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Lizardite versus Antigorite Serpentinite: Magnetite, Hydrogen, and Life(?)
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Bernard W. Evans
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Department of Earth and Space Sciences, Box 351310, University of Washington,
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Seattle, WA 98195-1310, U.S.A.
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ABSTRACT
The serpentinization of peridotite operates according to one or other, or a
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combination, of two end-member mechanisms. In low-temperature environments (50-
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300˚C), where lizardite is the predominant serpentine mineral, olivine is consumed by
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reaction with H2O but its composition (Mg#) remains unchanged. Mg-rich lizardite,
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magnetite and dihydrogen gas (± brucite) are products of the reaction. At higher
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temperatures (400-600˚C) rates of MgFe diffusion in olivine are orders of magnitude
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faster, with the result that the growth of Mg-rich antigorite can be accomodated by a
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compositional adjustment of olivine, eliminating the need to precipitate magnetite and
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evolve hydrogen. This latter end-member mechanism probably best reflects the situation
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in the forearc mantle wedge.
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INTRODUCTION
The serpentinization reaction is ordinarily written as conserving oxygen. Because
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it is accompanied by the oxidation of iron, the mass balance necessarily involves the
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liberation of dihydrogen gas. The precipitation of magnetite, uptake of ferric iron in
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lizardite (Whittaker and Wicks, 1970; Seyfried et al., 2007; Evans, 2008), and evolution
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of hydrogen (Thayer, 1966; Barnes et al., 1972; Neal and Stanger, 1983; Abrajano et al.,
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1990; Sherwood-Lollar et al., 1993) are features widely observed as accompanying the
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low-temperature serpentinization of peridotite. So for dunite the generalized equation is:
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MgFe-olivine + H2O = MgFe-serpentine + MgFe-brucite + magnetite + H2(aq)
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(1)
The inverse correlation between magnetic susceptibility and density that equation
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1 implies has been confirmed by measurements (Coleman, 1971; Toft et al., 1990; Oufi et
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al., 2002). Together with seismic velocities, these properties constitute valuable tools for
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geophysicists assessing the percentage of serpentinite in subducted and underthrust
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(wedge) mantle. Equally important, the evolution of hydrogen has been recognized as a
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source of energy for the potential production of abiotic hydrocarbons via Fischer-Tropsch
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reactions.
To the author’s knowledge, no serious attempt has been made to explain why the
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low-T serpentinization reaction should be an oxidizing one.
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Textbook metamorphic reactions involving solid solution minerals (e.g., Spear,
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1993) are generally written oxygen conserved and unaccompanied by oxidation of iron.
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For a long time it has been axiomatic in metamorphic petrology that reactions among
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solid-solution minerals tend to proceed so as to maintain the equivalence of the chemical
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potentials of components among the phases (in cation exchanges such as FeMg–1, Fe3+Al–
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1,
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according to this equilibrium principle. Coexisting solid solution minerals are represented
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in projection by intersecting T –X or P – X phase loops, and the loci of points on the loops
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conform to the exchange coefficient KD. Progress of such reactions, whether prograde or
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retrograde, requires the free exchange of components among the minerals.
KNa–1). PTX-relations, including pseudosections, are routinely calculated and graphed
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Ferromagnesian minerals, as in Barrovian metamorphism, are prime examples. Changes
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in the redox state of iron are not predicted. It seems to have escaped notice that the
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reaction describing the low-temperature formation of serpentine from peridotite does not
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follow this classic metamorphic model. Perhaps, quite reasonably, we have simply not
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had that expectation.
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At this point it is useful to draw a distinction between lizardite serpentinization,
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which is generally inferred to occur in the range 50-300˚C, and antigorite serpentinization
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in the range 400-600˚C. In both cases we can expect the first-formed layer of serpentine
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to be Mg-rich, a reflection of Fe/Mg partition between olivine (or orthopyroxene) and
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serpentine (KD ~ 0.4) and the chemical potentials of Fe and Mg (more specifically the
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Fe2+Mg-1 exchange potential) imposed on the thin film of fluid at the growth surface
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(Evans, 2008). Contrary to a seemingly popular view, the Fe-content of the serpentine is
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not limited by an intrinsic instability of more Fe-rich serpentine, as any global
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compilation of lizardite analyses can testify (e.g., Evans et al., 2009).
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THE DIFFUSION PROBLEM
Next we should question whether the reactant olivine truly “reacts” in the sense of
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changing its Fe/Mg ratio in accordance with reaction progress, or whether it merely
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dissolves at constant composition? Informative in this context are laboratory data on Fe-
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Mg diffusion in olivine, which is by far the slowest step in a full serpentinization process.
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At 100 MPa and 200˚C, DFeMg after extrapolation is –32 log units (m2/s) along [001], the
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fastest direction, and the characteristic diffusion distance √(Dt) over 1 my is 10-4 micron
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(Dohmen and Chakraborty, 2007). At 500 and 600˚C and the same pressure, DFeMg is 10-
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23.6
and 10-21.9 (m2/s) and the characteristic diffusion distances over 1 my are respectively
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8 and 60 microns (Fig. 1). Thus, in antigorite serpentinization, especially above 500˚C
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and over several my, we might expect reaction behavior to follow the textbook
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metamorphic model, there being no serious limitation on the export of Mg to the growing
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serpentine and consequent enrichment in iron of the olivine. As a result, there may be
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little or no growth of magnetite, possibly more brucite, and no yield of hydrogen. In
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lizardite serpentinization, olivine dissolves but it cannot react by changing its
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composition. In this case the precipitation of magnetite and evolution of hydrogen are
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assured. The system as a whole does not attain a condition of minimum Gibbs free energy
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because the mass-balance contribution of olivine interiors is not included.
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These rates of Fe-Mg exchange equilibration stand in marked contrast to the rate
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of serpentinization of olivine as measured hydrothermally in the laboratory (Martin and
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Fyfe, 1970; Wegner and Ernst, 1983). For example, at 310˚C and 100 MPa, 50˚C below
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the reaction curve, the hydration reaction is 50% complete after only 20 days. Thus, even
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at the low temperatures of lizardite serpentinization, the rate of the net-transfer reaction
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will be limited only by the rate of supply of H2O (MacDonald and Fyfe, 1985). At the
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higher temperatures of antigorite serpentinization, if hydration is facilitated by a copious
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supply of H2O, the time for diffusion may be curtailed, and reaction will behave more
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like lizardite serpentinization.
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LIZARDITE
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Consistent with the foregoing, microprobe analyses of low-T serpentinites show
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that the growth of Mg-rich lizardite from olivine is not accompanied by a corresponding
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enrichment in its fayalite content. The most frequent composition of lizardite in
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peridotites of mantle origin has an Mg# of 0.97-0.96 (Evans, 2009, Fig. 1). In many
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instances though, the Fe content is higher owing to its uptake of ferric iron. These
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compositions of lizardite reflect an attempt to equilibrate with the Fe2+Mg -1 exchange
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potentials of the primary olivine and orthopyroxene. Although olivine incrementally
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dissolves in the process, it is “non-reactive” in the sense of being unable to adjust its
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composition. For mass-balance another Fe mineral must form, such as magnetite or
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awaruite Ni3Fe. Although brucite tends to be iron rich, for example with Mg#s of 0.85-
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0.70, it is not present in sufficient quantities to satisfy the Fe mass-balance. Lizardite
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serpentinization may be viewed then as a “factory” that adds H2O to a semi-infinite
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supply of olivine to continuously yield products consisting of Mg-rich lizardite,
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magnetite, hydrogen, and in some cases brucite and FeNiCo alloys. Thus the
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mineralogical properties of most serpentinites are the consequence of the enormous
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differences in effective rates between the hydration reaction and MgFe-diffusion in
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reactant olivine.
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ANTIGORITE
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The growth of antigorite at the expense of olivine will reflect progress of one or
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more of four possible hydration reactions in the MSH system above about 300ºC: Fo +
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H2O → Atg + Brc, Fo + Tlc + H2O → Atg, Fo + En + H2O → Atg, and Fo + SiO2(aq) +
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H2O → Atg. In the CMSH system, antigorite can form along with diopside: Fo + Tr +
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H2O → Atg + Di, as inferred for the Happo ultramafic complex, Japan, believed to be
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mantle wedge (Nozaka, 2005). In the FMSH system, as argued above, the higher the
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temperature the less likely will be an accompanying precipitation of magnetite. These
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reactions could take place in a subducting slab if they are initiated by the introduction of
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H2O into anhydrous peridotite, much as anhydrous basalt is converted into a greenschist
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or subgreenschist facies assemblage. A likely alternative in the slab is antigorite growth
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at the expense of early-formed, oceanic lizardite and chrysotile, forming prograde
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antigorite serpentinite (O’Hanley, 1996). In this case, the antigorite serpentinite will
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likely inherit some magnetite. On the other hand, if antigorite grows directly from
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anhydrous FMSH peridotite at temperatures as high as 500-600ºC, we would predict little
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or no magnetite nor evolution of hydrogen. These conditions of alteration are to be
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expected in the mantle wedge above subducted ocean crust that is yielding aqueous fluid
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because of thermally driven dehydration reactions.
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There is in fact evidence from examples of antigorite serpentinization, notably
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from the mantle wedge, that the growth of Mg-rich antigorite from MgFe-olivine results,
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as anticipated, in Fe-enrichment of the olivine. Amounts of magnetite are small or even
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zero. The potential for magnetite and serpentine growth is in any case less than in
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lizardite serpentinization because olivine is stable above ~ 400˚C. The modal amount of
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antigorite is constrained by the whole-rock composition: unless silica is introduced, there
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will be no antigorite in a meta-dunite, and in a harzburgite (e.g., 80% olivine, 20%
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enstatite) there will be less than 50% antigorite. Estimates of the degree of
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serpentinization in the mantle wedge based on seismic velocities range widely: 15% to
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60% (Zhao et al., 2001; Bostock et al., 2002; Courtier et al., 2004).
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Smith (2010) described an antigorite-rich metaperidotite (sample N15-GN) from
the Green Knobs diatreme in the Navajo Volcanic Field that he inferred had formed by
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hydration of dunite above the Farallon plate. Modally the sample consists of 33% olivine
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(Mg# 0.884), 63% antigorite (Mg# 0.942), 0.5% diopside, 0.3% magnesite, 0.1%
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chromite, and 0.1% clinohumite. No magnetite is present. The whole-rock composition
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reconstructed from mineral and modal analyses has Mg# = 0.918, which suggests that the
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protolith olivine was in the range 0.92- 0.89 that is typical for unaltered Colorado Plateau
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xenoliths. “The olivine is more Fe-rich than that in any other Green Knobs
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peridotites…the Fe-rich composition may be attributed to formation of antigorite and
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partition of Fe into residual olivine” (Smith, 2010).
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In Mariana forearc serpentinite seamounts, Murata et al. (2009a,b) found that
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iron-rich olivine (0.86-0.88) formed at the margins of more Mg-rich, primary mantle
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olivine (0.89-0.93) in close spatial relationship to antigorite formation – with “only small
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amounts of magnetite”. Iron-enriched olivine also formed preferentially in stripes along
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(100) olivine “cleavages” (deformation zones). MnO is also enriched in olivine in the
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stripes, consistent with its partition with respect to antigorite (Trommsdorff and Evans,
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1974). Formation temperatures were estimated at 450-550ºC. These samples provide
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good evidence for the influence of MgFe pipe-diffusion along planar defects associated
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with dislocations in mantle olivine (Ando et al., ms), perhaps aided by ingress of water
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(Boudier et al., 2009).
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Forearc antigorite serpentinites of Neoproterozoic age from the Eastern Desert,
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Egypt (Khalil and Azer, 2007) contain highly strained olivine (Mg# > 0.89) of mantle
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derivation, and later unstrained olivine (0.84-0.85). The authors suggest a thermal event
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(granite intrusion) for the latter; nevertheless, whole-rock data show that the unstrained
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olivines grew in peridotite with an Mg# of 0.91.
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In samples of meta-peridotite from serpentinite mélange in Guatemala, antigorite-
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olivine serpentinite is reported in some cases to be free of magnetite (G. Harlow, personal
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communication, February, 2010). In two cases, two generations of olivine were found:
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Mg#s 0.90 and Mn-enriched 0.84 in one, and 0.91 and Mn-enriched 0.85 in the other.
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The absence of any indication of pre-existing lizardite, and association with sediment-
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fertilized HP blocks, supports a supra-subducted plate origin.
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Antigorite formed from olivine in the supra-subduction zone Trinity peridotite,
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California, above and close to the Trinity thrust (Peacock, 1987). Although magnetite is
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present along fractures in the olivine, antigorite growth (as blades penetrating the olivine)
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was interpreted as largely postdating this magnetite. Other minerals are tremolite,
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chlorite, talc, carbonate, and ferrit-chromite. The antigorite serpentinization was caused
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by infiltrating fluids at ~ 500 ºC and ~ 5 kbar derived from the subducting slab beneath.
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Field examples of the hydration of olivine directly to antigorite are far fewer than
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in the case of lizardite. Interpenetrating textures are typical (Wicks and Whittaker, 1977;
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O’Hanley, 1996). However, it is not clear from their examples how much magnetite was
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produced and whether or not the olivine underwent an increase in iron. In a good many
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serpentinites, there is textural evidence for the growth of antigorite at the expense of
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lizardite and chrysotile. The MSH phase diagram (Evans, 2004) suggests temperatures in
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excess of 300˚C for the growth of antigorite from lizardite.
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Texturally well-equilibrated metamorphic antigorite-olivine serpentinites
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ordinarily show the attainment of FeMg-exchange equilibrium between olivine and
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antigorite (Fig. 2). Excellent examples come from regionally metamorphosed Tethyan
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ophiolites in the Alps, which reached temperatures on the order of 400-550ºC or more
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over a metamorphic cycle of perhaps 10 my. Temperatures were evidently sufficiently
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high for olivine, either by recrystallization or solid-state diffusion, to adjust to the
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demands of the partition coefficient KD and modal mineralogy. These conditions favor
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the textbook hydration model for serpentinization in which iron is not oxidized. In some
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cases however, a prograde metamorphic path via lizardite serpentinite is recognizable:
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olivine Mg#s can be as high as 0.97-0.94, magnetite is abundant, and bodies of rodingite
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and ophicarbonate occur. In other samples (Trommsdorff and Evans, 1972; Worden et al.,
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1991; Scambelluri et al., 1991), olivine falls in the range 0.89-0.84, while whole-rock
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Mg#s are around 0.91-0.90. This suggests recrystallized olivine in FeMg-exchange
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equilibrium with antigorite, and minimal prior oxidation. These samples do contain ~ 1%
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spinel but it is generally in the ferrit-chromite range (Evans and Frost, 1975), sometimes
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in aggregates ~0.1 mm in diameter suggesting derivation by alteration of primary
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chromite.
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EXPERIMENTS AND MODELLING
Hydrothermal laboratory experiments on MgFe-olivine at low temperature have
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produced magnetite along with lizardite, and presumably some hydrogen (Moody, 1976;
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Janecky and Seyfried, 1986; Normand et al., 2002; Allen and Seyfried, 2003). Even at
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500 or 600˚C the characteristic MgFe diffusion distance in olivine over one year is so
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small that laboratory experiments on antigorite growth from MgFe-olivine are equally
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unlikely to follow the textbook model of reaction progress, and so we should expect them
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to produce magnetite. Conversely, at very low temperatures, e.g. 100-200˚C, it seems that
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magnetite formation is somewhat diminished because of the take-up of ferric iron in
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lizardite (Seyfried et al., 2007; Evans, 2008; Klein et al., 2009).
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Reaction-path modeling that assumes a constant composition of olivine
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(McCollom and Bach, 2009; Klein et al., 2009) is entirely appropriate for low-
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temperature serpentinization in nature; the “full equilibrium” model would allow reaction
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progress to enrich Ol, Srp and Brc in iron. However, model W/R ratios as high as 1:1
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probably overestimate the amounts of magnetite, antigorite, and H2 that result
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(irrespective of MgFe-diffusion in olivine) from internal oxygen and silica buffering
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down-T along the isobaric MFSH quasi-univariant curve Ol-Atg-Mag. If, as is likely in
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nature, fluid-absent conditions prevail during the decompression of mantle peridotite, the
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oxygen and silica buffer capacities are likely to endure during cooling until olivine finally
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succumbs to a major influx of fluid and low-T serpentinization ensues. Therefore, we can
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largely rule out much magnetite formation and hydrogen production during cooling prior
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to the major serpentinizing event, which ordinarily is related to faulting of some kind.
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At the isobaric MFSH quasi-invariant point Ol + Liz + Brc + Mag in dunite
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(reaction 1), the loss or armoring of olivine caused by hydration releases a buffer
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constraint that in turn allows ferroan brucite to accept silica from any source and be
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converted to serpentine, magnetite, H2 and H2O (e.g. Bach et al., 2006; Frost and Beard,
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2007; Beard et al., 2009). Large inputs of ocean or ground water will eventually
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overwhelm the oxygen-buffer capacity of the rock and yield more magnetite (or hematite)
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and hydrogen by oxidation of olivine, serpentine and brucite.
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ABIOGENIC METHANE
The hydrogen-rich fluids found issuing from serpentinite on land and on the ocean
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floor are often also enriched in methane. As many have pointed out, when hydrogen can
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be combined with CO2, potential abiotic hydrocarbons are possible via a natural
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equivalent of Fischer-Tropsch synthesis. Since the discovery of the serpentinite-hosted
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Lost City hydrothermal field, 15 km W of the Mid-Atlantic Ridge (Kelley et al., 2001;
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2005; Früh-Green et al., 2003), many workers, for example, Sleep et al. (2004), Schulte
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et al. (2006), Martin et al. (2008), Proskurowski et al. (2008), have seriously advanced
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the idea of a serpentinite origin for life on the planets. The necessary ingredients
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(peridotite or komatiite and water) were available together on the surfaces of the early
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planets.
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CONCLUDING REMARKS
The contrasting mechanisms of lizardite and antigorite serpentinization suggest
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that the production of H2, and consequently potential abiogenic CH4, may not be an
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accompaniment of mantle wedge serpentinization. Similarly, the amount of magnetite
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generated in mantle wedge serpentinite may be much less than found in ocean floor
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serpentinites. Therefore the magnetic anomalies these rocks induce (e.g., Hyndman and
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Peacock, 2003; Blakeley et al., 2005) may be quite different. There are certainly risks in
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taking serpentinized abyssal peridotites as models for the kind of serpentinite that occurs
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in the mantle wedge. On the other hand, the model for abiogenic hydrocarbons and
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potential life on planetary surfaces is secure because of the low temperatures involved. It
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is a sobering thought that we might owe our existence on earth to the sluggish low-
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temperature FeMg lattice diffusion in olivine in peridotite.
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ACKNOWLEDGMENTS
The author thanks N.I. Christensen, R. Dohmen, B.R. Frost, G.E. Harlow, P.B.
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Kelemen, H. Maekawa, O. Müntener, D. Smith and F.J. Wicks for helpful comments on
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the manuscript.
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Text Figures
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Figure 1. Characteristic diffusion distance in olivine (Mg# 0.9) at 100 MPa.
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Extrapolated from Dohmen and Chakraborty (2007)
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Figure 2. Roozeboom plot of iron-magnesium partitioning between antigorite and
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olivine in texturally equilibrated serpentinites. Data from the literature and
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unpublished work of the author. Note that roughly 10-30 % of the iron in antigorite is
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ferric iron. For this reason the trend of data-points does not pass through the origin.