Journal of Petrology Advance Access published online on October 16, 2008
Journal of Petrology, doi:10.1093/petrology/egn050
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Control of the Products of Serpentinization by the Fe2+Mg–1 Exchange Potential of Olivine and Orthopyroxene
Department of Earth and Space Sciences, Box 351310, University of Washington, Seattle, WA 98195-1310, USA
Received December 3, 2007; Revised typescript accepted September 18, 2008
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONTHE SERPENTINITE SYSTEMEXPERIMENTSWHY MAGNETITE?MICROSTRUCTURESFERROUS AND FERRIC IRON...FURTHER FIELD CONSTRAINTSOPEN-SYSTEM CONDITIONSCONCLUSIONSREFERENCES |
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It is argued that the high-Mg content (mg-number = 95 ± 3) of the serpentine minerals in serpentinized peridotite is a consequence of the environmental Fe2+Mg–1 exchange potential imposed on the system by the abundance of olivine and orthopyroxene. Mass balance in the serpentinization reaction then requires the precipitation of an iron-rich mineral that in most cases is magnetite. This causes hydrogen to be evolved in an oxygen-conserved reaction. The low-variance mineral assemblage Ol + Srp + Brc + Mag sets the chemical potentials of H2O, SiO2 and O2 internally at an early stage in the process, but the paragenetic assessment of serpentinites is rendered difficult by the variable and usually unknown Fe3+ content of the serpentine minerals, particularly lizardite. Whole-rock analyses of highly to completely serpentinized peridotites reveal Fe3+/
Fe ratios > 0·4, with an average value (0·69) similar to that of magnetite (0·67). This feature may be attributed to the presence of high-Fe3+ lizardite, as has been found in Mössbauer spectroscopy studies. Electron microprobe and scanning electron microcope analyses in the literature exhibit element trends (e.g. decreasing Si vs Fe a.p.f.u.) for olivine-pseudomorph lizardite and, with some exceptions, for bastite lizardite, that show a substitution of the cronstedtite component (Fe3+ charge-balanced on T and M sites). Cronstedtite substitution will be favoured at low temperature and/or low hydrogen fugacity, and in these circumstances less magnetite will be evolved during serpentinization, in some cases none at all. Some bastite lizardites from sea-floor settings show evidence of M-site vacancy substitution of Fe3+ for Fe2+. In the course of progressive serpentinization, micrometer to millimeter-scale variations in SiO2 potential may well be present, but their influence on Fe in lizardite seems to be limited to a few cases of lizardite associated with orthopyroxene. Chrysotile is on average more Mg-rich and less variable in Fe/Mg ratio than lizardite, facts that may be attributed to the greater Fe3+ content of lizardite. Chrysotile veins provide the best record available to us of the environmental Fe2+Mg–1 exchange potential in the pore fluid attending serpentinization. This potential serves as a robust control on serpentine and brucite compositions, although it may fail after olivine and orthopyroxene have been armoured or eliminated, and in more open-system environments (high water/rock ratio) such as on the sea floor or at serpentinite host-rock contacts. The default assumption in microprobe analyses that measured iron is all Fe2+ can lead to inappropriate petrological conclusions in the case of serpentinites. KEY WORDS: serpentinization; serpentine; magnetite; FeMg–1 exchange potential; lizardite; chrysotile; cronstedtite; open or closed systems
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONTHE SERPENTINITE SYSTEMEXPERIMENTSWHY MAGNETITE?MICROSTRUCTURESFERROUS AND FERRIC IRON...FURTHER FIELD CONSTRAINTSOPEN-SYSTEM CONDITIONSCONCLUSIONSREFERENCES |
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The recent paper by Frost & Beard (2007
) represents a step forward in our understanding of the processes governing the formation of serpentinite from peridotite. However, in emphasizing the role of silica activity control of magnetite formation in serpentinite, they overlook in my opinion a first-order observation that I believe is relevant to the general issue of serpentinization, specifically the precipitation of magnetite and resulting evolution of hydrogen gas, and in some cases abiogenic methane. Assuming some modest approach to equilibrium (which Frost & Beard did), serpentinite phase assemblages and compositions correspond to fixed or variably constrained values of the local chemical potentials of SiO2, H2O and O2 (or H2), and to the exchange potential of FeMg–1. A comprehensive examination of all four of these potentials should aim to find which, if any, has a potentially controlling influence on reaction products. A fundamental petrological question, surely, is why magnetite should form at all in serpentinites? It does so regularly, if not ubiquitously. The answer is critical to our understanding of the geological, geophysical and geobiological implications of serpentinization that have generated great interest in the last few years.
The vast majority of large-body (massif) serpentinites contain a small percentage of magnetite as well as serpentine minerals that are, with some exceptions, notably richer in Mg/Fe than the original olivine and orthopyroxene. I am not aware of the existence in the literature of any cogent explanation for this general observation. Indeed, it has led to two widespread misconceptions among some in the geological community with respect to the conditions that typically attend the serpentinization of peridotite. The first is that it involves the addition of oxygen from an oxidizing fluid. It is true that whole-rock ferric/ferrous ratios increase as a result of serpentinization and the growth of magnetite. However, the serpentinization process itself witnesses a lowering of oxygen fugacity to close to the IM (iron–magnetite) buffer (Frost, 1985), and there are good reasons for writing serpentinization reactions as oxygen conserved. It may be judged the more likely geochemical process (Greenwood, 1975) and is consistent with the evidence for hydrogen gas in spring waters and vents associated with the serpentinization of peridotite (e.g. Charlou et al., 2002; Sleep et al., 2004; Kelley et al., 2005; Bach et al., 2006; Seyfried et al., 2007). The addition of H2O and corresponding release of H2 combine hydration and hydrolysis, and accomplish petrographically the same thing as simply adding O2. This view is not meant to rule out possible oxygen addition in those instances of serpentinization that may have been driven by major infiltration of a near-surface fluid or one derived from more oxidized country rocks nearby.
The second misconception is that the growth of serpentine minerals richer in Mg/Fe than the original olivine and orthopyroxene results from the intrinsic instability of more Fe-rich serpentine. It is true that the replacement of Mg by Fe2+ in ordinary serpentines expands the M layer and worsens the misfit with T layers. However, the coupled, heterovalent substitutions of Al3+ and Fe3+ on M and T sites have both been shown to stabilize the lizardite structure (Caruso & Chernosky, 1979; Mellini, 1982; Mellini & Viti, 1994). Whereas XFe [= Fe/(Fe + Mg)] of mesh-pseudomorph serpentine is typically 0·02–0·08 and seldom exceeds 0·10, there are cases of vein and bastite serpentine, and antigorite, that exceed this value (up to at least 0·21); for example, in layered ultramafic intrusions and kimberlites. Furthermore, serpentine minerals with as much or more Fe as Mg (ferro-antigorite and greenalite, both modulated sheet silicates) occur in low-temperature ironstones and hydrothermal deposits.
In the following I will review relevant aspects of the current state of knowledge of serpentinization as a process, and conclude with the inference that the overriding chemical factor governing the identity and composition of reaction products is the Fe2+Mg–1 exchange potential 
µ(Fe2+Mg–1) in the local environment. This postulate accounts for the growth of magnetite and evolution of hydrogen. Large and variable amounts of Fe3+ occur in strongly serpentinized peridotite, and in the lizardite of which it is principally composed. We explore here the equilibria and possible controls on Fe3+ uptake in lizardite; they have relevance to the yields of magnetite and hydrogen accompanying serpentinization.
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THE SERPENTINITE SYSTEM |
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TOPABSTRACTINTRODUCTIONTHE SERPENTINITE SYSTEMEXPERIMENTSWHY MAGNETITE?MICROSTRUCTURESFERROUS AND FERRIC IRON...FURTHER FIELD CONSTRAINTSOPEN-SYSTEM CONDITIONSCONCLUSIONSREFERENCES |
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Serpentinization of dunite can be represented in the simple system MgO–SiO2–H2O by the reaction
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| (1) |
). The univariant assemblage expressed by the serpentinization reaction (1) should be depicted as an invariant point in any diagram at fixed P (cf. Frost & Beard, 2007).
More realistically for natural serpentinites, we should consider the four-component system MgO–FeO–SiO2–H2O. Now, at arbitrarily fixed P and T, there are five possible assemblages each with four phases. One of these assemblages is Ol + Srp + Brc + H2O. Iron/magnesium partitioning among the three minerals in this assemblage results in a divariant field on the phase diagram that probably has a width of some tens of degrees (Moody, 1976a).
In an idealized isobaric equilibrium model of falling temperature, the progressive serpentinization of Fa9–10 dunite would show a temperature interval where (in order of increasing XFe) serpentine, olivine, and brucite coexist. Mg/Fe partitionings are given in the literature (Trommsdorff & Evans, 1972; Evans & Trommsdorff, 1975, 1977; Dungan, 1977; Vance & Dungan, 1977; Worden et al., 1991). Because serpentine is considerably more abundant than brucite in the reaction, mass balance requires all minerals to increase in XFe as the reaction proceeds. Near its completion, the small amount of residual olivine might be as rich in iron as Fa17. The fractionation effect is potentially stronger in the case of meta-harzburgite because this composition does not produce the more iron-rich brucite along with the serpentine. Kunugiza (1982) showed that fractionation in the system along these lines was operative and the cause of zoning in olivine during 400–500°C prograde metamorphism of serpentinite in the Ryumon peridotite, Japan. Similarly, upgrade metamorphism of the antigorite + olivine Malenco serpentinite (Worden et al., 1991) produced progressively more Mg-rich olivine and antigorite (XFe = 0·144 
0·087 and XFe = 0·069 0·047, respectively).
Unsurprisingly, there is no evidence from nature that the retrograde serpentinization reaction proceeds at equilibrium along a fractionation path corresponding to the isobaric T – XFe phase loop (e.g. Kunugiza, 1982, fig. 4). It would require rates of element diffusion and mineral recrystallization that are extremely unlikely at the low temperatures of serpentinization. Stable-isotopic, petrological and geological evidence [reviewed by Evans (2004)] has shown that the reaction boundary (1) in many instances in nature is overstepped (down temperature) by 100 or 200°C or more. This fact alone does not necessarily signify that the reaction proceeds irreversibly, because we can reasonably allow for the activity of H2O in the reaction assemblage to fall with temperature (e.g. Sanford, 1981; Frost & Beard, 2007, fig. 4a) and thereby maintain equilibrium. However, the lack of any evidence that the reaction in nature is a sliding one (with a phase loop) requires us to conclude that it does take place irreversibly. The reaction assemblage at around 200°C implies not only a low pressure of H2O but also very reducing conditions and very low SiO2 activity (Hemley et al., 1977); for example, 10–3·5 relative to a quartz standard state (Frost & Beard, 2007). Peridotite at these temperatures is a significant sponge (sink) not only for H2O but also for SiO2. Although this is also true for O2, the absence of a potential gradient capable of transporting oxygen means that, to maintain O–H equilibrium, the introduction of H2O has to be offset by the evolution of H2, and so the progress of the basic reaction is best described (e.g. Thayer, 1966; Hostetler et al., 1966; Moody, 1976b), without stoichiometric coefficients, as a hydration + hydrolysis reaction:
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| (2) |
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EXPERIMENTS |
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TOPABSTRACTINTRODUCTIONTHE SERPENTINITE SYSTEMEXPERIMENTSWHY MAGNETITE?MICROSTRUCTURESFERROUS AND FERRIC IRON...FURTHER FIELD CONSTRAINTSOPEN-SYSTEM CONDITIONSCONCLUSIONSREFERENCES |
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Janecky & Seyfried (1986
) performed laboratory experiments at 200°, 300°C, and 500 bars reacting peridotites with seawater at 10 : 1 water/rock ratios. Their solid products included lizardite (XFe from 0·06 to 0·09), brucite, and magnetite. In a later experiment on lherzolite at 200°C and 500 bars with seawater in nearly 1 : 1 ratio with the rock (Seyfried et al., 2007), magnetite was barely detectable and appreciable ferric iron was found in the lizardite. Hydrothermal experiments on iron-bearing olivine (Fa7) at 2 kbar on the IM buffer by Moody (1976a) also produced magnetite along with serpentine and brucite.
A hydrothermal flow-through experiment by Normand et al. (2002) showed that air-saturated deionized water injected at 300°C and 300 bars into crushed San Carlos olivine (Fa9) underwent an initial rise followed by a dramatic loss of Si to below the detection limit in the first 100 hs of fluid flow; the latter was attributed to the incipient formation of serpentine in an irreversible hydration reaction. Steady-state values of Mg and pH in the fluid after 550 h were interpreted to reflect saturation with brucite as well as serpentine [as identified by X-ray diffraction, scanning electron microscopy (SEM) and high-resolution transmission electron microscopy (HRTEM)]. Octahedra of magnetite were also produced, signifying a major decline in fO2 in the fluid to that consistent with the reaction products. Energy-dispersive (EDS) analysis of the serpentine minerals (lizardite and chrysotile) showed a composition averaging 
96% Mg end-member. This experiment showed that reaction of Fe-bearing olivine with air-saturated water produces on a laboratory time-scale a detectable five-phase reaction assemblage (Ol + Mg-rich Srp + Brc + Mag + H2O), with attendant major changes in fluid chemistry in line with thermodynamic calculations. As shown below, this result is in every respect an accurate reproduction of the serpentinization reaction in nature, despite an experimental flow rate perhaps 5–7 orders of magnitude faster than fluid fluxes during serpentinization in nature (MacDonald & Fyfe, 1985; Normand et al., 2002), and an H2O pressure that was higher than is probably the case at the reaction site in nature.
The serpentinization of harzburgite was studied by Palandri & Reed (2004) at 25°C to 300°C in a series of calculated water–rock reaction simulations. Product minerals were serpentine, brucite and magnetite. At low water/rock ratios, the reacted water became highly alkaline and strongly reducing, and rich in dissolved calcium. These conditions are well-known accompaniments of serpentinization in the field, including the opportunity they provide to convert dyke-rocks into rodingite. Thus, the essence of the serpentinization reaction, the growth of serpentine and brucite, accompanied in most cases by magnetite, and the evolution of H2 gas, appears to be well supported by experiment and calculation.
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WHY MAGNETITE? |
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TOPABSTRACTINTRODUCTIONTHE SERPENTINITE SYSTEMEXPERIMENTSWHY MAGNETITE?MICROSTRUCTURESFERROUS AND FERRIC IRON...FURTHER FIELD CONSTRAINTSOPEN-SYSTEM CONDITIONSCONCLUSIONSREFERENCES |
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Why does serpentinization so commonly lead to the precipitation of magnetite, as generally seen in thin section and demonstrated by measurements of magnetic susceptibility in natural samples, and reproduced experimentally? If all the iron from the olivine can be accommodated in serpentine and brucite (the equilibrium phase-loop model), there is no need for another iron mineral to form. If we want to allow some Fe to occur in redox states other than 2+, we must now consider the system FeO–MgO–SiO2–H2O–H2. (H2 is chosen as a component rather than O2 because its partial pressure is experimentally measurable, and it demonstrably participates in reaction progress.) At fixed P and T this system allows five-phase assemblages, such as Ol + Srp + Brc + Mag (or iron) + H2O. However, this still does not tell us why in most cases we obtain magnetite (± iron), and sometimes not. Based on calculated semi-quantitative phase diagrams, Frost & Beard (2007
) argued that magnetite will form at very low silica activities accompanied by Mg-rich serpentine, whereas magnetite is not stable in the system at higher SiO2 activities, where the serpentine mineral may be more iron-rich. The critical constraining reaction for their hypothesis [Frost & Beard, 2007, reaction (23)] is
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| (3) |
Reaction (3) links three chemical species in the intergranular fluid. Theoretically, the reaction can be driven to the left to eliminate magnetite by increasing the potentials of SiO2 or H2O or by decreasing the potential of O2. However, the chemical potentials of all three fluid species are fixed when all five phases (Ol, Srp, Brc, Mag, fluid) are present in equilibrium and P and T are held constant. These potentials cannot change until reaction progress has responded to the influx or outflow of any or all of these species by the loss of a phase, and this might be brucite, olivine, or magnetite. We are free to choose a change in the concentration of any one of these species to bring this about. An increase in the chemical potential of SiO2, of course, is likely to eliminate brucite. An increase in the potentials of either SiO2 or H2O or both, according to reaction (3), can lead to the loss of magnetite and growth of serpentine richer in iron; in other words, to an increase in µ(Fe2+Mg–1). However, at the same time, µ(Fe2+Mg–1) of olivine (or orthopyroxene) will decrease according to the reaction
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| (4) |
). Thus equilibria (3) and (4) cannot be perturbed simultaneously without inducing exchange disequilibrium.
The averages of chemical analyses in the literature of serpentine from massif serpentinites show that for lizardite XFe = Fe/(Fe + Mg) = 0·058 ± 0·030 (n = 460) and for chrysotile XFe = 0·034 ± 0·018 (n = 198). Uncertainties are 1
. The frequency histogram for lizardite (not chrysotile) is asymmetric, with a more Fe-rich shoulder and a mode value (XFe = 0·04) smaller than the average. I use ‘massif’ as synonymous with the descriptive terms alpine-type, orogenic, ophiolitic, mantle, and abyssal. If the process of serpentinization is basically isochemical for Si, total Fe, and Mg [and there have been many studies to show this generally to be the case; see the review by O’Hanley (1996); also Shervais et al. (2004)], the growth of serpentine (with or without small amounts of brucite) from olivine and orthopyroxene with the typical XFe = 0·09–0·10 composition requires the precipitation of an additional Fe-rich mineral for mass balance. This is generally magnetite (or nickel–iron). Texturally equilibrated metamorphic olivine and antigorite in high-grade serpentinites (Trommsdorff & Evans, 1972; Dungan, 1977; Evans & Trommsdorff, 1977; Worden et al., 1991) has antigorite (XFe 
0·05) coexisting with olivine (XFe 0·10); the exchange KD [= (Fe/Mgsrp)/(Fe/Mgol)] is 0·45 (not allowing for unknown Fe3+ in the antigorite). We can reasonably expect partitioning with respect to the other serpentine minerals to be similar. This suggests that the growth of serpentine minerals in mantle peridotite in the range XFe 0·02–0·08 is not coincidental, but the result of an attempt to achieve the same Fe2+Mg–1 exchange equilibrium as in the olivine–antigorite pairs.
The progressive serpentinization of peridotite produces mostly SiO2-undersaturated minerals: not only serpentine and brucite, but also less common minerals such as hydrogrossular, Ti-andradite, perovskite, corundum, clinohumite, and native Fe–Ni–Co metals, as well as SiO2-undersaturated minerals in related rocks such as rodingites. The native metals (e.g. Rossetti & Zucchetti, 1988) and occasional graphite (Pasteris, 1981) also indicate very low redox states. This assembly of minerals is fully consistent with quantitative phase-equilibrium calculations, which show that the chemical potentials of the Fo and Fa components of reactive olivine in mantle peridotite induce conditions of low fO2 (approaching IM) and low a(SiO2) in low-temperature serpentinization (e.g. Hemley et al., 1977; Frost, 1985; Frost & Beard, 2007). Low SiO2 concentrations are also a feature of the fluid accompanying serpentinization in experiments (Hemley et al., 1977; Janecky & Seyfried, 1986; Normand et al., 2002).
A logical extension of the statements above is that the difference in the Fo and Fa potentials, the exchange potential µ(Fe2+Mg–1), also governs the mineral growth environment at the site of serpentinization. That is, µ(Fe2+Mg–1) exerts a control on the mole-fraction X(Fe2+) of the product ferromagnesian minerals serpentine and brucite.
The exchange potential is imposed on the site of alteration by all the surrounding, not-yet-altered olivine and orthopyroxene, and this control is likely to be most powerful at the outset when there is abundant olivine and orthopyroxene in the rock. Along with the phase equilibrium constraints on fO2 (or fH2), fH2O and aSiO2 supplied by the five-phase assemblage, the crystallization of magnetite (or native iron) is inevitable by this model. As long as the influence of the exchange potential prevails, the product serpentine will continue to grow with an Mg-rich composition, yielding more magnetite. This differs from the process of equilibrium fractional crystallization where the exchange potential would change with reaction progress. Sluggish kinetics apparently prevents this from happening; even close to the site of serpentinization, relict olivine seems always to retain its original composition. Olivine yields by incremental dissolution rather than by reaction; that is, it does not supply the forsterite component to make serpentine and brucite by becoming marginally enriched in fayalite. The bulk rock increases in the ratio Fe3+/Fe2+, utilizing oxygen from the added H2O, and H2 is liberated. Exchange equilibrium (or the KD effect) is the clue to the fundamental question posed at the head of this section.
It is almost axiomatic that an infiltrating aqueous fluid, especially if it is pervasive rather than channeled, will eventually take on the properties of the rock through which it has flowed. In the case of serpentinite these properties are likely to include pH, fO2, and silica activity (and other species), as well as the relative potentials of Fe2+ and Mg. Serpentinite-related fluids have such distinctive geochemical properties that they can be readily recognized in springs on land and on the ocean floor.
Reversal of the magnetite-forming serpentinization reaction in nature would require the unlikely reintroduction of H2. Thus, by this model the ferric iron is effectively sequestered in magnetite (and serpentine; see below). For this reason, metaperidotites that have formed by deserpentinization tend to contain abundant magnetite and olivines with less iron than the primary Fa9–10, down to as little as Fa2 (e.g. Gabrielse, 1963; Trommsdorff & Evans, 1969; Dungan, 1977; Pinsent & Hirst, 1977; Vance & Dungan, 1977; Blais & Auvray, 1990; Nozaka, 2003; Hattori & Guillot, 2007). With the exhaustion (or armoring) of original olivine and pyroxene, the µ(Fe2+Mg–1) evidently decreases in this process. Likewise, recrystallized and prograde metamorphosed serpentinites may have very magnesian serpentines. For example, in the completely serpentinized and recrystallized Cassiar serpentinite, British Columbia, the average content of the Fe2+-serpentine end-member is less than 1 mol % (O’Hanley & Dyar, 1993).
At the onset of a major serpentinization event, the potentials of O2, H2O, and SiO2, theoretically at least, are fixed at any given P and T by the phases already present: the potential of SiO2 by the coexistence of olivine and orthopyroxene, anthophyllite, or talc; the oxygen fugacity by the coexistence of MgFe-olivine and spinel (chromite, ferrit-chromite, or magnetite); and the H2O potential by the coexistence of olivine with incipient serpentine and brucite or talc. At constant P and T, as described by Frost & Beard (2007), the potentials of SiO2 and O2 (these being species in low concentration) will quickly converge on those for the assemblage Ol + Srp + Brc + Mag + H2O—unless the introduction of H2O-rich fluid is sufficiently massive to induce changes in rock composition by mass transfer of elements with the fluid (open-system behavior sufficiently strong to change the system variance). The open-system experiment of Normand et al. (2002) described above was not of sufficient duration to change the system variance (and eliminate a phase).
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MICROSTRUCTURES |
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TOPABSTRACTINTRODUCTIONTHE SERPENTINITE SYSTEMEXPERIMENTSWHY MAGNETITE?MICROSTRUCTURESFERROUS AND FERRIC IRON...FURTHER FIELD CONSTRAINTSOPEN-SYSTEM CONDITIONSCONCLUSIONSREFERENCES |
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Serpentinites are complex microstructurally (Wicks & Whittaker, 1977
; Cressey, 1979; Wicks & Plant, 1979; O’Hanley, 1996; Viti & Mellini, 1998; Auzende et al., 2002; Baronnet & Belluso, 2002; Rumori et al., 2004). Constituent minerals are very fine-grained (in some places amorphous), they vary in crystal quality, are rich in stacking and other defects, and are not always uniform in composition. Although this means that we should question any formal application of equilibrium phase relations, it does not imply that we should refrain from being guided by it. Evidence for multiple stages of alteration has been presented in a number of papers, and the reasonable suggestion made that they are associated with changes in fluid flux or fluid source (e.g. Wicks & Plant, 1979; Viti & Mellini, 1998; Bach et al., 2006). Dissolution and recrystallization of serpentine and brucite might also be expected, especially in mesh rims (Cressey, 1979; Rumori et al., 2004) and perhaps in bastite, making it hard to reconstruct the evolution of serpentine compositions. Olivine rims are patently the first parts to undergo alteration, but space/time relations can be affected by recrystallization and the influx of different fluids at various times. Disequilibrium may also be evidenced; for example, by a major change in apparent KD for Brc/Srp between the two events, as described by Bach et al. (2006). Several workers have adopted the not unreasonable view that serpentinization in space and time at any locality is a continuum between rock-dominant and fluid-rich stages. The former will be characterized by low-variance assemblages and volume expansion, and generally be associated with the presence of magnetite and low-Fe2+ serpentine, the evolution of hydrogen, and the development of rodingites. The latter can be ‘open system’ and in principle accompanied by significant mass transfer (of Si, for example), favoring the production of high-variance assemblages (in some cases monomineralic), more variable mineral compositions [relaxation of control by the µ(Fe2+Mg–1) of olivine] and less volume expansion (Evans, 2004). |
FERROUS AND FERRIC IRON IN SERPENTINE |
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TOPABSTRACTINTRODUCTIONTHE SERPENTINITE SYSTEMEXPERIMENTSWHY MAGNETITE?MICROSTRUCTURESFERROUS AND FERRIC IRON...FURTHER FIELD CONSTRAINTSOPEN-SYSTEM CONDITIONSCONCLUSIONSREFERENCES |
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A range of serpentine compositions is found in massif peridotites and serpentinites, but most of them have XFe well below 0·10. The small proportion of serpentines that have XFe larger than 0·10 (Frost & Beard, 2007
) tend to be associated with bastite after orthopyroxene and serpentine in veins. The serpentine forming the rims and much of the cores of mesh pseudomorphs of olivine is lizardite, but it is far from uniform in XFe. At this point, however, we have to recognize that a deeper understanding of the paragenetic relations of serpentinites, especially as they relate to the iron content of serpentine, has to confront the issue of a ferric iron component. A critical assessment of classical analyses of the serpentine minerals by Whittaker & Wicks (1970) showed that lizardite, as recognized by Page (1967), tends to be enriched in Fe3+. This fact was later confirmed by Mössbauer spectroscopy: O’Hanley & Dyar (1993) found that 21 lizardite samples from serpentinized peridotite in British Columbia, Quebec and Australia contained 0·33–0·88 Fe3+/Fe (average of 0·59), and Votyakov et al. (1993) showed that lizardite in 23 samples of metadunite and metaharzburgite in the Urals fell in the range 0·43–0·86 Fe3+/Fe (average of 0·67). Values of this ratio in excess of 0·67 (as in 40% of the O’Hanley & Dyar samples and 65% of the Urals samples) cannot be attributed to contamination by magnetite. Chrysotile in veins in the same serpentinites (O’Hanley & Dyar, 1998) was found to contain smaller proportions of ferric iron (Fe3+/Fe mostly 0·22–0·41). Thus, when we compare Fe in serpentine obtained by electron-microprobe analysis of samples from different locations, we cannot interpret the differences exclusively in terms of Fe2+. At most, the Fe content of serpentine sets an upper limit for Fe2+.
Whole-rock analyses in the literature of strongly serpentinized ultramafic rocks (with more than 10 wt % H2O) show ratios of Fe3+ to Fe in the range 0·4 to >0·9 (Fig. 1). The serpentinization of forsterite by reaction (1) produces an equimolar mix of serpentine and brucite that theoretically contains 13·9 wt % H2O, whereas a metaharzburgite without brucite will theoretically contain 12·4 wt % H2O. Figure 1 shows that typical proportions of ferric to total iron in highly serpentinized peridotite are virtually identical to that of ideal magnetite (0·67). This fact was already known to Thayer (1966, fig. 1), whose average of 60 analyzed serpentinites contained 1· 3% of CIPW normative hematite. The high Fe3+/Fe values are too robust to be explained by the occasional presence of accessory andradite, or by small amounts of Mg replacing Fe2+ in the magnetite, or by the failure to dissolve a per cent or so of accessory chromite.
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From the data used to plot Fig. 1, the average highly serpentinized massif peridotite contains 4·7 wt % Fe2O3 and 2·3 wt % FeO (equivalent to roughly Fa9 olivine). If all the Fe2O3 were contained in magnetite (giving 6·8 wt % modal magnetite), then the remaining 93% serpentine and brucite in the average rock would have only 0·2 wt % of the rock's FeO and thus would be virtually Mg end-member in composition. Figure 1 shows, on the other hand, that roughly 50% of highly serpentinized peridotites have Fe3+/
Fe ratios greater than that of magnetite. The excess of ferric iron can only stem from the serpentine mineral (principally lizardite) that modally overwhelms the small amounts of magnetite present. Thus, more than one-half of the rock's Fe2O3 is contained in the serpentine. This conclusion is entirely consistent with the Mössbauer data cited above. With this in mind, we can now correct the average lizardite XFe (total Fe) of 0·058 down to about 0·03 for the mol fraction of the Fe2+-serpentine component. Similarly, the empirical KD for olivine–antigorite pairs cited above (0·45) should probably be lowered to about 0·35 to correct for the presence of Fe3+ in the antigorite.
Lizardite and chrysotile in massif serpentinites, as cited above, differ strikingly in their averages and standard deviations of XFe. Although these minerals differ in their mode of occurrence and in their crystal chemistry, as well possibly as in their timing and temperature of formation, it seems very likely that the differences relate in large measure to the greater Fe3+ content of lizardite. It is well known that the lizardite structure can accommodate more R3+ cations such as Al and Cr than chrysotile, and that these cations help to minimize the misfit between the M and T sheets in the planar structure. Thus, if we want a monitor of the environmental exchange potential µ(Fe2+Mg–1) in serpentinites, we should focus on the composition of the chrysotile, which is mostly in veins. (Think of a nurse sampling the blood in your veins to determine chemical levels in the rest of your body.) Even then, the measurement will be high because chrysotile itself is not entirely free of ferric iron. Thus, the average mol per cent of the Fe2+-serpentine component in the chrysotile of massif metaperidotites will be slightly less than 0·034, much as we have proposed for lizardite.
If Fe3+ accounts for the high iron content of lizardite, and its substitution is charge-balanced between the M and T sites of the lizardite structure (Wicks & Plant, 1979), it can be described in terms of a ferri-Tschermak substitution. This substitution is represented by the vector Fe3+2(R2+Si)–1 in projection from H2O and O2 on the MgO–FeO–SiO2 diagram (Wicks & Plant, 1979, fig. 1). It may be expressed in terms of an equilibrium that involves only magnetite, fluid, and the iron-components of serpentine, namely ferroan lizardite and cronstedtite:
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| (5) |
µ(Fe2+Mg–1), this reaction can be represented by ‘equipotential’ lines on the MgO–FeO–SiO2 projection (Fig. 2). These kinds of lizardite, which have Si 
2·0 a.p.f.u., form the edge of a phase volume in the unprojected system that includes brucite, olivine, and magnetite. Increase in Fe3+ in them will not be occasioned by an increase in the SiO2 activity.
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The 1 : 2 correlation for variation towards the cronstedtite end-member [equation (5)] should ideally show as a line with a slope of –0·5 on a graph of Si a.p.f.u. vs
Fe a.p.f.u. Natural lizardites that have replaced olivine (mesh, hourglass, recrystallized) contain modest amounts of Al and very little Cr. Figure 3 shows the variation of Si + Al/2 vs Fe a.p.f.u. for a global population of olivine-pseudomorph lizardites taken from the literature. The ordinate assumes that Al is distributed 1 : 1 on the M and T sites. The negative slope of 0·39 in Fig. 3 shows the overriding importance of the cronstedtite substitution, which means that most of the measured variation in Fe is related to Fe3+ rather than to Fe2+. Detailed inspection of the data suggests that interlaboratory differences contribute in part to the scatter in Fig. 3. Excluded from Fig. 3 are analyses clearly contaminated by brucite as judged from their low anhydrous totals and low T/(T + M) ratios (Fig. 4). A trend line in Fig. 3 produced simply by contamination with brucite (e.g. XFe = 0·3) would have a slope of –1· 4; a less Fe-rich brucite contaminant (e.g. XFe = 0·15; Evans & Trommsdorff, 1975) would have a negative slope twice as steep. The most Fe-rich sample in our dataset is an analysis of a lizardite-1T mesh rim from the Lizard harzburgite (Wicks & Plant, 1979, FW-L-4), which has an anhydrous total of 88·15 wt % and XFe = 0·128, and a formula that can be expressed as |
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) lizardites averages 1· 1 (basically equal site partition, allowing for experimental uncertainty), whereas IVFe3+/VIFe3+ in the Votyakov et al. (1993) dataset is much smaller (0·28). The latter would require a contribution to overall charge-balance to come from elsewhere, such as from cation vacancy (see below) or from partial dehydrogenation (O = 5 + y, OH = 4 – y p.f.u.) (Fuchs et al., 1998). According to reaction (5), it would appear that a limit to cronstedtite substitution is reached when the supply of magnetite is exhausted. If the cronstedtite-rich serpentine is a direct product of the H2-liberating serpentinization reaction, it will be the olivine composition (plus a small contribution of Fe from altered chromite) that is limiting.
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Figure 4 examines the stoichiometry of lizardites from olivine-pseudomorphs in terms of total T- and M-site occupancies. It is constructed by assuming a constant ferrous iron content of 0·08 a.p.f.u. on M, so that all variation in iron is in Fe3+, which, like Al, is distributed equally between T and M. The stoichiometry overall shows no dependence on the total Fe content. Most samples fall between T/(T + M) = 0·39 and 0·41, straddling the ideal value of 0·4, but clearly some have higher values. These latter samples also plot above the best-fit line in Fig. 3 (in their absence a slope closer to –0·5 would have been obtained). Assuming that the analyses are accurate, these lizardites show M-site deficiency (Fe3+ on M not fully coupled with Fe3+ on T) as well as ferri-Tschermak substitution. This tendency shows up dramatically in some bastite lizardites (see below).
Equilibrium (5) may well explain the observations of recrystallization reactions among serpentine minerals in the Cassiar, Woodsreef, and Jeffrey serpentinites, which release magnetite in some cases and consume it in others (O’Hanley & Wicks, 1995; O’Hanley & Dyar, 1998). Because reaction (5) involves net volatile release on the left-hand side, higher temperatures of serpentinization should favour magnetite + low-iron serpentine. Experimental work on the serpentinite reaction by Moody (1976a, 1976b) found that magnetite occurred only at the higher temperatures of alteration. A magnetite-absent, high Fe-serpentine product will be favoured when H2 fugacity is low (or oxygen fugacity is higher). Temporal fluctuations in the fugacity of H2 at the reaction site can of course drive reaction (5) one way or the other. For example, an increase in effective water/rock ratio proportionally decreases the H2 content of the fluid (Allen & Seyfried, 2004, table 1), driving reaction (5) to the right.
The yield of hydrogen in the overall mass balance [equation (2)] for the serpentinization of dunite will depend on the composition of serpentine that forms. Incorporation of the cronstedtite component renders the serpentine poorer in Si, so that less brucite, a host for hydrogen, forms in equation (2), and more H2 is evolved. With the uptake of Fe3+ in lizardite, the yield of magnetite can be reduced to zero, as is sometimes observed in nature, but nevertheless H2 is released. In the case of (brucite-free) meta-harzburgite, the yield of H2 (and magnetite; Coleman & Keith, 1971) is basically high at the outset (Evans, 2004, p. 491).
Wicks & Plant (1979, fig. 1) also proposed that ferric iron might enter lizardite via a substitution involving M-site vacancy formation at constant Si a.p.f.u., as we have hinted at above. This produces a dioctahedral ferri-lizardite component, the ferrian analogue of kaolinite:
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| (6) |
; Le Gleuher et al., 1990; Shervais et al., 2004).
Most bastite lizardites (Fig. 5) in fact show T/(T + M) ratios (0·39–0·43) that are not very different from those of the olivine-pseudomorph lizardites (Fig. 4). However, with ocean-floor bastites separately identified (Uehara & Naka, 1983; Agrinier et al., 1988, 1996; Hébert et al., 1990; D’Antonio & Kristensen, 2004), we can see (Fig. 5) that among these there is a divergent trend towards increasing apparent non-stoichiometry. This trend is in part an artefact of the assumption in the plot that Fe3+ is equally distributed on T and M sites. Values of T/(T + M) 0·48 are absurdly high in comparison with 0·5 in end-member ferri-lizardite, which has 2 a.p.f.u. Fe3+. The divergence can be substantially lowered, but not totally eliminated, if the ferri-Tschermak substitution is withdrawn from the ordinate value in Fig. 5 and all Fe3+ is placed on the M sites. However, many of the bastite analyses of Agrinier et al. (1988, 1996) and D’Antonio & Kristensen (2004) have values of T/(T + M) that remain as high as 0·42–0·44 when this is done (and random high data-points rejected as talc contaminated). Antigorite has on average T/(T + M) = 0·414, so there is no possible confusion here. These observations indicate that the high Fe in some oceanic bastites is related to vacancy substitution [equation (6)] rather than to ferri-Tschermaks substitution [equation (5)]. Samples drilled through ultramafic clasts in the mud volcano of the South Chamorro Seamount, Marianas (D’Antonio & Kristensen, 2004), account for all the data-points above 0·38 Fe a.p.f.u. in Fig. 5. They follow the trend of the upper set of equipotential lines on the triangular FeO–MgO–SiO2 projection (Fig. 2). We should note that the actual values of Fe a.p.f.u. plotted in Fig. 5 for the non-stoichiometric bastite serpentines are strictly speaking incorrect in that they have been computed with a five-cation formula unit. The seven (anhydrous) oxygen formula contents of the M-site deficient serpentines cannot be calculated without separate analysis of Fe2+ and Fe3+. The analysis of the most Fe-rich bastite sample in Fig. 5, number 9-6-5 from D’Antonio & Kristensen (2004, table 1), can be recast to seven oxygens on the basis of either all ferrous or all ferric iron:
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The first requires all R3+ cations to reside on M sites as well as minor Si. Both formulae are M-cation deficient (0·085 and 0·291, respectively). The second is likely to be closer to the truth based on crystal chemical logic (T–M site misfit and cation deficiency). Most of the Chamorro bastite serpentines are described in thin section as having a green colour and high birefringence.
The occurrence of these cation-vacancy Fe3+-rich lizardites can probably be attributed to reaction with very low-temperature ocean water under conditions of high water/rock ratio at near-surface sites (e.g. Uehara & Naka, 1983; Agrinier et al., 1988; Evans & Baltuck, 1988; D’Antonio & Kristensen, 2004). The more standard cronstedtite trend may be associated with serpentinization taking place at greater depths (Mével, 2003) under lower water/rock ratios. According to reaction (6), ferri-lizardite substitution is encouraged by an increase in the SiO2 potential. Although this appears to support Frost & Beard's (2007) proposal regarding the influence of SiO2 activity on serpentine composition, equation (6) promotes Fe3+ in lizardite rather than Fe2+ [reaction (3)]. Compositional variation according to reaction (3) in the direction of ferroan lizardite would take the form of a horizontal line (constant Si a.p.f.u.) in Figs 3 and 6; Fig. 6 is discussed below. If what we see in these figures is a combination of reactions (5) and (3), then for geometric reasons the amount of additional Fe2+ involved (left to right across the diagram) could only be about one-fifth of the total amount of additional Fe3+. With the model proposed here, the slope of the best-fit lines in Figs 3 and 6 should ideally flatten sharply to zero below 0·08 Fe a.p.f.u. (all ferrous iron).
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Figure 6 includes only bastite lizardites that, according to Fig. 5, are approximately stoichiometric. To accommodate the higher Cr as well as Al in bastite as compared with olivine-pseudomorph lizardite, ordinate values are (Si + Al/2 + Cr/2) a.p.f.u. The population includes both continental and oceanic occurrences. The range in XFe is large, but now a trend toward cronstedtite enrichment (slope = –0·41) can be discerned. It is practically the same as for olivine-pseudomorph lizardites (Fig. 3). The possible influence of brucite contaminant on this slope is unlikely in the case of bastite lizardite. The appropriateness of the choice of ordinate variable is arguable, but the slope of the fit is little changed by another choice. This comparison shows that the environmental geochemical factors attending the serpentinization of orthopyroxene are, with the few exceptions described above, not detectably different from those accompanying olivine serpentinization.
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FURTHER FIELD CONSTRAINTS |
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TOPABSTRACTINTRODUCTIONTHE SERPENTINITE SYSTEMEXPERIMENTSWHY MAGNETITE?MICROSTRUCTURESFERROUS AND FERRIC IRON...FURTHER FIELD CONSTRAINTSOPEN-SYSTEM CONDITIONSCONCLUSIONSREFERENCES |
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A recurring petrographic observation in many natural serpentinites is that little or no magnetite seems to form in the beginning stages of alteration (Page, 1967
; Ikin & Harmon, 1983; Toft et al., 1990; O’Hanley & Dyar, 1993; Bach et al., 2006, fig. 3; Frost & Beard, 2007, fig. 6). Toft et al. (1990) noted that, at seven sub-sea-floor sites of abyssal peridotite, the early stages of serpentinization, as indicated by high grain-density measurements, are not accompanied by a corresponding increase in bulk magnetic susceptibility; that is, there is no formation of magnetite. Similarly, Oufi et al. (2002, fig. 11) noted that the Fe content of serpentine was high (5–7 wt % as FeO) in samples of high-density serpentinite (low modal per cent serpentine) from sea-floor drill sites, and lower (2–3·5 wt % FeO) in strongly serpentinized samples. Because the early formed iron-rich serpentine was physically replaced by later-formed iron-poor serpentine, they inferred progress of a magnetite-forming reaction. Figure 1 shows that serpentinization at the outset is accompanied by an increase in the whole-rock Fe3+/Fe ratio, notwithstanding any failure to form magnetite; evidently, the ferric iron is entering the serpentine. equation (5) can potentially provide a simple explanation for these observations if serpentinization is initiated at a comparatively low temperature. Some temperature increase in the course of serpentinization might be anticipated from the strongly exothermic nature of the serpentinization reaction ( 35 kJ/mol H2O), as discussed by many workers (e.g. MacDonald & Fyfe, 1985; Peacock, 1987b; Fyfe, 1990; Bach et al., 2002; Lowell & Rona, 2002; Früh-Green et al., 2003). The effectiveness of this heat source is limited in the case of hydrothermal systems with their substantial fluid mass-flow rates (Allen & Seyfried, 2004), but if water/rock ratios are 0·5 (Agrinier & Cannat, 1997) or less (close to the lower limit of 0·16 in the case of virtually fluid-absent conditions), a larger temperature increase is possible. A low-temperature start, with entry of Fe3+ into lizardite as a cronstedtite component, could thus explain the initial absence of magnetite. In our model this evolutionary process could take place under environmental conditions of constant µ(Fe2+Mg–1) (Fig. 2). The mole per cent of Fe-lizardite that is to be attributed to Fe2+ remains at approximately 2–3%, whereas the rest of the iron in lizardite is Fe3+.
However, early enrichment of Fe in lizardite at the expense of magnetite formation, followed by a decline in Fe, is apparently not universally the case. The Mössbauer work of Votyakov et al. (1993, fig. 1) clearly showed an increase in Fe3+/Fe ratio of lizardite with increasing degree of serpentinization. Frost & Beard (2007, fig. 6) showed that there are diverse patterns for the timing of magnetite growth.
The serpentinized peridotites of the Northern Apennines Ligurian units in Tuscany, Italy, studied by M. Mellini and coworkers, constitute a good type-example of massif serpentinite, for which extensive and thorough mineralogical, geochemical and textural characterization has been provided; and we have already used some of their data above. Observations from the Tuscany rocks reinforce and amplify some of the observations made above. They are mainly altered harzburgites with some lherzolites, so the SiO2 content of the protoliths was sufficiently high to prevent the growth of brucite alongside serpentine (Viti & Mellini, 1998). Chlorine contents up to 0·6 wt % in the serpentine minerals support the suggestion that the serpentinization was oceanic (Anselmi et al., 2000). Magnetite is abundant, and hydrogarnet is reported. Bodies of rodingite are present. Fe3+/Fe ratios in whole-rock analyses (Fe by X-ray fluorescence and Fe2+ by titration) of Tuscany serpentinites are similar to the global average of highly serpentinized peridotite (Fig. 1); namely, 0·69 (n = 18) (Viti & Mellini, 1998; Anselmi et al., 2000; Rumori et al., 2004).
XFe values of EDS and SEM analyses of lizardite and minor chrysotile (Viti & Mellini, 1998) in the mesh replacements of olivine in serpentinites on Elba are 0·050 ± 0·012 (rims) and 0·052 ± 0·023 (cores), and for other Tuscany harzburgites (Anselmi et al., 2000) the overall average mesh serpentine is 0·066 ± 0·020. The large standard deviations in XFe of serpentine reflect inter-sample variation rather than inhomogeneity within samples (Viti & Mellini, 1998). Although no measurements of Fe3+ and Fe2+ have been made on these serpentines, the whole-rock data and a trend line with a slope of –0·31 on the Si + Al/2 vs Fe plot (Fig. 3) indicate enrichment of Fe3+ at the expense of Si [equation (5)], despite the absence of brucite from the assemblage. One can speculate whether the inter-sample variation is due to varying temperatures of alteration, or to variations in H2 or H2O. Perhaps alteration at different times and temperatures is the best explanation? Notwithstanding their variable Fe3+ contents, most of the serpentines still have XFe less than the protolith olivine. I infer that the exchange potential of Fe2+Mg–1 (which are elements with relatively fast transport rates; Carlson, 2002) was spatially and temporally constant in the rock, and equivalent to serpentine with XFe2+ 0·03.
Corresponding data for bastite replacements of orthopyroxene in Elba and other Tuscany serpentinites are XFe = 0·050 ± 0·016 and 0·057 ± 0·029, respectively, and the total range is from 0·03 to 0·10. Thus, bastite lizardite in the Tuscany serpentinites is, within uncertainty limits, no richer in Fe than serpentine in the mesh pseudomorphs after olivine. The slope of the trend line in a plot analogous to Fig. 6 is –0·37. Frost & Beard (2007) correctly interpreted hydrogarnet as an indicator of low SiO2 activity. In the Tuscany serpentinites, some of the hydroandradite occurs in the bastite. Perovskite is another SiO2-undersaturated accessory mineral in some serpentinites. This and hydrogarnet occur in the less depleted, higher silica serpentinized harzburgites, lherzolites, and pyroxenites in Italy and elsewhere. Evidently, in these as well as dunites, serpentinization has brought the silica potential well down from that defined by olivine + orthopyroxene in the protolith (Frost & Beard, 2007, fig. 3). Even if the SiO2 potential were locally elevated in the vicinity of orthopyroxene, it was not responsible in Tuscany for any corresponding elevation in the average Fe content, or the T/(T + M) occupancy ratio of the bastite as compared with the mesh serpentine.
A similar lack of contrast between bastite and mesh serpentine compositions is seen in abyssal peridotites drilled from locations close to spreading centers (Oufi et al., 2002). XFe of serpentine after olivine was found to range from 0·02 to 0·08 and after orthopyroxene from 0·03 to 0·08 (Oufi et al., 2002, fig. 10; note their diagram plots wt % FeO rather than XFe).
A consistent difference between bastite and olivine-pseudomorph serpentine is the higher Al and Cr in the former. If bastite lizardite is not systematically different in average XFe from olivine pseudomorph lizardite, what explains the typical absence of magnetite from bastite (e.g. Dungan, 1979; Wicks & Plant, 1979; Shervais et al., 2004; Frost & Beard, 2007)? The R3+ elements have no doubt been inherited from the pyroxene. They are known to have sluggish intergranular transport rates (Carlson, 2002), and thus support local chemical potential gradients indefinitely. As suggested by Frost & Beard (2007), this is possibly true also for Si, but evidence for this in terms of the serpentine composition turns up in only a few cases.
The XFe of lizardite in veins in the Elba serpentinite (0·049 ± 0·003, Viti & Mellini, 1997) is similar to that of lizardite in the surrounding rock, whereas that of chrysotile (0·037 ± 0·004) is poorer in total Fe, a comparison that, as we have seen, is globally true, and probably related to their different uptake of Fe3+. Both are notably more homogeneous than in the rocks. Perhaps this reflects a one-time cracking event, as opposed to a protracted alteration history for the rocks during which temperature and H2 and H2O fugacity might have varied. One infers that the fluid that filled the cracks was locally derived, and, furthermore, broadly in exchange equilibrium with relict olivine in the wall-rocks. This might not be the case in other locations where the serpentinite is strongly tectonized and broken into small bodies, where the vein-filling fluid may be drawn from a foreign source, in which case more varied serpentine compositions (higher Al for example) can be expected, and perhaps more Fe2+Mg–1 substitution.
I see no reason to interpret the serpentinization of the Tuscany peridotites, apart for the introduction of H2O-fluid and addition of some minor elements, not to have been primarily a rock-dominated or closed-system event.
A caveat is appropriate here. Perhaps the compositional variability of both mesh and bastite serpentine (in Tuscany and elsewhere) signifies some randomness in the uptake of iron into material that is very poorly crystallized? At 150–350°C we can hardly expect the product minerals to reflect environmental chemical potentials with the same fidelity we have come to expect in prograde metamorphism at higher temperatures. We should also remember that the incremental delivery of H2O to the site of serpentinization is driven by a steep chemical potential gradient (MacDonald & Fyfe, 1985), and it is likely to be episodic and triggered by local strain events. It is probably unreasonable to expect the reaction rate faithfully to reflect the inflow rate and so maintain the steady-state or equilibrium H2O and H2 potentials. In this case, the Fe content of serpentine might vary correspondingly. The correlations expressed in Figs 3 and 6, however, suggest an orderly assembly of crystals with compositions controlled by intrinsic energetic factors.
Having stated at the outset that we shall assume some modest approach to equilibrium in the serpentinization process, we cannot afford to ignore instances where the formation of metastable products under unfavorable kinetic conditions has demonstrably taken place. For example, in a serpentinized harzburgite from Oman, Baronnet & Boudier (2001) described, using TEM, first-formed lizardite and brucite that had XFe similar to the reactant olivine. Further growth and/or recrystallization produced magnetite. This follows well-understood principles of metastable crystallization.
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OPEN-SYSTEM CONDITIONS |
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TOPABSTRACTINTRODUCTIONTHE SERPENTINITE SYSTEMEXPERIMENTSWHY MAGNETITE?MICROSTRUCTURESFERROUS AND FERRIC IRON...FURTHER FIELD CONSTRAINTSOPEN-SYSTEM CONDITIONSCONCLUSIONSREFERENCES |
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A reasonable model for massif serpentinites has, in my opinion, to be one that invokes the very slow diffusive passage of pervasive aqueous fluid along intergrain boundaries. This provides ample opportunity for the mineral assemblage in the rock to impose on the fluid its values of the chemical potentials of H2O, H2, SiO2 and Fe2+Mg–1. Rates of reaction (consumption of H2O and release of H2) and delivery of H2O will be very slow (e.g. Früh-Green et al., 2003
) and roughly in mutual balance, the water/rock ratio remaining low. These are the classic conditions of retrograde metamorphism that are sometimes called ‘water absent’ or ‘dry’. Periodic tectonic disturbances, particularly if they involve brittle fracture, such as in the formation of chrysotile veins, will tend to speed up both rates because of the arrival into grain boundaries of fluid that is richer in H2O and poorer in H2. This model clearly contrasts greatly with ‘hydrothermal’ metamorphism, a process that involves high fluid/rock ratios and is amenable to laboratory study (Moody, 1976a; Janecky & Seyfried, 1986; Normand et al., 2002; Seyfried et al., 2007). Depending on the amounts of fluid or its rate of supply, new chemical potentials can be imposed on the solids and the variance of the system may be increased (fewer phases). We might expect sub-sea-floor serpentinization to fall into this category. The first (‘dry’) model was implicit in the discussion above of the formation of mesh serpentine after olivine and much of the bastite serpentine in massif peridotites. More important is the fact that this model of serpentinization can deliver the maximum amounts of magnetite precipitate and evolved hydrogen. As we have seen, these amounts are subject to the whole-rock composition (modal orthopyroxene/olivine ratio) and how much of the cronstedtite component is incorporated in the serpentine.
The situation is less clear for some instances of bastite pseudomorphs after pyroxene. These can be composed almost entirely of lizardite—what looks like a local increase in phase rule variance. Le Gleuher et al. (1990) described the formation of bastite in a TEM study of enstatite in meta-pyroxenite from Pernes, France, in terms of system openness. Minor talc-like and brucite-like layers, in addition to chlorite, testify to metastable conditions among the products, even in the more open microsites in their sample. The XFe of the bastite resembles that of the original orthopyroxene (0·12). Its T/(T + M) ratio is high and can be reduced to 0·4 only if all Fe and two-thirds of the Al are placed on the M-site; perhaps there was some talc contamination? The common chloritization of biotite in granites might be a reasonable analogue for the bastite alteration of orthopyroxene to serpentine. In the former there is little evidence of significant volume expansion, despite major introduction of H2O, and loss of Si, K and Ti to the fluid (Veblen & Ferry, 1983). Similarly, bastite pseudomorphs, which usually involve a major component of topotactic lizardite, are not associated with significant volume expansion (Dungan, 1979; Wicks, 1986). Volume expansion accompanies the serpentinization of olivine because of the lack of diffusion gradients capable of removing Mg. Bastite alteration may well be accompanied in some cases by a locally higher SiO2 potential, as claimed by Frost & Beard (2007). However, does that drive reaction (3) or reaction (6)? It does not drive reaction (5). If the bastite environment is locally open to foreign fluid, then it might experience an elevated µ(Fe2+Mg–1). However, as noted above, Figs 3 and 6 can allow only about one-fourth of the excess Fe on the diagrams to be ferrous iron.
Where system openness is most obvious in serpentinite bodies is in marginal zones (e.g. Labotka & Albee, 1979; Kelemen et al., 2003), where the addition of Si promotes the formation of talc (± carbonate as listwanite), and tremolite from diopside + brucite (or olivine). Mass-transfer of nearly all elements on a smaller scale (centimeters to meters), producing monomineralic biotite, chlorite, and actinolite, is typically developed at lithological contacts of serpentinite bodies.
Active serpentinization processes in continental and oceanic environments have many geochemical features in common in that they release distinctive hydrothermal fluids that are notable for their high pH, relatively low SiO2 concentrations, and the presence of hydrogen and methane gases (e.g. Janecky & Seyfried, 1986; Beard & Hopkinson, 2000; Kelley et al., 2005; Charlou et al., 2002; Douville et al., 2002; Mével, 2003). These fluids are consistent with progress of the magnetite- and ferrian lizardite-forming serpentinization reaction, which would indicate that, overall, fluid/rock ratios are not so high as to destroy fundamentally closed-system behaviour with respect to modal mineralogy, change the abundances of major elements such as Si, Fe, Mg, and Al, or affect the internal control of µ(Fe2+Mg–1).
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CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONTHE SERPENTINITE SYSTEMEXPERIMENTSWHY MAGNETITE?MICROSTRUCTURESFERROUS AND FERRIC IRON...FURTHER FIELD CONSTRAINTSOPEN-SYSTEM CONDITIONSCONCLUSIONSREFERENCES |
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In evaluating environmental conditions controlling magnetite-present low-Fe serpentine vis-à-vis magnetite-free high-Fe serpentine, we are confronted (in addition to T and P) with a choice not only of the potential of SiO2 (Frost & Beard, 2007
), but also of the pressure of H2O, the oxygen (or hydrogen) fugacity, and the Fe2+Mg–1 exchange potential. These potentials are to a large degree pre-determined by the phase rule variance of the serpentine mineral assemblage. Their values can be changed if sufficient influx of fluid can locally eliminate a phase and increase the variance. The high frequency of serpentine compositions centered around XFe = 0·05 argues for the predominant influence of the prevailing of µ(Fe2+Mg–1) (the KD effect), otherwise we are looking at a curious coincidence. The large amounts of ferric iron in lizardite render equation (3) of the present paper (Frost & Beard, 2007) inappropriate for paragenetic analysis. Reaction (5) allows variability of XFe in serpentine under uniform levels of µ(Fe2+Mg–1). Rather than attribute the observed modal and compositional differences in serpentinites exclusively to the influence of local variations in silica activity, it is more appropriate to view them in terms of progress of temperature-dependent reactions such as (5) and possibly (6). In a generalized model of serpentinization of mantle peridotite (with Fa9–10 olivine), we can expect product lizardite to contain in solution about 3% of the Fe2+ end-member plus 2–7% of an Fe3+ end-member that is mostly cronstedtite, but in some instances a ferrian analogue of kaolinite, or both. The microprobe has been a valuable tool in petrological research, but unfortunately it can lead us unconsciously to ignore Fe3+ and draw petrological conclusions that in some cases are not warranted. Significant system openness to fluids from a geochemically foreign source is likely to be limited to serpentinite contacts with other rock-types and some, but not many, oceanic occurrences. The ferri-Tschermak compositional trend of lizardite is shared by continental and oceanic serpentinites, but more of the latter have in addition a trend of Fe3+ uptake via M-site vacancy formation, especially in the bastite pseudomorphs.
Much more could be learned from further study by Fe-Mössbauer spectroscopy of the Fe2+ and Fe3+ contents of lizardite in massif serpentinites, to the extent possible in their microstructural contexts. When possible, this should be combined with micro-Raman spectroscopy, which serves as a quick and reliable technique to check on the identity of the serpentine species (Groppo et al., 2006). The uptake of Fe3+ in lizardite involves predominantly the ferri-Tschermaks substitution, but can a stronger or more precise case be made in some samples for cation vacancies or hydrogen loss? Is the level of Fe2+ abundance in the lizardite in mantle meta-peridotite as uniform and low as suggested in this paper? In situ oxygen isotope thermometry (serpentine–magnetite) could throw light on the possible temperature dependence of Fe3+ in lizardite in serpentinites. Confirmation (or otherwise) of the micro-analytical trends for serpentine depicted in Figs 3 and 6 could be the objective of further careful microprobe study of specific occurrences. A examination of serpentinized olivine (Fa >10%) in large layered mafic intrusions might reveal sets of more ferroan equipotential trend-lines on the MgO–FeO–SiO2 diagram.
The composition of brucite was not discussed in detail here. Inasmuch as brucite has been reported in some serpentinites as yellow, brown, red, and rusty, and to contain amounts of the Fe(OH)2 end-member amakinite exceeding the constraints of Fe2+Mg–1 equilibrium, a study of their possible contents of ferric iron would seem to be desirable.
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ACKNOWLEDGEMENTS |
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I am grateful to D. K. Bird, M. D’Antonio, M. D. Dyar, J. M. Ferry, B. R. Frost, K. H. Hattori, S. M. Kuehner, T. C. Labotka, M. Mellini, E. Pope, and an anonymous reviewer for critical comments and other assistance.
*Corresponding author. E-mail: bwevans{at}u.washington.edu
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REFERENCES |
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