5. The buffering capacity of lithospheric mantle： implications for diamond formation

Contrib Mineral Petrol (2014) 168:1083DOI 10.1007/s00410-014-1083-6The buffering capacity of lithospheric mantle: implications for diamond formationRobert W. Luth · Thomas StachelReceived: 2 June 2014 / Accepted: 28 October 2014 / Published online: 8 November 2014 © Springer-Verlag Berlin Heidelberg 201450 ppm or less O 2 is required to shift a depleted man-tle peridotite the observed four log units of f O 2. Coupled with the observed distribution of samples at values of f O 2 intermediate between the most reduced (metal-saturated) and most oxidized (carbonate-saturated) possible values for diamond stability, these results demonstrate that peri-dotites are very poor sinks or sources of O 2 for possible redox reactions to form diamond. A corollary of the poor redox buffering capacity of cratonic peridotites is that they can be employed as faithful indicators of the redox state of the last metasomatic fluid that passed through them. We propose that diamond formation from CHO fluids is a pre-dictable consequence either of isobaric cooling or of com-bined cooling and decompression of the fluid as it migrates upward in the lithosphere. This establishes a petrological basis for the observed close connection between subcalcic garnet and diamond: based on high solidus temperatures of harzburgite and dunite effectively precluding dilution of CHO fluids through incipient melts, such highly depleted cratonic peridotites are the preferred locus of diamond formation. Due to a rapid increase in solidus temperature with increasing CH 4 content of the fluid, diamond forma-tion related to reduced CHO fluids may also occur in some cratonic lherzolites.Keywords Mantle petrology · Diamonds · Oxygen barometry · Mantle oxidation state · CHO fluidIntroductionCurrent models for the formation of natural diamond involve reaction of either a methane-bearing fluid or a car-bonate-bearing fluid (or melt) with the mantle (e.g., Deines 1980; Stachel and Harris 2009). The former mechanism canAbstract Current models for the formation of natural diamond involve either oxidation of a methane-bearing fluid by reaction with oxidized mantle, or reduction of a carbonate-bearing fluid (or melt) by reaction with reduced mantle. Implicit in both models is the ability of the man-tle with which the fluid equilibrates to act as an oxidizing or reducing agent, or more simply, to act as a source or sink of O 2. If only redox reactions involving iron are oper-ating, the ability of mantle peridotite to fulfill this role in diamond formation may not be sufficient for either model to be viable. Using the recent experimental recalibration of olivine–orthopyroxene–garnet oxybarometers of Stagno et al. (2013), we re-evaluated the global database of ~200 garnet peridotite samples for which the requisite Fe 3+/Fe 2+ data for garnet exist. Relative to the previous calibration of Gudmundsson and Wood (1995), the new calibration yields somewhat more oxidized values of Δlog f O 2 (FMQ), with the divergence increasing from <0.5 units of log f O 2 at ~3 GPa to as much as 1.5 units at 5–6.5 GPa. Globally, there is a range of ~4 log units f O 2 for samples from the diamond stability field at any given pressure. Most samples are sufficiently reduced such that diamond, rather than car-bonate, would be stable, and CHO fluids at these conditions would be H 2O-rich (>60 mol%), with CH 4 being the next most abundant species. To ascertain the capacity for mantle peridotite to act as a source or sink of O 2, we developed a new model to calculate the f O 2 for a peridotite at a given P , T , and Fe 3+/Fe 2+. The results from this model predictCommunicated by J. Hoefs.R. W. Luth (*) · T. StachelDepartment of Earth and Atmospheric Sciences,University of Alberta, Edmonton, AB T6G 2E3, Canada e-mail: robert.luth@ualberta.caContrib Mineral Petrol (2014) 168:1083 1083Page 2 of 12be represented by the reaction CH4+ O2= C + 2 H2O; the latter by the reactions CO2= C + O2 or CO32− (melt/ fluid) = C + O2− (melt/fluid) + O2. The stoichiometries of all these reactions have C:O2 of 1:1 on a molar basis. To what extent is the mantle capable of acting as a source or sink of this quantity of O2 as a necessary by-product of diamond formation?Arguably, the first mechanism could be expressed by CH4= C + 2 H2, in which hydrogen is liberated, and the second by CO32− (melt/fluid) + 2 H2= C + O2− (melt/ fluid) + 2 H2O, in which hydrogen is consumed. In these cases, the mantle must be able to act as a source or sink for hydrogen. Either hydrous minerals or nominally anhydrous minerals could fill this role, but there would need to be a coupled sink or source of oxygen, given that H dissolves in both hydrous and nominally anhydrous minerals as OH−(Keppler and Bolfan-Casanova 2006), which brings us back to the oxygen storage ability of the mantle.Oxidation–reduction reactions are usually consid-ered to be the mechanism by which the mantle acts as a source or sink of oxygen. These redox reactions involve elements potentially stable in multiple valence states at mantle conditions. Candidates previously discussed in the literature include carbon, iron, sulfur, and nitrogen (e.g., Frost and McCammon 2008; Haggerty 1986; Luth 1999; Wood et al. 1990). Much attention has focused on C and Fe; the low S content of the mantle argues for a limited role, although possibly significant in regions with local sulfur enrichment. Redox reactions involving nitrogen, although mentioned by previous authors (e.g., Haggerty 1986), have not been experimentally studied to date. This paper will focus on the coupling of C and Fe equilibria, and the implications of the limitations of Fe equilibria to influence reactions that might precipitate diamond from a fluid. We suggest that if only redox reactions involving iron are available, the mantle has an extremely limited capability to fulfill the required role of source or sink of oxygen.Buffering capacityIn the context of this paper, the “buffering capacity” of lithospheric mantle is its ability to act as a source or sink for oxygen; a rock with a high buffering capacity would have an ability to provide or absorb the oxygen required or produced during diamond-forming redox reactions with-out significant impact on the rock itself. Determining this capacity requires understanding the nature of the “buffer-ing” reactions in mantle lithologies, as well as how sensi-tive the redox state of mantle peridotite or eclogite is to addition or subtraction of oxygen—in other words, how much mantle must be involved in this process.“Buffering” by ferric/ferrous equilibria in the mantle may be most simply represented by the reaction 2 Fe2O3 (silicate) = 4 FeO (silicate) + O2. An obvious problem with this representation is that Fe2+ and Fe3+ do not substi-tute for each other in minerals for charge-balance reasons, and so more complicated reactions such asare required (or analogous ones for incorporation of fer-ric iron into pyroxene). These reactions can be considered as the mechanism by which oxygen is either supplied or consumed, as well as the equilibrium expressions that allow calculation of the f O2 of a sample. This perspective makes it clear that reactions between different minerals are required to supply or absorb oxygen. This in turn suggests the length scale over which reaction takes place, and the kinetics of these reactions, may play a role in any “buffer-ing capacity” scenario.In the reactions above, olivine is essential as well as pyroxene and garnet; it is not possible to write a balanced reaction without either olivine or SiO2—the first reac-tion above could be written as 2Fe3Fe3+2Si3O12(gt)+ 4SiO2(cs)=5Fe2Si2O6(px)+O2, for example. This stoichiometric requirement implies that bimineralic eclog-ite (cpx + gt) could supply oxygen only by precipitating olivine and could absorb oxygen only if coesite is also formed.Revised oxybarometryWe applied the recent calibration of the olivine–orthopy-roxene–garnet oxybarometer of (Stagno et al. 2013) to the extant database of cratonic garnet peridotite mantle samples that have the necessary compositional information, includ-ing measurements of the Fe3+/ΣFe of their garnets (Canil and O’Neill 1996; Creighton et al. 2009, 2010; Goncharov et al. 2012; Lazarov et al. 2009; Luth et al. 1990; McCam-mon and Kopylova 2004; Woodland and Koch 2003; Woodland and Peltonen 1999; Yaxley et al. 2012). We restricted our analysis to samples that had Fe3+/ΣFe meas-ured by Mössbauer spectroscopy, the EPMA flank method (Höfer and Brey 2007), or by XANES. Compared with the 2Fe3Fe3+2Si3O12(gt)=4Fe2SiO4(ol)+Fe2Si2O6(opx)+O22Ca3Fe3+2Si3O12(gt)+2Mg3Al2Si3O12(gt)+2Fe2Si2O6(opx)=2Ca3Al2Si3O12(gt)+4Fe2SiO4(ol)+3Mg2Si2O6(opx)+O22Ca3Fe3+2Si3O12(gt)+2Fe3Al2Si3O12(gt)=2Ca3Al2Si3O12(gt)+4Fe2SiO4(ol)+Fe2Si2O6(opx)+O2(Luth et al.1990)Contrib Mineral Petrol (2014) 168:1083 Page 3 of 12 1083 previous calibration (Gudmundsson and Wood 1995), thenew oxybarometer tends to produce more oxidized values,and the differences tend to be larger at higher pressures, aspreviously noted by the authors of the new oxybarometer(Stagno et al. 2013) (Fig. 1). The new oxybarometer alsoproduces a wider range of values at any given pressure.Plotting these data on a pressure (or depth)—Δlogf O2 (FMQ) diagram (Fig. 2) facilitates comparison withreactions that bound the likely oxidation state of the lith-ospheric mantle. On such a diagram, temperature alsoincreases with depth; the reactions on this diagram are cal-culated using the P–T covariance for a 40 mW/m2 geotherm(Hasterok and Chapman 2011). In Fig. 2, the iron–wüstitereaction (IW) can serve as a proxy for the redox conditions at which metal will begin to precipitate from a peridotitic mantle. Once metal begins to precipitate, carbon will dis-solve into the melt as carbide (Dasgupta and Hirschmann 2010; Lord et al. 2009; Rohrbach et al. 2014). Models (Frost and McCammon 2008; Rohrbach et al. 2007, 2011) predict metal saturation at ~250–300 km depth in the man-tle, but the absence of metal in lithospheric mantle samples, and the paucity of samples plotting near the IW curve in Fig. 2, suggests we can use the IW curve as a lower Δlog f O2 bound for likely values for lithospheric mantle.The EMOD and EMOG curves represent the reac-tions Mg2Si2O6+ 2 MgCO3= 2 Mg2SiO4+ 2 C + 2 O2 (Eggler and Baker 1982) involving diamond and graph-ite, respectively. At values of Δlog f O2 above these curves, crystalline carbonate, rather than elemental carbon, would be stable in the presence of olivine. (Stagno and Frost 2010) experimentally determined the conditions at which a carbonate-bearing melt, rather than crystalline carbonate, would provide an upper-f O2 bound to diamond stability; the SF10 curve (Fig. 2) illustrates where a carbonate melt with X(CO2) = 0.5 would become stable. As shown by (Stagno and Frost 2010) and (Stagno et al. 2013), the position of this curve shifts to lower f O2 with decreasing X(CO2) in the melt. The curve shown in Fig. 2 is therefore an upper bound to the stability of diamond relative to carbonate-bearing melt.From the location of these curves (Fig. 2), there is a range of approximately three log units f O2 from the most reduced to the most oxidized possible conditions at which diamond might crystallize in the lithospheric mantle. The observed range of values for natural samples is approxi-mately one log unit f O2 greater because of more oxidized samples from the Slave and Kaapvaal cratons that plot above the EMOD/EMOG and SF10 curves, and in which diamond/graphite would not be stable.Sensitivity of peridotite f O2 to oxygen additionor subtractionThe observed range in values prompted us to question how sensitive the calculated values of Δlog f O2 (FMQ) are toFig. 1 Difference between the values of Δlog f O2 (FMQ) calculated with the Stagno et al. (2013) and Gudmundsson and Wood (1995) oxybarometers (denoted as S13 and GW95, respectively) as a func-tion of pressureFig. 2 Values of Δlog f O2 (FMQ) for garnet peridotites calculated with the Stagno et al. (2013) formulation of the olivine-orthopyrox-ene-garnet oxybarometer as a function of pressure. Data from sam-ples with pressures <2.5 GPa are omitted. Reference reactions are calculated for a 40 mW/m2 cratonic geotherm (Hasterok and Chap-man 2011). IW is the iron–wüstite buffer reaction (Ballhaus et al. 1991). To plot the f O2 of this reaction relative to that of FMQ, the fayalite-magnetite-quartz formulation of Ballhaus et al. (1991) (from O’Neill 1987) was used. Gr–Dia is the graphite = diamond reac-tion and EMOG/EMOD are the enstatite + magnesite = olivine + C (G-graphite, D-diamond) reactions (Eggler and Baker 1982). The Gr–Dia and EMOG/EMOD reactions are calculated from the thermo-dynamic data of Holland and Powell (2011) (see Appendix 1). SF10 curve represents the analogous reaction to EMOG/EMOD involving a carbonate melt rather than crystalline magnesite from Stagno and Frost (2010). See text for data sources for samplesContrib Mineral Petrol (2014) 168:1083 1083Page 4 of 12the compositions of the garnet peridotites, in particulartheir bulk Fe2+/Fe3+, which is intrinsically dependent ontheir modal mineralogy. A simple way to pose this questionis how much oxygen would be required to shift a peridotitefrom IW to EMOD? To address this question, we need tocalculate how the Δlog f O2 (FMQ) of a garnet lherzolite ofa given bulk composition at a given pressure and tempera-ture changes with bulk Fe3+/ΣFe. The latter parameter canbe related to the amount of oxygen that could be supplied by the peridotite interacting with a fluid or melt; in the absence of metallic iron, a sample with a bulk Fe3+/ΣFe of zero has no ability to provide oxygen to a fluid or melt—but the maximum potential to absorb oxygen from that fluid/melt.Our calculations, modeled after those done by (Stagno and Frost 2010) and (Stagno et al. 2013), take into account temperature- and pressure-dependent partitioning of Mg and Fe2+ between the phases, as well as the Al- and Fe3+ partitioning between garnet and pyroxenes. Experimen-tal calibrations (Brey and Köhler 1990; Brey et al. 1990; Harley 1984; Krogh 1988; O’Neill and Wood 1979) and empirical correlations in elemental partitioning observed in our compiled database of garnet peridotites were used to constrain the partitioning behavior.We used this model to calculate the behavior of primi-tive mantle “pyrolite” (McDonough and Sun 1995) as a function of bulk Fe3+/ΣFe along a cratonic geotherm of 40 mW/m2 (Hasterok and Chapman 2011). In agree-ment with previous calculations (Ballhaus and Frost 1994; Creighton et al. 2009, 2010; Frost and McCammon 2008; Stagno and Frost 2010; Stagno et al. 2013; Wood et al. 1996), the Δlog f O2 (FMQ) of primitive mantle lherzolite decreases with increasing pressure along this geotherm by ~0.7 log units f O2 per GPa (Fig. 3). Increasing bulk Fe3+/ΣFe increases the value of Δlog f O2 (FMQ), with the magnitude of the effect decreasing as Fe3+/ΣFe increases, as can been seen by comparing the spacing between the 0.02 and 0.03 curves with that between the 0.06 and 0.07 curves in Fig. 3.The bulk Fe3+/ΣFe for a peridotite will depend on the Fe3+/Fe2+ of the constituent minerals as well as their rela-tive proportions, because the Fe3+/Fe2+ of the minerals differ—with significant amounts of Fe3+ in the pyroxenes and garnet (e.g., Canil and O’Neill 1996; Luth et al. 1990; Woodland 2009) but essentially none in olivine. This means that samples with the same bulk Fe3+/ΣFe may have con-siderably different values of Δlog f O2 (FMQ); conversely, samples with the same value of Δlog f O2 (FMQ) may have quite different bulk Fe3+/ΣFe depending on their modal mineralogy. To explore these relationships in a simplified situation, we varied the modal mineralogy of our primi-tive mantle pyrolite (Fig. 4), and re-calculated the Δlog f O2 (FMQ)—bulk Fe3+/ΣFe relationships for the new bulk compositions. Decreasing the amount of garnet and clinopy-roxene (GT02, GT05) shifts the curves to the left (reducing their bulk Fe3+/ΣFe). This change also steepens the curves,Fig. 3 Relative oxygen fugacities for a primitive mantle bulk com-position (McDonough and Sun 1995) calculated along a 40 mW/m2 geotherm (Hasterok and Chapman 2011) for varying values of bulk Fe3+/ΣFe (given by the numbers at the top of the curves). The iron–wüstite, graphite–diamond, EMOD/G, and carbonate liquid—carbon (SF10) reactions from Fig. 2 are plotted for referenceFig. 4 Values of Δlog f O2 (FMQ) calculated for various bulk com-positions derived from our primitive mantle pyrolite as a function of Fe3+/ΣFe. Pressure and temperature conditions were held fixed at those appropriate for the base of the lithosphere on a 40 mW/m2 geotherm (~200 km depth). The values of the IW and EMOD reac-tions at the same P–T conditions are shown for reference. Samples plotted are the original primitive mantle pyrolite composition (“MS95 pyrolite”—solid circles), the same composition with a higher Mg# of 92 rather than 90 (“HiMg pyrolite”—open circles), and five other compositions derived from the MS95 pyrolite by changing the modal mineralogy. GT05 (upright triangles) has 76 % ol, 14 % opx, 5 % cpx, 5 % gt. GT02 (inverted triangles) has 81 % ol, 15 % opx, 2 % cpx, 2 % gt. GT02LowOl has 55 % ol, 41 % opx, 2 % cpx, 2 % gt. Web1 and Web2 represent websterites, with Web1 (filled diamonds) having 10 % ol, 21 % opx, 35 % cpx, and 34 % gt, and Web2 having 10 % ol, 30 % opx, 50 % cpx, and 10 % gt. See text for discussionContrib Mineral Petrol (2014) 168:1083 Page 5 of 12 1083 meaning a smaller change in bulk Fe3+/ΣFe (i.e., a smallerchange in O2) is required to shift the Δlog f O2 (FMQ) agiven amount. Smaller amounts of garnet and clinopyrox-ene would be expected in a peridotite more depleted thanprimitive mantle. These depleted peridotites, which char-acterize cratonic lithosphere, would have a lesser ability toact as a source or sink for O2 than does primitive mantle. Incontrast, olivine-poor and pyroxene-rich lithologies, mod-eled by “Web1” and “Web2” in Fig. 4, require much largerchanges in Fe3+/ΣFe to effect the same change in Δlog f O2(FMQ), and hence would be much better sources or sinksfor O2, as long as they contain olivine required for the oli-vine–orthopyroxene–garnet reaction to occur.Depleted mantle also differs from the composition ofprimitive mantle in having a higher Mg#. We modeledthe effect of this difference by increasing the Mg# of ouroriginal composition from ~90 to 92. This change by itselfhas a very small influence on the Δlog f O2 (FMQ)—bulkFe3+/ΣFe curve (HiMg pyrolite vs. MS95 pyrolite curvesin Fig. 4). We conclude that the effect of changing modalmineralogy is much more significant than varying the Mg#.To apply this model to natural peridotites, we requiresamples for which the modal mineralogy, and Fe3+/ΣFefor garnet and (ideally) the pyroxenes, is available. Thefirst requirement narrows the field considerably, becausethe number of samples for which modal mineralogy isavailable is remarkably limited—in part because the smallsizes of many samples, combined with relatively coarsegrain size, render estimates of modal mineralogy problem-atic. There have been a number of studies that have meas-ured Fe3+/ΣFe in garnets from mantle-derived peridotitesbecause these data are essential to determining a value ofΔlog f O2 (FMQ) as outlined above. There are only a fewstudies that have determined Fe3+/ΣFe in clinopyroxenefrom garnet peridotites (Canil and O’Neill 1996; Woodland2009), and only one the authors are aware of that deter-mined Fe3+/ΣFe in orthopyroxene from garnet peridotites(Canil and O’Neill 1996). That study demonstrated thatthere is a good correlation between the Fe3+ in orthopyrox-ene and Fe3+ in clinopyroxene in garnet peridotites, a cor-relation that we can use to expand the number of samples towhich to apply this model. As noted by (Woodland 2009),the partitioning behavior of Fe3+ between clinopyroxeneand garnet in peridotites is more complex than previouslymodeled by (Canil and O’Neill 1996). In the absence ofa thermodynamic model for this partitioning, we used thecorrelation we observed between Fe3+/(Al + Cr + Fe3+)in coexisting clinopyroxene and garnet from the data of(Canil and O’Neill 1996) and (Woodland 2009) to infer Fe3+ contents in clinopyroxenes from the measured Fe3+ in the coexisting garnet. These approximations allow us to model the behavior of samples for which we have only gar-net Fe3+/ΣFe and modal mineralogy.Calculated trends in Δlog f O2 (FMQ) versus Fe3+/ΣFe space are shown in Fig. 5 for two samples from the Finsch mine (Lazarov et al. 2009) and one from Premier (PHN 5267; Boyd and Mertzman 1987; Canil and O’Neill 1996). The Premier sample had measured Fe3+/ΣFe for garnet andFig. 5 Δlog f O2 (FMQ) versus Fe3+/ΣFe trends calculated for sam-ples from Finsch and Premier compared with the trend for primi-tive mantle pyrolite from Fig. 4. Data sources: Finsch F-15 and 695 (Lazarov et al. 2009), Premier (PHN 5267) (Boyd and Mertzman 1987; Canil and O’Neill 1996). Pressure–temperature conditions of calculation same as in Fig. 4. The values of the IW and EMOD reac-tions are shown for referenceFig. 6 Amount of O2 required to shift a sample from IW to EMOD, plotted against the modal garnet in the sample. Finsch samples are cpx-bearing peridotites from (Lazarov et al. 2009) (their samples 695, F-1, F-3, F-5, F-6, F-8, F-11, F-12, F-14, F-15, F-16). Calculated at P–T of equilibration for each sample. Two values for MS95 pyrolite are shown; the higher one is calculated at the same conditions as in Figs. 4 and 5, the lower one is calculated at the same conditions as sample 695 from Finsch (5.6 GPa and 1,198 °C). For comparison to the curves in Fig. 5, samples 695 and F-15 are the most garnet-poor and most garnet-rich points, respectively, of the “Finsch sample” on this diagram. See text for discussion of the right-hand axes concern-ing CContrib Mineral Petrol (2014) 168:1083 1083Page 6 of 12both pyroxenes (Canil and O’Neill 1996); the Finsch sam-ples had measured Fe3+/ΣFe for garnet and clinopyroxene (F-15) or just for garnet (695).As expected from the results of our model calculations above, these variably depleted peridotites all have steeper trends than primitive mantle pyrolite, indicating less capa-bility of these samples to act as a source or sink of O2 for diamond-forming reactions. The near-vertical curve for the most depleted sample (Finsch 695, a clinopyroxene-bear-ing garnet harzburgite) reflects the low modal garnet and clinopyroxene in this sample, and shows that very little change in oxygen content would induce a very large change in Δlog f O2 (FMQ).We can now quantify this relationship and answer the question about how much oxygen is required to change the Δlog f O2 (FMQ) of these samples from IW to EMOD. The results of this calculation are plotted in Fig. 6 against the modal amount of garnet for a more extensive set of samples, mostly from the Finsch dataset of (Lazarov et al. 2009). The key point is that for all these samples, less than 200 ppm O2 is required to shift their redox state (relative log f O2) from IW to EMOD. For four of the sam-ples, 50 ppm O2 or less would be required. Obviously, the amount of O2 would be less for a smaller shift in redox state (Table 1).The stoichiometry of the diamond-forming reactions outlined in the Introduction can be used to calculate the amount of elemental carbon that would be produced by reaction with these amounts of O2. The right-hand axes in Fig. 6 provide these values in both ppm C and car-ats C per metric ton. To place these values in context, the concentration of diamond in diamond-bearing perido-tite xenoliths can be as high as 27,700 carats/metric ton (Viljoen et al. 2004), and garnet pyroxenite layers in the Beni Bousera peridotite massif contain up to 15 vol% gra-phitized diamond (Pearson et al. 1989), equivalent to about 800,000 carats per ton.Evolution of CHO fluidsIn order to compare these results with the redox capabilities of fluids or melts, thermodynamic models of the composi-tional evolution of these fluids or melts as a function of P, T, and f O2 are needed. Such a model for melts, especially carbonatitic melts at these pressures and temperatures, is lacking. There are models available for C–H–O fluids (Hui-zenga 2001, 2005; Zhang and Duan 2009, 2010), however, which provide a starting point for assessing the behavior of fluids. As discussed previously (e.g., Stachel and Har-ris 2009), C–H–O fluids may be responsible for diamond formation in harzburgites, which have higher hydrous solidus temperatures relative to lherzolites or eclogites. If we focus on such fluids, the first question is what are the relevant compositions as functions of pressure, tempera-ture, and oxidation state? This question has been exten-sively addressed previously in the literature (e.g., Frost and McCammon 2008; Huizenga et al. 2012; Zhang and Duan 2009), but we can now compare the changes in oxy-gen content of this fluid as it changes its oxidation state to the ability of lithospheric peridotite to supply or absorb this oxygen.In the diamond stability field, the C–H–O ternary con-tains a large field of diamond plus fluid; the composition of the C-saturated fluid is given by the red curve in Fig. 7. To a first approximation, the diamond-saturated fluids canTable 1 Calculated oxygen required to shift value of Δlog f O2 (FMQ) from IW to EMOD for primitive mantle “pyrolite” and samples from southern AfricaSample Pressure (GPa)Temperature (°C)Modal garnet (wt%)ppm O2 (IW–EMOD)ppm O2 per unit Δlog f O2Pyrolite 6.41,40317.2422141Pyrolite 5.61,19817.2377117Finsch 695 5.61,1980.83310Finsch F-1 5.71,20910.414846Finsch F-3 5.31,184 3.83310Finsch F-5 5.31,154 3.97022Finsch F-6 5.41,2077.412238Finsch F-8 5.61,200 3.33110Finsch F-11 5.51,225 4.07223Finsch F-12 5.91,198 6.912238Finsch F-14 5.61,191 3.15015Finsch F-15 5.51,19512.214646Finsch F-16 6.01,2109.318156PHN 5267 6.41,415 6.913646Contrib Mineral Petrol (2014) 168:1083 Page 7 of 12 1083be separated into CH4–H2O fluids under reducing condi-tions and H2O–CO2 fluids under oxidizing conditions. Plot-ting the location of the IW buffer and the EMOD reaction on the C-saturated fluid curve (Fig. 7) provides the same bounds for possible redox conditions that we used in the previous section. At IW, the fluid would be dominantly CH4 and H2O, and the fluid becomes more H2O-rich with increasing f O2, becoming an almost pure H2O fluid with minor CO2 and negligible CH4 at EMOD.The shape of the C-saturation curve constrains the abil-ity of fluids to precipitate diamond by interacting with mantle of differing oxidation state. Although very CO2-rich fluids could be stable at oxidizing conditions, fluids more CO2-rich than EMOD would precipitate magnesite by reacting with olivine in peridotitic mantle. Therefore, the EMOD point on the C-saturation curve is the most oxi-dized C-saturated fluid that can be present in peridotitic mantle. EMOD is slightly more CO2-rich and higher in f O2 than the point on the curve at which the C concentration of the fluid is lowest (which is approximately the “water-maximum” fluid composition (“m” in Fig. 8). It is only between EMOD and this point on the C-saturation surface that oxidized fluids would precipitate diamond by reacting with reduced mantle. For example, at the P, T conditions illustrated in Figs. 7 and 8, reduction of an oxidized fluid starting at EMOD would precipitate 0.27 wt% diamond while liberating 2.2 wt% O2. This oxygen would have to oxidize ferrous iron in the surrounding mantle—it cannot merely oxidize diamond, because that would produce CO2 and make a more oxidizing fluid. Further reduction of this fluid would dissolve diamond as the fluid becomes more C-rich. This evolution would require more O2 to be trans-ferred to the surrounding mantle. If that mantle could not absorb this amount of O2, the fluid could not change com-position and evolve to more reducing conditions. The total amount of O2 liberated as the fluid evolves from EMOD to the actual Δlog f O2 (FMQ) value for Finsch 695 would be 6.44 wt% O2 (g O2/100 g original fluid). That amount of O2 would oxidize 8.17 × 105 g of peridotite of Finsch 695 composition from IW to the measured value of Δlog f O2 (FMQ) for sample 695. Looked at a different way, this ratio corresponds to 26.6 g fluid per 106 g peridotite or 26.6 ppm fluid. If the peridotite was initially more oxidized than IW, even larger volumes of peridotite would have to be involved.If a fluid is liberated by exsolution from a crystallizing magma, it is likely to be a H2O-rich fluid with subordinate CO2, given the higher relative solubilities of H2O and CO2 compared with reduced species (e.g., Ardia et al. 2013; Foley 2011). This H2O + CO2 fluid would be undersatu-rated in diamond and would precipitate magnesite by react-ing with olivine if it is too CO2-rich. For such a fluid to reach the C-saturation surface at EMOD (to maximize the possible amount of diamond it could precipitate by further evolution), it would have to oxidize the surrounding mantle; it would need to get rid of 1.2 wt% O2 to reach the C-satu-ration surface from its original binary H2O + CO2 compo-sition. Once it has reached the C-saturation surface, con-tinued evolution of this fluid would precipitate 0.27 wt% diamond while producing an additional 2.2 wt% O2, asFig. 7 Molar C–H–O ternary diagrams illustrating the composition of fluids (red curves) in equilibrium with diamond at the P–T con-ditions of Finsch 695 (5.6 GPa, 1,198 °C). In the bottom figure, the numbers along the curve give the value of Δlog f O2 (FMQ) for the fluid. The top figure shows the location of fluids at Δlog f O2 (FMQ) corresponding to IW and EMOD. Calculated with GFluids (Zhang and Duan 2009, 2010). The Δlog f O2 (FMQ) for Finsch sample 695 (−3.34) is shown by the blue circle. See text for discussionFig. 8 Mole fraction of species in a CHO fluid in equilibrium with diamond at the same P, T as in Fig. 7. The values of Δlog f O2 (FMQ) for IW, Finsch sample 695, SF10 (S), and EMOD (E) are shown for reference. The “water maximum” is indicated by the diamond labeled “m”. Calculated with GFluids (Zhang and Duan 2009, 2010)。