thermal system that have undergone high-temperature sub-ocean-floor alteration. Mineralised rocks (pyrite c quartz c glaucophane c garnet c cummingtonite).
Int Journ Earth Sciences (1999) 88 : 219–235
Q Springer-Verlag 1999
ORIGINAL PAPER
I. Cartwright 7 A. C. Barnicoat
Stable isotope geochemistry of Alpine ophiolites: a window to ocean-floor hydrothermal alteration and constraints on fluid–rock interaction during high-pressure metamorphism
Received: 28 January 1999 / Accepted: 2 February 1999
Abstract The subduction of hydrated oceanic lithosphere potentially transports large volumes of water into the upper mantle; however, despite its potential importance, fluid–rock interaction during high-pressure metamorphism is relatively poorly understood. The stable isotope and major element geochemistry of Pennine ophiolite rocks from Italy and Switzerland that were metamorphosed at high pressures are similar to that of unmetamorphosed ophiolites, suggesting that they interacted with little pervasive fluid during highpressure metamorphism. Cover sediments also have oxygen isotope ratios within the expected range of their protoliths. In the rocks that escaped late greenschistfacies retrogression, different styles of sub-ocean-floor alteration may be identified using oxygen isotopes, petrology, and major or trace element geochemistry. Within the basalts, zones that have undergone highand low-temperature sub-ocean-floor alteration as well as relatively unaltered rocks can be distinguished. Serpentinites have d 18O and d 2H values that suggest that they were formed by hydration on or below the ocean floor. The development of high-pressure metamorphic mineralogies in metagabbros occurred preferentially in zones that underwent sub-ocean-floor alteration and which contained hydrated, fine-grained, reactive assemblages. Given that the transformation of blueschist-facies metabasic rocks to eclogite-facies assemblages involves the breakdown of hydrous minerals (e.g. lawsonite, zoisite, and glaucophane), and will thus liberate considerable volumes of fluids, metamorphic fluid flow must have been strongly channelled.
I. Cartwright (Y) Department of Earth Sciences, Victorian Institute of Earth and Planetary Sciences, Monash University, Clayton 3168, Australia Tel.: c61-3-99054879; Fax: 61-3-99054903 A. C. Barnicoat Department of Earth Sciences, University of Leeds, Leeds LS2 9JT, UK
High-pressure (quartzccalciteBomphaciteBglaucophaneBtitanoclinohumite) veins that cut the ophiolite rocks represent one possible channel; however, stable isotope and major element data suggest that they were not formed from large volumes of exotic fluids. Fluids were more likely channelled along faults and shear zones that were active during high-pressure metamorphism. Such strong fluid channelling may cause fluids to migrate toward the accretionary wedge, especially along the slab–mantle interface, which is probably a major shear zone. This may preclude all but a small fraction of the fluids entering the mantle wedge to flux melting. Additionally, because fluids probably interact with relatively small volumes of rock in the channels, they cannot “scavenge” elements from the subducting slab efficiently. Key words Fluids 7 Subduction 7 Oxygen isotopes 7 Western Alps 7 High-pressure metamorphism 7 Ophiolites
Introduction Subduction of hydrothermally altered oceanic lithosphere is an important geodynamic process that can potentially transport significant quantities of water and other volatiles (e.g. C, B, Cs, N, S) into the mantle (Bebout 1997). Documenting the fate of these volatiles during subduction is critical to our understanding of metamorphic, tectonic and igneous processes in convergent orogenic belts. Fluids released by devolatilisation can result in metasomatism of the mantle wedge and control island-arc magma generation (e.g. Davidson 1997; Bebout 1997). Relative to MORB, island-arc magmas are enriched in large-ion lithophile elements, notably Ba, Rb, and K, and are depleted in other elements such as Nb, La, and Ce (Davidson 1997). One explanation for this difference in chemistry is element transport in fluids introduced from the subducted material. Additionally, the presence of fluids may facilitate
220
deformation on faults and shear zones within the subducting slab, which may promote the break-up of the slab in the subduction zone. However, despite their potential importance, the role of fluids during high-pressure metamorphism, and the fate of water liberated during dehydration of the subducting crust, is poorly understood. Much fluid escapes from the subducting slab at relatively high crustal levels due to physical dewatering and lowtemperature dehydration reactions that break down clays or micas (e.g. Moore and Vrolijk 1992). These fluids interact with the rocks in the accretionary wedge. However, significant volumes of fluid are probably liberated at depth under blueschist- or eclogite-facies conditions (Ridley and Dixon 1984). Peacock (1990, 1993a, 1993b, 1997) proposed that fluids are able to migrate directly from the subducting slab into the mantle wedge resulting in wholesale hydration of the mantle. Tatsumi (1989) and Davies and Bickle (1991) suggested that this hydrated mantle is carried downwards by induced flow above the downgoing slab until melting occurred as the solidus of hydrous peridotite was intersected. By contrast, Philippot (1993) suggested that the interface between the subducting slab and the overlying mantle has limited permeability and focuses fluid along the top of the slab toward the accretionary wedge. Philippot (1993) considered that hydrous melts, rather than fluids, from the subducted metasediments causes metasomatism of the mantle wedge. The possibility that anhydrous rocks are hydrated and metasomatised within the subducting slab is also contentious. Coarse-grained gabbros locally contain hydrous eclogite-facies mineralogies; however, whether this reflects hydration during subduction (Meyer 1983; Wayte et al. 1989; Rubie 1990) or recrystallisation of rocks that were altered by oceanic circulation systems prior to subduction (Pognante 1985, 1989; Messiga and Scambelluri 1988; Barnicoat and Cartwright 1997) has been debated. In addition, high-pressure metamorphic rocks commonly contain veins that from their mineralogy are constrained to have formed during the highpressure event (Nadeau et al. 1993; Philippot and Selverstone 1991). That veins reflect the presence of fluids is largely accepted, but whether these veins were formed by large volumes of fluids infiltrating from external sources or by more local-scale flow or diffusion is not certain. Bebout (1991, 1997) showed that certain rocks from the Catalina Schist (California), which probably represent the deeper levels of an accretionary wedge, interacted with large volumes of fluid during high-pressure metamorphism. In particular, a variety of lithologies from melange zones have relatively homogeneous d 18O and d 2H values implying that they interacted with significant volumes of fluid. Away from the melange zones, the rocks have heterogeneous isotopic ratios and probably interacted with lower volumes of fluid. Barnicoat and Cartwright (1995, 1997) and Barnicoat and Bowtell (1995) have interpreted oxygen isotope ratios,
petrology, and major element geochemistry of metabasalts and metagabbros from the Western Alps (Switzerland–Italy), which underwent metamorphism significantly deeper within the subduction zone, to have been little modified during high-pressure metamorphism. Whereas these studies are not necessarily contradictory, they focussed on relatively small suites of rocks from which it is hard to make general conclusions. Herein we use oxygen isotope, geochemical, and petrological data from the Pennine ophiolite sequence in the Western Alps and its cover sediments to constrain the large-scale fluid flow patterns in the subducted crust. This incorporates both new data and that from smallerscale studies (Barnicoat and Bowtell 1995; Barnicoat and Cartwright 1995, 1997). To our knowledge, this represents the first attempt to constrain fluid flow in the deeper levels of subduction zones using such a wide range of rocks and, as such, provides a direct comparison with work on structurally higher rocks (e.g. the Catalina Schist).
Hydrothermal alteration in ophiolites Detailed investigations of ophiolites that have been obducted, and thus which have not undergone highpressure metamorphism, constrain the expected major element and stable isotope geochemistry of subducted oceanic crust (Gregory and Taylor 1981; Cocker et al. 1982; Muehlenbachs 1986; Schiffman and Smith 1988). Hydrothermal alteration below the ocean floor produces a characteristic modification of d 18O values from their original igneous values (F5.7‰). In the higher levels of the oceanic crust, interaction with sea water at low temperatures results in an elevation of d 18O values, whereas higher-temperature alteration at deeper levels causes d 18O values to be lowered (Fig. 1). Alteration is generally widespread within the sheeted dykes and pillow basalts but more localised (often around fractures) in the deeper gabbros and ultramafic rocks. The pattern of isotopic resetting is also dependent on whether the rocks are situated in the downflow parts of the system (where temperatures at any given depth are relatively cool) or in the warmer upflow zones (Fig. 1). Fluid flow also alters the major element chemistry of the rocks (Alt et al. 1986; Gillis et al. 1993; Gillis 1995). Basalts from the downflow zones are commonly enriched in Mg and depleted in Ca due to the formation of chlorite and the breakdown of plagioclase. By contrast, Ca- and Mn-rich as well as Mg- and Na-poor epidote-bearing rocks are common in the upflow zones. The upflow zones may also be the sites of base metal and/or copper mineralisation associated with black smokers. Oceanic sediments also have characteristic d 18O values (Hoefs 1997) varying from sandstones with relatively low d 18O values (F10–12‰) through pelitic rocks (d 18OF12–17‰) to calcareous rocks and cherts (d 18O 1 20‰).
221
Fig. 1 Hydrothermal alteration patterns below the ocean floor. Low-temperature alteration occurs in broad downflow zones leading to elevation of d 18O values and enrichment in Mg and depletion in Ca. By contrast, higher-temperature alteration in the relatively narrow upflow zones produce rocks which are Ca- and Mn rich, and Mg- and Na poor, with relatively low d 18O values. Sulphide mineralisation may occur on the ocean floor in blacksmoker-type deposits
weakly or unmetamorphosed Mesozoic–Cenozoic sediments and their basement rocks. Apulian rocks are preserved in the Austroalpine units of the Sesia zone and the Dent Blanche klippe (Fig. 2). These Apulian rocks overlie Piemonte oceanic crust of the Pennine zone, which occurs in the Zermatt-Saas zone in southern Switzerland and analogous units in northern Italy. Locally, the Piemonte zone ophiolites preserve blueschist- and eclogite-facies mineralogies (e.g. Barnicoat and Fry 1986; Pognante 1991). The Piemonte ophiolites overlie continental basement and associated cover rocks of the European plate (e.g. the Monta Rosa, Gran Paradiso, and Dora Mira massifs; Fig. 2). Alpine subduction was initially thought to be of midCretaceous age (120–100 Ma) with subsequent continental collision in the late Cretaceous to early Tertiary (80–60 Ma; Chopin and Moiné 1984). However, recent Sm–Nd and Ar–Ar geochronology (Barnicoat et al. 1993, 1995; Bowtell et al. 1994) suggest that subduction continued to at least the late Cretaceous. Later greenschist-facies overprinting of the high-pressure assemblages occurred at 45–30 Ma (Hunzinker 1970; Moiné 1985; Hurford et al. 1991; Becker 1993; Barnicoat et al. 1995), often associated with the reactivation of thrusts as extensional faults (Merle and Ballèvre 1992; Wheeler and Butler 1993).
Piemonte zone If widespread metamorphic fluid flow within the slab has occurred, some degree of homogenisation of stable isotopes is expected. Due to the larger volumes of fluid required to cause metasomatism (e.g. Barton et al. 1991), major element geochemistry may be less affected by metamorphic fluid flow. However, if the metamorphosed rocks preserve a range of stable isotope ratios and major element concentrations similar to those of the obducted ophiolites, limited pervasive fluid flow during high-pressure metamorphism is implied. In addition, if high-pressure veins were formed by infiltration of external fluids, they would probably have distinctly different d 18O values to those of their host rocks. Thus, examination of the stable isotope systematics, geochemistry, and petrology of high-pressure metamorphic rocks provides important constraints on fluid flow during subduction.
Geology of the Western Alps The Western Alps record the subduction of the Piemonte Ocean and subsequent collision of the southern Apulia continent with Europe (e.g. Coward and Dietrich 1989). There is a broad division of the Alps into: (a) Internal Zones that contain relics of oceanic crust and which locally preserve evidence of being metamorphosed at high pressures and temperatures; and (b) External Zones that are dominated by
The Piemonte zone comprises a lower unit dominated by ophiolites and an upper unit that contains abundant carbonate-rich metasediments. The Zermatt-Saas zone and analogous rocks in the Aosta region of northern Italy preserve a dismembered ophiolite sequence with all members recognised (Fig. 2b) including manganiferous cherts, pillow lavas, sheeted dykes, gabbros, and serpentinites (which probably represent altered ultramafic lithologies). Detailed descriptions of the petrology and geochemistry of these rocks have been presented by Bearth (1953, 1967, 1973), Dal Piaz (1965), Chinner and Dixon (1973), Dal Piaz and Ernst (1978), Bearth and Stern (1979), Meyer (1983), Martin and Kienast (1987), Tartarotti et al. (1986), Martin and Tartarotti (1986), Pfeifer et al. (1989), Bearth and Schwander (1981), Barnicoat and Fry (1986), Barnicoat and Bowtell (1995), and Barnicoat and Cartwright (1995, 1997), and only brief details are given herein. The ultramafic rocks comprise mainly serpentinite (antigorite) with minor magnetite. Locally relict augite and olivine are present. Titanoclinohumite, chlorite, secondary forsterite and diopside, chlorite and, locally, calcite are minor metamorphic phases. Talc–chlorite–actinolite schists that occur within, and especially at, the margins of larger serpentinite bodies may reflect alteration zones produced by late greenschist-facies fluid flow. Centimetre- to metre-scale lenses that contain coarse-grained (up to 10 cm diameter) epidote,
222 Fig. 2 a Geology of part of the western Alps showing major geological units and areas described in this paper. Ta Täschalp; Sa Saas-Fee; Sv Servette; Vt Valtournanche. b Detailed Map of Täschalp showing localities: P Pulwfe; H Hubeltini; M Mellachen; m1, m2 metasediment outcrops. (After Barnicoat et al. 1995)
grandite garnet, vesuvianite and/or cordierite are probably metamorphosed rodingites. As illustrated by the Allalin gabbro in the ZermattSaas region, the metagabbros commonly display only partial replacement by high-pressure minerals (Bearth 1967; Meyer 1983; Pognante 1989; Barnicoat and Cartwright 1997). A range of protoliths (including olivine gabbro, troctolite, and trondhjemite) are present and magmatic textures with grain sizes of up to 5 cm are variably preserved. There is a transition from rocks preserving magmatic mineralogies (plagioclase c augite c olivine) to those where olivine is replaced by talc c kyanite B chloritoid, plagioclase is replaced by omphaciteczoisiteckyanite, and augite is replaced by omphacitectalc. Glaucophane, paragonite and phengite locally replace the eclogite-facies minerals, whereas later greenschist-facies alteration produced talc, chlorite, actinolite and albite in some rocks. Metabasalts representing metamorphosed dykes, pillows, and flows contain a variety of mineral assemblages. The most common high-pressure minerals are omphacite, garnet, paragonite, clinozoisite, quartz, and glaucophane with minor rutile, apatite, dolomite, and
calcite. Lawsonite, now pseudomorphed by clinozoisitecparagonitecquartz, is also common, and talc, chloritoid and kyanite are locally present. Variations in high-pressure mineral assemblages partially reflect compositional variations (Barnicoat and Bowtell 1995). For example, glaucophane-bearing omphacite-absent blueschists are richer in Mg and poorer in Ca than the eclogites, whereas pillow lavas at Täschalp (Fig. 2b) have eclogite facies assemblages in their cores and blueschist-facies assemblages in their rims that is also a consequence of a difference in chemistry (Barnicoat 1988a). As discussed herein, the variation in chemistry often correlates with oxygen isotope ratios and most probably reflects sub-ocean-floor alteration prior to high-pressure metamorphism. Locally intense retrogression of the high-pressure mineral assemblages to greenschist-facies albitecchloritecactinolitecepidote assemblages has occurred adjacent to thrust planes and around albite veins (Barnicoat 1988b; Miller et al. 1998). In addition, many samples away from albite veins or thrusts show minor to moderate replacement of the high-pressure minerals by greenschist-facies assemblages.
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The ophiolitic rocks are in fault contact with metasediments (e.g. Bearth 1953; Debelmas 1980) that include: (a) pelites dominated by quartz, garnet and phengitic white mica, often with kyanite, chlorite and talc, and locally with minor calcite and epidote; (b) calcareous schists that are part of the schists lustrés or Bündnerschiefer that occur throughout the western European Alps, and which comprise quartz, calcite, and phengite together with chlorite; (c) quartzites that represent metamorphosed clastic sediments, rather than cherts, which are dominated by quartz, locally with millimetre-wide segregations of muscovite; and (d) albitecquartz-bearing gneisses that possibly represent metamorphosed arkoses or volcaniclastic rocks
Conditions of metamorphism Peak metamorphic conditions for the Täschalp region (Fig. 2) were constrained to be 550–600 7C at 1700–2000 MPa by thermobarometry and phase diagram considerations (Barnicoat and Fry 1986). At Lago Cignana, 22 km to the SW in the Valtournanche region (Fig. 2), cover sediments to the Zermatt-Saas ophiolite are coesite bearing, and the peak of metamorphism was probably at 2600–2800 MPa and 590–630 7C (Reinecke 1991). The Allalin gabbro, which occurs between Zermatt and Saas Fee (Fig. 2), underwent eclogite facies metamorphism at 550–650 7C and 1 2000 MPa (Meyer 1983; Barnicoat and Cartwright 1997). Figure 3 shows some reactions in the system Na2O–CaO–MgO–Al2O3–SiO2–H2O constructed using updated versions of Thermocalc (Powell and Holland 1989) and the Holland and Powell (1990) thermodynamic database (R. Powell, unpublished data). The reactions form part of a larger set of reactions between albite, chlorite, glaucophane, grossular, kyanite, lawsonite, paragonite, quartz, talc, tremolite, zoisite, and H2O for quartz-bearing assemblages. Since the focus of this paper is not the petrological evolution of the rocks, we have only depicted a few equilibria that describe reactions in the metabasic rocks and the invariant points at which they terminate. Reactions (2) and (3) account for the breakdown of early lawsonite to clinozoisitecparagonitecquartz assemblages, with the liberation of water. Albite may replace jadeitic pyroxene at reaction (1). Late growth of tremolitic amphibole and albite was possible by reactions such as (7), whereas later growth of chlorite may have been by reactions such as (4), (5), (6) or (8). The late growth of talc from higher-pressure assemblages may have been via reaction (9). Overall, these reactions may be intersected by P–T–time paths of the form proposed for the Zermatt-Saas region (Barnicoat and Fry 1986) suggesting that the mineralogical changes are part of a single metamorphic episode.
Fig. 3 Reactions in the system Na2O–CaO–MgO–Al2O3–SiO2–H2O that may explain mineral changes in the metabasalts. The mineral changes are consistent with metamorphism along a simple P–T–time path
Analytical techniques Samples for whole-rock stable isotope and XRF analyses were crushed to a fine powder in a steel mill. Mineral separates for stable isotope analyses were hand picked from cubes of rock (F1 cm on a side) that had been coarsely crushed and sieved. Stable isotope ratios were measured at Monash University. Oxygen isotope ratios of silicates were analysed following Clayton and Mayeda (1963) but using ClF3 as the oxidising reagent. Carbon dioxide was extracted from calcite by reaction with H3PO4 at 25 7C for 4–18 h in sealed vessels (McCrea 1950). The extracted gases were analysed as CO2 on a Finnigan MAT 252 mass spectrometer (Finnigan MAT, Bremen, Germany), and the results are expressed relative to V-PDB (C) and V-SMOW (O). Internal and international standards run at the same time as the samples from this study generally yielded d 18O and d 13C values within 0.2‰ of their accepted values. The long-term average d 18O value of NBS 28 at Monash is 9.55B0.11‰. Reproducibility based on replicated analyses of standards and samples is B0.15‰ (1s) for silicate d 18O, and B0.1‰ (1s) for carbonate d 18O and d 13C values. Hydrogen isotope ratios were analysed by a modification of the technique of Vennemann and O’Neil (1993) using a furnace (heatable up to 1500 7C) instead of an oxyacetylene torch. Conversion of H2O to H utilised Zn provided by Indiana University, with the precision and accuracy of Zn reduction monitored using water standards. d 2H
224
values are expressed relative to V-SMOW, and reproducibility is 2–3‰. Major element analyses were made on a Siemens SR S3000 (Siemens, Erlangen, Germany) XRF spectrometer at the University of Melbourne using fused beads: precision is B1% relative (1s).
Stable isotope geochemistry Oxygen isotope ratios of ophiolite rocks Figure 4a and Table 1 show the d 18O values of ophiolitic rocks from several localities in the Zermatt-Saas region and the Aosta Valley (Fig. 2). The samples were collected away from sites of intense late greenschistfacies alteration (such as albite veins or faults), although some show the effects of greenschist-facies overprinting. The protoliths of many of the rock types are relatively clear in the field; however, some basalts may represent dykes where the margins have not been recognised due to metamorphism, deformation or locally-poor outcrop. There is a distinct correlation of d 18O values with lithology: ultramafic rocks (serpentinites and rodingites) have d 18O values between 1.2 and 5.9‰ (average 2.8B1.5‰, np12); gabbros have d 18O values between 2.3 and 6.6‰ (average 5.2B1.0‰, np26); dykes have d 18O values between 4.4 and 5.7‰ (average 5.1B0.5‰, np4); homogeneous basalts have d 18O values between 4.0 and 9.5‰ (average 7.2B1.3‰, np58); and pillow lavas have d 18O values between 7.1 and 9.1‰ (average 7.6B0.7‰, np14). With the exception of the serpentinites, the d 18O values of individual rock types from each locality overlap. The d 18O values are very similar to those of certain ophiolite complexes that have escaped high-pressure metamorphism, such as Macquarie Island (Fig. 4b). Unlike ophiolites that have not undergone highpressure metamorphism, metamorphosed ophiolites have lost up to 90% of their original water contents (Philippot 1993; Peacock 1993a, 1993b); hence, the effects of metamorphic devolatilisation on d 18O values needs to be considered. Due to the small fraction of oxygen that is lost from the rock, and the small isotopic fractionations between phases at elevated temperatures, changes in oxygen isotope ratios due to metamorphic devolatilisation should be small. For example, removing 6% H2O during the transition from lawsonite blueschist (18% glaucophane, 28% lawsonite, 19% chlorite, 29% clinopyroxene, 6% quartz) to anhydrous eclogite (38% clinopyroxene, 51% garnet, 11% quartz) at 500 7C (mineral modes based on Peacock 1993b) changes d 18O values by ~0.2‰ (calculated using the practionations of Zheng 1993, and Matthews 1994). Similar calculations have been made for other metamorphic transformations in ophiolitic rocks (Barnicoat and Cartwright 1995), as well as for metamorphic rocks in general (Valley 1986); hence, closed-system devolatilisation is unlikely to significantly affect the d 18O values of these rocks.
Fig. 4 a d 18O values of ophiolitic rocks from the Zermatt-Saas region and the Aosta Valley (data from Table 1). There is a distinct correlation of d 18O values with lithology which is very similar to that observed in ophiolite complexes that have escaped high-pressure metamorphism, such as b that on Macquarie Island. (Data from Cocker et al. 1982)
As discussed previously, the variation in d 18O values through the oceanic crust (e.g. Fig. 4b) reflects the temperature of alteration by hydrothermal systems close to the mid-ocean ridges. Overall, the data in Fig. 4a suggest that the Alpine ophiolites preserve d 18O values that reflect alteration below the ocean floor. As discussed below, there are strong correlations between d 18O values, mineralogy, and geochemistry within the ophiolite profile which suggest that specific zones of sub-ocean-floor alteration can be recognised. The range of d 18O values in the Alpine rocks is broader than in typical ophiolite sections. This may be the result of samples being collected over a wide area and representing protoliths that have experienced variable subocean-floor alteration. Additionally, although we attempted to collect the freshest rocks, many of the samples show at least minor retrogression with late tremolitic amphibole and/or chlorite replacing some of the earlier phases suggesting that they have experienced some fluid flow during uplift. Whereas any fluid flow at that time was insufficient to obscure the overall pattern of d 18O values, it may have caused minor isotopic resetting.
225 Table 1 Stable isotope geochemistry of ophiolite rocks. Cc calcite; Cpx clinopyroxene; Gla glaucophane; Gnt garnet; Qtz quartz; WR whole rock; ctd chloritoid Sample
Wall rocks 18
d O WR
18
d O Cc
Da
Veins 13
d C Cc
18
d O Qtz
18
d O Gla
18
d O Cpx
18
d O Gnt
18
d O Qtz
18
d O Cc
Notes
13
d C Cc
Pillow lavas Mellichen (M) b JM96SA363 7.1 JM96SA364 7.6 JM96SA365 7.2 JM96SA366 7.4 JM96SA367 9.1 JM96SA368 9.0 JM96SA369 7.4 JM96SA370 7.7 JM96SA371 7.2 JM96SA372 7.1 JM96SA373 7.3 JM96SA374 7.2 JM96SA375 7.5 JM96SA376 8.0
Blueschist Blueschist Blueschist Blueschist Blueschist Blueschist Blueschist Blueschist Blueschist Blueschist Blueschist Blueschist Blueschist Blueschist
Metabasalts Mellichen (M) 96A9 7.4
8.1
6.0
96A10
6.9
8.0
6.3
4.6
96A11
5.8
8.1
0.0
8.5
5.2
4.3
96A17
4.4
7.6
0.0
96A18
5.7
9.9
0.0
96A21
8.7
96A22
8.5
Eclogite (r) f Eclogite (r) Eclogite Eclogite 8.2 10.5
5.3
5.2
7.6
–3.3
Eclogite Blueschist (r)
–2.2
Blueschist (r)
–2.2 –2.5
Blueschist (r) Blueschist Blueschist
12.0 10.7 96A23
7.6
96A24 96A25A 96A25 96A26 96A27A 96A27B 96A28 96A29 96A32 96A33 96A34 96A39 97A68/69 97A70 97A71 97A72 97A74
9.2 8.5 8.0 9.0 9.2 8.1 8.7 6.1 7.4 7.6 8.4 7.7 8.3 7.5
10.5 11.0
0.1 0.1
10.7
–0.1
10.1
–0.3
10.7
7.8
11.8
8.2
10.5 8.2 10.4 10.5
11.4 11.0 11.1
7.8 5.2 7.2
11.2 11.7 11.5 10.5 11.4
10.7
–0.5
10.5
–0.1
11.0 10.7
0.0 0.0
9.8
–0.3
0.6
12.0 9.8
–3.7
Blueschist (r)
–3.1 –3.8
7.5 11.7
11.5
–4.0
Blueschist (r) Eclogite (r) Eclogite (r) Eclogite (r) Eclogite (r) Eclogite (r) Blueschist (r) Blueschist (r) Blueschist (r) Blueschist (r)
–2.2 –2.5 –3.4
7.4
11.3
–0.7
Blueschist
Hubeltini (H) 96A51 7.3 96A56 6.7
Eclogite (r) Eclogite (r)
Plufwe (P) 83175 83176 85031 85019 85061 85065 85035
Eclogite Eclogite Eclogite Eclogite Talc-ctd Talc-ctd Talc-ctd
5.7 5.4 5.1 5.6 6.4 6.2 6.0
For continuation of Table 1 please see next page
226 Table 1 Continued Sample
Wall rocks
d 18O WR 85064 96A58 96A59 96A60 96A61 96A62 96A63 Saas (Sa) 97A32 97A33
6.6 7.7 7.8
Valtournanche 96A67 96A68/69 96A70 96A71 96A74 96A75 96A82 96A85 96A86 96A87 96A88
(Vt) 7.5 8.4
Servette (Sv) d 53229 52611 52612 91048 91049 91051 91052 53230 53231 53232
d 18O Cc
Da
Veins
d 13C Cc
d 18O Qtz
d 18O Gla
d 18O Cpx
d 18O Gnt
d 18O Qtz
d 18O Cc
d 13C Cc
10.0 10.2 10.8 10.3 10.4
7.8 7.7 8.1
–2.3 –2.4 –2.5 –2.7
6.5 6.4
7.3 7.4 7.6 7.2 9.5 7.0
14.2
9.9
0.4
0.2
6.8 5.9 5.2 4.8 4.9 5.3 5.2 4.6 8.0 8.1
10.5
7.9
10.2
7.0
9.5
6.9
5.4
11.5 8.4 10.4
10.8
0.3
10.4
0.7
9.5 9.2 9.7
8.8
0.2
Blueschist (r) Blueschist (r) Blueschist (r) Blueschist Blueschist Blueschist Eclogite (r) Eclogite (r) Eclogite (r) Blueschist Blueschist Blueschist Talcschist Talcschist Talcschist Blueschist Talcschist Mineralised Mineralised
Dykes Mellichen (M) 97A75 5.0 97A76 5.7 97A77 5.1 97A78 4.4
Saas (S): Allalin gabbro e 51292 6.6 51293 6.6 51600 6.0 51601 5.1 51606 5.3 51607 4.5 51608 7.6 51609 2.3 51610 4.8 51611 4.9 AL85/4 4.8 AL85/13 5.8 AL85/14 4.9
Talc-ctd Blueschist Blueschist Blueschist Blueschist Blueschist Blueschist (r) Eclogite (r) Eclogite (r)
7.1
Gabbros Mellichen (M) 97A61 6.1 97A62 5.6 97A64 5.1 97A65 8.5 97A66 8.3 97A67 8.1
Notes
Blueschist (r) Blueschist (r) Blueschist Blueschist
8.3
0.4
11.1
0.3
Altered Altered Altered Altered Altered Altered Relict magmatic Relict magmatic Relict magmatic Altered Altered Altered Altered Altered Altered Altered Altered Altered Altered
227 Table 1 Continued Sample
Wall rocks
d 18O WR AL85/18 AL85/20 AL85/24 AL85/27 AL85/28 AL85/33 AL85/44 AL85/47 AL85/53 AL85/55 AL85/56 AL85/57 D835
d 18O Cc
d 13C Cc
d 18O Qtz
d 18O Gla
d 18O Cpx
d 18O Gnt
d 18O Qtz
5.4 5.3 5.0 3.8 4.5 4.8 4.9 5.1 4.7 5.6 5.3 5.4 5.7
d 18O Cc
Notes
d 13C Cc Altered Altered Altered Altered Altered Altered Altered Altered Relict magmatic Relict magmatic Relict magmatic Relict magmatic Relict magmatic
Serpentenites and rodingites Valtournanche (Vt) 96A89 2.9 96A90 2.4 96A91 2.7 96A92 2.2 96A93 1.6 96A94 2.0 96A95 3.2 (–66) c 96A96 1.5 (–70) 96A97 2.2 (–71) 96A98 1.2 (–68) 96A99 1.9 (–75) 96A100 1.6 Mellichen (M) 97A57 5.6 97A58 5.8 97A73 4.5 97A79 5.8 97A80 5.9
Da
Veins
Rodingite Rodingite Rodingite Rodingite Rodingite Rodingite Serpentenite Serpentenite Serpentenite Serpentenite Serpentenite Serpentenite
11.0 9.6
0.3 0.5
10.0
–0.3
Serpentenite Serpentenite Serpentenite Rodingite Rodingite
a
d 18O(Vein Qtz)-d 18O(WR) Localities in Fig. 3 c 2 d H values in parentheses
d
b
e
Carbon isotope ratios of calcite in the ophiolite
values of 0.1–2.2‰ (Staudigel et al. 1997) metamorphism of such rocks would probably produce calcite with d 13C values similar to those in the Alpine rocks.
Calcite is a minor component in many of the ophiolite rocks (Table 1). The d 18O values of the calcite vary with whole-rock d 18O values; however, d 13C values are remarkably consistent (0.1B0.3‰) and similar to those of calcite in the Piemonte zone sediments. Calcite with similar d 13C values (0.0B0.4‰) occurs in the high-pressure veins (Table 1) and also within and around late greenschist- facies albite veins (–0.2B0.2; I. Cartwright, unpublished data). However, in most rocks calcite occurs as grains intergrown with the silicate phases, and it is not obviously associated with those vein sets. Calcite has a sedimentary carbon signature, suggesting that it was precipitated from fluids in equilibrium with the overlying calcareous sediments or directly from seawater. The calcite probably formed during the subocean-floor alteration that was discussed previously. Altered basalts from DSDP sites 417 and 418 contain measurable, and in some cases major, amounts of carbonate (CO2 contents from 0.4 to 32.1%), with d 13C
f
Data from Barnicoat and Cartwright (1995) Data from Barnicoat and Cartwright (1997) Sample showing minor greenschist retrogression
Cover sediments Figure 5 and Table 2 summarise the d 18O values of the Piemonte zone metasedimentary rocks that originally overlaid the ophiolite, and which are now in tectonic contact with it. The pelites have d 18O values of 8.3–16.1‰ (average 11.9B2.8‰, np7), micaceous quartzites have d 18O values of 12.3–14.3‰ (average 12.9B0.7‰, np7), and albitic gneisses, which are probably metamorphosed arkoses, have d 18O values of 8.8–12.5‰ (averagep11.0B1.3‰, np4), whereas calcareous schists have d 18O values of 19.5–24.4‰ (averagep21.7B1.7‰, np11). Calcite in the calcareous schists has d 18O and d 13C values of 19.8–26.2 and –1.2 to 2.3‰, respectively. Overall, the oxygen- and carbon isotope ratios of most of these rocks are similar
228 Table 2 Stable isotope geochemistry of cover sediments. Cc calcite; Qtz quartz; WR whole rock; Wt mica phengitic white mica Sample d 18O WR
Fig. 5 d 18O values of the Piemonte zone metasedimentary rocks that originally overlaid the ophiolite (data from Table 2). The oxygen isotope ratios of most of these rocks are similar to those expected for metasediments of these compositions (Hoefs 1997) suggesting that they have been little altered by interaction with fluids during metamorphism
to those expected for metasediments of these compositions (Hoefs 1997) suggesting that they have been little altered by interaction with fluids during metamorphism. Two samples of pelite from Täschalp have lower than expected d 18O values (8.3 and 8.9‰). Petrologically, these rocks are similar to adjacent ones that have more normal d 18O values, and their low d 18O values may reflect local resetting by fluids. High- and low-temperature ocean-floor alteration in basalts Figure 6 shows that within the basalt samples there is a close correspondence between the d 18O values and the degree and style of sub-ocean-floor alteration where the latter can be recognised (Table 1). In this grouping we have only included the freshest, least-retrogressed samples, to avoid possible complications caused by later low-temperature alteration. Samples were designated as undergoing little sub-ocean-floor alteration, high-temperature sub-ocean-floor alteration, and lowtemperature sub-ocean-floor alteration on the following basis (Barnicoat and Bowtell 1995; Barnicoat and Cartwright 1995): rocks with eclogite (garnetcomphacite-bearing) mineralogies have major element geochemistry that is most similar to MORB and are interpreted as experiencing little sea-floor alteration. By contrast, blueschists (glaucophane B garnet-bearing without omphacite) and talc–chloritoid rocks (garnet c talc c chloritoid c glaucophane c rutile c omphacite) have higher Mg but lower Ca, Mn, Na and K contents than the eclogites. These rocks have compositions similar to those in unmetamorphosed ophiolites (such as Troodos) that have experienced low-temperature alteration in the downflow regions of hydrothermal alteration cells. Clinozoisite-rich rocks and talcschists (talc c garnet c glaucophane c rutile) that are depleted in Si relative to the eclogites probably represent rocks from the upflow zones of the hydrothermal system that have undergone high-temperature
d 18O Cc
d 13C Cc
Pelites Taschalp (m1) a 97A14 12.0 97A15 13.7 Taschalp (m2) 97A25 8.3 97A26 8.9 97A28 10.5 Saas Fee (Sa) 97A36 16.1 97A39 13.4 Quartzites Taschalp (m2) 97A8 14.3 97A9 13.3 97A19 12.5 97A21 12.9 Saas Fee (Sa) 97A41 12.5 97A42 12.4 97A43 12.3 Calcareous schists Taschalp (m2) 97A12 21.9 97A13 21.9 97A17 97A18 19.8 Tashalp (m1) 97A22 23.3 97A23 24.1 97A24 20.3 97A29 19.5 97A30 23.8 Saas Fee (Sa) 97A34 23.2 97A35 20.6 97A38 24.4
d 18O Qtz
d 18O Wt mica
13.8 14.3
11.2 11.8
18.1 14.8
15.5 12.0
13.6 12.7
10.8 10.1
12.8 12.5
10.0 9.9
23.2 22.3
–0.4 –1.2
23.9 22.8
23.1
0.0
23.7
24.6 26.2 20.3 21.8 21.3
1.6 1.9 0.3 2.2 0.8
20.9
20.2 23.9
0.7 1.4
20.6 24.8
Albitic gneisses Taschalp (m1) 97A20 8.8 Saas Fee (Sa) 97A46 11.3 97A47 11.4 97A49 12.5 a
Localities from Fig. 3
sub-ocean-floor alteration. Mineralised rocks (pyrite c quartz c glaucophane c garnet c cummingtonite) that occur in association with talcschists at Servette (Fig. 2) probably represent fossil black-smoker-type deposits (Barnicoat and Bowtell 1995). The clinozoisite-rich and talcschist lithologies are less common than the blueschist-facies rocks, which accords with observations in unmetamorphosed ophiolites that fluid downflow occurs over broad areas, whereas upflow was much more focussed. The samples that have undergone high-temperature alteration generally have lower d 18O values, whereas those that underwent low-temperature
229
There is a greater, however, with many degree of scatter in calcite d 18O values, however, with many calcite–quartz pairs have fractionations within the expected range, suggesting that calcite d 18O values were also little reset following the peak of metamorphism. Overall, these data suggest that the minerals approached isotopic equilibrium at the peak of highpressure metamorphism and were preserved during exhumation. Minor greenschist facies alteration
Fig. 6 d 18O values of unretrogressed metabasalts that vary in their degree and style of sub-ocean-floor alteration (data from Table 1). The range of d 18O values is similar to that of equivalent unmetamorphosed ophiolitic rocks. For samples that have experienced late greenschist facies retrogression, the overall range of d 18O values is wider than the unaltered rocks and there is not the same clear division in d 18O values between the eclogites and the blueschists
alteration have higher d 18O values. As shown in Fig. 1, this is the pattern that is expected to develop during sub-ocean-floor hydrothermal alteration, suggesting that the major element and isotope geochemistry of these rocks has remained unaltered through high-pressure metamorphism. Mineral fractionations Figure 7 shows d 18O values of calcite, clinopyroxene, glaucophane and white mica from the ophiolite and the cover sequence, plotted as d 18O(mineral) vs d 18O(Qtz), data from Tables 1 and 2. We analysed mainly minerals from unretrogressed to moderately retrogressed rocks where high-pressure minerals were preserved. The lines show the expected 18O fractionations at 700–400 7C (calculated from the data of Zheng 1993, and Matthews 1994). Whereas stable isotope geothermometry is often imprecise as it uses d 18O values determined from aggregates of minerals (and thus depends on the purity of mineral separates, individual grains having identical d 18O values, and minerals not being appreciably zoned in their d 18O values), this approach is useful in identifying isotopic disequilibrium between minerals. Whereas there is some scatter in the data, most mineral pairs yield isotopic temperatures close to those of highpressure metamorphism (550–650 7C). The mafic minerals analysed in the ophiolite (clinopyroxene, garnet and glaucophane) and the white mica from the cover sequence are those that were stable at the peak of highpressure metamorphism, suggesting that quartz d 18O values have not been significantly reset subsequently.
Although care was taken to sample the freshest material, many of the rocks have undergone post-peak metamorphic retrogression (Table 1), typically involving partial replacement of the high-pressure minerals by tremolite, chlorite and albite. In all cases the rocks still contain relict high-pressure mineralogies and the degree of retrogression of these samples is relatively minor. More extensively retrogressed rocks do exist, but they are not included in this study. Locally, greenschist facies alteration occurs adjacent to albite veins (Barnicoat 1988a, 1988b; Miller et al. 1998); however, areas of patchy alteration apparently not associated with veins are also common. The range of d 18O values of these rock types is somewhat wider than the unaltered rocks, and there is not the same clear division in d 18O values between the eclogites and the blueschists (Fig. 6). Overall, however, there is little difference between the d 18O values of unaltered and altered metabasalts (although the latter may have slightly higher d 18O values). This suggests that either the volumes of fluids involved in retrogression were relatively minor (the volume of fluid required to change mineralogy is far less than that required to reset stable isotope ratios), or that the fluid was derived from the ophiolitic rocks themselves and was thus approximately in isotopic equilibrium with the rocks through which it flowed. The generation of fluids during uplift may be promoted as rocks decompress through dehydration reactions that have a positive slope in P–T space (e.g. Miller et al. 1998). It is possible that the fluids may have come from the surrounding rocks (e.g. the overlying metasediments that may also have generated fluids during decompression (cf. Bousquet et al. 1998). Extensively retrogressed samples from close to the contacts with the metasediments do have high d 18O values that approach those of the metasediments (I. Cartwright and A.C. Barnicoat, unpublished data) suggesting interaction with fluids that were derived from, or which had exchanged with, the metasediments. However, how far such fluids have penetrated into the ophiolitic rocks is not known. The “blurring” of d 18O values within the retrogressed rocks illustrates the importance of using the least-retrogressed samples to determine patters of sub-ocean-floor alteration.
230
Fig. 7 d 18O values of quartz and coexisting minerals from a–c the ophiolite rocks, d, e the cover sequence, and f the high-pressure veins in the ophiolite. The isotopic fractionations imply that many minerals were in mutual isotopic equilibrium at peak metamorphic temperatures (550–650 7C). Data from Tables 1 and 2; bars show analytical errors. Predicted fractionations from Zheng (1993) and Matthews (1994)
Timing of fluid flow in gabbros As discussed in more detail by Barnicoat and Cartwright (1997), the stable isotope geochemistry can help to constrain the timing of fluid flow in Alpine gabbros.
The Alpine gabbros have undergone only partial transformation to eclogite facies parageneses. Within coarsegrained gabbro units, such as the Allalin gabbro near Saas Fee (Fig. 2), there is a complete transition between gabbros preserving relict magmatic textures (where eclogite facies minerals are largely restricted to grain boundaries), and eclogite facies gabbros where: (a) olivine is replaced by zoned pseudomorphs with talcckyanite cores, chloritoidcomphacite outer cores, and garnet rims; (b) plagioclase is pseudomorphed by zoisite, omphacite and kyanite; and (c) augite is pseudomorphed by omphacite, talc and rutile (Bearth 1973; Chinner and Dixon 1973; Meyer 1983; Wayte et al.
231
1989; Barnicoat and Cartwright 1997). The eclogitic assemblages are hydrous, and fluid flow has been alternatively suggested to have occurred during subduction (Meyer 1983; Wayte et al. 1989; Rubie 1990) or below the ocean floor with recrystallisation at eclogite facies producing the current mineralogy (Pognante 1985, 1989; Messiga and Scambelluri 1988; Barnicoat and Cartwright 1997). As discussed by Barnicoat and Cartwright (1997), the eclogitic gabbros generally have lower d 18O values (as low as 2.3‰) than those that preserve magmatic relicts (Table 1). This is consistent with 18O depletion during high-temperature alteration below the ocean floor. Additionally, d 18O values of individual mineral sites within the eclogitic gabbros are similar to those of altered ocean-floor gabbros that have escaped high-pressure metamorphism, and the zoned olivine pseudomorphs have compositional variations that are similar to altered olivines in ocean-floor gabbros (Barnicoat and Cartwright 1997). Overall, hydration of the Allalin gabbro probably occurred prior to high-pressure metamorphism. The selective development of eclogite facies mineralogies in rocks that have undergone sub-ocean-floor alteration is likely due to those rocks containing hydrated, fine-grained and, hence, reactive mineralogies. Serpentinisation of ultramafics Metamorphosed ultramafic rocks in the Alps are generally serpentinised; however, such serpentinisation may reflect fluid flow in one of numerous settings: below the ocean floor, during high-pressure metamorphism, or during uplift and retrogression. Stable isotope geochemistry of serpentinites may be used to distinguish between fluid sources and temperatures of alteration (Wenner and Taylor 1973, 1974; Heaton and Sheppard 1977; Ikin and Harmon 1983; Burkhard and O’Neil 1988). Serpentinites from the ocean floor have relatively high d 2H (–35 to –65‰) and low d 18O (1–7‰) values (Fig. 8), reflecting their formation by interaction with ocean water at 300–400 7C. By contrast, serpentinites in some ophiolites have lower d 2H and higher d 18O values (Fig. 8) reflecting interaction with low-temperature (possibly ~150 7C; Barnes and O’Neil 1969) meteoric water during or after obduction. Serpentinites from Valtournanche (Fig. 2) have d 18O and d 2H values that plot within the oceanic chrysolite/lizardite field (as defined by Wenner and Taylor 1974), which are significantly different from the expected values for serpentinites formed by fluid flow in continental settings. We interpret these data to indicate that the original hydration occurred on, or below, the ocean floor rather than at any stage during metamorphism. The serpentinites were collected from large masses; other serpentinites from the Täschalp region that occur along fault planes may have formed or have been remobilised by interaction with metamorphic or meteoric fluids and may have a different isotopic signature. Burkhard and O’Neil
Fig. 8 d 18O and d 2H values of serpentinites from Valtournanche. These rocks plot within the oceanic chrysolite/lizardite field (as defined by Wenner and Taylor 1974), suggesting that hydration occurred on, or below, the ocean floor, rather than during metamorphism
(1988) describe serpentinites from the eastern Alps that have a wide range of d 18O (4–13‰) and d 2H values (–134 to –40‰) which probably formed by interaction with a wide range of fluids (meteoric, oceanic and metamorphic). Due to their susceptibility to alteration, ultramafic rocks are potentially sensitive indicators of fluid flow. Formation of high-pressure veins The metabasalts are cut by centimetre- to decimetrewide quartz veins that additionally may contain calcite, kyanite, glaucophane, rutile, garnet or omphacite. The mineralogy of these veins suggests that they were formed during high-pressure metamorphism. Because the veins occupy ~~1% of the outcrop at Täschalp and Valtournanche, they probably do not represent the products of a large-volume fluid flow system; however, discerning whether they are produced by external fluids or from fluids derived from the local rocks is important
232
in understanding fluid flow during high-pressure metamorphism. Figure 7f shows that coexisting quartz and calcite in the veins have d 18O values that imply that these minerals were in mutual isotopic equilibrium at F600 7C. These temperatures are close to those estimated by Barnicoat and Fry (1986) for the peak of metamorphism at Täschalp, which is consistent with the veins forming during the high-pressure event. Figure 9 shows that the d 18O values of quartz in the veins correlate broadly with the d 18O values of their immediate host rocks (collected 1–20 cm from the vein). D 18O(Qtz-WR) values are between 2.2 and 4.0‰ (average 2.8B0.6‰). For a “typical” eclogite comprising 15–20% garnet, 20–30% omphacite, 15–25% glaucophane, 10–15% paragonite, 5–10% zoisite and 5–10% quartz with mineral compositions as recorded for the Zermatt-Saas region (Barnicoat and Fry 1986), the expected D 18O(Qtz-WR) values at 600 7C are 2.5–3.0 (calculated from the data of Matthews 1994, and Zheng 1993). Minor variations in mineralogy or the assumed temperature of veining make relatively little difference to these values. That the D 18O(Qtz-WR) values are close to those predicted suggests that the veins approach isotopic equilibrium with their hosts, and were locally derived by diffusion or small-scale fluid flow. Further evidence for local vein derivation comes from observations that wallrocks within 20 cm of one of the kyanite-bearing veins is depleted in Al2O3 and SiO2 (the major components of the veins; Fig. 10; Table 3). Overall, we interpret the data to indicate that the high-pressure veins were formed from fluids that were derived from, or which had equilibrated with, the surrounding rocks. One plausible source of internally generated fluids during high-pressure metamorphism are those liberated by the breakdown of lawsonite (Fig. 3; e.g. Barnicoat and Fry 1986) or the breakdown of early formed glaucophane to form omphacite.
Fig. 10 Geochemical changes adjacent to a kyanitecquartzbearing high-pressure vein. Over 20 cm the host rock is depleted in Al2O3 and SiO2 (the major components in the vein) suggesting local derivation of the veins (data from Table 3)
Table 3 Major element compositions of rocks adjacent to highpressure vein
SiO2 TiO2 Al2O3 Fe2O3 MgO MnO CaO Na2O K2O SO3 P2O5 LOI Total
96A25A a
96A25B b
96A25C c
50.17 2.52 9.33 11.71 7.08 0.98 9.80 5.61 0.10 0.02 0.93 1.85 100.10
52.17 1.61 9.76 11.37 7.08 0.48 10.86 4.75 0.17 0.01 0.11 1.65 100.02
54.17 2.12 10.87 10.15 7.08 0.53 8.60 4.88 0.15 0.02 0.33 1.15 100.06
a
Distance from vein 1 cm Distance from vein 8 cm c Distance from vein 20 cm b
Discussion
Fig. 9 d 18O values of quartz and calcite in high-pressure veins. Vein d 18O values correlate broadly with those of their host rocks, suggesting that the veins approach isotopic equilibrium with their hosts and were locally derived by diffusion or small-scale fluid flow. Dpd 18O(Vein)Pd 18O(WR)
Overall, this study demonstrates that, in the rocks which escaped significant late greenschist facies retrogression, the stable isotope and major element geochemistry of Pennine ophiolites that were metamorphosed at high pressures is similar to that of unmetamorphosed ophiolites. Cover sediments also have oxygen isotope ratios within the expected range of their protoliths. Furthermore, different styles of sub-oceanfloor alteration may be identified using oxygen isotopes, petrology and major or trace element
233
geochemistry. For example, within the basalts, zones that have undergone high- and low-temperature subocean-floor alteration, as well as relatively unaltered rocks, can be distinguished. Serpentinites have d 18O and d 2H values that suggest that they were formed by hydration on or below the ocean floor. The development of high-pressure metamorphic mineralogies in metagabbros occurred preferentially in zones that underwent sub-ocean-floor alteration due to the development of hydrated, fine-grained, reactive assemblages by that early alteration. The petrology, major element variations and stable isotope geochemistry of the Pennine rocks thus largely reflects processes that predated high-pressure metamorphism. Oxygen isotope profile The overall d 18O profile reconstructed through the dismembered Pennine ophiolite sequence resembles that of the Macquarie Island ophiolite (Cocker et al. 1982). The transition between 18O-enriched and 18Odepleted rocks in that ophiolite occurs at the transition between the sheeted dykes and the basalts. By contrast, other ophiolites (e.g. Semail) show elevation of d 18O values in all rocks above the gabbro-sheeted dyke contact (Gregory and Taylor 1981). The oxygen isotope profiles through ophiolites reflect the sub-ocean-floor hydrothermal fluid flow systems, and often correlates with the setting in which the ophiolite formed. The Macquarie Island rocks were probably formed on a slow-spreading ridge with small magma chambers (Cocker et al. 1982), and the isotopic resetting may reflect the effects of two separate fluid flow systems (Cocker et al. 1982; Muehlenbachs 1996): (a) a shallow low-temperature system involving large fluid volumes; and (b) a deeper, higher-temperature, fluid flow system involving lower fluid volumes. If the stable isotope geochemistry of the Pennine ophiolites have been little altered during metamorphism, the tectonic setting and hydrothermal circulation system may have been similar to that which prevailed at Macquarie Island.
Implications for fluid flow during high-pressure metamorphism The subduction of oceanic lithosphere potentially transports large volumes of water initially bound up in hydrous minerals formed on the sea floor into the upper mantle. Given the substantial amount of dehydration that must have accompanied the transformation of the subducted rocks (especially the higher-level basalts that are commonly extensively hydrated) to eclogite facies assemblages (Philippot 1993; Peacock 1993a, 1993b, 1997), the lack of homogenisation of oxygen isotopes in particular suggests that fluid flow was not pervasive. For any given length scale, oxygen isotopes are generally reset by much smaller time-inte-
grated fluid fluxes (at least one to two orders of magnitude) than those required to reset the major element geochemistry; hence, even moderate amounts of pervasive fluid flow would be expected to reset d 18O values, even if rock chemistry remained unaltered. The rocks that show partial greenschist facies retrogression do have more scattered stable isotope values, which may reflect some resetting during later fluid flow. In order not to interact with large volumes of rock, the fluids generated during high-pressure metamorphism must have been strongly channelled. The highpressure veins represent one possible channel; however, the stable isotope and major element data presented in this study suggest that they were locally derived, or were derived from small volumes of fluid that were able to equilibrate isotopically with the rocks adjacent to the veins. Similar conclusions were also reached for high-pressure veins elsewhere in the Alps on the basis of stable isotope and fluid inclusion studies (Nadeau et al. 1993; Philippot and Selverstone 1991). The channels along which the fluids escaped have not yet been identified. One strong possibility is the tectonic contacts between the units. Within the ophiolite, pillow lavas are elongated to F5 : 1 aspect ratios (presumably from near-spherical forms); however, they are still recognisable, as are the sheeted dykes. The metagabbros appear to be little deformed, retaining their internal igneous textures. However, the ophiolite as a whole has been dismembered, and some of that dismembering may be due to deformation at high pressures. Shear zones or faults developed between units would have been zones of relatively high permeability (cf. McCaig 1997) which could have acted as fluid channels. If fluid flow in subduction zones is generally strongly channelled, the overall direction of fluid flow will depend on the orientation of the channels. If the channels are oriented subparallel to the subduction zone, the fluids may have been channelled towards the accretionary prism. The melange zone rocks from the Catalina schist that have interacted with large volumes of fluid (Bebout 1991, 1997) may represent suitable analogues for rocks within these channels at higher crustal levels. The slab–mantle interface may itself be a major shear zone (Philippot 1993), in which case fluids intersecting it are also likely to be channelled along the top of the slab. Such strong fluid channelling may preclude all but a few of the fluids entering the mantle wedge to flux melting; however, relatively small volumes of fluid (compared with that generated by the dehydration of subducted material) may be sufficient to account for the observed island-arc magmatism (Peacock 1990, 1993a, 1993b, 1997). Thus, the conclusions of this study accord with predictions derived from the chemistry of ocean island magmatic rocks. One consequence of the fluid flow pattern as inferred from this study is that the fluids are probably interacting with relatively small volumes of rock (in the channels). It is therefore not possible for those fluids to “scavenge”
234
elements as would be the case if the fluids were able to flow through the crustal pile as a whole. The flow of fluids in such channels will also facilitate deformation within the subducting slab, probably aiding its imbrication. This study has shown the importance of examining fluid flow during high-pressure metamorphism in order to constrain processes in subduction zones. Examination of ophiolites in other high-pressure terrains will potentially better constrain the volatile input to the mantle wedge during subduction. Acknowledgements This research was funded by Australian Research Grant A39531124 (to I.C.). We thank J. Miller for help in the field, M. Jane and M. Yanni (stable isotope analyses), and S. Reeves (XRF). Helpful comments from M. Engi and an anonymous reviewer helped improve the manuscript.
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