Open Access

Natural moissanite (SiC) – a low temperature mineral formed from highly fractionated ultra-reducing COH-fluids

  • Max W Schmidt1Email author,
  • Changgui Gao1, 2, 4,
  • Anastasia Golubkova1,
  • Arno Rohrbach3 and
  • James AD Connolly1
Progress in Earth and Planetary Science20141:27

https://doi.org/10.1186/s40645-014-0027-0

Received: 15 September 2014

Accepted: 30 October 2014

Published: 13 December 2014

Abstract

Natural moissanite (SiC) is reported from dozens of localities, most commonly from ultramafic rocks where it may be associated with diamond and iron silicides. Yet, formation conditions of moissanite remain in the realm of speculation. The key property of SiC is its extremely reduced nature. We have experimentally equilibrated SiC with olivine and orthopyroxene at 1300-1700°C, 2 and 10 GPa, to determine the oxygen fugacity of the C + orthopyroxene = SiC + olivine + O2 buffer (MOOC) and the equilibrium X Mg of coexisting mantle silicates. The experiments resulted in olivine and orthopyroxene with X Mg of 0.993-0.998 in equilibrium with SiC and iron silicides. Calculated oxygen fugacities are 5-6.5 log units below the iron-wustite (IW) buffer at 2-10 GPa. The experimental results concur with calculated phase relations for harzburgitic mantle under reducing conditions that include metal alloys, carbides and silicides. The extremely reducing character of SiC precludes coexistence with silicates with appreciable Fe2+, and hence excludes equilibrium with mantle phases with typical X Mg ’s of ~0.9. Calculated Fe-Mg diffusion lengths reveal that SiC grains of 1 mm would react with the Fe-component of olivine to iron carbide or metal and orthopyroxene within <1 Ma at temperatures above 800°C. We thus conclude that SiC forms through a low-temperature process (<700-800°C) where equilibrium is only reached at the grain scale. The most plausible formation mechanism is a strong fractionation of a C-O-H fluid from metamorphosed sediments originally rich in organic material. Such a fluid is initially saturated with graphite or diamond and is slightly more reduced than the H2O-C join. Fluid percolation in the mantle leads to H2O-sequestration by crystallizing hydrous phases (most likely serpentine, brucite or phase A), and hence O2-removal from the fluid causing its reduction and continuous graphite or diamond precipitation. A small, highly fractionated fluid fraction may then reach CH4 and H2 concentrations that allow SiC formation on grain boundaries without equilibration with the bulk rock on a larger scale. Such a mechanism is corroborated by the strongly negative δ13C of moissanites (-20 to -37), consistent with reduced fluids originating from metamorphosed organic carbon.

Keywords

MoissaniteSiCUltra-reducingIron silicides and carbidesCOH-fluidFluid-fractionation

Background

Moissanite is a refractory mineral stable at extremely reducing conditions with respect to typical Earth environments. Natural occurrences are manifold (Lyakhovich, [1980]; Trumbull et al. [2009]) but many fall into three categories: (i) peridotites, serpentinites (Xu et al., [2008]) and podiform chromitites (Bai et al., [2000]); (ii) kimberlites (Leung et al., [1990]; Shiryaev et al., [2011]); and (iii) metasomatic rocks (Lyakhovich, [1980]; Di Pierro et al., [2003]). In most cases, moissanite is recovered from heavy mineral separates with little textural and phase assemblage information available. Often Fe-silicides or metallic Si are attached to SiC grains (Trumbull et al. [2009]), there are also diamonds which contain moissanite inclusions (Gurney [1989]; Leung, [1990]).

The genesis and crystallization conditions of natural SiC remains unclear, nevertheless, agreement exists about the extremely reduced nature of moissanite. This property has led to hypotheses such as moissanite being a remnant from a primordial ultra-reduced mantle or from the core-mantle boundary as discussed by Mathez et al. ([1995]). The extreme oxygen fugacities required for moissanite formation have also led to propose extremely high temperatures and/or pressures of formation (Trumbull et al., [2009]), in part motivated by a not uncommon association with diamonds, in part probably motivated by the high temperature synthesis practiced for industrial SiC. As conditions in the upper mantle would not allow moissanite to be stabilized at the predominant oxygen fugacities near the iron-wustite (IW) reference equilibrium (Ballhaus [1995], Woodland and Koch [2003], Frost and McCammon [2008]), moissanite has been hypothesized to stem from the lower mantle (Mathez et al., [1995]).

The most basic fact pertinent to moissanite genesis is that the quartz + C = SiC + O2 equilibrium is at oxygen fugacities ≥5-7 log units below those of the iron-wustite buffer (Mathez et al. [1995]; Ulmer et al., [1998]). At such low oxygen fugacities, most Fe2+ should be reduced to metal. Consequently, coexisting mantle phases such as olivine and orthopyroxene should have unusually high X Mg values.

Inspiration for the low temperature hypothesis which forms the starting point of this study stems from a moissanite location in a Cenozoic basalt from North China, where SiC appears in mantle xenoliths infiltrated by carbonatites (Gao and Liu, [2008]); these composite xenoliths are in turn erupted by alkali basalts. Clinopyroxene, orthopyroxene and olivine in these xenoliths have average mantle X Mg values near 0.90 and abundant interstitial calcite with no indication for a reduced environment. Nevertheless, moissanite and a panoply of metals (Fe, FeNiSi-alloy, Ti) and Fe-silicides occur in these xenoliths but are constrained to small voids. The close spatial occurrence of an assemblage (i.e. calcite + mantle silicates) requiring oxygen fugacities close to the CCO buffer (graphite/diamond-CO-CO2, i.e.) with another assemblage requiring oxygen fugacities 10 log-units below this buffer motivates the formulation of a low-temperature hypothesis. In this model, the SiC in the voids would be precipitated from an ultra-reduced highly fractionated fluid that is not in equilibrium with the adjacent silicate and carbonate phases of the xenolith itself.

In this study, we demonstrate experimentally that olivine and orthopyroxene in equilibrium with SiC are almost purely magnesian, containing only 2-7‰ of their Fe-endmembers. Consequently, close spatial association of moissanite and silicates with X Mg ≤0.9 demonstrates chemical disequilibrium and indicates temperatures lower than those allowing for diffusional equilibration. Secondly, we combine thermodynamic data for Si- and Fe-carbides, stoichiometric FeSi-compounds and FeSiC alloys with silicate data to compute P-T-f O2 phase relations, phase compositions and hence the stability of moissanite for an Fe-Mg-Si-O-C peridotitic bulk composition.

Methods

High pressure experiments

One set of experiments reacted San Carlos olivine (X Mg  = 0.90) and a natural orthopyroxene from the Urals (X Mg  = 0.91) with SiC at 2 GPa, 1300 and 1500°C; and at 10 GPa, 1500 and 1700°C (Table 1). The 2 GPa experiments were performed in an endloaded, 14 mm bore diameter piston cylinder employing a talc–pyrex–graphite-MgO assemblage. The 10 GPa experiments were performed in a multi anvil employing 18 mm Cr2O3-doped MgO-octahedra on a 11 mm truncation. The multi-anvil assemblage contained a zirconia insulator, a stepped LaCrO3 furnace, inner MgO pieces, and a Mo ring and disc between the furnace and WC cube truncation faces. In both apparatus B-type PtRh thermocouples were used. In experiment G#1 we intended to monitor oxygen fugacity by using an Ir sensor (Woodland and O’Neill, [1997]), but a ternary Fe-Si-Ir alloy resulted for which thermodynamic data are not available.
Table 1

Experimental phase compositions and calculated oxygen fugacity

 

P [GPa]

T [°C]

t [h]

X Mg olivine

X Mg opx

Metal or silicide [wt%]

Δ log f O 2 I W SiC-C-ol-opx

      

Fe

Ni

Si

Ir

  

G#1

2

1300

48

---*

0.982(5)

8.1(26)

0.24(9)

59.4(43)

20.3(20)

+quartz

---

G#2

2

1500

15

0.997(1)

0.998(1)

77.4(13)

3.3(4)

18.0(4)

---

Fe2Si

-6.4

G#3

10

1500

72

---*

0.995(1)

60.3(10)

2.7(8)

35.0(17)

---

FeSi

(-4.9)**

      

46.6(21)

0.9(5)

49.5(16)

---

FeSi2

 
      

30.8(6)

1.3(2)

64.4(20)

---

FeSi4

 

G#6

10

1700

54

0.993(2)

0.995(1)

75.6(11)

2.6(9)

20.8(4)

---

Fe2Si

-5.6

*Rims or silicide-free exsolution zones of previously San Carlos olivine too small to measure by EMPA.

**From equilibrium (1), the FeSi is most likely formed metastably in the early stage of the experiment.

A second set of experiments intended to synthesize SiC first at low temperature from tetrakis-silane (((CH3)3Si)4Si) or stearic acid (C18H36O2) and olivine + serpentine or talc + MgO at 500°C, 0.2 GPa. These experiments did not result in SiC or equilibrium assemblages. We then aimed at synthesizing SiC + magnesio-silicates at 1000-1600°C, 2-10 GPa employing talc + MgO and tetrakis-silane as starting materials. At 6 GPa, 1300°C we obtained SiC + olivine + opx, but the assemblage was unequilibrated with strong compositional zoning in the silicates and further experiments in this direction were abandoned.

Thermodynamic data

Thermodynamic databases suited for calculating phase equilibria in Earth sciences mostly contain only silicates and oxides. In order to perform calculations in mantle materials at ultra-reduced conditions, alloys, carbides and silicides were combined with silicate data from Holland and Powell ([2011]). The data sources employed were: Cementite (Fe3C): calorimetric data from Gustafson ([1985]), volume data from Li et al. ([2002]), Wood et al. ([2004]); Fe7C3: Djurovic et al. ([2011]), Mookherjee et al. ([2011]), Nakajima et al. ([2011]), Moissanite (SiC): Lacaze and Sundman ([1991]), Miettinen ([1998]), Aleksandrov et al ([1989]), Li and Bradt ([1987]). The FeSiC-alloy and the iron silicides are based on the 1-atm phase diagram by Lacaze and Sundman ([1991]) and on the volume EoS of Brosh et al. ([2009]). Details of the adoption of the ternary metal, carbides and silicides are given in a forthcoming paper (Golubkova et al. [2014]). Phase relations were computed by free energy minimization with the Perple_X software package (Connolly JAD [2009]).

Results

Phase equilibria and phase compositions

Experiments at 2 GPa, 1500°C (G#2) and 10 GPa, 1700°C (G#6) resulted in equilibrated olivine + orthopyroxene + SiC plus iron silicides (Table 1, Figure 1). The experiment at 10 GPa, 1500°C (G#3) resulted in coarse opx + SiC plus three iron silicides (FeSi, FeSi2, and FeSi4) and olivine that did not develop equilibrium rims large enough for measurement. These charges did not contain any residual silicates from the starting material, but in both the 1500°C and 1700°C experiments at 10 GPa, most of the San Carlos olivine disintegrated to an almost Fe-free olivine containing many submicron iron silicide inclusions (Figure 1b). These inclusions are too small to determine their stoichiometry. The equilibrium X Mg of olivine and opx in these experiments are 0.993-0.998 (Table 1). Initially, we added Ir metal as internal oxygen fugacity monitor (Woodland and O’Neill, [1997]; Stagno and Frost, [2010]), this experiment (G#1) at 1300°C, 2 GPa yielded opx (X Mg = 0.982) and quartz. However, abundant residual starting material San Carlos olivine testifies for incomplete equilibration questioning whether the X Mg of the opx of this experiment corresponds to equilibrium with SiC. We have thus only used experiments #2, #3 and #6 for oxygen fugacity calculations.
Figure 1

BSE images of run products from (a) experiment G#2 (1500°C, 2 GPa) and (b) experiment G#6 (1700°C, 10 GPa). Olivine and orthopyroxene have XMg > 0.99 and cannot be well distinguished from each other. Note the graphite flakes at 2 GPa. At 10 GPa diamonds result mostly in plug-outs during preparation and are hence rarely visible. Note that in (b) the original San Carlos olivine forms almost pure forsterite with many Fe-rich inclusions which stoichiometry cannot be determined due to their small size.

Oxygen fugacity (f O2 ) in the experiments

For the experiments, which yielded iron silicides and equilibrium compositions of olivine and orthopyroxene, oxygen fugacities were calculated from equilibrium (1) and (2), employing the olivine and orthopyroxene activity formulations from O’Neill and Wall ([1987]).
1 ferrosilite = 2 FeSi + 3 O 2
(1)
1 fayalite = 1 F e 2 S i + 2 O 2
(2)
At equilibrium,
0 = 2 G FeSi P , T + 3 G O 2 P , T G fs P , T R T ln a O 2 3 a fs opx ,
(3)
FeSi being a pure phase and
0 = G F e 2 S i P , T + 2 G O 2 P , T G fay P , T + R T ln a O 2 2 a fay olivine ,
(4)
Fe2Si being a pure phase. This leads to
log f O 2 = 2 G FeSi P , T + 3 G O 2 P , T G fs P , T 2.3025 R T + log a fs o p x / 3 ,
(5)
using the measured opx composition, log f O 2 = − 11.2 (5) or Δ log f O 2 I W = 4.9 results for G#3 at 10 GPa, 1500°C. For the Fe2Si equilibrium (4)
log f O 2 = G F e 2 S i P , T + 2 G O 2 P , T G fay P , T 2.3025 R T + log a fay olivine / 2 ,
(6)
measured olivine compositions yield log f O 2 = − 13.3 (10a, 10b) or Δ log f O 2 I W = 4.6 for G#2 at 2 GPa, 1500°C; and log f O 2 = − 11.9 (4) or Δ log f O 2 I W = 5.8 for G#6 at 10 GPa, 1700°C. For the two experiments G#2 and G#6, which resulted in equilibrated olivine and orthopyroxene, oxygen fugacities were also calculated from
1 ferrosilite + 1 graphite / diamond = 1 fayalite + 1 SiC + O 2
(7)
At equilibrium
0 = 1 G SiC P , T + 1 G O 2 P , T + 1 G fay P , T 1 G fs P , T 1 G graph / diam P , T + R T ln a O 2 a fay olivine a fs o p x
(8)
and
log f O 2 = G fs P , T + G graph / diam P , T G fay P , T G SiC P , T G O 2 P , T 2.3025 R T log a fay olivine + log a fs o p x .
(9)

SiC and graphite or diamond being pure phases. The resulting oxygen fugacities are log f O 2 = − 15.1. (6) or Δ log f O 2 I W = 6.4 . at 2 GPa, 1500°C (G#2), and log f O 2 = − 10.6 (4) or Δ log f O 2 I W = 5.6 at 10 GPa, 1700°C (G#2).

Calculation of T-f O2 sections

Reactions governing reduced phases in the mantle

To understand the succession of phase assemblages with decreasing oxygen fugacity, we calculated f O 2-temperature diagrams (Figure 2) and phase compositions (Figures 3 and 4) for a model harzburgite at 2 and 10 GPa. For reference, we also give the quartz-fayalite-magnetite (QFM), the graphite/diamond-CO2-CO equilibrium (CCO), and the iron-wustite buffer (IW), even if these do not occur as reactions in the harzburgite phase diagram. The position of the iron-wustite buffer is calculated from Campbell et al. ([2009]), which yields about half a log-unit lower values than the formulation of O'Neill ([1988]).
Figure 2

Calculated f O 2 -temperature diagrams for carbon-bearing harzburgitic peridotite simplified in Mg-Fe-Si-C-O 2 at (a) 2 GPa and (b) 10 GPa. The saturation hierarchy at moderately oxidizing conditions is chosen as olivine > orthopyroxene > graphite/diamond > magnetite. The FMQ, CCO and IW buffers are given for reference. Squares are calculated oxygen fugacities for the experiments, red fill for f O2 ’s calculated from the iron silicide involving equilibria (1) and (2), black fill for f O2 ’s calculated from the SiC-involving reaction (7).

Figure 3

Isopleths of olivine composition in terms of X Mg in the same harzburgitic bulk composition as in Figure 2 for (a) 2 GPa and (b) 10 GPa. Within two log-units of the metal forming reaction, X Mg olivine increases to 0.99, when reaching the SiC-buffer, olivines are essentially Fe-free. X Mg olivine is almost constant in the stability fields of magnesite or graphite/diamond. The grey field corresponds to olivine + orthopyroxene saturation and only additional phases are indicated.

Figure 4

Isothermal composition – f O2 sections at (a) 2 GPa, 1100°C and (b) 10 GPa, 1300°C. At oxygen fugacities below the carbide- or metal–forming reaction, X Fe olivine drops rapidly to per mill values. The grey field corresponds to olivine + orthopyroxene saturation and only additional phases are indicated. With decreasing oxygen fugacity the FeSiC-alloy becomes more and more Si-rich, leading to the fcc-bcc phase transition and then to iron silicide stability. Note that the Si-alloy formed orthopyroxene and olivine breakdown is essentially pure Si with 1 mol% C and almost no iron. The calculations are done at discrete f O2 -values represented by symbols; green: X Mg olivine , black: X Mg opx , red: C content in metal alloy, blue: Si-content in metal alloy.

In the carbon-bearing upper mantle, the iron-wustite equilibrium does not occur, its equivalent leading to the appearance of reduced Fe0 are the iron carbide forming reactions
6 olivine + 2 graphite = 3 o p x + 2 cementite F e 3 C + 3 O 2
(10a)
and
14 olivine + 3 diamond = 7 o p x + 2 F e 7 C 3 + 7 O 2
(10b)
which are only stable at relatively low temperatures, i.e. 1140 to 1260°C at 2 to 10 GPa. At higher temperatures, carbides do not form but a metal alloy results from
2 olivine + 2 x graphite / diamond = 1 o p x + 2 + 2 x FeSiC alloy + 1 O 2
(11)
where FeSiCalloy is a ternary metal solid solution. The stoichiometry of reaction (11) depends on the alloy composition. The molar fraction of C in the alloy ranges from 0.13 to 0.33 at 2-10 GPa, 1150-1700°C. At oxygen fugacities of initial metal formation, Si-fractions are <0.1 mol% in the alloy. For a graphite/diamond saturated mantle with a bulk X Mg of 0.9, reactions (10a, 10b) and (11) occur at 2 GPa about one log-unit below the iron-wustite buffer and about one log-unit above IW at 10 GPa (Figure 2a,b). About 5-9 log-units below the carbide-forming reactions, increasing chemical potential of Si leads to the replacement of Fe3C or Fe7C3 by the FeSiC-alloy then containing significant Si (Figure 5b,d). Silicon contents in the alloy increase with decreasing oxygen fugacity leading to a phase transition from a face centered to a body centered metal structure (fcc and bcc in Figure 5).
Figure 5

Calculated isopleths of metal composition for the ternary FeSiC alloys at 2 and 10 GPa. (a) and (c) C-content in mol-fraction; (b) and (d) Si-content in mol-fraction. Note the phase transition from the fcc to bcc alloy structure with increasing Si-content when decreasing oxygen fugacity. At the metal-forming reaction, C-contents increase with temperature and decrease with pressure, Si contents are below the lowest increment of the calculations. At 2 GPa, calculated C- and Fe-contents of the Si-alloy stable at ultra-reducing conditions are <1 mol%. Grey fields correspond to olivine+opx saturation.

Silica carbide forms principally through a reaction involving moissanite, olivine, orthopyroxene and elemental carbon as graphite or diamond (moissanite-olivine-opx-carbon: MOOC)
1 o p x + 1 graphite / diamond = 1 olivine + 1 SiC + 1 O 2
(12)

at 2 GPa at 7.7-4.8 log-units (1100-1700°C) below equilibria (10a, 10b) and (11). At 10 GPa this difference is 9.9-6.5 log units (1100-1700°C). This corresponds to a Δlogf O2 of IW-9 to IW-6, the difference decreasing with temperature. Along a mantle adiabat, moissanite forms at IW-8.0 at 2 GPa and at IW-6.7 at 10 GPa. Reaction (12) was previously identified as being responsible for the appearance of moissanite in the upper mantle (Mathez et al. [1995]; Ulmer et al. [1998]).

Further reduction causes the FeSiC-alloy to be replaced by stoichiometric iron silicide FeSi. At 2 GPa this happens about ½ a log-unit below the SiC-buffer (12), while at 10 GPa, the FeSi-forming reaction intersects reaction (12) leading to an invariant point. At about 3 and 5 log-units below the SiC-forming reaction (12), opx and then olivine (at these conditions pure enstatite and forsterite, respectively) become unstable and a ternary metal alloy of almost pure Si forms through
1 o p x + 0.01 SiC = 1 olivine + 1.01 S i alloy + 1 O 2
(13)
1 olivine + 0.01 SiC = 1 periclase + 1.01 S i alloy + 1 O 2
(14)

where the Si-alloy has 1 mol% C. Completion of reactions (13) and (14) would leave no oxidized Si. The calculations do not indicate any Fe in the Si metal and only 0.2 to 1.3% C. Reaction (13) is relevant in nature as pure Si-metal attached to SiC has been reported by Trumbull et al. ([2009]).

Summarizing, a harzburgitic mantle would have the following succession of reduced phases: graphite/diamond → (Fe3C/Fe7C3) → Fe-rich FeSiCalloy → SiC → FeSi → Sialloy. At pressures ≥10 GPa, FeSi may form before moissanite.

Note that our calculations do not include liquids and that at high temperatures FeSiCalloy and FeSi become metastable with respect to liquid. This is of particular relevance at oxygen fugacities just below the metal-forming reaction (11) where a de facto binary FeC-alloy is almost Si-free. The Fe-C eutectic is located at 1150°C, 1 atm (Chipman [1972]) and 1210°C, 10 GPa (Hirayama et al. [1993]; Rohrbach et al. [2014]), although this location is not entirely unambiguous as Lord et al. ([2009]) report the eutectic at 10 GPa at 1420°C. In a graphite/diamond-saturated system, the ternary FeSiC-alloy would then be replaced by a metallic melt phase. FeSi has a temperature of congruent melting of 1410°C at 1 atm (Lacaze and Sundman [1991]), but to our knowledge its high pressure melting is unknown.

Mineral compositions at reduced conditions

The above reactions are discontinuous reactions that delineate the various stability fields. For the silicate mantle the continuous reactions consuming the fayalite and ferrosilite components in olivine and orthopyroxene are more dramatic. As illustrated in Figures 3 and 4, X Mg in olivine is almost invariant in each field of coexistence with magnesite or graphite/diamond. Immediately below the metal forming reaction, Fe2+ becomes strongly reduced, within only 2 log-units, the X Mg of olivine (and similarly opx) in equilibrium with metal or Fe3C/Fe7C3 and graphite/diamond increases to ~0.99 (Figures 3 and 4). When oxygen contents corresponding to the oxygen fugacities of SiC are reached, the equilibrium X Mg in the silicates is >0.999, almost invariant with pressure and temperature (at least for 1100-1700°C, 2-10 GPa). This confirms the experimental findings. Concomitant with the increase in X Mg of the silicates, an increase of the Si-content of the FeSiC-alloy is predicted (Figures 4 and 5), mirrored by a decrease of X Fe metal from ~0.90 to 0.70 at 2 GPa and from 0.76 to 0.66 at 10 GPa.

Silica carbide stability with temperature and pressure

A major question concerning the stability of SiC in the mantle regards the evolution of the difference in oxygen fugacity between the metal or Fe-carbide forming (10a, 10b, 11) and the SiC-forming reactions (12). This difference, Δlogf O 2 (12)-(10a,10b,11), decreases with increasing temperature but at near-adiabatic temperatures only by -0.1 log-units per 100°C. With increasing pressure there is a slight increase of Δlogf O 2 between the moissanite forming reaction and the equivalent of IW in the mantle such that along a mantle adiabat Δlogf O 2 (12)-(10a,10b,11) remains almost constant at ~ -4.5.

Discussion

Oxygen fugacities calculated from the experiments

In Figure 2a and b we plot the oxygen fugacities calculated for the three experiments with olivine and/or orthopyroxene equilibrated with SiC. The oxygen fugacities calculated from the experiments through the moissanite-forming equilibrium (7) are in perfect agreement with the calculated phase diagram. They are also in good agreement with the values derived from Ulmer et al. ([1998]) from experiments on a Mg-Si-C-O system that cross-compared equilibrium (12) at 1500°C to other common oxygen buffers.

Oxygen fugacities from the iron silicide involving equilibria (red symbols) were calculated without involving SiC and may hence be taken as independent confirmation of the f O2 -value of the SiC-forming reaction. Nevertheless, the interpretation of the silicides in the two experiments at 1500°C is not straightforward. Congruent melting of Fe2Si occurs at 1212°C at 1 atm, but the dP/dT slope of this reaction is unknown. It is thus well possible that the Fe2Si in the 2 GPa experiment G#2 crystallized upon quench from a FeSiC-liquid. At 10 GPa, the silicides should be stable as solids. Nevertheless, the experiment at 1500°C (G#3) contains three silicides, FeSi, FeSi2 and FeSi4. Together with opx, diamond and SiC, these are six phases, one phase too much for a divariant field of a five component system (Mg-Fe-Si-C-O), but possible if Ni is also accounted for as a component. Nevertheless, one of the silicides may well form in a transient stage, when the FeO of the fayalite component is reduced. We interpret the f O2 -value calculated from equilibrium (1) as reflecting a transient stage at the beginning of the experiment.

X Mg of silicates coexisting with SiC

The results from the experiments in terms of X Mg of silicate minerals coexisting with SiC agree very well with the calculated phase compositions (Figures 3 and 4). Together with the constraints on oxygen fugacity, a coherent picture forms: moissanite can only be expected in a mantle or any rock with silicates essentially devoid of Fe2+. This unsurprising finding has strong implications for the origin of SiC in mantle rocks. It demonstrates that any significantly large domain of mantle at oxygen fugacities of SiC-stability (i.e. void of Fe2+ in silicates and oxides) has not yet been identified.

Both the experiments and the calculations demonstrate that for plausible mantle compositions with SiC stable, the X Mg of coexisting olivine or orthopyroxene is >0.993. Hence SiC cannot be in equilibrium with typical mantle that has olivine and pyroxenes with X Mg of 0.88-0.92. The experiments only constrain the X Mg of olivine and orthopyroxene in equilibrium with SiC. Nevertheless, in nature, X Mg of serpentine is generally higher than that of coexisting orthopyroxene, which is similar to that of clinopyroxene and X Mg of olivine is lowest among these phases. One can thus safely assume that for all silicates characteristic for peridotites (except of garnet) similarly high X Mg of >0.993 are to be expected.

As SiC is found in mantle rocks and kimberlites from which anomalous X Mg values for the silicate phases are not reported, the formation mechanism of SiC should explain how average X Mg mantle silicates and SiC persist in the same rock volume.

Diffusion limits on temperatures of SiC preservation

Considering a SiC grain within a matrix of FeMg-silicates, one can calculate the time scale for diffusive equilibration between SiC and the FeMg-silicates. With time, SiC would react with the fayalite component of olivine to form opx and metal or carbide until the SiC grain is completely consumed. Taking a putative SiC grain in a matrix of otherwise undisturbed mantle silicates, Fe-Mg diffusion coefficients of the silicates allow for calculation of the time-temperature conditions for which a given grain size of SiC would react with olivine until its complete destruction. Assuming spherical symmetry, the amount of Fe2+ needed to oxidize all Si of a SiC grain in an olivine (Fo90) matrix occurs within 2.3 radii. Fe-Mg volume diffusion in olivine and orthopyroxene are characterized by diffusivities of D0 = 10-7 m2/s and an activation energy E A of 250 kJ/mol (Brady and Cherniak, [2010]). We have calculated the relation between time and temperature for a given grain size of radius r employing the characteristic diffusion length 2.3r. This calculation yields the time necessary to diffuse a sufficient amount of Fe2+ towards a SiC grain, then being consumed and forming orthopyroxene (and metal), for a given temperature and grain size (Figure 6). These calculations show that mm-sized grains would disappear through diffusive equilibration within a million years at temperatures above 800-900°C. Preservation of SiC in a mantle mineral matrix with XMg = 0.9 would not be possible at adiabatic mantle temperatures, even SiC grains of 3-10 cm would only last for 10’000 years. This conclusion remains robust when employing the experimentally determined range of D 0 E A pairs (for summary of diffusion data see Brady and Cherniak, [2010]).
Figure 6

Illustration of the diffusion time – temperature – grain size relation for SiC in an olivine matrix with X Mg = 0.9. For a given grain size, this graph illustrates the time of SiC decomposition through diffusive equilibration with the Fe-component in olivine, effectively delimiting the temperature at which SiC could reside in a mantle with X Mg  = 0.9. For details see text.

These time spans could be prolonged if SiC forms in veins of essentially Fe-free silicates, but such mineral compositions are not reported for peridotites or kimberlites. Regarding the putative high-temperature origin of SiC, at temperatures of the asthenospheric mantle, SiC grains of initially >50 cm would be necessary to persist for >1 million years. It is thus unlikely that SiC would survive any period of sustained high temperatures in the mantle.

A model for SiC formation in nature

Terrestrial environments are typically too oxidizing to permit SiC formation, thus exceptional circumstances need to be explored. One plausible scenario starts with a slightly reduced graphite/diamond-saturated COH-fluid that evolves towards an ultra-reduced composition through the formation of first carbonates or spinel solid solution and then hydroxide minerals, all from anhydrous silicates. In this model, the fluids react with its immediate surroundings while percolating upwards, leaving newly formed carbonates, spinel or hydroxides behind. The fluid-mediated formation of magnetite component in spinel from mantle silicates drives any fluid to more reducing compositions (away from the O2-corner in Figure 7). The fluid-mediated formation of carbonates (from mantle silicates) leads to more reduced residual fluids when the starting COH-fluid has a composition to the reduced side of the CO2-H2O tie-line (Figure 7). Fluids from sediments rich in organic material are expected to already have H:O-ratios slightly to the H-rich side of the H2O-C tie-line (Connolly JAD [1995]; Figure 7c). Nevertheless, as long as they are saturated in graphite, also fluids slightly to the O-rich side of H2O could be suitable precursors (Figure 7b). The following steps would then lead to a fluid that may mediate SiC formation:
Figure 7

COH-diagrams for fluids illustrating the graphite saturation surface (a) and with calculated graphite-saturation surface and oxygen fugacitiy values and isopleths for 2 GPa, 900°C (b) and 2 GPa, 700°C (c). The solid line traversing the diagram represents the graphite-saturation surface, possible fluid compositions plot below. The CCO-buffer locates per definition where the graphite-saturation surface intersects the C-O side of the triangle. Note that in nature, we do not expect fluid compositions with significantly more oxygen than the H2O-CO2 tie-line, rendering the field of oxidized geological fluid compositions extremely narrow, in fact graphically almost not resolvable at 700°C. The red numbers give oxygen fugacities (in log-units) on the graphite saturation surface (a C = 1), the blue numbers on the H2O-CO2 tie-line approximated by a C-activity of 10-6, the green numbers those between H and H2O. Almost the entire width of fluid compositions on the graphite-saturation surface spans a relatively small range in oxygen fugacities equivalent to CCO to IW. More reducing fluids are all on the H-rich side of possible fluid compositions, essentially within a X O fluid of 0.01. The star in (b) indicates a putative fluid from an organic rich sediment that also has some CO2; to drive this composition to the reducing side, CO2 has to be removed from the fluid in carbonates. Magnesite and dolomite are only stable to oxygen fugacities of the H2O-maximum (the inflection in the graphite-saturation surface), but calcite until an approximate logf O2 of -15 (900°C) or -19 (700°C). Fluids from carbonate-free organic-rich sediments are expected to lie on the graphite/diamond saturation surface on the H-side of the H2O-C tie-line, as indicated by the star in (c). These fluids move away from H2O, on the graphite-saturation surface co-precipitating graphite or diamond, if reaction with peridotite produces hydrous phases (phase A at high pressure, serpentine or brucite at low pressure). Eventually such fluids would become more reduced than the IW buffer and would possible reach a H:O ratio of >1000 necessary to precipitate SiC. Note that there is no direct proportionality between oxygen fugacity and X O fluid .

1a) Starting from fluids with some CO2, the precipitation of carbonates through reaction of the fluid with silicates (e.g. olivine + CO2 = opx + magnesite; or cpx + olivine + CO2 = calcite + opx) removes oxygen from the fluid, causing the fluid composition to evolve away from CO2, the star in Figure 7b representing a hypothetical starting composition.

1b) Alternatively, fluid percolation in mantle rocks may lead to the formation of magnetite component in spinel and hence consumption of oxygen from the fluid. This would lead to co-precipitation of graphite/diamond and the rising fluid would evolve towards the left along the graphite/diamond saturation surface of Figure 7. In both cases, the fluid could only evolve to more reducing compositions if it leaves the zones of carbonate or spinel formation behind.
  1. 2)

    If the initial fluid composition is at H:O <2, steps 1a or b would need to fractionate the fluid composition to H:O >2. Any fluid initially already to the H-rich side of the H2O-graphite/diamond tie-line or H2O-maximum (fluids are mostly H2O at the inflection of the graphite saturation surface, Figure 7) would evolve towards more methane-rich compositions by removal of H2O from the fluid. This can easily be achieved by the precipitation of hydrous phases when the fluid percolates through mantle rocks. In peridotitic compositions, such hydrous phases could be hydrous alphabet phases at higher pressure, otherwise humites, serpentine, brucite or talc. Removal of H2O from the fluid forces the fluid to evolve along the graphite/diamond saturation surface continuously precipitating graphite/diamond concomitantly with the formation of hydrous magnesium silicates. Note that step 2 requires temperatures within the stability fields of these phases, which, in equilibrium with olivine, for the most part do not exceed 700-800°C (Fumagalli and Poli, [2005]).

     
  2. 3)

    Finally, still fractionating e.g. serpentine or brucite, such fluids could reach oxygen fugacities of IW-5 to IW-8 where SiC and Fe-silicides become stable. Note that the compositional change in the fluid from moderately reducing conditions near the IW-buffer to ultra-reducing conditions that allow for SiC formation is small in terms of X O , and is graphically not resolvable in the C-O-H ternary. This final step of fluid fractionation occurs very close to the intersection of the graphite-saturation surface with the C-H sideline (Figure 7).

     

The simplest case enabling SiC formation is to start with an aqueous fluid that contains more methane than CO2 (molar), i.e. with a bulk composition more reducing then the H2O-maximum. Sequestration of oxygen from the fluid by hydroxide minerals would then directly lead to extremely reducing conditions. In mantle compositions, serpentine or brucite are only stable at relatively low temperatures (<700°C) as could be expected for fluid-percolation in supra-subduction mantle or during orogenic metamorphism. In fact, SiC has been reported from serpentinites (Xu et al. [2008]). Moissanite accounting for 8% of a rock with a bulk X Mg  = 0.998 has been reported from a beach pebble which matrix is brucite-dominated (Di Pierro et al. [2003]). This latter rock contains minor (remnant) chromite and orthopyroxene of typical mantle composition and may represent a completely metasomatized peridotite. Otherwise, moissanite is rare in most rocks and mostly described from heavy mineral concentrates. Phases grown with or onto moissanite are Fe-silicides, Si- and other metals. Unfortunately, the textural context and relations to silicates remain mostly unknown providing little evidence for the type of reactions leading to moissanite.

Step 3 is likely to occur through grain scale equilibria where new hydrous minerals form in equilibrium with the fluid but where the bulk rock is in disequilibrium with this fluid. Alternatively, discrete metasomatic zones could host SiC and hydrous phases as in the rock from Turkey described by Di Pierro et al. ([2003]). As SiC is a rare mineral, it appears likely that in most cases either reduced fluids are consumed before extreme fractionation or fluid:rock ratios in percolation zones are too high to change fluid-compositions much. However, when such prolonged fractionation occurs, C-O-H fluids may fractionate to H:O ratios sufficiently high for SiC-precipitation.

Evidence for reduced zones in the mantle wedge

Although the mantle wedge is generally thought to be rather oxidized, evidence for a fairly reduced mantle wedge has been reported for xenoliths from the Avacha volcano, Kamtchatka (Ishimaru et al. [2009]). The bulk of minerals in these peridotite xenoliths equilibrated at 870-1040°C and oxygen fugacities of ΔQFM = -0.2 to +1.5, and most silicates have XMg values characteristic for the mantle. Nevertheless, some xenoliths contain small amounts of Ni-, Fe- and possibly Ti-metal and Ti-rich Fe silicides mostly along grain boundaries. Serpentine is absent in these xenoliths. The sporadic occurrence of metal and silicides are ascribed to a reducing fluid phase with oxygen fugacities ≤ IW that causes precipitation of the reduced phases under disequilibrium with the host rock. Exact temperatures could not be determined but are suggested to lie below the bulk equilibration temperature and serpentine stability (Ishimaru et al. [2009]), probably between 700 and 900°C.

Carbon isotope composition

Natural moissanites have a light carbon isotopic composition of δ13C = -20 to -37‰ (Mathez et al. [1995], Trumbull et al. [2009]). Generally, the main source material for such light compositions is organic matter typically reaching down to -30‰. Upon burial and heating through subduction or orogenesis such organics will generate fluids with low δ13C. However, the relation between the original fluid composition and the δ13C of precipitated SiC is not straightforward: The fluid evolves to ultra-reduced compositions through removal of H2O by hydrous phases, this process is necessarily accompanied by graphite or diamond precipitation (see Figure 7). Poulsen ([1996]) calculated a fractionation factor Δ13Cgraphite-CH4 of +1 to +5‰ (750 to 500°C) and hence graphite fractionation renders an already reduced fluid lighter. On the other hand mixing of the fluid with mantle carbon of δ13C ~ -5‰ (Deines [2002]) would lead to heavier compositions. Nevertheless, Deines ([2002]) points out that the light carbon mode of mantle samples of δ13C ~ -20 to -25‰ is not necessarily solely related to surface carbon, also the fractionation factor Δ13CSiC-CH4 remains unknown. Hence a quantitative model is difficult, but the light isotopic carbon composition of moissanite supports an involvement of initially reduced fluids probably originating from organic material.

Kimberlites

Moissanite is reported from kimberlites (e.g. Shiryaev et al. [2011]) which contain light 13δ C diamonds (Trumbull et al. [2009]), but their direct paragenetic association remains uncertain (Gurney [1989]). In one case, SiC was described as part of a complex inclusion in diamond (Leung [1990]), the latter effectively halting diffusive equilibration with the surroundings. Moissanite could survive millions of years in the sub-lithospheric cratonic mantle at low temperatures (<800°C). In principle, the same fluid or melt that leads to diamond formation could also precipitate SiC in a later more fractionated stage, which should occur at shallower pressures and lower temperatures. Regardless of these observations, as kimberlites are CO2-rich magmas whose eruptions are propelled by CO2, it is not possible that the kimberlite magma would be equilibrated with SiC. Hence, moissanites in kimberlites represent xenocrysts, the kimberlites sampling cold and reduced parts of lithospheric mantle keels.

Oxygen-content and mass balance

Ultra-reduced mantle

To appreciate the amount of oxygen to be removed from a given mantle volume to attain oxygen fugacities of the SiC-buffer, the following mass balance is at order: To move a graphite-bearing mantle (with 8.0 wt% FeO) from just below the CCO-buffer (logf O2  = -8.5 or ΔIW = +3.3; all values for 2 GPa, 1200°C) to the metal-forming reaction (logf O2  = -12.9 or ΔIW = -1.1) where a few per mill of metal form, about 0.05 wt% O2 have to be removed from the rock. To further move a mantle volume to just above the SiC-buffer (logf O2  = -20.1 or ΔIW = -8.3), where essentially all Fe2+ is reduced to Fe0 and where the metal contains about 26 mol% Si, a total of 2.82 wt% O2 have to be removed. This is equivalent to 6.4 wt% relative of the total oxygen contained in the mantle volume considered, making it hard to imagine how such a process could possibly operate on a large scale. To remove such an amount of oxygen, an extremely high fluid/rock ratio would be required (as ultra-reduced fluids have little oxygen); this is most easily realized if only very small rock volumes in a zone of high fluid throughput are reduced to oxygen fugacities of the SiC-buffer.

Highly fractionated fluid

In the mantle, there is little change in oxygen content between oxygen fugacities just below the olivine-opx-magnetite equilibrium to the olivine-opx-metal equilibrium (i.e. the equivalents to the magnetite-quartz-fayalite (QFM) and iron-wustite (IW) buffers in peridotites), followed by a large change in oxygen content with reduction to the olivine-opx-C-SiC buffer. Exactly the contrary is true for COH-fluid. A fluid from graphite-saturated sediments could have e.g. 20 mol% CO2 and a X O fluid (=molar O/(H + O)) of roughly 0.75 (Figure 7). Its oxygen fugacity would be 0.6 to 1.3 log-units below QFM or ΔIW = +5 to +3.5, (all values for 700-900°C, 1 GPa). In order to evolve to an oxygen fugacity just above the IW-buffer which locates at X O fluid of 0.004 to 0.011 on the graphite-saturation surface, almost 99% of the fluid’s oxygen needs to be removed. To further move to a fluid composition that would be in equilibrium with SiC, X O fluid becomes 0.001, which composition-wise is only a small step.

In summary, oxygen fugacity has no simple proportionality to the oxygen content of a rock. In particular, oxygen fugacity is not a parameter that can be freely varied, one has to carefully consider which changes in oxidation state (mostly of Fe and C, but in this case also Si and Cr) and bulk composition of the rock volumes under consideration are involved.

Conclusions

Mantle silicates in equilibrium with natural moissanite have X Mg values of >0.993, SiC forming only 4.5-6 log units below the IW reference buffer. Large volumes of peridotite with such mineral compositions are unknown, hence, SiC forms locally in disequilibrium with the bulk rock and does not reflect large scale modifications of oxygen contents in the mantle. SiC cannot be in equilibrium with minerals that have typical average mantle X Mg of 0.88-0.92. As Fe-Mg diffusion in silicates is generally fast, SiC could not survive at adiabatic temperatures in the mantle for any reasonable time span (i.e. >100’000 yrs) hence precluding any high temperature (>1000°C) origin of natural moissanite. Consequently, any primordial, core-mantle boundary or lower mantle origin for SiC can be ruled out on the basis of the temperatures involved. Nevertheless, natural moissanite does yet not directly indicate any particular pressure of formation.

We conclude that SiC forms through a relatively low temperature mechanism (<700-800°C) and identify this mechanism as percolation of ultra-reduced fluids. Such temperatures in mantle rocks are only realistic in the subduction realm or during orogenesis. Suitable fluids may originate from sediments rich in organic matter. These fluids evolve from moderately reduced compositions through progressive removal of the oxidized species H2O through hydrous magnesium silicates, which form through reaction with the host rock during fluid percolation. After >99% fractionation, such fluids may reach H:O ratios of ~1000 that stabilize SiC. Such extreme fractionation is obviously rare but may occur occasionally. The initial fluids are expected to form upon heating of sediments, this could occur in the subduction realm, but also during orogenic burial, possibly precipitating SiC in ophiolitic peridotites upon orogenesis.

Abbreviations

CCO: 

Graphite/diamond-CO-CO2, Oxygen fugacity buffers

IW: 

Iron-wustite

MOOC: 

Moissanite-olivine-orthopyroxene-graphite/diamond

QFM: 

Quartz-fayalite-magnetite.

Phases: 

bcc: Body-centered cubic structure of Fe-alloy

graph: 

Graphite

fcc: 

Face-centered cubic structure of Fe-alloy

FeSiCalloy

Alloy of ternary composition with dominant iron

mgt: 

Magnetite

ol: 

Olivine

opx: 

Orthopyroxene

Sialloy

Alloy of ternary Si-Fe-C composition, almost pure Si

Declarations

Acknowledgements

The experiments of this study were performed while C.Gao visited ETH with the State Sponsored Study Abroad Program by the Chinese Scholarship Council. A. Golubkova was supported through SNF grants 200020-140541/1 and 200020-153112/1 to MWS.

Authors’ Affiliations

(1)
Department of Earth Sciences, ETH
(2)
State Key Laboratory of Geological Processes and Mineral Resources, School of Earth Science, China University of Geosciences
(3)
Institut für Mineralogie, Westfälische Wilhelms-Universität Münster
(4)
School of Ocean and Earth Science, Tongji University

References

  1. Aleksandrov IV, Goncharov AF, Stishov SM, Yakovenko EV: Equation of state and Raman scattering in cubic BN and SiC at high pressures. Jetp Letters 1989, 50: 127–131.Google Scholar
  2. Bai W, Robinson PT, Fang Q, Yang J, Yan B, Zhang Z, Hu XF, Zhou MF, Malpas J: The PGE and base-metal alloys in the podiform chromitites of the Luobasa ophiolite, southern Tibet. Can Mineral 2000, 38: 585–598. 10.2113/gscanmin.38.3.585View ArticleGoogle Scholar
  3. Ballhaus C: Is the upper mantle metal-saturated? Earth Planet Sci Lett 1995, 132: 75–86. 10.1016/0012-821X(95)00047-GView ArticleGoogle Scholar
  4. Brady JB, Cherniak DJ: Diffusion in minerals: an overview of published experimental diffusion data. Rev Mineral Geochem 2010, 77: 899–920. 10.2138/rmg.2010.72.20View ArticleGoogle Scholar
  5. Brosh E, Makov G, Shneck RZ: Thermodynamic analysis of high-pressure phase equilibria in Fe-Si alloys, implications for the inner core. Phys Earth Planet Int 2009, 172: 289–298. 10.1016/j.pepi.2008.10.012View ArticleGoogle Scholar
  6. Campbell A, Danielson L, Righter K, Seagle CT, Wang Y, Prakapenka VB: High pressure effects on the iron-iron oxide and nickel-nickel oxide oxygen fugacity buffers. Earth Planet Sci Lett 2009, 286: 556–564. 10.1016/j.epsl.2009.07.022View ArticleGoogle Scholar
  7. Chipman J: Thermodynamics and phase diagram of the Fe-C system. Metal Trans 1972, 3: 55–64. 10.1007/BF02680585View ArticleGoogle Scholar
  8. Connolly JAD (1995) Phase-diagram methods for graphitic rocks and application to the system C-O-H-FeO-TiO2-SiO2. Contrib Mineral Petrol 119:94–116View ArticleGoogle Scholar
  9. Connolly JAD (2009) The geodynamic equation of state: what and how. Geochem Geophys Geosys 10:, doi:10.1029/2009GC002540Google Scholar
  10. Deines P: The carbon isotope geochemistry of mantle xenolioths. Earth Sci Rev 2002, 58: 247–278. 10.1016/S0012-8252(02)00064-8View ArticleGoogle Scholar
  11. Di Pierro S, Gnos E, Grobety BH, Armbruster T, Bernasconi SM, Ulmer P: Rock-forming moissanite (natural α-silicon carbide). Am Mineral 2003, 88: 1817–1821.View ArticleGoogle Scholar
  12. Djurovic D, Hallstedt B, von Appen J, Dronskowski R: Thermodynamic assessment of the Fe-Mn-C system. CALPHAD: Comput Coupling Phase Diagrams Thermochemistry 2011, 35: 479–491. 10.1016/j.calphad.2011.08.002View ArticleGoogle Scholar
  13. Frost DJ, McCammon CA: The Redox State of Earth’s Mantle. Annual Rev Earth Planet Sci 2008, 36: 389–420. 10.1146/annurev.earth.36.031207.124322View ArticleGoogle Scholar
  14. Fumagalli P, Poli S: Experimentally determined phase relations in hydrous peridotites to 6.5 GPa and their consequences on the dynamics of subduction zones. J Petrol 2005, 46: 555–578. 10.1093/petrology/egh088View ArticleGoogle Scholar
  15. Gao C, Liu Y: Moissanite-bearing carbonatite xenoliths from Cenozoic basalt, North China: products of ancient oceanic crust subduction ? Geochim Cosmochim Acta 2008, 72: A292.Google Scholar
  16. Golubkova A, Schmidt MW, Connolly JAD (2014) Simulating reducing and ultra-reducing conditions in the mantle: application to the occurrence of SiC in podiform chromitites. Submitted Contrib Mineral PetrolGoogle Scholar
  17. Gurney JJ: Diamonds. In Kimberlites and Related Rocks, Vol. 2 Edited by: Ross J. 1989, 935–965.Google Scholar
  18. Gustafson P: A thermodynamic evaluation of the Fe-C system. Scand J Metallurgy 1985, 14: 259–267.Google Scholar
  19. Hirayama Y, Fujii T, Kurita K: The melting relation of the system iron and carbon at high pressure and its bearing on the early stage of the Earth. Geophys Res Lett 1993, 20: 2095–2098. 10.1029/93GL02131View ArticleGoogle Scholar
  20. Holland TJB, Powell R: An improved and extended internally consistent thermodynamic dataset for phases of petrological interest, involving a new equation of state for solids. J Metam Geol 2011, 29: 333–383. 10.1111/j.1525-1314.2010.00923.xView ArticleGoogle Scholar
  21. Ishimaru S, Arai S, Shukuno H: Metal-saturated peridotite in the mantle wedge inferred from metal-bearing peridotite xenoliths from Avacha volcano, Kamchatka. Earth Planet Sci Lett 2009, 284: 352–360. 10.1016/j.epsl.2009.04.042View ArticleGoogle Scholar
  22. Lacaze J, Sundman B: An assessment of the Fe-C-Si system. Metall Transact A 1991, 22A: 2211–2223. 10.1007/BF02664987View ArticleGoogle Scholar
  23. Leung I: Silicon carbide cluster entrapped in a diamond from Fuxian, China. Am Mineral 1990, 75: 1110–1119.Google Scholar
  24. Leung I, Guo W, Friedman I, Gleason J: Natural occurrence of silicon carbide in a diamondiferous kimberlite from Fuxian. Nature 1990, 346: 352–354. 10.1038/346352a0View ArticleGoogle Scholar
  25. Li Z, Bradt RC: Thermal expansion of the hexagonal (4H) polytype of SiC. J Am Ceramic Soc 1987, 70: 445–448. 10.1111/j.1151-2916.1987.tb05673.xView ArticleGoogle Scholar
  26. Li J, Mao HK, Fei Y, Gregoryanz E, Eremets M, Zha CS: Compression of Fe 3 C to 30 GPa at room temperature. Phys Chem Minerals 2002, 29: 166–169. 10.1007/s00269-001-0224-4View ArticleGoogle Scholar
  27. Lord OT, Walter MJ, Dasgupta R, Walker D, Clark SM: Melting in Fe-C system to 70 GPa. Earth Planet Scie Lett 2009, 284: 157–167. 10.1016/j.epsl.2009.04.017View ArticleGoogle Scholar
  28. Lyakhovich VV: Origin of accessory moissanite. Geol Rev 1980, 22: 63–74. 10.1080/00206818209466961View ArticleGoogle Scholar
  29. Mathez EA, Fogel RA, Hutcheon ID, Marshintsev VK: Carbon isotopic composition and origin of SiC from kimberlites of Yakutia, Russland. Geochim Cosmochim Acta 1995, 59: 781–791. 10.1016/0016-7037(95)00002-HView ArticleGoogle Scholar
  30. Miettinen J: Reassessed thermodynamic solution phase data for ternary Fe-Si-C system. CALPHAD: Comput Coupling Phase Diagrams Thermochemistry 1998, 22: 231–256. 10.1016/S0364-5916(98)00026-1View ArticleGoogle Scholar
  31. Mookherjee M, Nakajima Y, Steinle-Neumann G, Glazyrin K, Wu XA, Dubrovinsky L, McCammon C, Chumakov A (2011) High-pressure behavior of iron carbide (Fe7C3) at inner core conditions. J Geophysical Res-Solid Earth 116, B04201Google Scholar
  32. Nakajima Y, Takahashi E, Sata N, Nishihara Y, Hirose K, Funakoshi K, Ohishi Y (2011) Thermoelastic property and high-pressure stability of Fe7C3: implication for iron-carbide in the Earth’s core. Am Mineral 96:1158–1165View ArticleGoogle Scholar
  33. O’Neill H, Wall VJ: The olivine-orthopyroxene-spinel oxygen geobarometer, the nickel precipitation curve, and the oxygen fugacity of the Earth’s upper mantle. J Petrol 1987, 28: 1169–1191. 10.1093/petrology/28.6.1169View ArticleGoogle Scholar
  34. O'Neill HS (1988) The system Fe-O and Cu-O: thermodynamic data for the equilibria Fe-“FeO”, Fe-Fe3O4, “FeO”-Fe3O4, Cu-Cu2O and Cu2O-CuO from emf measurements. Am Mineral 73:470–486Google Scholar
  35. Poulsen SR: Equilibrium mineral-fluid stable isotope fractionation factors in graphitic metapelites. Chem Geol 1996, 131: 207–217. 10.1016/0009-2541(95)00153-0View ArticleGoogle Scholar
  36. Rohrbach A, Ghosh S, Schmidt MW, Wijbrans CH, Klemme S: The stability of Fe-Ni carbides in the Earth's mantle: evidence for a low Fe-Ni-C melt fraction in the deep mantle. Earth Planet Science Lett 2014, 388: 211–221. 10.1016/j.epsl.2013.12.007View ArticleGoogle Scholar
  37. Shiryaev AA, Griffin WL, Stoyanov E: Moissanite (SiC) from kimberlites: polytypes, trace elements, inclusions, and speculations on origin. Lithos 2011, 122: 152–164. 10.1016/j.lithos.2010.12.011View ArticleGoogle Scholar
  38. Stagno V, Frost DJ: Carbon speciation in the asthenosphere: experimental measurements of the redox conditions at which carbonate-bearing melts coexist with graphite or diamond in peridotite assemblages. Earth Planet Sci Lett 2010, 300: 72–84. 10.1016/j.epsl.2010.09.038View ArticleGoogle Scholar
  39. Trumbull RB, Yang JS, Robinson PT, Di Pierro S, Vennemann T, Wiedenbeck M: The carbon isotopic composition of natural SiC (moissanite) from the Earth’s mantle: new discoveries from ophiolites. Lithos 2009, 113: 612–620. 10.1016/j.lithos.2009.06.033View ArticleGoogle Scholar
  40. Ulmer GC, Grandstaff DE, Woermann E, Gobbels M, Schonitz M, Woodland A: The redox stability of moissanite (SiC) compared with metal-metal oxide buffers at 1773K and at pressures up to 90 kbar. Neuses Jahrb Min Abhand 1998, 172: 279–307.Google Scholar
  41. Wood IG, Vočadlo L, Knight KS, Dobson DP, Marshall WG, Price GD, Brodholt J: Thermal expansion and crystal structure of cementite, Fe3C, between 4 and 600K determined by time of flight neutron powder diffraction. J Appl Crystallogr 2004, 37: 82–90. 10.1107/S0021889803024695View ArticleGoogle Scholar
  42. Woodland AB, Koch M: Variation in oxygen fugacity with depth in the upper mantle beneath the Kaapvaal craton, Southern Africa. Earth Planet Sci Lett 2003, 214: 295–310. 10.1016/S0012-821X(03)00379-0View ArticleGoogle Scholar
  43. Woodland AB, O’Neill HS: Thermodynamic data for Fe-bearing phases obtained using noble metal alloys as redox sensors. Geochim Cosmochim Acta 1997, 61: 4359–4366. 10.1016/S0016-7037(97)00247-0View ArticleGoogle Scholar
  44. Xu A, Wu W, Xiang W, Yang J, Chen J, Ji S, Liu Y: Moissanite in serpentinites from the Dabie Shan Mountains in China. Mineral Mag 2008, 72: 899–908. 10.1180/minmag.2008.072.4.899View ArticleGoogle Scholar

Copyright

© Schmidt et al.; licensee Springer. 2014

This article is published under license to BioMed Central Ltd. This is an Open Access article distributed under the terms of the Creative Commons Attribution License (http://creativecommons.org/licenses/by/4.0), which permits unrestricted use, distribution, and reproduction in any medium, provided the original work is properly credited.

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