American Mineralogist; November-December 2009; v. 94; no. 11-12;
p. 1731-1734; DOI: 10.2138/am.2009.3301
© 2009 Mineralogical Society of America
Magnetite-free, yellow lizardite serpentinization of olivine websterite, Canyon Mountain complex, N.E. Oregon
Bernard W. Evans1,*,
Scott M. Kuehner1 and
Anastasia Chopelas2
1 Department of Earth and Space Sciences, Box 351310, University of Washington, Seattle, Washington 98195-1310, U.S.A.
2 Institute of Geophysics and Planetary Physics, University of California Los Angeles, Box 951567, Los Angeles, California 90095-1567, U.S.A.
Correspondence: * E-mail: bwevans{at}u.washington.edu
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ABSTRACT
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We document an example of serpentinization of olivine and orthopyroxene that produced virtually no magnetite, but instead relatively Fe-rich yellow-colored lizardite (XFe = 0.08 to 0.17), and the native Fe-Ni-Co metals, awaruite and wairauite. Lizardite’s identity was confirmed by micro-Raman spectroscopy, although peaks are broad. Electron microprobe analyses of the lizardite yield a continuous compositional trend of formula contents suggestive of the progressive uptake of Fe3+ exclusively on M sites, where it is charge balanced by vacancies. Although these observations are unusual, this secondary mineral assemblage can be explained in terms of the likely intensive variables T, fH2O, fH2, and aSiO2 attending the alteration. The absence of magnetite in serpentinization does not signify a lack of oxidation. By forming the hydrated phase-component ferri-lizardite instead of magnetite from the fayalite and ferrosilite components, the yield of hydrogen is reduced by two-thirds. The usual inverse correlation of rock density with magnetic susceptibility is unlikely to be the case in this kind of serpentinization.
Key Words: Serpentinite • ferrian lizardite • olivine-websterite • micro-Raman • hydrogen • magnetic susceptibility
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INTRODUCTION
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The serpentinization of olivine and orthopyroxene by low-temperature aqueous fluids normally produces serpentine that is noticeably richer in Mg (higher Mg/Fe ratio) than the primary silicates. Aside from the addition and release of volatiles and trace elements, serpentinization has been shown in most cases to be broadly isochemical (O’Hanley 1996). Consistent with the MFSHO phase diagram, mass balance then results in the precipitation of a small modal percentage of magnetite, typically as a dusting at the margins of serpentine mesh pseudomorphs after olivine.
Following up a suggestion made in Evans (2004), it was later argued (Evans 2008) that the Mg-rich nature of serpentine reflects the ambient chemical potentials of Fe and Mg components in the grain-boundary regions of a peridotite undergoing serpentinization, with the understanding that this control will weaken as the modal amounts of primary olivine and orthopyroxene decline, or when the system is infiltrated by large amounts of aqueous fluid. This is the same internal thermodynamic control that establishes the low oxygen and silica activities that the mineralogy of typical serpentinites reflects (Frost 1985; Frost and Beard 2007).
Frost and Beard (2007) pointed out instances when the product serpentine is more iron rich, in some cases with an Mg/Fe ratio comparable to that of the primary silicates. They proposed the presence of higher silica activities than elsewhere, such as might be the case in the vicinity of orthopyroxene grains, as an explanation for the higher Fe content of the serpentine.
Unlike chrysotile, whose iron content expressed as XFe [=Fe/(Fe+Mg)] shows a sharp frequency maximum at 0.03 in a histogram of analyzed samples from the literature, the XFe of lizardite in serpentinized mantle peridotite has a broad maximum, ranging in XFe values from 0.02 to 0.08 (Fig. 1
). Primary olivine and orthopyroxene in such rocks have XFe ratios mostly in the range 0.08 to 0.12. In the vast majority of examples of lizardite in both mesh and bastite pseudomorphs, microanalyses (Evans 2008, Figs. 3 and 6) show that increasing total Fe apfu in the lizardite is accompanied by decreasing Si apfu (or Si+Al/2, or Si+Al/2+Cr/2). This correlation indicates that low levels of Fe2+ become augmented by increasing amounts of Fe3+ that are taken up in a ferri-Tschermaks substitution involving a cronstedtite component: (Fe22+Fe3+)(Fe3+Si)O5(OH)4. This conclusion is consistent with the Fe3+/total-Fe ratios of highly and fully serpentinized peridotites (Thayer 1966; Evans 2008, Fig. 1), most of which are in the range 0.5 to 0.95, with an average close to that of ideal magnetite (0.67), and with Mössbauer spectroscopy studies (O’Hanley and Dyar 1993; Votyakov et al. 1993). Thus, the shape of the lizardite histogram reflects the presence of both ferrous and ferric iron, and we infer that the higher XFe values correspond to larger Fe3+ contents.
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FIGURE 1. Frequency histogram of the total Fe contents of lizardite from serpentinized mantle peridotites; data taken from the literature.
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Whereas the cronstedtite substitution accounts for ferric iron in most high-Fe lizardites, there are also examples where the uptake of iron is accompanied by what appears to be an increase in Si apfu (or Si+Al/2). In these cases, the ratio of total T to total T+M cations rises to values well above the ideal stoichiometric 0.40, for example, Evans (2008, Fig. 5), where Fe3+ was divided equally between T and M sites. But rather than signifying an increase in T cations, this trend may be attributed to ferric iron substitution coupled with vacancies on the M site, whence total formula cations will be <5.0 and the excess of T cations is an artifact of the 5-cation assumption. This substitution may well be expected in bastite pseudomorphs after orthopyroxene, although it is certainly not always the case (Evans 2008, Fig. 6). Chemographic relationships (e.g., Evans 2008, Fig. 2) suggest that the vacancy-coupled Fe3+-substituted lizardite should be found in meta-harzburgite when talc (alternatively tremolite) is present. In contrast to the development of cronstedtite (cf. Evans 2008, reaction 5), the ferrous and ferric components of lizardite in this case are linked by a potential equilibrium that explicitly involves SiO2 as a species:
 | (1) |
Equilibrium 1, like the cronstedtite equilibrium, may be shifted at constant T and P by changes in the fugacities of H2O and H2 (or O2). In both cases, the side containing the ferric-iron lizardite components will be favored at low temperatures because it incorporates fewer moles of free volatiles. A tentative general conclusion might therefore be that ferrian lizardite is associated with very low temperatures of serpentinization. However, we need not necessarily invoke "progress" during serpentinization of reaction 1 as such (see reactions 2, 3, and 4 below).
In all probability, ferric iron may be present in natural lizardites through a combination of these two substitution mechanisms. A third possible substitution mechanism for ferric iron invokes dehydrogenation (Fuchs et al. 1998). In this case, there is no change in the formula proportions of cations from the ideal composition, and so it cannot be detected by electron-probe analysis.
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OLIVINE WEBSTERITE
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This letter describes a drill-core sample of olivine websterite taken from the Canyon Mountain ophiolite complex, N.E. Oregon, which is inferred to exhibit M-site coupled Fe3+-vacancy substitution in the lizardite. The sample (CA26C) was obtained through the courtesy of N.I. Christensen, who included it in a study of the seismic velocities of ophiolite sections (Christensen 1978, Table 1). The olivine and orthopyroxene have been patchily altered to a lizardite that, under the microscope, is feebly pleochroic yellow in plane-light and up to high first-order in birefringence. Its identity was confirmed by micro-Raman spectroscopy (Groppo et al. 2006) although the peaks are broad, suggestive perhaps of some stacking disorder and as well as compositional variability (Fig. 2). Lizardite is present as anastomosing veinlets in olivine and as bastite pseudomorphs after orthopyroxene. Only a few tiny specks of magnetite associated with the lizardite were found with the electron microprobe. On the other hand, awaruite (Ni-Fe alloy) and wairauite (Co-Fe alloy) are present in grains several mircometers across embedded in the lizardite. Relics of chromite are present elsewhere in the sample. Small amounts of calcic amphibole have formed at the expense of clinopyroxene. No talc was found.
The primary silicate minerals (olivine Fo84Fa16, orthopyroxene Ca1.5Mg84.3Fe14.2, clinopyroxene Ca47.8Mg47.4Fe4.8) are relatively homogeneous (Table 1). The clinopyroxene was not serpentinized, although, perhaps at an earlier stage, it yielded a small amount of magnesiohornblende (Table 1). The XFe of the serpentine varies from 0.08 to 0.17 (Fig. 3), overlapping that of the olivine (0.16) and the orthopyroxene (0.145). Thus, there is a rough mass-balance for Mg/Fe except for a few more Mg-rich lizardite spots. The most Fe-poor lizardite replaced olivine; other than this, the XFe values of lizardite after olivine and orthopyroxene overlap. However, minor elements are more distinct (Table 1); lizardite formed from olivine was found to contain 0.02 to 0.34% NiO, barely detectable Cr, and 0.4% Al2O3, whereas that clearly replacing orthopyroxene has zero Ni, 0.04 to 0.36% Cr2O3, and 1.1% Al2O3. Based on a formula content of 5 total cations, the tetrahedral cations Si+Al/2+Cr/2 in lizardite (Al+Cr is assumed to be charged balanced on T and M; potential tetrahedral Fe3+ is not included) increase from 2.00 to 2.08 with increasing Fe. Similarly, the relative proportions of T to T+M cations increase with increasing Fe (Fig. 4) from 0.40 to 0.418. The y-axis values in this plot, unlike Figure 3, are not influenced by the choice of cation normalization. These compositional trends correspond not to the Fe2+Mg –1 vector, but to the vacancy-balanced substitution of ferric iron, as shown by the calculated line in Figure 4 that starts at a zero Fe3+ composition of 0.15 Fe apfu; this model-based starting point is justified below. If we were to allow some ferric iron on T (as in a ferri-Tschermaks substitution), it would increase T/(T+M) ratios and steepen the slope of the data points.
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FIGURE 3. Number of T cations per 5 total formula cations plotted against Fe/(Fe+Mg) for orthopyroxene, lizardite, and olivine in olivine websterite from the Canyon Mountain complex.
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FIGURE 4. Ratio of T cations (Si+Al/2+Cr/2) to T+M cations vs. total Fe apfu of the lizardite. The calculated line starts from Fe2+ = 0.15 and is for perfect vacancy-coupled substitution of Fe3+.
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The average awaruite on an atom% basis is 70 Ni, 29 Fe, and 1 Co; wairauite is 38 Co, 61 Fe, and 1 Ni. A literal interpretation of the Ni-Fe phase diagram (Kubaschewski 1982) would suggest a formation temperature above about 250 °C. The presence of wairauite presumably reflects the higher Co/Ni whole-rock ratio of pyroxenite as compared to dunite or harzburgite.
Two meta-harzburgites from the Stillwater Complex, Montana (samples 18479, 18480), with primary olivine and orthopyroxene similar in composition to those described here, were reported by Wicks and Plant (1979) to contain yellow, high-Fe lizardite in mesh and bastite locations. Contents of Si+Al/2+Cr/2 in the lizardite vary from 1.89 to 2.15 apfu (from their Tables 2 and 3), which suggests the presence of both cronstedtite and vacancy-balanced substitutions in these samples. When 20–30 sample spots are analyzed, with a sufficient spread in XFe values, it is much easier to distinguish between Fe2+Mg–1 substitution (constant levels of T and M atoms) and Fe3+ substitutions (Figs. 3
and 4).
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ESTIMATE OF FERROUS IRON IN LIZARDITE
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If we assume that the Fe2+/Mg ratio of the lizardite in the sample is appropriate for Fe2+-Mg exchange equilibrium with the olivine and orthopyroxene from which it formed, then we can use partitioning data (KD) from metamorphic antigorite-olivine pairs to obtain a first approximation estimate of the ferrous iron content of the lizardite. The possibility seems remote that the KD for lizardite is greatly different from that for antigorite. It turns out that KD is a function of XFe of the olivine (Evans, in preparation), and the dependence is even more pronounced than for olivine-orthopyroxene and olivine-cummingtonite pairs. For Fa16 olivine, the KD is about 4.0. In this case, the serpentine would have XFe2+ = 0.05, so its formula content of ferrous iron would be 0.15 (Fig. 4). Lizardite compositions with Fe > 0.15 apfu in our sample would then represent increasing substitution of a Fe3+ end-member. Thus, the most Fe-rich lizardite measured, 0.50 Fe apfu, would have a Fe3+/total-Fe ratio of 0.70. This figure is quite normal for natural lizardites in serpentinites. The sloping line in Figure 4 starts at 0.15 Fe apfu, and shows the ideal trend of Fe3+ vacancy-coupled substitution in terms of T/(T+M). Although it is a "model" construction, it fits the independent analytical data quite well.
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CONCLUDING REMARKS
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Natural lizardite in the relatively high-Fe range (~6 to at least 17% Fe end-member) contains appreciable amounts of ferric iron. This is introduced via one or both of two kinds of coupled substitution: (1) ferri-Tschermaks or cronstedtite (charge balanced across T and M sites), and (2) M-site substitution coupled with M-site vacancies. Both substitutions are favored when the environment of serpentinization is a low-temperature one, the fugacity of H2 is low (or fO2 is high), and the fugacity of H2O is high. The latter may vary substantially and even be driven to low values as a result of the serpentinization reaction itself (Sanford 1981). Substitution 2, as in the present case, is likely to be attended by higher values of aSiO2 than substitution 1; conversely, an increase in aSiO2 cannot be what drives substitution 1. Only modestly higher values of fO2 can be entertained in case 2, since these serpentinites contain native Fe-Ni-Co alloys. The compositional correlations shown in Figures 3 and 4 are subtle and certainly best resolved when only data from a single analytical laboratory are included, and potential interlaboratory differences in apfu calculations resulting from standardization, ZAF corrections, and sample preparation are excluded.
Only small amounts of magnetite, if any, may be produced alongside these high-Fe lizardites. Nevertheless, iron from the primary silicates is partially oxidized to Fe+3 and so, if as customary we assume that oxygen is conserved, the reaction will involve hydrolysis as well as hydration, and hydrogen will be evolved. However, only one-third as much hydrogen per mole of added H2O will be released, as a comparison of Equations 2, 3, and 4 shows for fayalite. The same proportions of H2 and H2O hold when ferrosilite component is the reactant.
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Simultaneously operative during serpentinization will be reactions among the Mg end-member minerals and Fe2+-Mg exchange equilibria. Magnetic susceptibility is likely to be low, with the result that usual inverse correlation of density with magnetic susceptibility (e.g., Toft et al. 1990, Fig. 1; Oufi et al. 2002, Fig. 9; Frost and Beard 2007, Fig. 6) will be poor to non-existent. In fact, the diversity of trends found in this diagram may well reflect spatial and temporal differences in the uptake of Fe3+ in lizardite in the course of serpentinization.
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ACKNOWLEDGMENTS
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We thank N.I. Christensen for supplying the sample from the Canyon Mountain complex. Reviews by B.R. Frost, M. Mellini, and F. Wicks helped us to sharpen this paper.
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Footnotes
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MANUSCRIPT HANDLED BY BRYAN CHAKOUMAKOS
MANUSCRIPT RECEIVED May 26, 2009;
MANUSCRIPT ACCEPTED July 8, 2009
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REFERENCES CITED
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Abstract
Introduction
