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Journal of Petrology Volume 42 Number 6 Pages 1119-1140 2001
© Oxford University Press 2001
New Constraints on the P–T Evolution of the Alpe Arami Garnet Peridotite Body (Central Alps, Switzerland)
MINERALOGISCHES INSTITUT, UNIVERSITÄT HEIDELBERG, IM NEUENHEIMER FELD 236, D-69120 HEIDELBERG, GERMANY
Received April 6, 2000; Revised typescript accepted September 30, 2000
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESTEXTURE AND MINERAL COMPOSITIONSMETAMORPHIC EVOLUTION AND...DISCUSSION AND CONCLUSIONSREFERENCES |
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The metamorphic evolution of the garnet peridotite body of Alpe Arami, Central Alps, is a matter of current controversy. In this paper, the inter- and intragrain distribution of major and trace elements obtained by electron and ion probe microanalyses is used to better constrain the P–T evolution of this peridotite. Using the compositions of homogeneous porphyroclast cores, peak metamorphic conditions of 1180 ± 40°C and 5·9 ± 0·3 GPa are estimated, based on consistent results from the application of several independent thermometers (Fe–Mg exchange between garnet, pyroxenes and olivine, Ni exchange between garnet and olivine, Co and Ni exchange between orthopyroxene and clinopyroxene), the Al-in-orthopyroxene barometer and the Ca–Cr systematics of garnet. Orthopyroxene and clinopyroxene porphyroclasts are, however, not in equilibrium with respect to some elements with low diffusivities, such as Ca, Ti, Cr, V and Sc. This disequilibrium appears to be the main cause for the lower P–T values suggested by some of the previous workers. On the other hand, there is no evidence for an ultradeep (>200 km) origin of the Alpe Arami body as postulated recently. Chemical zonation profiles across mineral grains suggest that during retrograde evolution a near-isothermal decompression was followed by accelerated cooling.
KEY WORDS: Alpe Arami; Central Alps; garnet peridotite; ultrahigh-pressure metamorphism; geothermobarometry; secondary ion mass spectrometry (SIMS)
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESTEXTURE AND MINERAL COMPOSITIONSMETAMORPHIC EVOLUTION AND...DISCUSSION AND CONCLUSIONSREFERENCES |
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The cores of collisional belts often contain (ultra)high-pressure peridotites and eclogites (e.g. Medaris & Carswell, 1990
; Wang et al., 1995). The tectonometamorphic evolution of these rocks as deduced from their mineral assemblages and textures can potentially provide important constraints on orogenic processes (e.g. Van der Wal & Vissers, 1993; Zhang et al., 1994; Altherr & Kalt, 1996; Van Roermund & Drury, 1999). One of the most intriguing examples of a subduction-related ultrahigh-pressure peridotite–eclogite association is exposed in the Central Alps at Alpe Arami (AA), Ticino, Switzerland (Grubenmann, 1908; O’Hara & Mercy, 1966; Möckel, 1969; Rost et al., 1974; Ernst, 1977, 1978, 1981; Evans & Trommsdorff, 1978; Trommsdorff, 1990; Pfiffner & Trommsdorff, 1998).
Whereas the age of the ultrahigh-pressure metamorphism that affected the AA rocks is now reasonably well constrained at 43–35 Ma (Sm–Nd on garnet–clinopyroxene–whole rock: Becker, 1993; SHRIMP U–Pb on zircon: Gebauer, 1996, 1999), the P–T evolution of the AA garnet peridotite has been the subject of a continuing controversy (Ernst, 1978, 1981; Evans & Trommsdorff, 1978; Becker, 1993; Dobrzhinetskaya et al., 1996, 1999; Brenker & Brey, 1997; Green et al., 1997a, 1997b; Risold et al., 1997; Ulmer & Trommsdorff, 1997; Pfiffner & Trommsdorff, 1998; Bozhilov et al., 1999; Nimis et al., 1999; Paquin et al., 1999a, 1999b; Trommsdorff et al., 2000; Nimis & Trommsdorff, 2001). Early estimates of the maximum P–T conditions experienced by the AA peridotite were 830–950°C and 2·5–4·2 GPa (Ernst, 1978, 1981; Evans & Trommsdorff, 1978; Becker, 1993). Brenker & Brey (1997) recognized for the first time the existence of chemical disequilibrium between the phases of the AA peridotite. By applying a combination of the Fe–Mg garnet–olivine exchange thermometer (O’Neill & Wood, 1979, 1980) and the Al-in-orthopyroxene barometer (Brey & Köhler, 1990) on mineral core compositions, they obtained values of 1120 ± 50°C and 5·0 ± 0·2 GPa for the peak metamorphic conditions. These values are in accord with Ca–Cr systematics of garnet (Brenker & Brey, 1997). Similarly high P–T values of 
1095°C and 5·2 GPa were derived by Medaris (1999). However, thermometers based on the enstatite–diopside solvus (Brey & Köhler, 1990) and the Fe–Mg exchange between clinopyroxene and garnet (Krogh, 1988) yielded apparent temperatures 300°C lower. A combination of these thermometers with the Al-in-opx barometer resulted in apparent P–T values of 800–900°C and 3·0–3·5 GPa (Brenker & Brey, 1997). Mineral rim compositions yielded still lower values of 720 ± 50°C and 2·0 ± 0·3 GPa. Ignoring the chemical disequilibrium among the mineral phases found by Brenker & Brey (1997), Nimis et al. (1999) and Nimis & Trommsdorff (2001) proposed peak metamorphic conditions of 840°C and 3·2 GPa for the AA garnet peridotite, based on the following geothermobarometers: (1) Al between opx and grt (Taylor, 1998); (2) Cr between cpx and grt (Taylor & Nimis, 1998); (3) Fe–Mg between cpx and grt (Ai, 1994); (4) Ca–Mg between cpx and opx (Taylor, 1998).
Textural relationships between exsolved ilmenite and host olivine as deduced from observations by transmission electron microscopy led Dobrzhinetskaya et al. (1996) and Green et al. (1997a, 1997b) to propose that the ilmenite represents former FeTiO3 perovskite exsolved from wadsleyite. As a consequence, an extreme depth of origin (>300 km) of the AA garnet lherzolite body was inferred. Risold et al. (1997), however, showed that ilmenite inclusions in olivine are typical for all garnet peridotites from the principal three locations in the Cima Lunga unit of the Central Alps (Fig. 1) and the prograde character of garnet at least for Cima di Gagnone has been proven as it overgrows pre-existing folds with amphibole and spinel with an upper pressure limit of 3 GPa (Niida & Green, 1999). Furthermore, Ulmer & Trommsdorff (1997) presented data on the solubility of Ti in olivine and argued for the genesis of the FeTiO3 rods in olivine during decompression and cooling. Nevertheless, on the basis of one experiment, Dobrzhinetskaya et al. (1999) claimed that the solubility of Ti in olivine coexisting with rutile at 1400 K and 5 GPa is very low (<0·2 wt % TiO2). On the basis of their experimental solubility data for higher P–T conditions that are in conflict with those of Ulmer & Trommsdorff (1997), they suggested again that the AA peridotite body was exhumed from a depth corresponding to a pressure of 10 GPa.
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Bozhilov et al. (1999) reported on exsolved Ca-poor clinopyroxene lamellae in diopside grains of the AA peridotite that were first described by Yamaguchi et al. (1978). The orientation, crystallography and microstructures of these lamellae in conjunction with the phase relationships in the Mg0·9Fe0·1SiO3 system led these workers to suggest that the AA peridotite was exhumed from a minimum depth of 250 km. Arlt et al. (2000), however, have shown that relative to the phase boundaries for Mg0·9Fe0·1SiO3, the high-pressure (HP) C2/c–Pbca phase boundary is shifted 1 GPa towards lower pressure and the high-temperature (HT) C2/c–P21/c phase boundary is shifted 700°C towards lower temperature if the reported composition of the clinoenstatite lamellae in AA cpx is taken into account. Thus, an origin of the exsolved clinopyroxene lamellae by inversion from the HT C2/c polymorph during near-isothermal decompression cannot be ruled out.
In this paper, we describe a comprehensive set of new microanalytical data on the AA lherzolite. We use the inter- and intragrain distribution of major and trace elements obtained by electron probe microanalysis (EPMA) and secondary ion mass spectrometry (SIMS) to better constrain the P–T evolution of these rocks. By applying a number of widely tested and most up-to-date thermobarometers based on major and trace elements we will show that complete recrystallization and equilibration of the AA peridotite occurred at about 1180°C and 5·9 GPa. Our new data do not provide any evidence for an exhumation from greater depths (>200 km) as postulated by Dobrzhinetskaya et al. (1996, 1999) and Bozhilov et al. (1999).
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GEOLOGICAL SETTING |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESTEXTURE AND MINERAL COMPOSITIONSMETAMORPHIC EVOLUTION AND...DISCUSSION AND CONCLUSIONSREFERENCES |
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The Alps are the result of the continuing convergence between the European and the Adriatic plates. Recent geodynamic models of the Alps (e.g. Schmid et al., 1996
; Channell & Kozur, 1997; Gebauer, 1999) include the creation and consumption of three oceanic domains, namely the Meliata–Hallstatt, Piemontese–Ligurian (South Penninic) and Valais (North Penninic) oceans. The Penninic nappe stack of the Central Alps is interpreted to represent the transition from the southern European continental margin to the Valais and Piemontese–Ligurian oceans (e.g. Schmid et al., 1996). The AA peridotite–eclogite body is located near the Insubric Line and forms part of the Cima Lunga nappe, which correlates tectonically with the Adula nappe (Fig. 1). Both nappes are thought to represent the southernmost part of the European margin at a time before the closure of the Valais ocean in the Eocene period (Schmid et al., 1996). The lithologies of both nappes are characterized by continental basement rocks, metamorphosed Mesozoic sediments and numerous basic and ultrabasic lenses, most of which contain high-pressure mineral assemblages (e.g. Trommsdorff, 1990; Meyre & Puschnig, 1993; Grond et al., 1995; Meyre et al., 1997; Pfiffner & Trommsdorff, 1998). For some of the high-pressure rocks from the Cima Lunga unit, an origin from ocean-floor metamorphic protoliths (serpentinites, rodingites) has been demonstrated (e.g. Evans & Trommsdorff, 1978; Evans et al., 1979, 1981; Trommsdorff, 1990). As in many other high-pressure terranes world-wide, only very few relics of high-pressure metamorphism have been reported within the felsic country rocks (Adula nappe: Meyre et al., 1999). Such relics are as yet unknown from the Cima Lunga nappe (Grond et al., 1995). At Cima di Gagnone, however, the eclogite–peridotite suite is accompanied by high-pressure metamorphosed ophicarbonate rocks (Pfiffner & Trommsdorff, 1998). It has repeatedly been emphasized that the P–T conditions of the high-pressure metamorphism in the Adula–Cima Lunga nappe system increase from 450–550°C and 1·0–1·3 GPa in the north to 750–900°C and 1·8–3·5 GPa in the south (Heinrich, 1986; Trommsdorff, 1990; Meyre & Puschnig, 1993; Pfiffner & Trommsdorff, 1998; Trommsdorff et al., 2000). Among the ultramafic lenses, several occurrences of amphibole–garnet peridotites have been found (e.g. Cima di Gagnone, Mt. Duria; Fig. 1). As amphibole occurs as a prograde phase, partly included in garnet, there is general consensus that maximum pressures did not exceed 3·3 GPa, i.e. the upper-pressure limit of pargasitic amphibole in lherzolite (e.g. Niida & Green, 1999). Recent P-T estimates for the peak of metamorphism are 740°C and 3·2 GPa for Cima di Gagnone and 830°C and 3·0 GPa for Mt. Duria (Nimis et al., 1999; Nimis & Trommsdorff, 2001). However, prograde amphibole has yet to be found in the garnet peridotite at AA. Therefore, maximum P–T conditions experienced by this body may potentially be significantly higher than those deduced for the other two occurrences. |
ANALYTICAL TECHNIQUES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESTEXTURE AND MINERAL COMPOSITIONSMETAMORPHIC EVOLUTION AND...DISCUSSION AND CONCLUSIONSREFERENCES |
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Major element abundances were determined using a CAMECA SX51 electron microprobe equipped with five wavelength-dispersive spectrometers. Operating conditions were 15 kV accelerating voltage, a beam current of 20 nA and 10 s counting time for each element (except for Ti, Mn and Cr in grt, opx and cpx: 100 s; and Ni in ol: 30 s). A PAP matrix correction was applied to the raw data. Natural and synthetic silicates and oxides were used for calibration. In addition to numerous individual point analyses, many traverses through mineral grains were performed.
Sc, V, Co, Ni and in part also Ti were measured by SIMS with a modified CAMECA IMS-3f ion microprobe equipped with a primary beam mass filter. For the transition elements we employed the high mass resolution technique with a mass resolution M/
M of 8000 at 10% to discriminate element peaks from interfering molecular ions. Further details have been given by Seitz et al. (1999). Stepscan ion microprobe analyses were performed along the same profiles as investigated by electron microprobe.
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TEXTURE AND MINERAL COMPOSITIONS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESTEXTURE AND MINERAL COMPOSITIONSMETAMORPHIC EVOLUTION AND...DISCUSSION AND CONCLUSIONSREFERENCES |
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The AA body consists of a small (
1 km in length) lens of peridotite locally bordered by kyanite-bearing eclogite (Fig. 1), both showing variable degrees of low-pressure hydrous alteration (Möckel, 1969). Fresh garnet peridotite from the core of the AA body displays a porphyroclastic texture with porphyroclasts of garnet (grt I), orthopyroxene (opx Ib) and rare clinopyroxene (cpx Ib) and olivine (ol Ib) within a matrix of ol II, opx II and cpx II. Although we investigated about 30 thin sections from several samples, we never found opx Ib and cpx Ib in contact with each other, but they may both occur in contact with garnet. The large (up to 0·5 cm) grt I grains may have thin (<170 µm) kelyphitic coronas (amphibole + spinel ± opx/chl) and contain scarce inclusions of olivine (ol Ia), orthopyroxene (opx Ia) and clinopyroxene (cpx Ia). There is a textural transition from opx Ib porphyroclasts to opx II matrix grains. As recognized earlier (Yamaguchi et al., 1978; Green et al., 1997a, 1997b), there is a marked textural contrast between cpx Ib and cpx II. Whereas cpx Ib forms large (up to 4 mm) porphyroclasts, cpx II grains of the matrix are much smaller (<0·8 mm) and show interlobate and amoeboid shapes indicating textural disequilibrium. Both cpx Ia and Ib contain oriented exsolution lamellae of clinoamphibole (am) that were described in detail by Yamaguchi et al. (1978). Sometimes, these secondary amphiboles occur at the margins of larger cpx Ia inclusions in garnet. Furthermore, lamellae of low-T clinoenstatite (space group P21/c) were found in cpx Ia inclusions in garnet and cpx Ib grains adjacent to garnet (Yamaguchi et al., 1978; Bozhilov et al., 1999). On the basis of TEM observations these lamellae were interpreted to have exsolved as former high-pressure clinoenstatite (space group C2/c), implying exhumation from depths of >250 km (Bozhilov et al., 1999). Cpx Ia and Ib also contain small (
5 µm) inclusions of ilmenite and exsolution lamellae of chromite (Bozhilov et al., 1999), which were never found in cpx II. Larger olivine Ib porphyroclasts contain variable amounts of rod-shaped precipitates of ilmenite parallel to [010] of olivine and tabular chromite precipitates parallel to (010) (Dobrzhinetskaya et al., 1996; Risold et al., 1997). Three samples (AA-3P1, AA-R9, AA-7a) taken from the core of the AA garnet peridotite body were studied in detail for the compositions of their minerals. Only sample AA-3P1 contains rare cpx Ib porphyroclasts. Representative mineral analyses are given in Table 1 and overall chemical variations in the various minerals are provided in Table 2. With the exception of very minor compositional differences between the garnets (see below), all three samples are characterized by similar mineral compositions.
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Typical chemical zoning profiles across mineral grains are presented in Figs 2–5. Garnet I shows significant chemical zoning (Fig. 2 and Table 2) with an increase in Fe and Mn, and decreases in Mg, Ca, Ti, Sc and Ni towards the rims of the grains. In the same direction, mg-number [= 100 x Mg/(Mg + Fe)] decreases from 82 to 77. Except for minor differences in the abundances of Cr, Ca, Ti, V and Sc, all grt I grains display similar core compositions (Fig. 2) and incorporate detectable amounts of Na (0·04 wt %; Table 1). Most garnet grains display W-shaped Cr and V zoning profiles [e.g. grt I (1) in Fig. 2], but other profile shapes were also found [e.g. grt I (3) in Fig. 2]. Garnet compositions adjacent to cracks are often similar to those of the garnet rims [e.g. Ni, Cr, V in grt I (3); Fig. 2].
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Compositional profiles across large orthopyroxene Ib porphyroclasts reveal homogeneous cores that are often surrounded by heterogeneous outer zones whereby the relative widths of both depend on the grain size and on sectioning effects (Fig. 3 and Table 2). In the thin (<150 µm) outer zones, a progressive increase in Al, Cr, Ti, V, and sometimes also in Ca, away from the cores of the grains, is abruptly followed by a strong depletion in these elements. From the core to the rim of the grains, mg-number decreases gradually from about 91·3 to 90·3, Ni and Co also decrease, but Sc increases slightly (Fig. 3). Adjacent to cracks, opx Ib grains have compositions that are in most cases similar to those of the rims (e.g. opx Ib from sample AA-3P1 in Fig. 3). The Ca contents in opx Ib cores are somewhat variable and in all cases low.
Except for higher values of mg-number (92·9–92·2 vs 91·6–90·3), opx Ia inclusions in garnet have compositions that are similar to those of the cores of opx Ib porphyroclasts (Figs 3 and 4).
Typical opx II grains, with diameters of <700 µm, are characterized by increasing abundances of Al, Fe and Sc, and decreasing abundances of Ca, Mg, Cr, Ti, Ni, Co and V from core to rim (Fig. 3). Likewise, mg-number decreases slightly from about 91·5 to 90·2. In the cores of opx II grains, the abundances of elements with low diffusivities, such as Al, Cr, Ti and Sc, are intermediate between those observed in the cores and intermediate zones of opx Ib grains. A notable exception is V, which is more abundant in opx II (Fig. 3).
Cpx Ib porphyroclasts show significant chemical zoning whereby Al, Ti and Na decrease and Sc increases slightly towards the rims of the grains (Fig. 5). Ca contents of the cores are somewhat variable and relatively high (Table 1). The compositions of cpx Ia inclusions in grt I are variable and depend on grain size. Large (>1 mm) cpx Ia grains are compositionally similar to cpx Ib porphyroclasts (Fig. 5), but smaller grains have higher mg-number. Compared with cpx Ib, cpx II grains are characterized by lower abundances of Al, Fe, Na, Ti and V, and higher abundances of Ca, Mg and Sc (Table 1, Fig. 5). All cpx II grains display significant chemical zonation with Al, Na, Ti, Fe and V decreasing, but Ca, Mg and Sc increasing towards the rim (Fig. 5).
Regardless of their textural position and size, olivine grains are nearly homogeneous. There are, however, significant differences among the different types (Tables 1 and 2). Whereas ol Ib and ol II have similar abundances of Mg and Fe with an average mg-number of 90·3 ± 0·7 (1
), ol Ia is more Mg rich (mg-number = 92·0–92·6). The abundances of Ni in the three olivine types are similar and the average value is 2901 ± 251 ppm (1). Brenker & Brey (1997) reported low Ca contents (30–40 ppm) for the cores of large olivine Ib grains and a narrow zone of enhanced Ca concentration (up to 120 ppm) at the rims. Their interpretation of the increasing Ca content was that it was probably caused by a late-stage heating and/or decompression event.
The clinoamphibole occurring in cpx Ia and Ib as exsolution lamellae and as small grains at the margins of cpx Ia shows only moderate chemical variation with an average composition of (Na0·75K0·02)(Na0·10Ca1·75Fe2+0·15)(Al0·48Ti0·06Cr0·16Mg4·00Mn0·01Fe2+0·29)[Al1·35Si6·65O22(OH)2]. This composition is similar to that reported by Yamaguchi et al. (1978) and corresponds to edenite using the classification of Leake et al. (1997). Clinoamphibole from the thin kelyphite corona around grt I grains is even more Al rich and corresponds to pargasite: (Na0·82K0·01)(Ca1·88Fe2+0·12)(Al0·82Ti0·02Cr0·22Mg3·58Mn0·01Fe2+0·28Fe3+0·07)[Al2·01Si5·99O22(OH)2]. Spinel from the kelyphite has an average composition of Fe2+0·22Mg0·78Fe3+0·02Cr0·13Al1·85O4.
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METAMORPHIC EVOLUTION AND THERMOBAROMETRY |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESTEXTURE AND MINERAL COMPOSITIONSMETAMORPHIC EVOLUTION AND...DISCUSSION AND CONCLUSIONSREFERENCES |
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Evaluation of zonation profiles
Chemical zonation in mineral grains of a rock is caused by a change in P, T and effective bulk composition (X) and by diffusion during and after growth. The extent of chemical heterogeneity within a particular grain is controlled by the range of partition coefficients of the different elements as a function of P, T and X, the rates of change in these variables and the diffusion rates. In any case, an abrupt change in the slope of a zonation profile, as observed for some elements in mineral grains of the AA peridotite (Figs 2 and 3), signifies a complex evolution not fully recognized by previous workers (e.g. Brenker & Brey, 1997
; Nimis et al., 1999; Nimis & Trommsdorff, 2001). By using experimental and empirical evidence of P- and T-dependent element partitioning among minerals in peridotite systems and of the relative diffusivities of elements we will now evaluate our observed zonation profiles.
Garnet
On the basis of the relative lengths of zoning profiles for various elements in pyrope-rich mantle garnet it has been suggested that the diffusion coefficients, over the temperature range 1200–1400°C, decrease in the order Ni 
Mg, Fe, Mn > Ti > Ca > Cr, V (Griffin et al., 1996). Brenker & Brey (1997) and Griffin et al. (1999) have pointed out that Cr in peridotitic grt is a primary indicator for the Cr/Al ratio of the bulk rock. Given the relative immobility of Cr in garnet (e.g. Griffin et al., 1996, and references therein), we suggest that the variable Cr profiles observed in our AA garnets (Fig. 2) reflect variable effective bulk compositions during garnet growth. Thus, the slightly higher Cr content of grt I from sample AA-7a relative to garnet I in sample AA-3P1 (Fig. 2) reflects a more depleted nature of this sample. This hypothesis is corroborated by lower abundances of Ti and higher abundances of Sc in grt from sample AA-7a. At the moment, we have no convincing explanation for the W-shaped patterns of the Cr profiles that are most commonly observed in garnets from the chemically more fertile samples (e.g. AA-3P1). It is worth noting that V appears to mimic the behaviour of Cr.
All chemical zonation profiles across grt I grains suggest outward diffusion of Mg and inward diffusion of Fe and Mn (Fig. 2) as a result of cooling after initial equilibration (e.g. Medaris & Wang, 1986; Brey et al., 1990). This interpretation is supported by the observed decrease in Ni towards the rims of all AA garnets investigated (Fig. 2). Whereas the range of Ni content in olivine from mantle peridotites is small and independent of T, the abundances of Ni in coexisting garnet vary from 10 ppm at 600°C to 120 ppm at 1400°C (Griffin et al., 1989; Canil, 1994, 1999; Ryan et al., 1996). A comparison of the profiles for Ni and Mg in AA garnets supports the view of Griffin et al. (1996) that, at elevated temperatures, the diffusivity of Ni is equal to or slightly higher than that of Mg.
In summary, the observed zonation profiles of garnet I grains may be interpreted as follows. The variations in Cr, V, Ti and Sc originate from variable effective bulk compositions during garnet growth. At some P and T, all garnets were equilibrated with respect to their Mg, Fe, Mn, Ni and Ca contents. This stage is still reflected in the constant garnet core compositions. Subsequent cooling caused outward diffusion of Mg and Ni and inward diffusion of Fe and Mn.
Olivine
The chemical homogeneity and similarity of olivine Ib and II grains may be explained by the fact that Fe–Mg diffusion in olivine is about two orders of magnitude faster than that in pyroxenes and in garnet (Ganguly & Tazzoli, 1994; Brenker & Brey, 1997). However, as convincingly argued by Brenker & Brey (1997), the partitioning of Fe and Mg between olivine and the cores of garnet grains in peridotitic rocks may still reflect peak metamorphic conditions because (1) olivine is the most abundant mineral and forms a large reservoir, implying that any change in the Fe/Mg distribution coefficient between olivine and garnet will be reflected only in the composition of garnet, whereas olivine remains almost constant, and (2) the Fe/Mg ratio in the interior parts of large garnet grains is frozen in at high temperatures because Fe–Mg diffusion in garnet is slow compared with a much faster diffusion in olivine (Smith & Wilson, 1985; Chakraborty & Ganguly, 1991; references given by Griffin et al., 1996). The Fe and Mg zonation profiles across grt I and the homogeneity of ol Ib grains from the AA peridotite are in line with this argumentation. Furthermore, the higher values of mg-number observed in ol Ia inclusions in garnet relative to those in ol Ib and ol II (Table 1) clearly demonstrate the dependence of olivine compositions on the volume of the exchange reservoirs. Similar arguments may be made for the Fe–Mg exchange between grt and opx (Figs 2–4).
Pyroxenes
To evaluate the chemical zonation profiles of the AA pyroxenes, it is useful to compare their Ca contents with those of opx–cpx pairs equilibrated experimentally at various P–T conditions. At constant pressure, the partitioning of Ca between opx and cpx equilibrated at a range of temperatures should follow a near-linear trend whereby Ca in opx decreases and Ca in cpx increases with falling temperature (Fig. 6). With increasing pressure, the trend lines are shifted towards lower Ca contents (Brey & Köhler, 1990; Brey et al., 1990; Brey, 1991). The Ca contents of both pyroxene Ib and II pairs are not in equilibrium. In particular, the cores of opx Ib grains have Ca contents that are too low to be in equilibrium with cpx at any P. The exsolution of abundant clinoamphibole lamellae probably did not influence the compositions of cpx Ia and Ib grains because careful electron probe scans across clinoamphibole lamellae clearly demonstrate that an increase in Ca and Si and a decrease in Al in the host cpx is restricted to <10 µm from the exsolved clinoamphibole lamellae (Yamaguchi et al., 1978: Fig. 2 and Table 1). Furthermore, as observed zonation profiles for Ca in opx Ib grains (Fig. 3) do not suggest an outward diffusion of Ca (except for the outermost zones; Fig. 3), the Ca contents in the cores of both pyroxene Ib porphyroclasts are thought to result from differences in effective bulk compositions rather than from P–T controlled changes along the enstatite–diopside solvus. If so, the abundances of other elements in the cores of both pyroxenes could still reflect equilibrium conditions. In particular, the diffusivities of Mg and Fe in both types of pyroxenes are higher than those of Ca and Al (Brady & McCallister, 1983; Sautter & Harte, 1990; Smith & Barron, 1991; Ganguly & Tazzoli, 1994; Brenker & Brey, 1997; Dimanov & Sautter, 2000). Therefore, the abundances of Mg, Fe, Ni and Co in the cores of both pyroxenes could represent equilibration with the large olivine Ib–II reservoir and with the cores of grt I porphyroclasts. Furthermore, the fact that Al contents in the cores of pyroxene Ib porphyroclasts are nearly constant and similar to those of pyroxene Ia inclusions in garnet (Table 1, Figs 3–5) suggests equilibration with garnet at some P and T. As the Al content in opx coexisting with grt increases with rising T and decreasing P (e.g. Wood & Banno, 1973; MacGregor, 1974; Harley, 1984; Brey & Köhler, 1990; Brey et al., 1990), the zonation present in the outer zones of some opx Ib grains (Fig. 3) could be due to a near-isothermal decompression or a decompression accompanied by heating (Al increasing) followed by accelerated cooling (Al decreasing). In contrast, Al zonation patterns of cpx Ib porphyroclasts are characterized by continuously decreasing Al concentrations from core to rim (Fig. 5) suggesting a somewhat different P–T evolution and supporting our interpretation that both pyroxenes Ib were not in chemical equilibrium with respect to Ca and Al.
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Peak metamorphic conditions
On the basis of the considerations presented above, we will use different thermobarometers to constrain the P–T evolution of the AA peridotite body: (1) Al-in-opx coexisting with grt (Brey & Köhler, 1990); (2) Fe–Mg exchange between grt and ol (O’Neill & Wood, 1979, 1980); (3) Fe–Mg exchange between grt and opx (Brey & Köhler, 1990); (4) Fe–Mg exchange between grt and cpx (Krogh Ravna, 2000); (5) Ni partitioning between grt and ol (Canil, 1994, 1999; Ryan et al., 1996); (6) Ca–Cr systematics in grt coexisting with cpx (Brenker & Brey, 1997); (7) the partitioning of transition elements (Sc, Ti, V, Cr, Mn, Co, Ni) between opx and cpx (Seitz et al., 1999). For illustrative purposes, we will also calculate apparent equilibrium temperatures based on (8) the enstatite–diopside solvus (Brey & Köhler, 1990) and (9) the amount of Ca in opx coexisting with cpx (Brey & Köhler, 1990). Resulting temperature and pressure estimates are summarized in Table 3 and graphically presented in Fig. 7.
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Fe–Mg exchange thermometry between grt, ol, opx and cpx is generally compromised by uncertainties in the Fe3+/Fetot ratios of both natural phases and the phases from which the thermobarometers were calibrated in the original experiments (e.g. Canil & O’Neill, 1996). For natural garnet peridotites, it has been shown that (1) the Fe3+/Fetot ratio increases in the order ol (0·00), grt (0·03–0·12), opx (0·03–0·10), cpx (0·2–0·4), (2) the Fe3+/Fetot ratio in grt increases with both T and P, whereas that in cpx remains approximately constant, and (3) Fe3+opx 
0·55 Fe3+cpx (Canil & O’Neill, 1996). It is now widely accepted that the experiments used to calibrate the Fe–Mg exchange thermometers produced grt, opx and cpx with substantial Fe3+ contents, similar to those of the natural samples (e.g. Brey & Köhler, 1990; Canil & O’Neill, 1996; Smith, 1999, and references therein). Fe3+/Fetot values for minerals of the AA peridotite as determined by Mössbauer spectrometry are 0·034–0·036 for garnet (Fett, 1989; Nimis & Trommsdorff, 2001) and 0·14 for cpx (Nimis & Trommsdorff, 2001). Using the relationship Fe3+opx 0·55 Fe3+cpx (Canil & O’Neill, 1996) a value of 0·032 for opx can be calculated.
Peak metamorphic conditions may be derived by applying the Al-in-opx barometer (Brey & Köhler, 1990) in combination with thermometers based on the Fe–Mg exchange between grt and ol (O’Neill & Wood, 1979, 1980) and between grt and opx (Brey & Köhler, 1990) to the cores of grt I, ol Ib and opx Ib grains. In doing so, we will first ignore the influence of Fe3+ and calculate temperatures with all Fe treated as Fe2+. This yields consistent results with average temperatures of 1180 ± 40°C and pressures of 5·9 ± 0·3 GPa (Table 3). If Fe3+ concentrations are considered, grt–ol and grt–opx temperatures rise by 40°C and 30°C, respectively. These P–T conditions are slightly higher than, but in general agreement with those presented by Brenker & Brey (1997). The use of opx Ia inclusions in grt for barometry results in slightly lower pressure estimates of 5·2 GPa. This is probably due to the fact that the small opx Ia grains have lost Fe to grt during cooling (compare mg-number of opx Ib and Ia; Tables 1 and 2), resulting in an increase in KD for grt–opx (lower T), which, when combined with the Al-in-opx barometer, yields a lower pressure.
Although the partitioning of Ca between opx and cpx does not represent equilibrium (Fig. 6), the partitioning of Mg and Fe between grt and cpx (Krogh Ravna, 2000) may give some information about the thermal history of the AA peridotite. In contrast to grt, opx and ol, cpx in peridotitic and pyroxenitic systems is generally characterized by relatively high Fe3+/Fetot ratios (Canil & O’Neill, 1996). However, the calibrations of Krogh (1988) and Krogh Ravna (2000) do not consider possible Fe3+ contents and Brey & Köhler (1990) and Brey et al. (1990) demonstrated that, in peridotitic systems, the thermometer of Krogh (1988) reproduced experimental results within 70°C (in most cases better than 50°C) even when it is assumed that Fetot equals Fe2+. Applying the more recent calibration of Krogh Ravna (2000) to the compositions of grt I and cpx Ib cores yields temperatures of 1188 ± 34°C and 1133 ± 29°C for Fetot = Fe2+ (Table 3). Considering Fe3+ results in slightly lower temperature values of 1120°C and 1070°C, respectively.
Further important constraints on the reliability of the calculated equilibration temperatures come from the partitioning of Ni between garnet and olivine. Garnet I grains show a marked zoning with decreasing Ni contents towards the rim indicating cooling (Fig. 2). Garnet cores still show narrow plateaux with constant Ni abundances of 67 ppm (Fig. 2 and Table 3). Applying the Ni-in-grt thermometer of Canil (1999) results in temperatures of 1120 ± 20°C (Table 3). With the thermometer version of Ryan et al. (1996) slightly higher temperatures of 1135 ± 20°C are obtained.
Still another independent test for the peak metamorphic conditions is provided by the Ca–Cr systematics of garnet I grains. As shown by Brenker & Brey (1997) the Cr content in lherzolitic garnet is a function of the effective bulk composition only and does not depend on P and T. For constant Cr in garnet, Ca in garnet decreases with increasing pressure and rising temperature, and at constant P and T, Ca increases linearly with Cr. The Ca–Cr systematics of garnets from the AA peridotite and calculated isolines for various P–T conditions are presented in Fig. 8. Most of the data plot between the isolines for 1100°C and 5·0 GPa, and 1240°C and 6·3 GPa, representing the most extreme conditions found by Fe–Mg and Ni exchange thermometry in combination with Al-in-opx barometry. Even the most Ca-rich garnet composition observed in the AA peridotite would still require minimum conditions of 1030°C and 4·5 GPa. Lower P–T conditions of 840°C and 3·2 GPa as suggested by Nimis et al. (1999) and Nimis & Trommsdorff (2001) are clearly inconsistent with the garnet compositions. To match low temperatures of 840°C with the observed Ca–Cr systematics requires pressures between 5·9 and 9·2 GPa, and, correspondingly, extremely high temperatures between 1205 and 1645°C are needed to allow for a low equilibration pressure of 3·2 GPa. The very high pressures (>10 GPa) postulated by Dobrzhinetskaya et al. (1999) are also not supported by the garnet compositions (Fig. 8). The observed Ca zonation profiles across garnet (Fig. 2) are not indicative of inward diffusion of Ca that would be expected for a decrease in P and/or drop in T.
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In conclusion, the peak metamorphic conditions experienced by the AA peridotite body can safely be constrained to temperatures of 1100–1240°C and pressures of 5–6 GPa. Our preferred estimate for the peak P–T conditions is 1180 ± 40°C and 5·9 ± 0·3 GPa.
The partitioning of transition elements between pyroxene porphyroclasts
Both the zonation profiles of Mg and Fe across pyroxene Ib porphyroclasts and the consistent temperature estimates derived from the Fe–Mg exchange thermometers (grt–ol, grt–opx, grt–cpx) suggest that all porphyroclasts were once mutually equilibrated with respect to Fe and Mg, i.e. elements with a high diffusivity. This hypothesis is corroborated by temperatures calculated from the partitioning of Co and Ni (Seitz et al., 1999) between the cores of opx Ib and cpx Ib. Applying the exchange thermometer for Co we obtained temperatures of 1085 ± 41°C (Table 3), and measured partition coefficients for Ni [ln DNi(opx/cpx) 0·75] are also consistent with high temperatures around 1150°C [compare fig. 3 of Seitz et al. (1999)]. The apparent partition coefficients of Ti, Cr, V and Sc, however, yield inconsistent temperatures (Table 3) that are tentatively explained by disequilibrium at peak conditions as a result of the relatively low diffusivities of these elements, especially in cpx (Dimanov & Sautter, 2000).
Rim compositions of pyroxene porphyroclasts are controlled by exchange with olivine. This is clearly seen by the zonation patterns for Ni and Co in opx Ib (Fig. 3). The retrograde loss of Co and Ni from the outer zones of and along cracks in opx Ib grains is the opposite of what would be expected for opx–cpx exchange equilibria; with falling temperature, the values of DNi(opx/cpx) and DCo(opx/cpx) should increase (Seitz et al., 1999), which is the opposite to that observed for pyroxene Ib grains (compare Figs 3 and 5).
Retrograde evolution
As discussed above (see ‘Evaluation of zonation profiles’) the observed Al zonation in the outer parts of opx Ib grains is compatible with a decompression at near-isothermal conditions followed by rapid cooling. Cooling is also indicated by the zonation profiles for Fe, Mg and Ni across grt I and opx Ib grains (Figs 2 and 3, respectively). Indeed, apparent temperatures obtained for the rim compositions of the porphyroclasts (calculated for 6 GPa, to facilitate comparison with peak temperatures) are significantly lower than the peak equilibration conditions inferred above from the cores of the grains. Minimum temperatures obtained with the Fe–Mg exchange thermometers are 796°C for grt–ol (O’Neill & Wood, 1979, 1980), 818°C for grt–opx (Brey & Köhler, 1990) and 839°C for grt–cpx (Krogh Ravna, 2000). Measured Ni contents at garnet rims are 42 ppm (Fig. 2) resulting in apparent temperatures (Canil, 1999) of 1030°C.
The compositions of ol Ia, opx Ia and cpx Ia inclusions in grt I are variable and depend on grain size. With decreasing size of the inclusion, temperatures calculated with the various Fe–Mg exchange thermometers fall from 1175°C for large cpx Ia grains to 800°C for small ol Ia inclusions (Table 3). These values are in agreement with what is to be expected from diffusion data. The diffusion of Fe and Mg in olivine is about two orders of magnitude faster than in orthopyroxene (Ganguly & Tazzoli, 1994; Griffin et al., 1996).
Apparent temperatures derived from the observed partitioning of transition elements between the cores of opx II and cpx II pairs (Seitz et al., 1999) are similar to those obtained for pyroxene Ib pairs (Table 3). The fact that TNi and TCo are 1110°C indicates that the formation of pyroxene neoblasts occurred at high temperatures shortly after peak metamorphic conditions. This interpretation is supported by temperatures in excess of 1050°C derived from Fe–Mg partitioning between coexisting grt I and cpx II grains (Table 3). Apparent temperatures for the other transition elements indicate disequilibrium between opx II and cpx II, even though they frequently occur in contact. This disequilibrium suggests rapid formation of pyroxene neoblasts followed by rapid decompression and cooling of the AA peridotite body.
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DISCUSSION AND CONCLUSIONS |
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The data of this study indicate that the garnet peridotite body of Alpe Arami experienced P–T conditions that were significantly higher than those of the Cima di Gagnone and Mt. Duria bodies for which peak conditions of about 800°C and 3·3 GPa have been derived (Ernst, 1978
, 1981; Evans & Trommsdorff, 1978; Nimis et al., 1999; Nimis & Trommsdorff, 2001). Our preferred estimate of 1180 ± 40°C and 5·9 ± 0·3 GPa for the metamorphic peak conditions of the AA peridotite is based on consistent results from the application of several independent thermometers (Fe–Mg exchange between garnet, pyroxenes and olivine, Ni exchange between garnet and olivine, Co and Ni exchange between opx and cpx), the Al-in-opx barometer and the Ca–Cr systematics of garnet on chemically homogeneous cores of porphyroclasts. We also demonstrated that the opx Ib and cpx Ib grains (which were not found in contact with each other) are not in equilibrium with respect to the partitioning of Ca, Al, V, Ti, Cr and Sc. By this observation, we can explain the lower P–T values (830–950°C and 2·8–4·0 GPa) suggested by other workers (Ernst, 1978, 1981; Evans & Trommsdorff, 1978; Becker, 1993; Nimis et al., 1999; Nimis & Trommsdorff, 2001). Furthermore, the partitioning of Li among the AA mineral phases and Li zoning patterns in mineral grains indicate a metasomatic overprint of the garnet peridotite in the early part of the exhumation path. Both cpx Ib and cpx II grains are characterized by unusually high abundances of Li, suggesting partial recrystallization (Paquin et al., 1999b; Paquin & Altherr, 2000a, 2000b).
Our data on the composition of matrix pyroxenes and of the rims of porphyroclasts do not suggest a stagnation phase at about 700–800°C and 2–3 GPa (Brenker & Brey, 1997) during exhumation. Instead, we believe that the rim compositions of porphyroclasts (grt I, opx Ib, cpx Ib) and neoblasts (opx II, cpx II) were frozen in at various temperatures during rapid cooling.
The P–T conditions for the AA peridotite obtained in this study are significantly higher than those estimated for the other garnet peridotite bodies of the Cima Lunga unit (see ‘Geological setting’). Such a situation is not uncommon in collisional belts. In the Variscan Vosges and Schwarzwald, for example, ultrahigh-pressure (UHP) garnet peridotites and websterites are closely associated with prograde garnet–spinel and spinel lherzolites (Kalt et al., 1995; Altherr & Kalt, 1996; Kalt & Altherr, 1996). During the formation of collisional belts, HP and UHP mantle rocks may become tectonically introduced into continental materials at various stages (e.g. Brueckner & Medaris, 2000).
What remains to be discussed is the hypothesis of an ultradeep (10 GPa) origin of the AA body that has been repeatedly advanced by Green and co-workers (Dobrzhinetskaya et al., 1996, 1999; Green et al., 1997a, 1997b; Bozhilov et al., 1999). We will not repeat the conclusive arguments that led other workers to question this hypothesis (Ulmer & Trommsdorff, 1997; Risold et al., 1997; Arlt et al., 2000) but we will raise some additional points that may argue against an ultradeep origin of the AA peridotite. First, we have shown that all garnets were equilibrated with their effective bulk compositions at inferred peak conditions (1180°C and 5·9 GPa). At pressures higher than 6 GPa garnets coexisting with pyroxenes start to contain a detectable amount of a majorite component, and at 10 GPa such a component should be significant (i.e. Si 3·12 c.p.f.u.; Irifune, 1987). During decompression, such garnets would either retain their majoritic compositions or exsolve pyroxene. As there is no textural evidence for exsolution of pyroxene from AA garnets and our microprobe analyses leave no room for a significant majoritic component, an ultradeep origin of the AA peridotite would imply that the rock volumes containing former majorite and pyroxene porphyroclasts had recrystallized completely at 1180°C and 5·9 GPa producing almost homogeneous grains of grt I and pyroxene Ib with the compositions preserved in the cores of these mineral grains. If this were true, it is difficult to imagine how the thin clinopyroxene exsolution lamellae in the cpx Ia and cpx Ib grains, as well as the rod-shaped ilmenite exsolution bodies in olivine Ib were produced at the same time.
The composition of low-I clinopyroxene (P21/c) exsolution lamellae in cpx Ia and Ib grains as reported by Bozhilov et al. (1999) is Mg1·57Fe0·25Na0·03Al0·02Ca0·14Si2O6. According to experimental data on the pyroxene solvus (Brey & Köhler, 1990; Brey et al., 1990) such a pigeonitic composition is not stable at temperatures of 1200°C, but instead, requires excessive temperatures of 1800°C (at 6 GPa) and 1480°C (at 1 GPa). Such temperatures are, however, inconsistent with the compositions of cpx Ib, opx Ib and grt I grains. This problem needs further investigation.
The peak metamorphic conditions experienced by the AA peridotite as derived from the microanalytical data presented in this study are only slightly higher than those given by Brenker & Brey (1997). Our preferred retrograde P–T path, however, differs from the one suggested by those workers (Fig. 9) for the following reasons: (1) chemical zonation profiles across AA garnet I and pyroxene Ib grains are best explained by a near-isothermal decompression followed by rapid cooling; (2) rim compositions of AA minerals probably do not reflect a discrete partial re-equilibration stage allowing combination of thermometers with the Al-in-opx barometer, but rather were frozen in during rapid cooling; (3) temperatures of 500–600°C suggested by Brenker & Brey (1997) for the formation of narrow shear zones containing the stable parageneses ol + spl + opx + amphibole ± cpx are most probably too low, as the formation of amphibole signals the infiltration of hydrous fluids and in the presence of H2O the assemblage opx + ol + spl will be replaced by Mg-rich chlorite at temperatures below 800°C (Jenkins & Chernosky, 1986; Schmädicke, 2000).
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ACKNOWLEDGEMENTS |
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We thank H.-P. Meyer, T. Ludwig and H.-M. Seitz for technical assistance during electron and ion microprobe measurements, and I. Fin and U. Geilenkirchen for preparing numerous polished thin sections. Motivating discussions with F. Brenker, G. Brey, A. Kalt, B. Olker, H.-M. Seitz and A. B. Woodland are appreciated. We are thankful for the constructive and helpful reviews by D. A. Carswell, L. G. Medaris, Jr and V. Trommsdorff. This study is part of the Dr. rer. nat. thesis of J.P., which is financially supported by the Studienstiftung des Deutschen Volkes e. V.
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FOOTNOTES |
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*Corresponding author. Telephone: +49-6221-54-4651/8207. Fax: +49-6221-54-4805. E-mail: jpaquin{at}min.uni-heidelberg.de
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