The Redox State of Earth s Mantle

Transcription

1 ANNUAL REVIEWS Further Click here for quick links to Annual Reviews content online, including: Other articles in this volume Top cited articles Top downloaded articles Our comprehensive search Annu. Rev. Earth Planet. Sci : First published online as a Review in Advance on February 12, 2008 The Annual Review of Earth and Planetary Sciences is online at earth.annualreviews.org This article s doi: /annurev.earth Copyright c 2008 by Annual Reviews. All rights reserved /08/ $20.00 The Redox State of Earth s Mantle Daniel J. Frost and Catherine A. McCammon Bayerisches Geoinstitut, University of Bayreuth, Germany; Dan.Frost@uni-bayreuth.de Key Words oxygen fugacity, oxythermobarometry, silicate perovskite, fluids, diamond, disproportionation, lower mantle, core formation Abstract Oxygen thermobarometry measurements on spinel peridotite rocks indicate that the oxygen fugacity at the top of the upper mantle falls within ±2 log units of the fayalite-magnetite-quartz (FMQ) oxygen buffer. Measurements on garnet peridotites from cratonic lithosphere reveal a general decrease in fo 2 with depth, which appears to result principally from the effect of pressure on the controlling Fe 3+ /Fe 2+ equilibria. Modeling of experimental data indicates that at approximately 8 GPa, mantle fo 2 will be 5 log units below FMQ and at a level where Ni-Fe metal becomes stable. Fe-Ni alloy and an Fe 2 O 3 -garnet component will be formed as a result of the disproportionation of FeO, which is experimentally demonstrated through observations of high Fe 3+ / Fe ratios in minerals in equilibrium with metallic Fe. In the lower mantle, the favorable coupled substitution of Al and Fe 3+ into (Fe,Mg)SiO 3 perovskite results in very high perovskite Fe 3+ / Fe ratios in equilibrium with metallic Fe. As a result, the lower mantle should contain approximately 1 weight% metallic Fe formed through FeO disproportionation, if the bulk oxygen content is the same as the upper mantle. Loss of disproportionated metallic Fe from the lower mantle during core formation could explain the higher Fe 3+ / Fe ratio of the present-day upper mantle when compared to that expected during core formation. The influence of pressure on mantle fo 2 has important implications for the speciation of C-O-H-S volatile phases in Earth today and during its early evolution. 389

2 1. INTRODUCTION Oxygen fugacity is a key parameter in controlling the physical and chemical style of interactions between reservoirs within the mantle and between the mantle and surface (Frost 1991). The oxidation state of the mantle, for example, controls the speciation of C-O-H-bearing fluids and melts, which in turn can influence the mantle solidus and properties of the resulting liquids (Taylor & Green 1988, Ballhaus & Frost 1994, Holloway & Blank 1994). It therefore has an important role in magma genesis, magma degassing, and metasomatic processes (Kadik 1997, Holloway 1998). Through its influence on H 2 O activity, the mantle redox state will also affect the partitioning of H 2 O between fluids/melts and minerals, and it is a major parameter defining the stability of C-bearing phases in Earth, such as diamonds and carbonates. In addition, by controlling mineral Fe 3+ and OH contents, it will have a strong influence on most transport properties, such as electrical conductivity and creep (Ryerson et al. 1989). Although not so apparent, the redox state of the deeper mantle also has important implications for the geochemistry and geophysics of Earth as a whole. First, some melts, such as kimberlites, likely come from sources deep in the mantle (>200 km) and their volatile-rich nature implies deep redox processes. Seismic observations also suggest the presence of volatile-induced melting in the deeper mantle (Revenaugh & Sipkin 1994). Similarly, some rare diamonds apparently contain lower mantle mineral inclusions, implying that diamond growth is not confined to the lithospheric mantle (Stachel et al. 2000). In the deep mantle, it is possible that for crystal chemical reasons, rocks become more reduced with depth and may contain metallic Fe (O Neill et al. 1993b, Ballhaus 1995, Frost et al. 2004). This may influence the cycling of volatiles through the mantle, as reduced CH 4 -rich fluids may tend to rise out of the deep mantle. The presence of metallic Fe in the mantle may have had important implications for core formation and the early evolution of the mantle s redox state (Frost et al. 2004, Wood et al. 2006). The redox state of the early mantle would have influenced the style of mantle degassing and the development of the hydrosphere and atmosphere (Kasting et al. 1993). Here, the major lines of evidence for the redox state of the upper mantle are reviewed along with arguments and experimental data as to the redox state of the deeper interior. These studies indicate that the convecting mantle becomes more reduced with depth as a result of the effect of pressure on Fe 3+ /Fe 2+ equilibria that control redox state. The evidence and implications of this process on volatiles and the historical evolution of the mantle s redox state are then discussed. 2. UPPER MANTLE OXYGEN FUGACITY Although the upper mantle is the only part of the Earth s interior where the oxygen fugacity can be directly measured, this region is made complex by partial melting and by the fact that most of the available samples come from the subcontinental lithosphere and have also been influenced by metasomatism. A good understanding 390 Frost McCammon

3 of the Fe 3+ content of the mantle and how it relates to oxygen fugacity must, however, be obtained for the upper mantle before we can consider the likely redox state of the deeper mantle. This means considering the redox state of magmatic rocks formed by partial melting of the mantle, such as basalts, and mantle rocks that are either brought to the surface as xenoliths in volcanic magmas or are exposed in peridotite massifs, which are sections of the mantle that have become tectonically emplaced into the crust Basalts The redox state of the mantle is reflected in the oxidation state of mid-ocean ridge basalt (MORB). Analyses of the rapidly quenched glassy rims of pillow basalts from most spreading centers reveal a relatively narrow range of Fe 3+ / Fe ratios, between 8% and 15%, with a global average of 12 ± 2% (Bézos & Humler 2005). This is significantly higher than the average of 7 ± 3% obtained in a previous global study of MORB glasses (Christie et al. 1986); potential reasons for this difference are discussed by Bézos & Humler (2005). Both studies are in agreement, however, that Fe 3+ / Fe ratios do not vary significantly regionally and there is no relationship between Fe 3+ / Fe and ratios that discriminate source enrichment such as chondrite normalized La/Sm. In addition, rapidly chilled basalts sampled close to the vent of Mauna Loa (Rhodes & Vollinger 2005) also have average Fe 3+ / Fe ratios similar to that of global MORB measurements (11.6 ± 6%), implying that the oxidation state of plume-type Hawaiian magmas is similar to MORB. Using an empirical calibration, the fo 2 of MORB magmas can be calculated from their Fe 3+ / Fe ratio and oxide mole fractions (Sack et al. 1980, Jayasuriya et al. 2004). The global average from MORB glasses is 0.41 ± 0.43 log units relative to the fayalite magnetite quartz (FMQ) oxygen buffer (for buffering reactions and notations see Frost 1991). The fo 2 of MORB magmas at their mantle source can be calculated from the 1 bar fo 2 using data on the partial molar volumes of the Fe 3+ /Fe 2+ melt components and their compressibilities (Kress & Carmichael 1991). Assuming the system is closed to oxygen, this gives an fo 2 of FMQ for MORB at conditions of partial melting. As discussed in Section 2.2, this is at the more oxidizing end of the range displayed by abyssal peridotites (between FMQ and FMQ 2.5) that are considered to be MORB mantle residues. The absence of large Fe 3+ / Fe variations in MORB glasses is an important observation with respect to the redox reactions occurring during partial melting. Parameters sensitive to the degree of partial melting, such as Na 8.0 (Klein & Langmuir 1987), have little correlation with Fe 3+ / Fe ratios, even though the samples cover an apparent range of partial melting of between 6% and 20% (Bézos & Humler 2005). If Fe 2 O 3 behaved simply as an incompatible element during partial melting [it has been proposed to have a mineral melt partition coefficient of 0.1 (Canil et al. 1994], then Fe 3+ / Fe ratios could never stay constant over these partial melting intervals. If partial melting to create MORB took place in the presence of graphite, however, MORB: mid-ocean ridge basalt FMQ: Fayalite = Magnetite + Quartz oxygen buffer, 3Fe 2 SiO 4 + O 2 = 2Fe 3 O 4 + 3SiO 2 Na 8.0 : Na 2 O content of a magma corrected along a crystal fractionation trend to 8 wt.% MgO The Redox State of Earth s Mantle 391

4 CCO: oxygen buffer, C + O 2 = CO 2 then the melt Fe 3+ / Fe ratio could be buffered through the reaction (Kadik 1997, Holloway 1998): C + 2Fe 2 O 3 = 4FeO + CO 2. (1) graphite melt melt The oxygen fugacity is then a function of the melt CO 2 solubility, which in turn depends on pressure, temperature, and melt composition. The relationship between the degree of melting and CO 2 solubility could potentially keep Fe 3+ / Fe ratios relatively constant during melting. However, if MORB genesis occurs at approximately FMQ, as recent measurements would suggest (Bézos & Humler 2005), then this is 1 log unit above the CCO buffer and graphite would not be stable. Other mechanisms may, therefore, need to be sought for keeping Fe 3+ / Fe ratios in a narrow range during partial melting. Alteration processes render ancient basalts and peridotites unsuitable for fo 2 determinations based on Fe 3+ / Fe measurements, so a number of studies have investigated the behavior of other trace elements, such as V, Cr, and Eu, that have variable valence states and, therefore, have mineral-melt partition coefficients that are a function of fo 2 (Canil 1997, Delano 2001, Wadhwa 2001). The concentrations of these elements in magmatic rocks, which cannot be perturbed by degassing and are not easily affected by alteration processes, can be used to estimate the fo 2 of the melting process from knowledge of the source rock concentration and partition coefficients. Over ranges of fo 2 exhibited by terrestrial magmas, V can exist in 3+, 4+, and 5+ oxidation states (Canil 1999) and the V mineral melt partition coefficient (D V ) has been well calibrated as a function of fo 2. The effects of magma differentiation on the oxidation state can be accounted for if the V/Sc ratio is used (Canil 1999, Li & Lee 2004, Canil et al. 2006) because variations in the degree of partial melting and olivine fractionation affect both elements similarly. An important outcome is that the V/Sc ratios of peridotites and basalts from the Archean cover a similar range to presentday samples, with the inference that upper mantle fo 2 could have risen by no more than 0.3 log units over the last 3.5 Gyrs of Earth s history (Li & Lee 2004). Delano (2001) reached a similar conclusion by studying Cr. The examination of V/Sc ratios in ophiolite massifs also implies that mantle melting occurs over a narrow range of fo 2 (Canil et al. 2006) Spinel Peridotites Between depths of 30 and 60 km the mantle is predominantly composed of spinel peridotite rocks, although the mineral spinel can be stable to deeper levels in subcratonic lithosphere owing to higher Cr contents. Samples are available from a wide range of tectonic settings. Typical spinel peridotites contain between wt% Fe 2 O 3 (Fe 3+ / Fe 1% 3%) in the whole rock, which is quite evenly distributed between all minerals except olivine, which contains negligible Fe 2 O 3 (Canil & O Neill 1996, Woodland et al. 2006). Spinels have Fe 3+ / Fe ratios in the range of 15% 34%, which can fall to 5% 15% in peridotites that have undergone extensive melt extraction. Although spinels have high Fe 2 O 3 contents, they generally make up less than 3% of the rock, and their contribution to the rock s total Fe 2 O 3 content is therefore subequal to ortho- and clinopyroxenes (Canil & O Neill 1996, Woodland et al. 2006). 392 Frost McCammon

5 Two equilibria have been widely employed for oxygen thermobarometry measurements of spinel peridotites. Nell & Wood (1991) have calibrated the equilibrium, 6Fe 2 SiO 4 + O 2 = 3Fe 2 Si 2 O 6 + 2Fe 3 O 4, (2) olivine opx spinel where fo 2 is calculated through the relation log f O2 = Go (2) ln(10)rt + 3 log a opx Fe 2 Si 2 O log a spinel Fe 3 O 4 6 log a olivine Fe 2 SiO 4, (3) Annu. Rev. Earth Planet. Sci : Downloaded from where R is the gas constant; a spinel Fe 3 O 4 is, for example, the activity of the Fe 3 O 4 spinel component; and G o (2) is the standard state Gibbs free energy change for Equation 2. Ballhaus et al. (1991) performed an empirical calibration of this equilibrium. Alternatively the FMQ equilibrium, 3Fe 2 SiO 4 olivine + O 2 = 2Fe 3 O 4 +3SiO 2, (4) spinel was calibrated by O Neill & Wall (1987), where the activity of SiO 2 is calculated using the equilibrium Mg 2 SiO 4 olivine + SiO 2 = Mg 2 Si 2 O 6. (5) opx The three calibrations for Equations 2 and 4 (O Neill & Wall 1987, Ballhaus et al. 1991, Nell & Wood 1991) predict fo 2 s that are generally within 0.5 log units of each other (Wood 1991), although under particular conditions (e.g., low temperatures, high pressures, high Cr contents in spinel) the calculated values can differ by up to two log units (McCammon & Kopylova 2004). Equations 2 and 4 have been applied to rocks from a number of tectonic settings with the conclusion that fo 2 in the spinel peridotite facies generally falls between FMQ-2 and +2 log units, but exhibits significant variations between settings (Figure 1). In addition, most settings record a range of fo 2 of between 1 and 2 log units, inferring that the fo 2 of the mantle is heterogeneous on quite a small scale. Seafloor abyssal peridotites, which are MORB mantle residues, record some of the lowest mantle oxygen fugacities observed for spinel peridotites. Peridotite massifs, such as Ronda, record similar levels of fo 2, supporting a common asthenospheric character. Both abyssal and massif peridotites have fo 2 s that overlap at the more oxidized limit with those measured for pristine MORB glasses; however, their median values are 1 log unit below MORB at its source when recent glass measurements are employed (Bézos & Humler 2005). Massifs in the Pyrenees and xenoliths from the subcontinental lithospheric mantle record fo 2 s that are slightly higher than abyssal peridotites. A possible explanation for this is that lithospheric metasomatism has raised the fo 2 by, for example, adding Fe 2 O 3 (McGuire et al. 1991), CO 2 or carbonates. Analysis of light rare earth elements in samples from particular settings, however, does not reveal a correlation between fo 2 and modally or cryptically metasomatized mantle (Ionov & Wood 1992, Woodland et al. 1996). However, more oxidized samples from several continental localities contain metasomatic hydrous minerals, such as amphiboles (McGuire et al. 1991, Woodland et al. 1996). The high fo 2 s recorded in subduction-related settings The Redox State of Earth s Mantle 393

6 Grenada Simcoe Japan Lihir British Columbia Subduction xenoliths Annu. Rev. Earth Planet. Sci : Downloaded from San Carlos Beni Bousera KLB Central Asia Pyrenees Lherz Ronda Abyssal peridotite MORB at 1.5 GPa MORB glasses Log fo 2 ( FMQ) Continental xenoliths Peridotite massifs Oceanic Figure 1 Oxygen fugacities (range and median value) calculated for spinel peridotite assemblages from a number of tectonic settings. Measurements from the same setting are grouped within the shaded regions, which are, from the bottom up, oceanic (Bryndzia & Wood 1990), peridotite massifs (Woodland et al. 1992; Woodland et al. 1996, 2006), xenoliths from the continental lithosphere (Ionov & Wood 1992), and xenoliths from subduction settings (Wood & Virgo 1989, Canil et al. 1990, Brandon & Draper 1996, McInnes et al. 2001, Parkinson et al. 2003). Included within the oceanic group are fo 2 measurements of MORB glasses (Bézos & Humler 2005) and the same samples corrected to source region conditions of 1.5 GPa (Kress & Carmichael 1991). also seem to be related to metasomatic processes and are coupled in some instances with the formation of hydrous minerals (Wood & Virgo 1989, Brandon & Draper 1996). These observations fit the very general model that the lithosphere is initially as reduced as the asthenosphere from which it formed, but over time becomes oxidized as a result of further metasomatic interaction with the asthenosphere. Although many peridotites show major element trends that reflect differing degrees of melt extraction, the degree of fertility of a rock does not correlate 394 Frost McCammon

7 simply with fo 2 (Ionov & Wood 1992). Fe 2 O 3 is moderately incompatible and major element melt extraction trends are coupled to bulk rock Fe 2 O 3 depletions, but a crystal chemical effect decouples fo 2 from these trends. Melt extraction causes an increase in the spinel Cr/(Cr+Al+Fe 3+ ) ratio, but because Fe 3+ -Cr substitutions are energetically more favorable than Fe 3+ -Al substitutions, this leads to an increase in spinel/pyroxene Fe 2 O 3 partition coefficients. The result is a positive correlation between spinel Cr/(Cr+Al+Fe 3+ ) ratio and Fe 2 O 3 concentration (Canil & O Neill 1996, Woodland et al. 1992), which acts in opposition to the melt extraction trend and obscures the relationship with fo 2. This effect will presumably also influence the Fe 3+ / Fe ratios of the ensuing partial melts but it is hard to quantify without mineral/melt Fe 2 O 3 partition coefficients. A further aspect of melt extraction is that it decreases the modal proportion of spinel and clinopyroxene in peridotites, making fo 2 more sensitive to increases in Fe 2 O 3. It is, therefore, particularly easy for metasomatism to raise the fo 2 of depleted rocks, such as harzburgites, as only small increases in Fe 2 O 3 are necessary (Woodland et al. 2006). The lack of correlation between melt extraction trends and fo 2 in many settings probably also reflects the ease by which the oxidation state of depleted peridotites can be reset. More significant levels of metasomatism, such as those observed in recent or past subduction environments, can result in the formation of clinopyroxene, spinel, and hydrous minerals that all have significant Fe 2 O 3 contents, making the oxidation state of the peridotite far less susceptible to perturbations (Wood 1991, Woodland et al. 2006). Above FMQ+1, however, larger amounts of Fe 2 O 3 need to be added to a peridotite to raise its fo 2 by even a small amount, which acts as a quasi- fo 2 buffer and explains why very few rocks exhibit fo 2 values above FMQ+1 (Woodland et al. 2006) Garnet Peridotites Garnet-peridotite rocks, which are mainly found as xenoliths in kimberlite magmas, come on average from greater depths than spinel-peridotites but display similar ranges of whole-rock Fe 2 O 3 contents. The garnet Fe 3+ / Fe ratios in peridotite xenoliths are typically between 2% and 14% (Canil & O Neill 1996). Geothermometry and barometry reveals a trend of increasing garnet Fe 3+ / Fe ratio with temperature and pressure of equilibration, which is not coupled to a whole-rock increase. An increase in the garnet/pyroxene Fe 2 O 3 partition coefficient with temperature most likely explains this trend (Canil & O Neill 1996, Woodland & Koch 2003). Oxygen fugacities for garnet peridotites can be determined using the equilibrium 2Fe 3 Fe 3+ 2 Si 3O 12 garnet = 4Fe 2 SiO 4 olivine + 2FeSiO 3 + O 2, (6) opx which involves the Fe 3 Fe 3+ 2 Si 3O 12 (skiagite) garnet component (Gudmundsson & Wood 1995; see also Woodland & Peltonen 1999 regarding the correction of a typographic error). The oxygen fugacity is calculated with the equation log fo 2 = Go (6) gt + 2 log afe ln(10)rt 3 Fe 2 Si 3 O 12 2 log a opx FeSiO 3 4 log a olivine Fe 2 SiO 4. (7) The Redox State of Earth s Mantle 395

8 1 0 Equation 6 Annu. Rev. Earth Planet. Sci : Downloaded from Log fo 2 ( FMQ) Abyssal peridotites Slave Kaapvaal Fennoscandian USSR Pressure (GPa) Figure 2 Oxygen fugacities relative to FMQ calculated for garnet peridotite rocks from cratonic lithosphere as a function of pressure. The xenoliths are from the Kaapvaal craton, South Africa (Luth et al. 1990, McCammon et al. 2001, Woodland & Koch 2003); the Slave Craton, Canada (McCammon & Kopylova 2004); the Fennoscandian Shield, Finland (Woodland & Peltonen 1999); and the Vitim Plateau, USSR. Equation 6 is the fo 2 calculated for the end-member Equation 6 in the text, which appears to control the pressure dependence of the measurements. The volume change for Equation 6 is positive (8.6 cm 3 mol 1 ), which means pressure favors the stability of Fe 3 Fe 3+ 2 Si 3O 12. Therefore, for a fixed garnet peridotite composition, the calculated oxygen fugacity will have a tendency to be driven to lower levels with increasing pressure (Gudmundsson & Wood 1995). Oxygen fugacities, calculated using Equation 7, for garnet-bearing rocks show a general trend of decreasing fo 2 with depth (Figure 2). A comparison with the slope calculated for the end-member Equation 6 along a similar geotherm shows that this decrease results principally from the volume change of the equilibrium. Owing to interphase Fe 2 O 3 partitioning, garnet Fe 3+ / Fe ratios, and consequently Fe 3 Fe 3+ 2 Si 3O 12 activities, are increasing with pressure and temperature, but any effect of this on fo 2 is overridden with depth by the volume change of Equation 6. The sheared, more fertile peridotite rocks that are representative of some of the deepest samples [e.g., garnet peridotite PHN 1611 of Nixon & Boyd (1973)], therefore, display some of the highest garnet Fe 3+ / Fe ratios but record the lowest fo 2 s. At a constant pressure, however, higher fo 2 s do couple to some extent with temperature. 396 Frost McCammon

9 3. DEEP UPPER MANTLE AND TRANSITION ZONE The deepest garnet peridotite xenoliths come from approximately 6 GPa or 200 km, but the fo 2 at greater depth can be estimated if we can assume that the mantle retains, on average, a constant Fe 3+ / Fe ratio or more precisely a constant O/Fe ratio If this is the case, then the fo 2 calculated for a mantle assemblage using heterogeneous Fe 3+ /Fe 2+ equilibria will decrease with depth (O Neill et al. 1993b, Ballhaus 1995, Rohrbach et al. 2007). This occurs for two reasons: (a) the volume changes of Fe 3+ /Fe 2+ equilibria used to calculate fo 2 favor the stability of Fe 2 O 3 components with depth, and (b) the activity of Fe 2 O 3 components decreases as they become diluted in minerals that increasingly come to dominate the mantle assemblage with depth. This second point is easily illustrated with respect to Equation 2. In the (Mg,Fe)Al 2 O 4 spinel stability field olivine, the most abundant mineral has a negligible Fe 2 O 3 content and Fe 2 O 3 is instead concentrated in minerals such as spinel and clinopyroxene that are of relatively low abundance. The calculated fo 2 is relatively high because, as seen in Equation 3, it is proportional to the activity of the Fe 2 O 3 -bearing component. In the deeper mantle, however, garnet, and then the high-pressure olivine polymorphs wadsleyite and ringwoodite that dominate the assemblage, can all dissolve significant amounts of Fe 2 O 3.Fe 2 O 3 component activities will therefore be low, resulting in a low calculated fo 2. Pristine mantle whole-rock Fe 3+ / Fe ratios can be estimated from xenoliths that show minimal evidence of metasomatism or partial melting. In both the spinel and garnet peridotite facies, whole rock Fe 3+ / Fe ratios fall in a relatively narrow range in the region of 2% (Canil & O Neill 1996). Figure 3 shows a calculation for the oxygen fugacity of a bulk silicate Earth (BSE) (McDonough & Sun 1995) composition as a function of depth, normalized against the iron wüstite oxygen buffer (IW). The calculation is made using Equation 7 with mineral compositions for a fourphase assemblage calculated using experimental Fe-Mg partitioning data (Frost 2003) and accounting for the pyroxene-majorite and orthopyroxene-clinoenstatite (HCPX, C2/c) transitions. Although there are uncertainties in Fe 2 O 3 partitioning, they cannot change the general conclusion, which is that the fo 2 of a mantle assemblage with a constant Fe 3+ / Fe ratio will continue to drop with increasing pressure (Ballhaus 1995). This is principally a result of the volume change of Equation 6, but an additional effect arises from the dilution of the Fe 3 Fe 3+ 2 Si 3O 12 skiagite component owing to the pyroxene-majorite transformation. At approximately 8 GPa, the curve calculated using Equation 7 crosses the nickel precipitation curve (NiPC), which is also marked on Figure 3. The NiPC marks the fo 2 where NiO, mainly in olivine, will be reduced to form Ni-rich metal in a mantle peridotite assemblage (O Neill & Wall 1987). Metallic Fe will also partition into the alloy as it forms. The fo 2 of the NiPC is calculated with the reactions and Ni 2 SiO 4 olivine Fe 2 SiO 4 olivine = 2Ni metal + SiO 2 + O 2 (8) = 2Fe + SiO 2 + O 2, (9) metal BSE: bulk silicate Earth IW: iron wüstite oxygen buffer, 2Fe + O 2 = 2FeO The Redox State of Earth s Mantle 397

10 5 4 FMQ EMOG/D 3 Log fo 2 ( IW) 2 1 Kaapvaal craton Annu. Rev. Earth Planet. Sci : Downloaded from Figure NiPC [1] Pressure (GPa) The oxygen fugacity calculated for a four-phase garnet peridotite assemblage assuming a Fe 3+ / Fe ratio of 2% shown along a cratonic geotherm relative to the IW buffer. The fo 2 of the Ni precipitation curve (NiPC) calculated using Equations 8 and 9 for a peridotite assemblage is indicated with values for the Ni content (mol%) of the precipitating metal shown along the curve. Line 1 is the fo 2 calculated for garnet peridotite using Equation 6. Line 2 is the fo 2 calculated for a garnet peridotite assemblage once the NiPC curve is crossed and Ni-Fe metal forms. Line 3 is the metastable extrapolation of Equation 6 that assumes no metal precipitation. The EMOG/D buffer is given by Equation 15 in the text. where at equilibrium, log fo 2 = Go (8) ol + log ani ln(10)rt 2 SiO 4 log a SiO2 2 log a metal Ni. (10) A similar equation can be written for Equation 9. The a Si O2 is determined with Equation 5. The composition of the precipitating metal is calculated by simultaneously solving both Ni and Fe versions of Equation 10 at the same fo 2 and assuming typical Fe and Ni contents of olivine and orthopyroxene in the mantle, i.e., XNi olivine = (O Neill & Wall 1987). As shown in Figure 3, the NiPC is just slightly below the IW buffer. The Ni content of the metal precipitated at the curve, which is indicated on the curve, decreases with increasing pressure because the volume change for Equation 8 is smaller than Equation 9. This effect is well known from Ni partitioning experiments between silicate and metal at high pressure (Li & Agee 1996). At approximately 8 GPa or 250 km, where the fo 2 of a peridotite assemblage crosses the NiPC, the metal formed will contain 60 mole% Ni. Equations 8 and 9 are driven to the right as a result of Equation 6 being driven to the left with increasing pressure. [2] [3] Frost McCammon

11 This can be more simply described by the reactions NiO + 2FeO = Ni + Fe 2 O 3 (11) olivine metal garnet and 3FeO = Fe + Fe 2 O 3. (12) olivine metal garnet Equation 12 describes the disproportionation of FeO to produce phases where Fe is in both higher and lower oxidation states. As Ni-Fe metal precipitates, the Fe 3+ / Fe ratio of the assemblage increases, raising the activity of the Fe 3 Fe 3+ 2 Si 3O 12 garnet component and causing the fo 2 to diverge from that calculated, assuming a constant Fe 3+ / Fe ratio, and more closely mirror the NiPC (Figure 3). The effect of pressure on Equation 6 will force more metal to precipitate with depth, but only a small amount needs to form to keep the fo 2 of the assemblage close to the NiPC. By combining the effects of Equations 6, 8, and 9 in a peridotite assemblage it can be shown that by 14 GPa, i.e., the base of the upper mantle, between wt% metal would form and the fo 2 is calculated to be 0.3 log units below the NiPC. At higher pressure, the transformation of olivine to wadsleyite should drive Equation 6 to the right and raise the calculated fo 2 ; however, as discussed below, the activity of the Fe 3 Fe 3+ 2 Si 3O 12 garnet component will be lowered further in the transition zone because Fe 2 O 3 will also partition into wadsleyite. The prediction that Fe 2 O 3 -bearing garnet within peridotite rocks should coexist with Fe-Ni metal at conditions compatible with the lower region of the upper mantle is supported by high-pressure and -temperature experiments performed to determine the minimum Fe 3+ / Fe ratio of garnet (O Neill et al. 1993a, Rohrbach et al. 2007). In the experiments of O Neill et al. (1993a), samples of (Mg,Fe) 4 Si 4 O 12 majorite garnet were synthesized in equilibrium with metallic Fe and minor amounts of SiO 2 or (Mg,Fe) 2 SiO 4 polymorphs. The fo 2 is then described by the reaction Fe 4 Si 4 O 12 garnet = 4Fe + 4SiO 2 +2O 2. (13) metal stishovite This defines the lower limit of stability of Fe-bearing majorite garnet with respect to fo 2 and the Fe 2 O 3 content of garnet will, therefore, be at its lowest achievable level. For Fe/(Fe+Mg) = 0.15 in majorite, the fo 2 calculated from Equation 13 is more than 1 log unit below IW; however, majorite garnet at these conditions has an Fe 3+ / Fe ratio of 7% (Figure 4). This is a clear change in behavior when compared with low pressure minerals such as (Mg,Fe)Al 2 O 4 spinel, which have negligible Fe 2 O 3 contents at conditions of IW (Ballhaus et al. 1991). Volumetric effects favoring Fe 3+ over Fe 2+ components with pressure (as in Equation 6) drive this change. In the recent experimental study of Rohrbach et al. (2007), high garnet Fe 3+ / Fe ratios were found within peridotite assemblages equilibrated with metallic Fe up to 14 GPa. At conditions equivalent to depths >250 km, garnets in equilibrium with Fe-metal contained more Fe 2 O 3 than is believed to be available in average mantle. This again argues for the formation of metallic Fe-Ni alloy through Equations 8 and 9 to supply sufficient Fe 2 O 3 to satisfy the observed garnet content at these reducing conditions. Disproportionation of FeO: 3FeO = Fe 2 O 3 + Fe The Redox State of Earth s Mantle 399

12 a (Mg,Fe) 2 SiO 4 + Fe metal Ringwoodite; O'Neill et al. (1993), Frost et al. (2001) Wadsleyite; O'Neill et al. (1993) b (Mg,Fe)SiO 3 garnet + Fe metal O'Neill et al. (1993) McCammon & Ross (2003) Frost & McCammon (unpublished) Annu. Rev. Earth Planet. Sci : Downloaded from c Fe 3+ /(Σ Fe) Fe 3+ /(Σ Fe) Fe/(Fe + Mg) (Mg,Fe)SiO 3 perovskite + Fe metal Lauterbach et al. (2000) McCammon (1998) Terasaki et al. (2007) Fe/(Fe + Mg) d Fe 3+ /(Σ Fe) Fe 3+ (per 3 oxygen formula unit) Fe/(Fe + Mg) (Mg,Fe)(Al,Si)O 3 perovskite + Fe metal Frost et al. (2004) Lauterbach et al. (2000) Solidus; Frost et al. (2004) Al (per 3 oxygen formula unit) Figure 4 (a c) Minimum Fe 3+ / Fe ratios for transition zone and lower mantle minerals determined in equilibrium with metallic Fe [or Fe-S liquid in the case of Terasaki et al. (2007)] as a function of Fe/(Fe+Mg). (d ) The Fe 3+ versus Al 3+ content in atom formula units of (Mg,Fe)(Al,Si)O 3 perovskite in equilibrium with Fe metal. The vertical dashed line indicates the Al content of perovskite in a typical BSE composition. Most samples were synthesized between C and GPa, but those marked solidus are from the peridotite solidus at approximately 2200 C. 400 Frost McCammon

13 In the transition zone, wadsleyite and then ringwoodite (>17 GPa) replace olivine as the stable (Mg,Fe) 2 SiO 4 polymorph. Although olivine dissolves negligible Fe 2 O 3, experiments show the minimum Fe 3+ / Fe ratios of both (Mg 1.8,Fe 0.2 )SiO 4 wadsleyite and ringwoodite to be approximately 2% when synthesized in equilibrium with stishovite and Fe metal. At mid-transition zone conditions, the mantle is comprised of just wadsleyite and majoritic garnet in the approximate molar ratio 60:40. Using data in Figure 4, the minimum possible bulk rock Fe 3+ / Fe ratio is 3%. This is higher than the 2% estimated for pristine mantle, which implies that if the Fe 3+ / Fe ratio of the transition zone is the same as the upper mantle, Fe-Ni metal must form to provide enough Fe 2 O 3. This means that the transition zone must have a similar metal content and fo 2 as the base of the upper mantle, i.e., 0.1 wt% Fe-Ni metal with an fo 2 just below the IW buffer. 4. LOWER MANTLE (Mg,Fe)(Al,Si)O 3 perovskite, CaSiO 3 perovskite, and (Mg,Fe)O ferropericlase comprise the bulk mineralogy of the lower mantle to a depth of 2700 km where perovskite is likely replaced with a so-called postperovskite phase of similar composition (Murakami et al. 2004). The fo 2 of at least the upper portion of the lower mantle can be estimated if it is assumed that the bulk oxygen content of the lower mantle is the same as that of the upper mantle. Seismic tomography appears to show slabs entering the lower mantle (e.g., van der Hilst et al. 1997), which, if correct, means there must be an upward return flow of material mixing with the upper mantle. If the lower mantle were strongly enriched in Fe 2 O 3 then mixing with the upper mantle would have surely perturbed melt fo 2 proxies such as V/Sc over time (see Section 2.1). Regions of strong enrichment could exist, but they must be isolated. If, on the other hand, the lower mantle is slightly enriched in Fe 2 O 3, then as discussed below, this is unlikely to significantly affect the well-buffered fo 2 of the lower mantle. The fo 2 of the lower mantle can be assessed by examining the minimum Fe 3+ / Fe ratios of lower mantle minerals that occur when they are in equilibrium with metallic Fe. Samples of Al-free (Mg,Fe)SiO 3 perovskite synthesized in equilibrium with metallic Fe and either stishovite or ferropericlase at C(Figure 4c) have higher Fe 3+ / Fe ratios than any other mantle silicate at the same redox conditions. However, as shown in Figure 4d, the perovskite Fe 2 O 3 content is also a strong function of Al 2 O 3 concentration. The Fe 3+ / Fe ratio of perovskite in equilibrium with metallic Fe is more than 50% for a typical mantle Al 2 O 3 content. Perovskite Fe 2 O 3 contents at the peridotite solidus (2200 C) are slightly lower, but rather than an effect of temperature, this is more likely a result of the lower Fe/(Fe+Mg) ratios of samples at or just above the solidus. The strong preference of (Mg,Fe)(Al,Si)O 3 perovskite for Fe 2 O 3 most likely occurs because of the energetically favorable coupled substitution of Fe 3+ on to the eightfold coordinated Mg site charge balanced by Al substitution on to the sixfold coordinated Si site. The size mismatch caused by this coupled substitution is approximately the same for both sites (Kesson et al. 1995), and several computer simulations The Redox State of Earth s Mantle 401

14 have shown this substitution to be the most favorable for Fe 3+ and Al incorporation in perovskite (Richmond & Brodholt 1998, Zhang & Oganov 2006). Mineral/melt trace element partitioning experiments indicate that the perovskite Al 2 O 3 content is also coupled to a range of trivalent trace element concentrations (Liebske et al. 2005). An alternative mechanism, however, is to substitute both Al and Fe 3+ on to the Sisite and charge balance with the creation of an oxygen vacancy, which, for example, would invoke the end-member MgFe 3+ O 2.5. There is evidence that both Fe 3+ and Al take part in such an oxygen vacancy substitution mechanism particularly at low total trivalent cation concentrations (Navrotsky 1999, McCammon 1998, Navrotsky et al. 2003, Lauterbach et al. 2000). A number of studies indicate, however, that increasing pressure and trivalent cation concentration favor a coupled substitution mechanism, making FeAlO 3 and AlAlO 3 the most important trivalent perovskite components throughout the lower mantle (Richmond & Brodholt 1998, Frost & Langenhorst 2002, Walter et al. 2006, Zhang & Oganov 2006). Experiments on Al-bearing perovskite performed in the presence of ferropericlase and metallic Fe can be described by the equilibrium AlAlO 3 pv/gt +3FeO = Fe + 2AlFeO 3. (14) fper metal pv/gt The high concentration of the FeAlO 3 component in perovskite in equilibrium with Fe-metal implies that this equilibrium is shifted significantly to the right at lower mantle conditions, and if the bulk oxygen concentration of the lower mantle is the same as the upper mantle, FeO will be forced to disproportionate and precipitate metallic Fe to stabilize the FeAlO 3 component. On the basis of data, such as that in Figure 4d, Frost et al. (2004) calculated that of the order of 1 wt.% metallic Fe would be forced to precipitate if the lower mantle has a typical BSE composition. The metallic phase would comprise approximately 88 wt% Fe, 10 wt% Ni and 1 wt% S, and would strongly partition siderophile elements. The fo 2 in the lower mantle can be calculated as approximately 1.4 log units below IW, using the compositions of the metal and coexisting ferropericlase and employing ferropericlase activity coefficients (Frost 2003). The fo 2 of the lower mantle will, therefore, be very well buffered over a quite narrow range of values between IW and IW 1.5. Significant increases in the Fe 2 O 3 content of lower mantle would be necessary to raise the fo 2 to the IW buffer where, at the elimination of the metal phase, perovskite would have a Fe 3+ / Fe ratio of 50%. However, it is unlikely that the fo 2 would descend significantly below IW-1.5, as this would require large amounts of FeO to disproportionate. The formation of metallic Fe within and around grains of Al-bearing and Al-free perovskite has been reported in several experimental studies (Miyajima et al. 1999, Lauterbach et al. 2000, Frost et al. 2004, Kobayashi et al. 2005). This appears to be also evidence for disproportionation. In multianvil experiments, however, it is difficult to prove that Equation 14 actually formed this metal, as the multianvil does not provide a closed system with respect to oxygen and reduction could have occurred by some other agent. Although the diamond anvil cell (e.g., Miyajima et al. 1999) should be a closed system, laser heated experiments could conceivably create regions of oxidation and reduction in response to strong thermal gradients. To demonstrate the occurrence 402 Frost McCammon

15 AlAlO 3 Fe 3+ AIO 3 24 GPa Annu. Rev. Earth Planet. Sci : Downloaded from Figure 5 Molar volume (cm 3 mol -1 ) GPa MgSiO 3 mol% T 2 O 3 (T = Al, Fe 3+ ) Molar volumes of perovskites synthesized along the joins MgSiO 3 -Al 2 O 3 (Walter et al. 2004, 2006) and MgSiO 3 -Fe 3+ AlO 3 (Nishio-Hamane et al. 2005). The shaded region indicates the range of possible linear fits through the two MgSiO 3 -FeAlO 3 data points that give end-member Fe 3+ AlO 3 molar volumes in the range cm 3 mol 1. of FeO disproportionation high Fe 2 O 3 contents need to be measured within samples that are clearly in equilibrium with metallic Fe. Evidence for Equation 14 comes mainly from multianvil experiments performed at 25 GPa, i.e., only the very top of the lower mantle. Whether Fe metal remains stable throughout the entire Al-perovskite stability field will depend on the volume change of Equation 14. If it is negative, metal and the FeAlO 3 perovskite component will become more stable with pressure. The limited available data to assess this are shown in Figure 5. From measurements of perovskite volumes on the MgSiO 3 -AlAlO 3 join, the volume of the AlAlO 3 end-member (25.76 cm 3 mol 1 ) can be estimated if a linear relationship is assumed. Two measurements exist for FeAlO 3 -bearing perovskites synthesized at 24 and 51 GPa (Nishio-Hamane et al. 2005). The 51 GPa composition falls almost perfectly on the MgSiO 3 -FeAlO 3 join, whereas the 24 GPa composition also has a significant AlAlO 3 component (5%), although why it has a larger volume is unclear. If the volume of the FeAlO 3 perovskite end-member is estimated using linear extrapolations through both the 24 GPa and 51 GPa data points, the calculated volume change for Equation 14 is either 0 or 2 cm 3 mol 1, respectively. The available data, therefore, indicate either no volume change or a volume change that favors disproportionation. Experimental evidence exists for several pressure-induced transitions that could potentially shift Equation 14 toward the left- or right-hand sides at pressures >25 GPa. Between 50 and 60 GPa, for example, metallic Fe transforms from The Redox State of Earth s Mantle 403

16 γ (fcc) to ε (hcp) structures (Kubo et al. 2003), which will stabilize components on the right-hand side of Equation 14. However, the decrease in volume (Tsuchiya et al. 2006) caused by the change in the electronic spin state of Fe 2+ from high spin to low spin in ferropericlase at GPa (Badro et al. 2003) and perovskite at GPa (Badro et al. 2004) may potentially shift Equation 14 to the left, destabilizing Fe 2 O 3 in perovskite and Fe metal. The effect of pressure on Fe-Mg partitioning between perovskite and ferropericlase will also influence Equation 14 because, as shown in Figure 4c, the perovskite Fe 2 O 3 content also depends on the Fe/(Fe+Mg) ratio. The available experimental data seem to indicate that there is little effect of pressure on Fe-Mg partitioning between perovskite and ferropericlase (Kesson et al. 1998, Kobayashi et al. 2005, Murakami et al. 2005), which also provides at least some evidence that spin transitions may not significantly influence Fe-bearing equilibria. The existing data, however, are not comprehensive or entirely consistent. Some ab initio calculations also indicate that Equation 14 should stabilize metallic Fe in the lower mantle throughout the entire perovskite stability field (Zhang & Oganov 2006). Toward the bottom of the lower mantle perovskite transforms to the postperovskite phase at a depth that may coincide with the top of the D region (Murakami et al. 2004). Preliminary diamond cell measurements indicate that Fe 2 O 3 has a high solubility in the postperovskite phase (Sinmyo et al. 2006), with measured Fe 3+ / Fe ratios of more than 60%, whereas more recent experiments containing ferropericlase show significantly lower Fe 3+ / Fe values (<15%) (Sinmyo et al. 2007). Although the postperovskite Fe 3+ / Fe ratio in equilibrium with metallic Fe has not been measured, the ab initio calculations (Zhang & Oganov 2006) indicate that an Fe 3+ -Fe 3+ coupled substitution is energetically favored in postperovskite, and that this can drive the precipitation of metallic Fe through FeO disproportionation. Although further studies are necessary to determine the volume change of Equation 14 and to measure the Fe 2 O 3 content of perovskite in equilibrium with metallic Fe at very high pressures, the existing experimental data and ab initio calculations are not inconsistent with Fe metal and Fe 2 O 3 -rich perovskite and postperovskite coexisting throughout the entire lower mantle. Within the present-day lower mantle, this metal will be solid and will not separate from the solid lower mantle assemblage, although the metal may be molten within the thermal boundary layer at the core mantle boundary. As material downwells into the lower mantle, Equation 14 will shift to the right, precipitating metallic Fe. If, however, upwelling mantle leaves the Al-perovskite stability field, the equilibrium will move to the left, consuming both Fe metal and Fe 2 O 3 to the level at which they coexist in the overlying transition zone. At the core-mantle boundary the Fe-metal within the lower mantle may be liquid and could potentially separate to the core, enlarging the core over time and resulting in a net increase in the oxygen content of the mantle. Several mechanisms may prevent this, however. First, as shown in Figure 4d, at higher temperatures the Fe 2 O 3 content of perovskite (and potentially postperovskite) decreases and the proportion of precipitated metal will, therefore, also decrease. Second, it has been proposed based on high-pressure experiments that the Fe liquid of Earth s core may be undersaturated with respect to the typical FeO content of the mantle (Takafuji et al. 2005, Asahara et al. 2007) and may thus dissolve FeO from the mantle at the core mantle boundary. 404 Frost McCammon

17 This would leave an oxygen depleted region (i.e. low FeO and Fe 2 O 3 ) at the core mantle boundary that, based on rates of cation diffusion (Holzapfel et al. 2005) may be only a few meters thick but would effectively shield the remaining mantle from Fe metal loss. Annu. Rev. Earth Planet. Sci : Downloaded from 5. THE SPECIATION OF C-O-H-S-BEARING PHASES The Fe 3+ / Fe ratio of the mantle and the speciation of C-O-H-S phases are intimately connected and may both control mantle fo 2 under certain conditions. C-H-O-S species may control mineral and melt Fe 3+ / Fe ratios during melting, degassing, and some metasomatic processes, such as those that produce diamonds. The narrow range of fo 2 measurements in certain mantle samples has raised the suggestion that fo 2 may be buffered by, for example, the CCO buffer (Blundy et al. 1991). Most possible buffering reactions are sliding buffers where at least one side of the reaction potentially involves a solution. In many instances, reactions are therefore likely to adjust to changes in fo 2 rather than buffering them (Woodland et al. 2006). Although mantle C phases can be pure and are, therefore, potentially the most plausible redox buffers, it has also been questioned (Canil et al. 1994) as to whether typical mantle C contents are high enough to buffer perturbations arising from other components. Increasing pressure favors the stability of Fe 2 O 3 components, which lowers fo 2 with depth, leading to the reduction of C-O-H-S volatile species. The additional Fe 2 O 3 produced by volatile reduction will act to impede the decrease in fo 2 with pressure caused by Fe 2+ /Fe 3+ equilibria (shown in Figure 3), but this impedance will be small for typical mantle volatile contents (O Neill et al. 1993b). Fe 2+ /Fe 3+ equilibria, therefore, exert a directional control on the fo 2 with depth, but this does not mean they buffer fo 2, as it can still be perturbed by incoming components C-O-H Fluids and Melts The speciation of a pure C-O-H fluid as shown in Figure 6 can be used as a guide to the types of volatile species likely to be stable in the mantle, although these species will predominantly exist as components in melts and minerals. The fluid speciation in Figure 6a is calculated along an average mantle adiabat at an fo 2 determined using Equation 7 on a model fertile peridotite assemblage with a bulk Fe 3+ / Fe ratio of 2%, which results in Fe-Ni metal formation at approximately 8 GPa. At oxygen fugacities where graphite or diamond become stable, the speciation of a C- O-H fluid becomes a function of fo 2 and can be calculated from equation of state estimates of fluid fugacities using the method described by Holloway (1987). The calculations in Figure 6 are based on computer simulations of pure fluid properties (Belonoshko & Saxena 1992) and no high-pressure and -temperature experimental data exist to compare these determinations. Absolute proportions of volatile species calculated in Figure 6 are uncertain; however, the general trends in which pressure and temperature drive fluid speciation are likely to be correct. During MORB melting, Figure 6a predicts that CO 2 and H 2 O are the major volatile species, which will be dissolved in the silicate melt. No hydrous minerals The Redox State of Earth s Mantle 405

18 a 1 C-O-H fluid speciation along an adiabat in the upper mantle Graphite stable b 1 H 2 O and CH 4 speciation in the lower mantle H 2 O 25 GPa GPa GPa CH 4 Annu. Rev. Earth Planet. Sci : Downloaded from Mole fraction in fluid CO 2 + H 2 O stable CCO CO 2 Diamond stable H 2 CO NAM: nominally anhydrous mineral LVZ: low-velocity zone Pressure (GPa) Figure 6 CH 4 H 2 O Mole fraction in fluid GPa Expected fo Log fo 2 ( IW) The speciation of the C-O-H fluid phase in equilibrium with graphite/diamond calculated (a) as a function of pressure in the upper mantle along an adiabat with a potential temperature of 1200 C and at an fo 2 defined by curves 1 and 2 in Figure 3, and (b) at lower mantle pressures (25 and 50 GPa at 1600 C) as a function of oxygen fugacity relative to the IW buffer. At oxygen fugacities below the CCO buffer, indicated by the vertical line in (a), graphite/ diamond are unstable and mixtures of CO 2 and H 2 O are stable at any fo 2. The vertical dotted line in (b) shows the expected lower mantle fo 2. The calculations assume ideal mixing in the gas phase and use the equations of state of Belonosko & Saxena (1992). (e.g., phlogopite or amphibole) are stable along an adiabat (Frost 2006), although H 2 O will also partition into NAMs such as clinopyroxene (Hirschmann 2006). The depth where the peridotite solidus is crossed and melting commences will depend on the H 2 O and CO 2 contents of the mantle (Dasgupta et al. 2007). Typical MORB H 2 O contents, in the range ppm, can dissolve completely into NAMs with increasing pressure (Hirschmann 2006), but as carbonate minerals are unstable at adiabatic temperatures below 3 GPa, the peridotite solidus will be depressed by CO 2 (Luth 1999, Dasgupta et al. 2007) and a small-degree of volatile-rich melt may be produced at significant depths (>100 km or 3 GPa). Above 3 GPa, however, graphite will become stable and the concentration of CO 2 in melts will decrease with increasing pressure as the fo 2 decreases. In this way, the beginning of melting may be controlled by the elimination of graphite in upwelling mantle as the fo 2 increases, which has been termed redox melting (Taylor & Green 1988, Ballhaus & Frost 1994). The high pressure limit of the LVZ in the asthenospheric mantle (if it is caused by partial melting) could be controlled by redox melting processes in addition to the increase in H 2 O partitioned into NAMs with pressure (Mierdel et al. 2007). At lower temperatures, along a continental geotherm for example, carbonate minerals can be stable in the upper mantle, but their presence will be controlled by the 406 Frost McCammon

19 EMOG/D buffer: MgSiO 3 + MgCO 3 = Mg 2 SiO 4 + C + O 2. (15) enstatite magnesite olivine graphite/diamond As shown in Figure 3, the fo 2 of cratonic mantle crosses the EMOG/D buffer into the graphite stability field at approximately 3 GPa. At higher pressures, analogous reactions to Equation 15 show that the fo 2 in the transition zone and lower mantle is at least 3 log units below the carbonate stability field. Solid carbonates do not exist along an adiabat as they are unstable with respect to melts at lower pressure and to graphite and diamond at higher pressures. Therefore, CO 2 -rich magmas, such as kimberlites that originate from >4 GPa, must have sources more oxidized, i.e., with a higher Fe 3+ / Fe ratio, than typical mantle. H 2 O and graphite/diamond are predicated to be the dominant C-O-H species in the mantle at approximately 5 GPa (Figure 6a). At these conditions, the H 2 O budget of all but the most H 2 O-rich mantle sources can most likely be stored in NAMs (Hirschmann 2006). Above 6 GPa, however, CH 4 becomes an increasingly important fluid component. Studies at lower pressures (<3 GPa) indicate that CH 4 has a low solubility in silicate melts ( Jakobsson & Holloway 1986, Taylor & Green 1988). If this is also the case at higher pressures, CH 4 will not induce peridotite melting and NAMs and diamond may continue to be the main hosts of C and H. By 8 GPa, however, CH 4 is predicted to be the dominant fluid component. Measurements of the solubility of H 2 O in NAMs are performed, generally, by equilibrating minerals with pure H 2 O, although silicates dissolve in the fluid at high temperatures (Kohlstedt et al. 1996). Such measurements may be inappropriate for the deeper mantle if fluids are dominated by CH 4. In the deep upper mantle, NAMs may be forced to dehydrate to produce a CH 4 -dominated fluid with a low H 2 O activity, consistent with observations of low water contents in mantle-derived omphacites from eclogite xenoliths (Koch-Müller et al. 2004). If CH 4 -rich fluids do not contain a significant dissolved silicate component, they will retain a low density and be mobile. Such fluids should rise out of reduced regions of the mantle and be oxidized in the shallower mantle below 6 GPa. Such a mechanism may act to focus C-O-H fluids in the asthenosphere, maintaining the LVZ against volatile loss into the lithosphere and potentially keeping the Fe 3+ / Fe ratio of the upper mantle quite low. As shown in Figure 6b, with respect to fo 2, the H 2 O stability field in equilibrium with diamond is predicted to expand with increasing pressure. Unless the fo 2 of the mantle decreases relative to IW, fluids are predicted to evolve to become more H 2 O- rich with pressure. Although CH 4 may become an important fluid species in the transition zone and top of the lower mantle as the H 2 O stability field continues to expand with pressure in the deeper lower mantle, the prediction is that fluids will become H 2 O-dominated. EMOG/D: enstatitemagnesite-olivine-graphite/ diamond; MgSiO 3 + MgCO 3 = Mg 2 SiO 4 + C + O Diamonds If Fe-Ni rich metal precipitates in the mantle as proposed (Ballhaus 1995, Frost et al. 2004, Rohrbach et al. 2007), diamond should become unstable with respect to either The Redox State of Earth s Mantle 407

20 Fe 3 C or C dissolved in the Fe metal (Wood 1993). Rare occurrences of native Fe inclusions in diamonds would seem to contradict this (Sobolev et al. 1989, Stachel et al. 1998), although in some instances Fe metal seems to be separated from the diamond by another phase (Stachel et al. 1998) and there are also examples where Fe 3 C and Fe metal occur in diamond inclusions ( Jacob et al. 2004). Macroscopic diamonds in xenoliths represent large local concentrations of C formed through metasomatic processes. These processes must involve either reduction of CO 2 -rich fluids/melts or oxidation of CH 4 -rich fluids, both of which will perturb the fo 2 of the local environment. The likely involvement of a fluid influx means that mineral or fluid inclusions trapped in diamonds do not necessarily record typical mantle fo 2 s. It is also possible that FeS-rich liquids could directly precipitate dissolved C without a redox reaction (Litvin et al. 2002). CO 2 -bearing fluids or melts leaving a subducting slab would reduce to form diamonds in the ambient mantle at any pressure >4 GPa, whereas, as shown in Figure 6a, CH 4 -rich fluids rising from the deeper mantle toward the H 2 O maximum at 6 GPa would be oxidized to produce diamond and H 2 O. An increase in fo 2 within the upper regions of the lower mantle as Al-bearing perovskite becomes unstable may also produce conditions favorable to diamond growth. Diamonds might potentially exsolve from Fe-Ni alloy at this depth. The compositions of rare mineral inclusions in diamonds that appear to be lower mantle assemblages support the idea of diamond growth in the very top of the lower mantle (McCammon 2005) S-Bearing Phases As fo 2 decreases with pressure, the (Ni,Fe)/S ratio of the mantle sulfide assemblage should increase to produce more Fe-Ni-rich liquids. Above 8 GPa, Fe-Ni metal should coexist with a sulfide liquid because the solubility of S in solid metal is low. As shown in Figure 7 at 10 GPa where the mantle temperature is approximately 1400 C, Fe-rich solid metal will coexist with an Fe-S liquid that is far more Fe rich than the FeS liquid, which occurs at shallow mantle conditions. At lower mantle conditions, however, the S solubility in solid Fe is higher and the entire S budget of the mantle can be dissolved in Fe metal (Li et al. 2001), which should contain on the order of 1 wt% S. H 2 S should also become a significant volatile component in the deep mantle, although its speciation is currently hard to determine as models for the fs 2 of the mantle are lacking and its equation of state is poorly constrained at high pressure. 6. REDOX STATE IN SUBDUCTION ZONES There is potentially a flux of oxidized components returned to the mantle in subduction zones (Lécuyer & Ricard 1999). Iron in oceanic crust reacts with seawater before being subducted; for example, olivine reacts at hydrothermal temperatures to produce magnetite (Shanks et al. 1981), which is also a product of serpentinization reactions involving peridotite that take place at mid-ocean ridges (Bach et al. 2004). Other oxidized species, such as carbonate, may enter subduction zones through 408 Frost McCammon

21 Fe-FeS GPa Annu. Rev. Earth Planet. Sci : Downloaded from Figure 7 Temperature ( C) Fe 3 GPa Metal + liquid Fe mole % Liquid Metal + sulphide Sulphide + liquid FeS The Fe-FeS join at 3 and 10 GPa (Brett & Bell 1969, Fei et al. 1997). Circles show the approximate composition of sulfide liquids at each pressure. At 10 GPa, Fe-rich metal will coexist with an Fe-rich sulfide melt as shown by the horizontal dashed line. sediments (Rea & Ruff 1996), but may not reach the deep mantle if they are lost from the slab during dehydration and melting reactions. Oxidized components may be transferred to the overlying mantle through melts and fluids (e.g., Wood et al. 1990), although exact details of the oxidizing mechanism(s) are still disputed. Fe 2 O 3 is a potential candidate, for example, through mineral/melt partitioning during partial melting (e.g., Amundsen & Neumann 1992, Mungall 2002) leading to metasomatism (McGuire et al. 1991); however, H 2 O is not because, as shown in Figure 6a, the fo 2 of the ambient mantle is compatible with the presence of H 2 O until 5 GPa and there is no tendency for it to reduce. Only at deeper conditions in regions where CH 4 and H 2 are in equilibrium with the mantle fo 2 could the release of H 2 O lead to mantle oxidation, although this would create mobile reduced volatiles that are likely to ascend and reduce the mantle at shallower depths. Fe 2 O 3 could also be lost from the subducting slab through the enhanced solubility of magnetite in fluids in the presence of chloride (Chou & Eugster 1977), which is likely present in subduction-zone fluids based on observed trace element and isotopic enrichment patterns (Keppler 1996). One possibility is that subduction-related Fe 2 O 3 recycling may be restricted to the lithospheric mantle wedge environment, which could account for the constancy of upper mantle fo 2 since the Archean (Li & Lee 2004). Some oxidation of the asthenospheric mantle, on the other hand, may be required to balance mantle degassing of CO 2 and SO 2. Alternatively, slabs may retain oxidized components and recycle them into the lower mantle (Lécuyer & Ricard 1999), although this does not seem to be The Redox State of Earth s Mantle 409

22 reflected in the oxidation state of magmas and massifs associated with mantle plumes (Wood et al. 1990, Bézos & Humler 2005, Rhodes & Vollinger 2005). Mantle oxygen fugacity has a strong influence on the nature of subduction redox reactions; hence the possibility that if the mantle were more reduced in its early history, it could have been oxidized by H 2 O recycled by subduction (Kasting et al. 1993). If after core formation the upper mantle were at conditions below IW, for example, the release of H 2 O from a slab would cause oxidation because the mantle would be in equilibrium with an H 2 - and CH 4 -bearing fluid or melt. One argument against this, however, is that temperatures in the early Hadean and Archean mantle were likely higher, causing oxidizing components, such as CO 2,H 2 O and Fe 2 O 3,to be released into the mantle at much shallower depths where they would have strongly partitioned into melts that ultimately ended up back in the lithosphere. High degrees of melting brought on by injecting H 2 O into the early mantle may have eliminated spinel and clinopyroxene from the mantle assemblage, which are the main hosts of Fe 2 O 3, therefore limiting the oxidation of the residual asthenospheric mantle. 7. EVOLUTION OF THE MANTLE S REDOX STATE As described in Section 2.1, concentrations of Cr, V, and the V/Sc ratio of the oldest known rocks indicate that the redox state of the upper mantle has remained at approximately FMQ over the past 3.5 Gyrs (Delano 2001, Canil 2002, Li & Lee 2004). At 4.5 Gyrs during core formation, the silicate portion of Earth must have been in equilibrium with Fe-Ni rich metal, which would have resulted in an upper mantle fo 2 of 4.5 log units below FMQ. The mantle, therefore, must have been oxidized by some process during the first 1 Gyrs of Earth history. The increase in mantle redox state is more clearly evident when we consider the Fe 2 O 3 content of the upper mantle. Solid or liquid silicates in equilibrium with Fe-rich metal at upper mantle conditions contain essentially no Fe 2 O 3, whereas the current upper mantle contains at least 0.2 wt% Fe 2 O 3 (Canil & O Neill 1996). Most models attribute the oxidation of the mantle to the reduction of H 2 O releasing H 2, which was ultimately lost to space. H 2 O may have either been incorporated in the material forming the mantle, and oxidation therefore occurred as differentiation proceeded (Ringwood 1977, O Neill 1991), or once Earth had cooled, subduction of hydrated lithosphere may have oxidized the mantle (Kasting et al. 1993). The first possibility, however, must be seen in light of new insights into the redox state of the Martian mantle, which although potentially a more volatile-rich planet than Earth, seems to have a much lower upper mantle fo 2 (Wadhwa 2001, Herd et al. 2002); arguments against the second possibility were raised in the previous section. The potential for metallic Fe to form by disproportionation mainly in the lower mantle, however, raises another possible explanation for the increase in mantle oxidation state (Mao & Bell 1977, Frost et al. 2004, Galimov 2005). If the lower mantle formed from silicates that had been in equilibrium with metallic Fe and therefore contained no Fe 2 O 3, then as Al-bearing perovskite crystallized, FeO in mantle silicates must have disproportionated to produce Fe 2 O 3 in perovskite and Fe metal. If some of this precipitated metal separated to the core, then the residual silicate would have been 410 Frost McCammon

23 left with a greater O/Fe ratio. If approximately 10% of the disproportionated metal entered the core, then the remaining lower mantle would contain enough excess Fe 2 O 3 to explain the present-day upper mantle Fe 3+ / Fe ratio after convection had homogenized the entire mantle (Frost et al. 2004). This mechanism has been referred to as an oxygen pump (Galimov 2005, Wood et al. 2006), as oxygen is pumped into the mantle. It is interesting to note that the more reduced nature of the Martian mantle (Wadhwa 2001, Herd et al. 2002) may be consistent with the lack of a significant perovskite-bearing lower mantle to drive an oxygen pump (Wood et al. 2006). A number of scenarios could explain the loss of some disproportionated metal from the lower mantle to the core, thus raising the mantle O/Fe ratio. Many models propose that during the later stages of accretion, core-forming metal descended through the solid lower mantle as diapirs (Li & Agee 1996, Rubie et al. 2003), which could have entrained disproportionated metal as they descended. Alternatively, as Fe 2 O 3 concentrations are high in perovskite even at the peridotite solidus (Frost et al. 2004), then if perovskite crystallized from a deep magma ocean during accretion, some metallic Fe would have also been formed and settled, to be incorporated into separating metallic diapirs. Another possibility, particularly if a magma ocean extended deep into the lower mantle, is that large amounts of disproportionated metal were continually lost from the lower mantle to the core, but the rise in Fe 2 O 3 was always tempered by the addition of further Fe metal from accreting bodies that reduced Fe 2 O 3 back to FeO (Wood et al. 2006). To explain the depletions of certain weakly siderophile elements from the mantle, some models argue for an initial reduced phase of core formation where very little FeO (<1 weight %) would have been left in Earth s proto mantle (Wänke 1981, O Neill 1991, Wood et al. 2006). Disproportionation of FeO followed by loss of Fe to the core could not, however, have raised the fo 2 of the mantle from these very reducing conditions as this requires the addition of FeO to the mantle rather than an increase in the O/Fe ratio. The presence of metallic Fe in the lower mantle during core formation raises the possibility of whether this metal could explain the apparent overabundance of highly siderophile elements (e.g., Pt, Ir, Rh) in the mantle. It is likely, however, that the lower mantle formed from material that had already experienced core-mantle differentiation so that disproportionating metal would have formed from material already stripped of its siderophile elements. Only Fe metal that had equilibrated with undifferentiated silicates, which could be termed core-forming metal, would have high siderophile element concentrations. One possibility is that the highly siderophile element concentration in the mantle results from core-forming metal that was also trapped in the lower mantle along with disproportionated metal toward the end of accretion. This is reminiscent of inefficient core formation models ( Jones & Drake 1986) to explain siderophile element overabundance, except that while in these previous models the agent to oxidize the metal remained unclear (O Neill 1991), here the metal is in equilibrium with its own oxidant, i.e., perovskite Fe 2 O 3, which oxidizes the metal as material leaves the lower mantle. By adjusting the proportions of disproportioned and core-forming metal initially trapped in the lower mantle, the overabundance of highly sideriophile elements could be explained. The Redox State of Earth s Mantle 411

24 SUMMARY POINTS 1. Spinel peridotite rocks record upper mantle fo 2 in the range FMQ ± 2 log units. Samples with closer links to the asthenosphere fall at the more reduced end of this range, whereas samples with long histories in the lithosphere or from subduction zones appear more oxidized. 2. The fo 2 recorded by garnet peridotite rocks decreases with depth owing to the effect of pressure on the controlling Fe 3+ /Fe 2+ equilibria. 3. At low pressures, mineral Fe 3+ / Fe ratios drop to insignificant levels as fo 2 approaches IW. Pressure stabilizes mineral Fe 2 O 3 components, however, and many high-pressure minerals, including garnet, have significant Fe 3+ / Fe ratios at and below IW. 4. By extrapolating the effect of Fe 3+ /Fe 2+ equilibria on mantle rocks to higher pressure it can be shown that mantle fo 2 will drop to levels consistent with the precipitation of Ni-rich Fe metal at approximately 8 GPa. The formation of Ni-Fe metal results from the disproportionation of FeO to produce metal and Fe 2 O 3 -bearing garnet. 5. At the base of the upper mantle, CH 4 and H 2 O will be the dominant C-O-H fluid species. Such fluids may tend to rise out of the deeper mantle and be oxidized in the shallow asthenosphere. 6. Experimental results imply that Fe-rich metal should be stable throughout the transition zone and is likely present in quantities close to 1 weight% in at least the top of the lower mantle. The available experimental evidence suggests that Fe-Ni metal may be present throughout the lower mantle but the data are inconclusive. 7. Loss of disproportionated metallic Fe from a solid lower mantle during core formation would have raised the bulk oxygen content of the remaining mantle. After mantle homogenization by convection, the fo 2 of the upper mantle would have been raised to its present-day value toward the end of core formation. This would be one explanation as to why even the oldest rocks from the mantle (3.5 Gyrs) seem to record oxygen fugacities similar to the present-day mantle. FUTURE ISSUES 1. The Fe 3+ / Fe contents of MORB glasses define a very narrow range, but the mechanism that controls this needs to be resolved. 2. In upwelling mantle beneath spreading ridges, the conditions at which diamond or graphite is oxidized to carbonate liquid are unconstrained but could possibly control the beginning of MORB partial melting. 412 Frost McCammon

25 3. The behavior of reduced C-O-H-S fluids requires further investigation, such as their effect on the deep mantle solidus and equilibrium mineral OH contents. Virtually no experimental data exist on the equation of state or behavior of CH 4 or H 2 S at high pressure. 4. Fluxes of oxidized species into and out of the mantle need tighter constraints. 5. The minimum Fe 3+ / Fe ratios of high-pressure minerals over a range of pressures, temperatures, and compositions still need to be measured. In particular, the Fe 3+ / Fe ratio of magnesium silicate perovskite as a function of fo 2 needs to be examined over conditions covering the entire lower mantle. Annu. Rev. Earth Planet. Sci : Downloaded from DISCLOSURE STATEMENT The authors are not aware of any biases that might be perceived as affecting the objectivity of this review. ACKNOWLEDGMENT We are very grateful to A. Woodland for his comments on an early version of this manuscript. LITERATURE CITED Amundsen HEF, Neumann E-R Redox control during mantle/melt interaction. Geochim. Cosmochim. Acta 56: Asahara Y, Frost DJ, Rubie DC Partitioning of FeO between magnesiowüstite and liquid iron at high pressures and temperatures: implications for the composition of the Earth s outer core. Earth Planet. Sci. Lett. 257: Bach W, Garrido CJ, Paulick H, Harvey J, Rosner M Seawater-peridotite interactions: first insights from ODP Leg 209, MAR 15 N. Geochem. Geophys. Geosys. 5:Q09F26 Badro J, Fiquet G, Guyot F, Rueff JP, Struzhkin VV, et al Iron partitioning in Earth s mantle: toward a deep lower mantle discontinuity. Science 300: Badro J, Rueff JP, Vanko G, Monaco G, Fiquet G, et al Electronic transitions in perovskite: possible nonconvecting layers in the lower mantle. Science 305: Ballhaus C Is the upper mantle metal-saturated? Earth Planet. Sci. Lett. 132:75 86 Ballhaus C, Berry RF, Green DH High pressure experimental calibration of the olivine-orthopyroxene-spinel oxygen geobarometer: implications for the oxidation state of the upper mantle. Contrib. Mineral. Petrol. 107:27 40 Ballhaus C, Frost BR The generation of oxidised CO 2 -bearing basaltic melts from reduced CH 4 -bearing upper mantle sources. Geochim. Cosmochim. Acta 58: The Redox State of Earth s Mantle 413

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33 Contents Annual Review of Earth and Planetary Sciences Volume 36, 2008 Annu. Rev. Earth Planet. Sci : Downloaded from Frontispiece Margaret Galland Kivelson xii The Rest of the Solar System Margaret Galland Kivelson 1 Abrupt Climate Changes: How Freshening of the Northern Atlantic Affects the Thermohaline and Wind-Driven Oceanic Circulations Marcelo Barreiro, Alexey Fedorov, Ronald Pacanowski, and S. George Philander 33 Geodynamic Significance of Seismic Anisotropy of the Upper Mantle: New Insights from Laboratory Studies Shun-ichiro Karato, Haemyeong Jung, Ikuo Katayama, and Philip Skemer 59 The History and Nature of Wind Erosion in Deserts Andrew S. Goudie 97 Groundwater Age and Groundwater Age Dating Craig M. Bethke and Thomas M. Johnson 121 Diffusion in Solid Silicates: A Tool to Track Timescales of Processes Comes of Age Sumit Chakraborty 153 Spacecraft Observations of the Martian Atmosphere Michael D. Smith 191 Crinoid Ecological Morphology Tomasz K. Baumiller 221 Oceanic Euxinia in Earth History: Causes and Consequences Katja M. Meyer and Lee R. Kump 251 The Basement of the Central Andes: The Arequipa and Related Terranes Victor A. Ramos 289 Modeling the Dynamics of Subducting Slabs Magali I. Billen 325 vii