionating, the middle rare earths (MREEs) from one another. Similar observations can be made for ions of other charge, such as Th4+ and U4+, substituting into garnet and clinopyroxene. Th4+ is larger than U4+ and, in peridotitic garnet, E LEMENTS Application of the principles discussed above to identifying and quantifying the presence of garnet in the source regions of igneous rocks is not always unequivocaI. If garnet was present as a residual phase, we would expect fractionation of MREEs from HREEs, and this has frequently been reported. For example, Shen and Forsyth (1995) showed that Sm/Yb ratios in MORB correlate with other geochemical parameters that are believed to indicate melting in the deeper parts of the asthenosphere. After correcting for variable enrichment of incompatible elements, they concluded that the areas with the strongest “garnet signature” are the areas with the thinnest crust. In these regions, the relatively cold lithosphere inhibits shallow melting; thus a greater fraction of the observed MORB was, they argued, produced deep, in the garnet stability field (Shen and Forsyth 1995), and melting ceased relatively deep. In contrast to this result, Bourdon et al. (1996) showed that some MORBs have 230 Th/238U activity ratios above unity, implying melting in the presence of residual garnet. More importantly, these high 230 Th/ 238U activity ratios were observed at ridges with thick crust (Bourdon et al. 1996). This identification of the isotopic “garnet signature” in areas with thick crust is directly opposite to the inferences of Shen and Forsyth (1995), who identified the “garnet signature” in areas of thin crust. One possible way to reconcile these confl icting results is through the melting of a heterogeneous source that is a mixture of spinel peridotite and garnet pyroxenite (Hirschmann and Stolper 1996), but it is apparent that there is no clear unequivocal interpretation of the “garnet signature” in mid-ocean ridge basalts. Experimentally measured crystal–melt partition coeffi cients for garnet (Grt) and clinopyroxene (Cpx) from a single experiment at 3 GPa plotted as a function of ionic radius (Pertermann et al. 2004). Solid lines are lattice strain fits to the data using the model of Blundy and Wood (1994). The filled circles indicate the “best-fit” radius ro. 425 FIGURE 7 D ECEMBER 2013 CONCLUDING REMARKS We have shown that garnet is an extremely important phase at high pressures and temperatures in the mantle. Garnet is denser than other low-pressure ferromagnesian silicates because of the relatively high coordination numbers of Mg2+ and Fe2+ (both 8-coordination) and Al3+ (6-coordination). This means that once garnet becomes stable in peridotite (>85 km depth), it increases in modal amount, from ~10% to ~35%, by dissolving the low-pressure pyroxenes (which have the same number of cations per 12 oxygens) into its structure. A similar but more pronounced change in modal proportions is found in eclogite compositions in which the garnet–clinopyroxene assemblage becomes stable at about 70 km depth (Green and Ringwood 1967) and progressively changes from ~30% to ~95% garnet as clinopyroxene dissolves into the garnet and the latter becomes increasingly majoritic in composition. 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Contributions to Mineralogy and Petrology 124: 185-208 E LEMENTS The relatively high density and cation coordination numbers of garnet mean that the pyroxene–garnet transformations discussed above have potentially strong seismic “signatures” and could enable identification of lateral compositional heterogeneities in the mantle. They also mean that garnet can have a profound effect on the trace element concentrations of melts with which it is in equilibrium, providing the potential for distinguishing magmas by depth of melting. As yet, however, other unquantified effects mask these two potential applications. ACKNOWLEDGMENTS We thank E. Baxter, W. Griffi n and J. Van Orman for reviews and comments that improved the content and clarity of this article, and K. Macdonald and Z. Spetsius for the use of their images. We also acknowledge support from a European Research Council grant 267764. 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