Many oceanic serpentinites consist only of two phases: serpentine and magnetite. Determining the metamorphic evolution of these rocks is difficult, because fluid-mineral phase relationships cannot always be employed. The serpentinites from ODP Leg 209 are remarkable in this regard, as they are highly variable in their mineralogical composition and in the extents of serpentinization. Talc replaces serpentine at Site 1268, brucite is present in the variably serpentinized dunites from Sites 1271 and 1274, where orthopyroxene has reacted to talc, brucite and iowaite are developed in rocks from Site 1272, and a continuous transition from relatively fresh to completely serpentinized fresh harzburgites and dunite can be studied at Site 1274. The following discussion will evaluate the observed phase relationships in the light of experimental [Seyfried and Dibble, 1980; Janecky and Seyfried, 1986; Allen and Seyfried, 2003] and theoretical [Palandri and Reed, 2004; Wetzel and Shock, 2000] constraints for seawater-peridotite interactions and under consideration of peridotite-hosted hydrothermal vent fluid compositions from the Logatchev, Rainbow, and Lost City hydrothermal sites [Charlou et al., 2002; Douville et al., 2002; Kelley et al., 2001].
4.1. Serpentinite Phase Petrology
 Previous studies of oceanic serpentinites and associated gabbros from various settings [Cannat et al., 1992; Mével et al., 1991; Hébert et al., 1990; Bideau et al., 1991; Früh-Green et al., 1996; Dilek et al., 1997; Prichard, 1979; Mével and Stadoumi, 1996; Agrinier et al., 1995; Agrinier and Cannat, 1997; Schroeder et al., 2002] have revealed complex metamorphic evolution paths that include high-temperature (>300°–400°C) serpentinization and rodingitization and multiple lower-temperature overprints related to uplift and exposure of the mantle at the seafloor. In terms of their major components serpentinites are remarkably uniform, consisting of serpentine polytypes+magnetite±brucite±talc±tremolite [e.g., Mével, 2003].
 Temperature-pressure stabilities of serpentine phases are not well constrained. Thermodynamic calculations predict that antigorite is the stable serpentine phase above 200°–300°C and pressures below 2 kbar and that chrysotile and lizardite form from antigorite during low-temperature recrystallization [e.g., Evans, 1977]. An early oxygen isotope study of serpentinites is consistent with this view. Wenner and Taylor  suggest that lizardite and chrysotile form at temperature below 235°C. However, more recent stable isotope investigations [Agrinier et al., 1995; Früh-Green et al., 1996; Agrinier and Cannat, 1997], field studies [O'Hanley and Wicks, 1995; Wicks and Whittaker, 1977], and hydrothermal experiments [Normand et al., 2002; Janecky and Seyfried, 1986; Allen and Seyfried, 2003] indicate that lizardite and chrysotile form directly from olivine at temperatures above 300°C.
 The nature of the serpentine polytype is likely to be a function of temperature, fluid composition, Ptotal, P(H2O), and kinetics. While thermodynamical calculations predict that (at Ptotal = P(H2O) = 1kbar) antigorite forms at T > 300°C, lizardite forms at T < 200°C, and chrysotile forms at intermediate temperatures [e.g., Evans, 1977], stable isotope studies have shown that lizardite and chrysotile form at temperatures of 350–400°C [Agrinier and Cannat, 1997; Agrinier et al., 1995; Früh-Green et al., 1996]. Surface energy minimization effects may explain the metastable replacement of olivine by lizardite, followed by partial replacement of lizardite by chrysotile [Normand et al., 2002]. Sanford  and O'Hanley  argue that the stability fields of lizardite and chrysotile may extend to higher temperatures if P(H2O) < Ptotal. Actively serpentinizing systems within the oceanic lithosphere consume water, produce methane and hydrogen, and are likely at sub-lithostatic pressure, suggesting that P(H2O) is likely significantly smaller than Ptotal. The use of serpentine mineralogy in constraining the evolution of oceanic serpentinites is therefore somewhat problematic.
 The presence of incompletely serpentinized dunites and minor harzburgites at Sites 1271 and 1274 provides the opportunity to examine potential reaction paths of serpentinization. Field and experimental studies suggest that serpentinization at low to moderate temperature (<250°C) is a nonequilibrium process, during which dissolution of olivine proceeds at rates faster than precipitation of talc and serpentine [e.g., Nesbitt and Brickner, 1978; Martin and Fyfe, 1970].
 Figure 5a displays a Mg-Ca-Si-O-H mineral-fluid phase diagram for 200°C and 500 bars. The fluid composition will be driven toward the stability fields of serpentine and brucite, which represent the lowest energy mineral assemblage [e.g., Hemley et al., 1977]. The reaction path proposed by Nesbitt and Brickner  (black arrow in Figure 5a) is strongly curved, owing to rapid dissolution of olivine and consumption of acidity, followed by a shoaling of the trend due to dissociation of orthosilicate at high pH. The last part of the proposed reaction path is controlled by the forsterite dissolution boundary. If pH = 10 and silica concentrations of seawater are assumed (similar to the Lost City vent fluids [Kelley et al., 2001]), the calculated silica activity at 200°C (the estimated maximum temperature of Lost City vent fluids [Allen and Seyfried, 2004]) is near the serpentine-brucite boundary in Figure 5a. However, the in situ pH at 200°C for the Lost City fluids (room temperature pH values around 10.5 [Kelley et al., 2001]) is only 7.7. Still the Lost City fluid could be multiply saturated in olivine, brucite, and serpentine phases if the Mg+2 activity is ≤1 μmol/kg. Also shown in Figure 5a is a schematic reaction path for fluids from a harzburgite-seawater interaction experiment [Seyfried and Dibble, 1980] that trends toward the chrysotile-brucite boundary. In contrast, fluids from a lherzolite-seawater experiment at 200°C and 500 bars [Janecky and Seyfried, 1986] do not evolve toward brucite saturation, because the high abundance of pyroxenes keeps the fluid pH low and the silica activity of the fluids high (Figure 5a). The fluid reaction path is therefore a function of both the reaction temperature (i.e., the relative rates of olivine and pyroxene alteration) and the modal composition of the peridotite.
Figure 5. (a) Mineral-fluid phase diagram for the system Mg-Ca-Si-O-H at 200°C and 500 bars (right), constructed using thermodynamic data from Johnson et al.  and assuming Log(aCa2+/a2H+) = 6. The black trend is a hypothetical evolution path of fluids dissolving olivine at a higher rate than hydrous mineral are precipitated. The drop in silica activity is solely the result of the decrease in the orthosilicate activity coefficient as a function of increasing pH (calculated with SUPCRT92 [Johnson et al., 1992]). The blue square represents the hydrothermal fluids from the Lost City vent site [Kelley et al., 2001], respeciated for 200°C [Allen and Seyfried, 2004]. The in situ pH value of Lost City fluids at 200°C is 7.7; silica concentration is similar to seawater [Kelley et al., 2001]. The Lost City fluids could be multiply saturated in olivine, brucite, and serpentine, if the Mg+2 activity is ≤1 μmol/kg. The gray field encompasses fluid compositions from a lherzolite-seawater reaction experiment at 200°C and 500 bars conducted by Janecky and Seyfried . Fluids from a harzburgite-seawater interaction experiment [Seyfried and Dibble, 1980] evolve along a trend indicated by the green arrows toward the chrysotile-brucite boundary. The black arrow is a hypothetical evolution path of a solution that dissolves olivine faster than it precipitates serpentine [cf. Nesbitt and Brickner, 1978]. The red arrow represents a schematic fluid evolution path for seawater heated to 200°C followed by a drop in Mg and an increase in pH. (b) Mineral-fluid phase diagram for the system Mg-Ca-Si-O-H at 400°C and 500 bars. The gray field labeled “R&L” represents fluid compositions from the Rainbow and Logatchev hydrothermal sites, respeciated for 400°C and 500 bars [Charlou et al., 2002]. In situ pH values at these systems range between 4.5 and 5. Silica activities are 6 to 7 mmol/kg, and Mg2+ activity is assumed to be on the order of 0.1 mmol/kg.
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 The effect of pyroxene on solution chemistry is even stronger at higher temperatures (>250°–300°C), where pyroxenes react faster than olivine [Martin and Fyfe, 1970; Allen and Seyfried, 2003]. Figure 5b demonstrates that high-temperature (365°C) black smoker fluids from the Logatchev and Rainbow peridotite-hosted hydrothermal systems [e.g., Charlou et al., 2002], when respeciated at 400°C and 500 bars, are pinned by fluid saturation in pyroxene and that talc and possibly tremolite should form at the expense of pyroxene. As long as pyroxenes are present (and dissolve fast) the pH of the interacting fluids is low and the silica activity high [cf. Allen and Seyfried, 2003]. Only when the fluid is no longer pyroxene saturated can it evolve to higher pH and lower silica activity.
 These mineral-fluid phase relationships are relevant for the interpretation of the serpentine-brucite and olivine-amphibole-talc assemblages observed in cores from Sites 1271 and 1274. A possible interpretation of the relatively fresh dunites with rare, talc-altered orthopyroxene pseudomorphs is that these rocks reacted with hydrothermal fluids at high temperatures (>300°C), where orthopyroxene reacts faster than olivine. The proximity of mafic material (amphibolites) to intervals of unaltered dunite may suggest that the presence of pyroxene and amphibole kept the silica activity of the interacting fluids high so that serpentinization of olivine was inhibited. Formation of talc after olivine is thermodynamically favored (Figure 5b) but kinetically sluggish [e.g., Nesbitt and Brickner, 1978]. Replacement of orthopyroxene by talc, on the other hand, is commonly observed in alpine and abyssal serpentinites [e.g., Aumento and Loubat, 1971; Hostetler et al., 1966]. As the system moves toward lower silica activities after pyroxenes are exhausted, talc that had formed during initial alteration will react to serpentine. These processes explain the scarcity (or complete lack) of talc as part of the background alteration in many completely serpentinized rocks.
 The brucite-bearing dunites (and minor harzburgites) require a different genetic model (see Figure 5a). The phase relations indicate that these lithologies reacted with fluids of high pH and low silica activities. Such fluids are generated when alteration is controlled by rapid olivine dissolution and relatively sluggish hydrous mineral precipitation. Thermodynamically and kinetically, conditions for brucite-serpentine formation are favorable at low temperatures and in the absence of pyroxenes. The abundance of brucite at Site 1271 may suggest that the circulating fluids are dominantly high pH and low silica activity, which could be explained if the basement consists dominantly of dunite.
 At Site 1274, magnetite is rare where serpentinization is incomplete (as low as 60%), suggesting that significant magnetite formation does not begin until more than 60% of the primary rock is serpentinized. At Site 1274 brucite is most common in dunites, but does also occur in harburgites. Brucite is not expected to be associated with the breakdown of pyroxene, because the high silica activity of fluids reacting with pyroxene would prevent brucite formation and would cause previously formed brucite to react to serpentine according to reaction:
 The occurrence of both brucite and partially serpentinized pyroxene may hence indicate local disequilibria during the main stage of serpentinization. Alternatively, the formation of serpentine+brucite+magnetite after olivine took place at a later stage, at temperatures below 250°C. The following reaction can possibly account for this observation:
 The molar volumes of brucite, magnetite, and chrysotile are 26.43 cm3, 44.52 cm3, and 108.5 cm3, respectively [Helgeson et al., 1978], so that the stoichiometry of this reaction corresponds to a mode of 83% serpentine, 12% brucite, and 5% magnetite (83:12:5).
 While here magnetite accompanies serpentine and brucite and appears to form directly from olivine breakdown, little magnetite is produced during initial serpentinization. Allowing for Fe to substitute for Mg in serpentine and brucite, e.g.,
will increase the proportion of brucite and decrease the proportion of magnetite (83:15:2). The formation of magnetite and H2,aq maybe omitted altogether if Fe-rich brucite and serpentine are formed:
Hence the amount of magnetite, brucite, and H2,aq formed depends upon the degree of Fe-incorporation into secondary phases.
 To explain density-magnetic susceptibility relationships in variably serpentinized harzburgites from ophiolites, Toft et al.  suggested that initial serpentinization produces Fe-rich serpentine and Fe-rich brucite, and that magnetite forms during recrystallization of these phases as serpentinization proceeds. We tentatively propose a similar explanation for the observation that little magnetite is formed during initial serpentinization at Site 1274. However, we note that analyses of recrystallized serpentine from MARK serpentinites [Dilek et al., 1997] have much higher Fe contents (5.2 ± 1.0 wt.% FeO) than pseudomorphic serpentine (2.8 ± 1.2 wt.% FeO). Our observation of direct replacement of olivine by serpentine, brucite AND magnetite that, as argued above, likely occurred subsequent to the initial serpentinization and possibly represents lower reaction temperatures provides an alternative explanation for the increased magnetite formation during advanced stages of serpentinization. The presence of fresh clinopyroxene in rocks in which olivine is completely serpentinized (Figure 4f) can also best be explained by main stage alteration at temperatures at which olivine reacts much faster than pyroxene (<250°C).
4.2. Redox Conditions
 While the sequence of alteration reaction observed in Hole 1274A and previous studies [e.g., Toft et al., 1990] suggest that magnetite formation peaks when most of the rock is serpentinized, ongoing fluid-rock reactions after serpentinization is complete may lead to magnetite destruction. This is clearly demonstrated by the rocks recovered from Site 1268 that show reaction of serpentine and magnetite to talc and hematite along with a decrease in magnetic susceptibility [Shipboard Scientific Party, 2003]. The succession of oxide/sulfide minerals in Hole 1268A indicates the development of the rock-fluid system to more oxidizing conditions with time (Figure 6a). The presence of magnetite in veins as well as the occurrence of reduced Fe and Ni sulfides (relict pyrrhotite/pentlandite and partial replacement by heazlewoodite) in a moderately altered orthopyroxenite suggest fairly reducing conditions during early serpentinization and veining. Phases indicating even lower oxygen fugacities, found in peridotites from Hess Deep and MARK area [Alt and Shanks, 1998, 2003] (Figure 6b), were not identified in the course of our limited shipboard studies. Increasing modal percentages of hematite in the late stage veins are associated with decreasing H2S(aq) and H2(aq) activities of the interacting hydrothermal solution (Figure 7). Such a trend can be expected in long-lived hydrothermal systems where the reducing capacity of the rock is exhausted by oxidation of seawater sulfate [e.g., Seyfried and Ding, 1995]. This scenario is consistent with the observation that the peridotites in Hole 1268A underwent multiple episodes of intense and pervasive alteration under variable redox and pH conditions. This is generally similar to serpentinites from the Iberian Abyssal Plain [Beard and Hopkinson, 2000; Hopkinson et al., 2004].
Figure 6. Phase diagram for the Fe-Ni-S-O system at 300°C and 2 kbar. Phase boundaries of Fe phases are marked by thick lines, those of Ni phases and Ni-Fe alloys are thin lines, and others are dashed lines. Fe phases are hematite (Hem, Fe2O3), pyrite (Py, FeS2), magnetite (Mt, Fe3O4), pyrrhotite (Po, Fe1−xS), and Ni- and Ni-Fe phases are vaesite (Va, NiS2), polydymite (Pd, Ni3S4), millerite (Mi, NiS), heazlewoodite (Hz, Ni3S2), taenite (Ta, δFeNi), awaruite (Aw, Ni3Fe), and kamacite (Ka, αFeNi). Oxide-sulfide phase boundaries simplified after Frost ; others are calculated with SUPCRT92 [Johnson et al., 1992]. (a) Fields mark oxide-sulfide assemblages observed in Hole 1268A, and the red arrow is the inferred evolution path. (b) Fields mark oxide-sulfide assemblages observed in Hess Deep and MARK areas [Alt and Shanks, 1998, 2003].
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Figure 7. Mineral fluid phase diagram for the system Fe-H-O-S at temperatures of (a) 400°C, (b) 300°C, and (c) 200°C and a pressure of 1 kbar. The dashed line labeled AAC is for the reaction anhydrite + 3 anorthite + 4 H2(aq) = 2 clinozoisite + H2S(aq) + 2H2O. Anorthite activity is (a) 0.7, (b) 0.5, and (c) 0.3. Note that the presence of anhydrite excludes hematite at temperatures >300°C, while anhydrite and hematite may coexist at lower temperatures. Gray lines are for equal activities of HSO4− and H2S(aq) (continuous) and for equal activities of HS− and SO42− (dashed). The latter sulfide-sulfate species are more relevant at higher pH. The continuous gray lines are for calculated in situ pH values for Rainbow fluids (200°C: 3.0; 300°C: 3.9; 400°C: 4.6) and Lost City (200°C: 7.7; 300°C: 7.5) based on 25°C pH values of 2.8 (Rainbow [Charlou et al., 2002]) and 10.5 (Lost City [Kelley et al., 2001]). Note that the sulfate-sulfide boundary shifts to the left more rapidly than the hematite-magnetite boundary as temperatures decrease. Also note that in the presented temperature range sulfate and hematite may not coexist at high pH values relevant for Lost City. All calculations are based on the SUPCRT92 database [Johnson et al., 1992].
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 The late-stage veins are dominated by hematite and not by pyrite like in other locations [Alt and Shanks, 1998, 2003] (Figure 6b). Seyfried and Ding  suggest that the reaction anhydrite + 3 anorthite + 4 H2(aq) = 2 clinozoisite + H2S(aq) + 2H2O (AAC) keeps the system fairly reducing until plagioclase has been reacted to albite (and anorthite activity is small). This model can explain the lack of hematite in reaction zones of basalt-hosted hydrothermal systems [e.g., Alt, 1995]. In peridotite-hosted systems a Ca-Al silicate/anhydrite buffer may not be affective. In any case, high Ca concentrations (20–65 mM [Charlou et al., 2002; Kelley et al., 2001]) in fluids venting from peridotites indicate that enough Ca is released from the incompletely serpentinized rocks to keep sulfate level low. One way of simulating the effect of intense water rock reaction on redox equilibria is assuming a drop in anorthite activity along with decreasing temperatures of fluid-rock interaction as the crust is uplifted. The consequence of a coupled drop in temperature and anorthite activity (Figures 7a–7c) is a shift of the AAC equilibrium relative to hematite-magnetite-pyrite such that below 300°C hematite may coexist with anhydrite. Another way of explaining the shift in redox equilibria is independent of a possible CaAl-silicate/anhydrite buffer. Seawater has insufficient Ca contents (10 mM) to precipitate all the sulfate (28 mM) as anhydrite upon heating. In basalt-hosted hydrothermal systems sulfate is near-zero, because Ca is rapidly released to the fluid in the downwelling limb of a convection cell and sulfate is controlled by anhydrite solubility. However, Ca release from peridotite may be diminished if the protolith is strongly melt-depleted and/or the rock is completely serpentinized and Ca-leached. If such a rock undergoes continued reaction with seawater, sulfate may remain in the fluid at higher concentration levels and hematite may become stable. The temperature-dependent shift of the sulfate/sulfide equilibria relative to that of the hematite/magnetite boundary (Figure 7) indicates that acidic fluids with significant sulfate contents may coexist with hematite at temperature of 300°C and below. Slightly alkaline fluids (e.g., in situ pH of 7.5–8), such as at the Lost City site, cannot equilibrate with hematite (dashed, gray line in Figure 7) even if sulfate is present. This is consistent with our earlier suggestion that olivine reacts to serpentine, brucite, and magnetite at high fluid pH that is expected to evolve when olivine reacts faster than pyroxene. At Site 1268, olivine is completely reacted out. Therefore circulating seawater will not become alkaline and may produce hematite upon interaction with serpentinite. These examples illustrate the complex relations between water-silicate reactions, fluid chemistry, and redox equilibria in peridotite-hosted hydrothermal systems.
4.3. Si Metasomatism of Serpentinites at Site 1268
 The most unusual aspect of the hydrothermal alteration at Site 1268 is the widespread replacement of serpentine by talc under static conditions. Replacement of serpentinized peridotite and gabbro by talc has been reported before. Escartín et al.  documented the widespread presence of amphibole-chlorite-talc schists on rift valley walls in the 15°20′N area that represent a late stage of high-temperature syntectonic alteration along detachment faults. Früh-Green et al.  mention talc-rich lithologies from the detachment fault surfaces at Atlantic Massif at the MAR 30°N. D'Orazio et al.  describe talc-rich serpentinite and gabbro breccias and talc schists from the St. Paul and Conrad fracture zones in the Atlantic. All these occurrences have been attributed to large volume fluid flux along faults leading to Si metasomatism. The unusual aspect about the talcous rocks from Site 1268 is complete replacement of serpentine by talc under static conditions.
 Another remarkable observation is the lack of Ca metasomatism (rodingitization) of the gabbros and gabbroic veins encountered in Hole 1268A. In previously drilled and dredged sections of upper mantle [Dilek et al., 1997; Früh-Green et al., 1996; Bideau et al., 1991] gabbros are Ca-metasomatized, and it has been suggested that Ca released during serpentinization is fixed in Ca silicates replacing primary minerals of the gabbro. This static Ca metasomatism of gabbros embedded in peridotite is common elsewhere, but it is not developed at Site 1268, where, instead, serpentinites have undergone pronounced Si metasomatism.
 It is well documented that talc forms after orthopyroxene under static conditions during prograde serpentinization and is replaced by lizardite either during retrograde serpentinization [Wicks, 1984] or due to a combination of increased H2O activity and decreased silica activity in the fluid [Frost, 1985]. As discussed in the previous section the formation of noticeable amounts of early talc and tremolite (after orthopyroxene) and brucite (after olivine) during prograde serpentinization is consistent with field observations, experimental results and vent fluid chemistry. Hydrothermal experiments (at 400°C and 500 bars) confirm that fluids are talc-saturated as long as fresh pyroxene is left in the serpentinizing peridotite [Allen and Seyfried, 2003]. Similarly, brucite is believed to be a product of incipient serpentinization which, due to reaction with aqueous SiO2 released by the breakdown of orthopyroxene to lizardite, forms lizardite and magnetite [e.g., Toft et al., 1990].
 There is no relict talc, tremolite, or brucite in serpentinites from Hole 1268A, suggesting that serpentinization reached an advanced stage, at which only serpentine and magnetite were present. Another indication of the advanced degree of serpentinization in Hole 1268A is the general scarcity of pseudomorphic mesh textures and the common development of transitional hourglass and ribbon textures [O'Hanley, 1996]. Furthermore, interlocking textures, indicating recrystallization of mesh-textured serpentine, are common and serrate chrysotile veins formed at the expense of early lizardite.
 The replacement of serpentine by talc at Site 1268 is a metasomatic process. Fresh Harzburgite (80%olivine, 20% orthopyroxene) has a molar (Mg+Fe)/Si of 1.8. Serpentine has a molar (Mg+Fe)/Si of 1.5 so that some transfer of Mg, Fe, and Si must be inferred unless the rock contains abundant brucite [O'Hanley, 1996]. The molar (Mg+Fe)/Si of talc is 0.75, and hence the complete replacement of serpentine by talc in long sections of core from Hole 1268A indicates substantial mass transfer of Mg, Fe, and Si. Mg solubility in hydrothermal fluids is extremely low in the presence of hydrous Mg-silicates [e.g., Saccocia et al., 1994]. We calculate that at 150°–350°C and 1 kbar, the Mg concentration of 3.2 wt.% NaCl solution saturated with talc is less than 5 ppm at pH ≥ 6. To explain the replacement of serpentine (∼40 wt.% Mg) by talc (∼29 wt.% Mg) with Mg loss would require unreasonable high water-to-rock ratios >20,000. However, the solubility of Mg is strongly pH-dependent, and Mg solubility in talc-saturated seawater at 200°C is ∼100 ppm at pH = 5 and ∼10,000 ppm at pH = 4 (at an assumed silica activity of 1 mM). It is conceivable that seawater recharging into highly Ca-depleted peridotite cannot leach sufficient Ca+2 from the rock to balance the loss in Mg+2, which is then in part balanced by the release of protons to the fluid. If the pH was indeed ≤5, Mg loss may attribute to the replacement of serpentine by talc. Marcasite is expected to form at pH ≤ 4–5 [Murowchick and Barnes, 1986], and ongoing studies of Hole 1268A sulfide petrology will focus on the identification of marcasite in talc-altered rocks.
 One plausible reaction between serpentine and fluid to form talc is
which proceeds to the right if the silica activity of the fluid is increased and/or if the water activity is lowered. In some cases, the formation of talc-magnesite rocks after serpentinite have been a consequence of lowering the water activity of the fluid by adding CO2 [e.g., Peabody and Einaudi, 1992]. However, we did not observe magnesite (or other carbonates) in detectable amounts and conclude that carbonatization cannot account for the abundant development of talc after serpentine. It is hence more likely that an increase in silica activity is responsible for the conversion of serpentine to talc. Talc alteration may therefore more likely be a result of Si metasomatism.
 The close spatial association of talc alteration and gabbro intrusions suggests that gabbro emplacement and talc alteration are related. Gabbro intrusion and dike formation in serpentinite could revive hydrothermal circulation by providing a heat source and creating permeability. However, hydrothermal circulation long after gabbro emplacement could also transport SiO2 from gabbro into host peridotite. Theoretical geochemical models suggest that, at 350°C and 500 bars, mafic rock-seawater reactions form hydrothermal fluids that many orders of magnitude higher in silica and lower in pH than hydrothermal fluids produced by reaction of seawater with ultramafic rocks [Wetzel and Shock, 2000]. Hydrothermal reaction between gabbro and seawater under greenschist-facies conditions transform pyroxene and plagioclase in the gabbros to chlorite, amphibole, and talc, releasing silica and acidity and causing Si metasomatism and talc formation in the serpentinites. Serpentinization of olivine at 350°C and 500 bars is associated with extremely low silica concentrations (<0.01 mM) in the interacting fluids [Wetzel and Shock, 2000]. However, Allen and Seyfried  have demonstrated that ultramafic rock alteration under conditions under which pyroxenes react faster than olivine may generate fluids with moderate Si concentrations (4 mM). Breakdown of pyroxene during harzburgite-seawater reactions is therefore an alternative (or additional) source of silica. Similarly, the abundance of hydrothermal sulfides in Hole 1268A may suggest input of hydrothermal sulfide derived from a gabbro-driven hydrothermal system as proposed by Alt and Shanks  for serpentinites from the MARK area.
 While serpentinites at Site 1268 show extreme Si metasomatism, the gabbroic dikes lack rodingitization. Rodingitization commonly takes place when gabbros and ultramafic rocks undergo the same metamorphic history and Ca-rich fluids generated during serpentinization metasomatize the gabbros replacing pyroxene and plagioclase by diopside, tremolite, clinozoisite, prehnite, and hydrogrossular [e.g., Schandl et al., 1989; O'Hanley, 1996]. The lack of rodingitization (even in small gabbroic veins) may also relate to the Ca-poor nature of the host peridotite.
4.4. Late-Stage Fluid-Rock Interactions
 The latest stages of veining in the Hess Deep, MARK, and Lost City areas is the formation of aragonite veins under ambient conditions at or near the seafloor [e.g., Früh-Green et al., 2003; Blusztajn and Hart, 1996; Dilek et al., 1997]. There is also abundant late-stage aragonite in serpentinites from the Iberian margin drill cores [Hopkinson et al., 2000]. A variety of late-stage, low-temperature reactions is recorded in ODP Leg 209 drill core. The late-stage talc overprint and sulfide/oxide veining at Site 1268 likely took place at somewhat elevated temperatures. Aragonite veins were not identified in core from Hole 1268A, suggesting that present-day veining and seawater ingress may be minimal, or that exposure at the seafloor was recent. At other sites (in particular 1271 and 1274), aragonite veining and low-temperature oxidative alteration of relict primary phases is common.
 The discovery of iowaite in core from Hole 1272A is the first reported occurrence of this phase from a mid-ocean ridge setting. It has been inferred that iowaite may be formed from iron-bearing brucite under oxidizing conditions [Heling and Schwarz, 1992]. This process involves the oxidation of Fe2+ in brucite to Fe3+ generating a charge imbalance, which is accommodated by incorporation of Cl− between brucite layers. In a model suggested for iowaite formation in serpentinite muds in the Izu-Bonin forearc this reaction is proposed to take place immediately below the seafloor due to infiltration of ambient seawater [Heling and Schwarz, 1992]. Conversely, occurrences of iowaite in serpentinites from the Iberian margin (ODP Site 897) have been related to circulating low-temperature, seawater derived Cl-rich brines [Gibson et al., 1996]. Here, iowaite is restricted to a zone with elevated Cl-concentrations in bulk rock analyses (up to 1 wt%) and it is inferred that alteration is this area was probably independent of the earlier serpentinization, sulfidization and pyrite formation, and later surficial oxidative alteration.
 The observation of abundant iowaite distinguishes Hole 1272A from all other holes drilled during Leg 209. A distinctive stage of alteration, postdating serpentinization, may have occurred in this area which may have been associated with fluid flow along the major fault zones observed in this hole. While the involvement of a brine in the formation of iowaite is questionable, iowaite does indicate oxidizing conditions. It is remarkable to find long sections of basement that have been oxidized given the fact that fluids actively venting from oceanic core complexes and ultramafic massifs in ophiolites are enriched in aqueous H2 [Kelley et al., 2001; Neal and Stanger, 1983; Abrajano et al., 1988]. It has been proposed that H2 is generated under a range of temperatures by hydrolysis on ferrous oxide in olivine [Neal and Stanger, 1983; Holm and Charlou, 2001]. The hydrogen will react (possibly catalyzed by microbial activity) with dissolved oxygen, nitrate, and sulfate. As long as fresh olivine is present, the interacting fluids should hence be anoxic. There are small amounts of fresh olivine present in the interval of iowaite occurrence in Hole 1272A. Maybe the basement Site 1272 is effectively hydrated so that circulating fluids do not react with sufficient amount of fresh material to produce significant H2,aq.
 The presence of nontronite in serpentine muds recovered from Hole 1274A also indicates that water-rock reactions continued at low temperatures and under oxidizing conditions. This is also suggested by the development of aragonite veins with oxidation halos in the uppermost 90 m of Hole 1274A. The aragonite veins disappear abruptly below the first fault gouge at 95 mbsf. Either the fault gouge represents a hydrogeological barrier that prevents cold seawater from penetrating deeper into the basement, or the fault zone accommodates the strain so that fracturing and circulation of cold seawater is limited to the hanging wall of the gouge.