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Abstract

Garnet peridotite forms a large portion of the upper mantle but also occurs not uncommonly in association with crustal rocks in collisional mountain belts. The examples described here are all of the latest type, with the emphasis placed upon texture and metamorphism. All these examples are from the Central Alps and from the Eastern Alps, i.e., areas that have not been dealt with in previous chapters.

Garnet peridotite forms a large portion of the upper mantle but also occurs not uncommonly in association with crustal rocks in collisional mountain belts. The examples described here are all of the latest type, with the emphasis placed upon texture and metamorphism. All these examples are from the Central Alps and from the Eastern Alps, i.e., areas that have not been dealt with in previous chapters.

Central Alpine domain

The known outcrops of garnet lherzolite in the Central Alps all occur in the same tectonic zone, namely the upper part of the Adula-Cima Lunga unit (Fig. 1). Palaeogeographically, this unit has been considered part of the former European continental margin (Schmid et al., 1990) and thus it forms the uppermost nappe complex of the lower Penninic system. Heinrich (1982, 1986) mapped mineral assemblages in mafic and pelitic lithologies of the Adula nappe and established isograds of a regional high-pressure metamorphism predating the classic, Lepontine isograd belt (Trommsdorff, 1966; Jager et al., 1967) of the Central Alps. The peak conditions of the high-pressure metamorphism increase from north to south from 1 GPa and 500 °C to over 2-5 GPa and 800 °C, based on thermobarometry of eclogites, metapelites and meta-ophicalcite rocks (Heinrich, 1982, 1986; Partzsch, 1996; Meyre et al., 1997, 1999; Pfiffner, 1999).

Fig. 1.

Eocene high pressure metamorphism in the Adula–Cima Lunga Nappe (modified after Frese et al., 2003). Quartz eclogites (shaded dashed isograds and values in squares) indicate an increase in P–T from north to south (Heinrich, 1982). Eclogites and garnet peridotites show values around 800 °C and 3 GPa; some values for garnet peridotites (dots) are given after Nimis & Trommsdorff (2001). The isograds of the later amphibolite facies regional metamorphism (Oligocene decompression) crosscut the nappe boundaries.

Fig. 1.

Eocene high pressure metamorphism in the Adula–Cima Lunga Nappe (modified after Frese et al., 2003). Quartz eclogites (shaded dashed isograds and values in squares) indicate an increase in P–T from north to south (Heinrich, 1982). Eclogites and garnet peridotites show values around 800 °C and 3 GPa; some values for garnet peridotites (dots) are given after Nimis & Trommsdorff (2001). The isograds of the later amphibolite facies regional metamorphism (Oligocene decompression) crosscut the nappe boundaries.

Garnet lherzolites were found in the highest P-T region of the Adula-Cima Lunga nappe at three localities. These are, from east to west (Fig.1): Monte Duria (MD; Fumasoli, 1974), Alpe Arami (AA; Grubenmann, 1908; Möckel, 1969), and Cima di Gagnone (CdG; Evans & Trommsdorff, 1978). At all three localities, garnet lherzolites show a more or less pronounced layering determined by variations in olivine and pyroxene contents. At MD, garnet lherzolites occur in over 20 individual ultramafic bodies, which form 10-100 m boudins within migmatitic gneisses. Garnet lherzolites and gneisses are folded around a steeply plunging megafold having an amplitude of at least 4 km.

At AA, garnet lherzolites form the core of a 1 km × 400 m chlorite peridotite boudin surrounded by steeply south-dipping migmatitic gneisses. A discontinuous layer of eclogitic rocks separates the peridotite body from the country gneisses. At CdG, garnet lherzolite occurs in one of numerous ultramafic lenses of tens to hundreds of metres in size, surrounded by pelitic and semipelitic, migmatitic, gneisses, kyanite eclogites, marbles, and meta-ophicalcite rocks (Pfiffner & Trommsdorff, 1998). The peridotites show transitions to eclogite. Within the ultra-mafic rocks, the presence of metarodingite boudins, interpreted as former mid-ocean ridge basalt (MORB) dykes, testifies to an early serpentinite stage of the metaperidotites (Evans et al., 1979, 1981). The ultramafic–mafic–carbonate suite at CdG has been interpreted as derived from an ocean basin near a continental margin (Pfiffner & Trommsdorff, 1998), with the ultramafic rocks representing former subcontinental mantle that had been exhumed during oceanic rifting.

The chemical compositions of the garnet lherzolites from the Central Alps (O’Hara & Mercy, 1966; Fumasoli, 1974; Rost et al., 1974; Ernst, 1978; Evans & Trommsdorff, 1978; Pfiffner, 1999) are all remarkably similar and close to that of fertile mantle, which is typical for the subcontinental lithosphere in the Alpine realm and Liguria (Nicolas & Jackson, 1972; Piccardo et al., 1990; Menzies & Dupuy, 1991; Müntener, 1997). Isotope geochemical investigations of minerals from AA and CdG garnet lherzolites yielded consistent, Eocene ages (~ 40Ma), based on garnet–clinopyroxene–whole rock Sm–Nd isochrons (Becker, 1993) and of 43–35 Ma, based on the U–Pb method and SHRIMP analysis of zircons (Gebauer et al., 1992; Gebauer, 1996). These ages are consistent with the recent dating of Late Eocene high-pressure metamorphism in the Western Alps (Froitzheim et al., 1996; Gebauer et al., 1997). In agreement with the prograde metamorphism of the eclogite sequence mapped in the Adula-Cima Lunga unit by Heinrich (1982, 1986), most of the garnet peridotites from the Central Alps show evidence of prograde metamorphism, as documented by Evans & Trommsdorff (1978) for CdG. The garnet peridotite at CDG was interpreted by these authors, for the first time, as subduction zone garnet peridotite.

Poikiloblastic and porphyroclastic garnet peridotites in the Central Alps

Detailed petrographic descriptions of garnet lherzolites from MD, AA, and CdG have been given by Fumasoli (1974), Möckel (1969) and Evans & Trommsdorff (1978), respectively. Additional information regarding the textures is given here. The garnet lherzolites in the Central Alps exhibit two main textural types: poikiloblastic and porphyroclastic. The poikiloblastic garnet peridotites (Fig. 2b) occur at CdG and, locally, at MD.

Fig. 2.

Photomicrographs of HP metaperidotites from the Central Alps, after Frese et al. (2003). Lineations and trace foliation are parallel to the long edge of the images. Crossed polarisers. All mineral abbreviations after Kretz (1983). (a) Cima di Gagnone garnet peridotite (sample Mg160Tro). The foliation is defined by elongated and slightly flattened grains of olivine, enstatite and garnet. (b) Prograde poikiloblastic pyrope-rich garnet (dark shading) with inclusions of enstatite, olivine and amphibole from Cima di Gagnone (sample Mg160-4-8) overgrowing an older foliation. (c) To the left: oriented inclusions of ilmenite rods elongated parallel to [010] olivine and forming palisades parallel to (001) of olivine (Mg160-98-1). To the right: schematic representation of the crystallographic orientation of the rods and palisades after Risold et al. (2001). (d) Porphyroclastic texture with two generations of olivine from Alpe Arami. The older olivine I forms large, mm-sized grains, surrounded by a mortar textured matrix of recrystallised olivine II. (e) Poikiloblastic pyrope-rich garnet (dark shading) from Alpe Arami (sample AA4; courtesy of P. Nimis).

Fig. 2.

Photomicrographs of HP metaperidotites from the Central Alps, after Frese et al. (2003). Lineations and trace foliation are parallel to the long edge of the images. Crossed polarisers. All mineral abbreviations after Kretz (1983). (a) Cima di Gagnone garnet peridotite (sample Mg160Tro). The foliation is defined by elongated and slightly flattened grains of olivine, enstatite and garnet. (b) Prograde poikiloblastic pyrope-rich garnet (dark shading) with inclusions of enstatite, olivine and amphibole from Cima di Gagnone (sample Mg160-4-8) overgrowing an older foliation. (c) To the left: oriented inclusions of ilmenite rods elongated parallel to [010] olivine and forming palisades parallel to (001) of olivine (Mg160-98-1). To the right: schematic representation of the crystallographic orientation of the rods and palisades after Risold et al. (2001). (d) Porphyroclastic texture with two generations of olivine from Alpe Arami. The older olivine I forms large, mm-sized grains, surrounded by a mortar textured matrix of recrystallised olivine II. (e) Poikiloblastic pyrope-rich garnet (dark shading) from Alpe Arami (sample AA4; courtesy of P. Nimis).

The poikiloblastic garnet peridotite is weakly foliated, sometimes folded, with completely anhedral, often elongated garnet crystals concentrated along pyroxene-rich layers. Poikiloblastic garnet is often stuffed with abundant, sometimes folded inclusions of orthopyroxene and, less commonly, olivine, clinopyroxene and rare, pale brownish to greenish, rounded, magnesian Ca-amphibole. Olivine occurs in two generations: the older generation (olivine I) is by far major in quantity and was formed during high-pressure (HP) dehydration of a hydrous protolith, whereas the younger generation (olivine II) is minor and was formed during post-HP partial recrystallisation. Olivine I consists of up to 2 mm large, elongated grains containing numerous rod-shaped micro-inclusions of ilmenite (Fig. 2c). The rods are elongated parallel to [010] of olivine and form palisades aligned within the (001) plane of olivine (Risold et al., 2001). Especially along more pyroxenitic layers, pseudomorphs of olivine with inclusions of wormy ilmenite after titanian clinohumite are commonly found, with occasional relics of (OH)-titanian clinohumite (Evans & Trommsdorff, 1978). The breakdown of titanian clinohumite and the presence of planar OH defects with ilmenite palisades in olivine I indicate prograde metamorphism at high H2O activity (Trommsdorff et al., 2001). In addition to olivine, the matrix around the garnets is formed by an idioblastic mosaic of orthopyroxene, clinopyroxene and Ca-amphibole with a typical grain size of 0.5-1 mm.

At CdG, garnet has overgrown pre-existing, isoclinal folds involving all the matrix minerals, including Ca-amphibole and olivine + ilmenite pseudomorphs after titanian clinohumite (Pfiffner & Trommsdorff, 1998) (Fig. 3). Ca-amphibole grains enclosed in garnet have a distinctly higher K2O content (0.7 wt%) than those in the matrix (0.15 wt%). Thus there is a large body of evidence that, in the poikiloblastic type peridotites, garnet formed during prograde metamorphism. Evans & Trommsdorff (1978) interpreted the poikiloblastic peridotites from CdG as “subduction zone garnet peridotites” derived from a partially serpentinised, hydrous protolith. At CdG, prograde chlorite-amphibole peridotites are commonly found to be isofacial with the garnet peridotites. The simultaneous occurrence of both rock types can be explained in terms of variable bulk composition and/or H2O activity (Trommsdorff, 1990; Pfiffner, 1999).

Fig. 3.

Photomicrograph of garnet peridotite from CdG. Pokiloblastic garnet (grt) overgrows a preexisting microfold with diopside (dio), enstatite (en) and pargasitic amphibole (amph).

Fig. 3.

Photomicrograph of garnet peridotite from CdG. Pokiloblastic garnet (grt) overgrows a preexisting microfold with diopside (dio), enstatite (en) and pargasitic amphibole (amph).

The second textural type, a porphyroclastic garnet lherzolite, dominates at AA. In comparison with CdG, the AA garnet peridotite is composed of a four-phase assemblage (garnet, olivine, enstatite, diopside) which has been described in detail by Möckel (1969). Its microtexture is characterised by more or less equant, anhedral porphyroblasts of pyrope-rich garnet in a porphyroclastic matrix with a mortar texture (Fig. 2d). Garnet grew, most probably, poikiloblastically (Fig. 2e): its present, often porphyroclastic texture was induced by a late deformation. Garnet is mostly rounded, with a grain size up to 10 mm. Olivine, enstatite, clinopyroxene and, rarely, chromian spinel inclusions in garnet (Möckel, 1969; Nimis & Trommsdorff, 2001) indicate that garnet growth was prograde. Chromian diopside, up to 5 mm in diameter, is frequently concentrated near garnet. The matrix around garnet is composed of two generations. A porphyroclastic generation of olivine (olivine I), consists of large grains up to 2 mm in size, showing deformation-induced undulating extinction, kinks and subgrains, and containing rod-shaped ilmenite inclusions. A second generation (olivine II) displays small, subhedral grains forming the mortar texture between the porphyroclasts. The latter generation recrystallised during deformation along the Southern Steep Belt of the Central Alps (Trommsdorff et al., 2000). Similarly, enstatite and diopside form two generations (Paquin & Altherr, 2001). Apoikiloblastic microtexture (Fig. 2e), analogous to CdG, has been reported by Möckel (1969, sample G01-A) in one titanian clinohumite-bearing garnet peridotite block near AA: this texture also rarely occurs in the main AA body (Nimis, pers. comm.) (Fig. 2e). Recent detailed mapping (Bay, 1999) has confirmed that the locality of Alpe Stuello, where this block was sampled, represents remnants of a moraine deposit derived from the AA ultramafic body (see also Geological Map of Switzerland 1:25000, sheet 1313, Bellinzona; Bächlin et al., 1974).

Oriented ilmenite rods in olivine

The principal argument for an ultradeep origin of the AA rocks (Dobrzhinetskaya et al., 1996) was based upon the occurrence of topotactic FeTiO3 rods in olivine, originally described by Möckel (1969) as brown rutile needles, oriented parallel to [010]ol. On the basis of TEM electron diffraction patterns, Dobrzhinetskaya et al. (1996) re-determined the needles as ilmenite and three additional, previously unknown, polymorphs of FeTiO3 which they interpreted as intermediate between the high-pressure phase FeTi-perovskite (Leinenweber et al., 1991) and ilmenite. By means of optical measurements Dobrzhinetskaya et al. (1996) determined the bulk quantities of 1-3 vol% FeTiO3 rods in olivine corresponding to about 0.7 to over 2.0 wt% TiO2. These numbers have later been revised to lower values ranging from 0.1-0.9 vol% FeTiO3 (0.07 to 0.6 wt% TiO2 in olivine) with a mean of 0.53 vol% (Green et al., 1997). By comparison, the highest reported TiO2 content in natural olivine (Hervig et al., 1986) does not exceed 0.05 wt% (500 ppm). Maximum TiO2 solubility in olivine from experiments at oxygen fugacities equivalent to mantle conditions at 10 GPa/1400 °C (Ulmer et al., 1998) and 14 GPa/1600 °C (Gudfinnsson & Wood, 1998) yield 0.13 wt% and 0.11 wt% TiO2 in olivine, respectively. By contrast, Dobrzhinetskaya et al. (1996) determined, at uncontrolled oxygen fugacity, considerably higher solubility values, i.e. > 1 wt%, at 12 GPa/1700 K. Based on the very high Ti content and the determination of unknown polymorphs, Dobrzhinetskaya et al. (1996) inferred that the titanate rods in AA olivine exsolved as FeTiO3 perovskite during the polymorphic transformation of wadsleyite, which has a higher Ti solubility (Gudfinnsson & Wood, 1998), to olivine.

The existence of the three new FeTiO3 polymorphs has been questioned by Risold et al. (1997) on the basis of theoretical considerations and by Hacker et al. (1997) and Risold et al. (1997) because the published diffraction patterns can be explained by dynamical diffraction between ilmenite and the olivine matrix. The bulk quantities of TiO2 in Arami olivine published by Dobrzhinetskaya et al. (1996) have been questioned by Hacker et al. (1997) on the basis of EPMA measurements by Risold et al. (1996) and Reusser et al. (1998) on the basis of UV-laser ablation ICP-MS: all these authors agree on a TiO2 content in olivine of approximately 350 ppm (0.035 wt%). These results are still 2 to 20 times lower than the revised values determined optically by Green et al.(1997).

Ilmenite rods in olivine, identical in size, composition and quantity to those of Alpe Arami, have been detected in over ten peridotite localities of the Adula-Cima Lunga nappe (Fig. 1); among these are CdG and MD. Additional occurrences of ilmenite rods in olivine have been reported from the Nonsberg region in the Eastern Alps (Risold et al., 1999) and in the Sulu terrane, China (Hacker et al., 1997).

If the ilmenite exsolutions are indicative of an ultradeep origin, the geological evidence for prograde metamorphism mentioned above (see Nimis & Trommsdorff, 2001) would require that Central Alpine peridotites had to be subducted to more than 300 km and then exhumed again within a single collision cycle. However, based on new optical microscope and TEM observations, a mechanism has been proposed for the formation of the FeTiO3 rods in olivine that does not require mantle transition zone pressure, nor ultrahigh pressure (Risold et al., 2001): TEM investigation of the rods at AA, CdG and MD demonstrates that the titanate inclusions correspond to a Mg-bearing ilmenite with exclusively a trigonal crystal structure. No other FeTiO3 polymorphs have been identified. The shape of the ilmenite inclusions concurs with minimum topotactic misfits between the lattices of olivine and Mg-bearing ilmenite (Risold et al., 2003). Optical microscopy reveals that the ilmenite rods are not randomly distributed within olivine but have, at all localities, a preferred alignment parallel to (001)ol; Figure 2c, (Risold et al., 2001). TEM images of the ilmenite palisades at AA and CdG show that they are controlled by the presence of (001) planar defects in olivine along which the rods nucleated. The defects have been characterised by high resolution TEM and identified as 4.4 Å wide humite-type layers, i.e., (Mg2SiO4)-Mg(OH,F)2, intergrown within olivine. In such layers substitution up to 50 mol% MgH2 by Ti4+ is possible. Assuming Ti-saturated humite slabs, the maximum density of (001) faults observed at AA, approximately 1900 mm−1, can accommodate about 170 ppm TiO2 in olivine. This represents about half the values obtained by UV-laser ablation ICP-MS and EPMA analysis. Evidence for the annealing of the defects, however, suggests that only parts of the defects are still preserved. The formation of ilmenite rods is considered to be a consequence of the breakdown of isolated humite layers in olivine. The apparent TiO2 content in olivine at AA, CdG and MD is regarded as a function of the Ti-humite defect density.

Lattice preferred orientation (LPO) of olivine from AA and CdG

Orogenic peridotites (Den Tex, 1969) play a fundamental role in the understanding of the composition, geophysical properties and geodynamic processes of the Earth’s upper mantle. Forsterite-rich olivine dominates mantle composition and, being a relatively weak mineral, controls mantle rheology (e.g. Drury & Fitzgerald, 1998). During convection in the asthenosphere, deformation by intracrystalline slip, accompanied by dynamic recrystallisation, produces lattice preferred orientation (LPO) of olivine. Crystal LPO, with slip planes and slip directions parallel to the shear plane and shear direction, respectively, leads to an optimisation of the resolved shear stress on active slip systems, and thus to geometric softening of the aggregate. Because of a large elastic anisotropy of olivine crystals, the olivine LPO is the primary cause of seismic anisotropy in peridotites. The relationships between flow geometry, LPO and seismic anisotropy in olivine polycrystals have been repeatedly discussed for naturally and experimentally deformed aggregates (e.g. Nicolas & Christensen, 1987; Tommasi et al., 2000; Bystrickyet al., 2000).

Most published LPOs of olivine were obtained from ophiolites or xenoliths. They generally show strong preferred orientations of (010) or (0kl) parallel to foliation, and [100] or [u0w] parallel to lineation (Ben Ismail & Mainprice, 1998). This is in good agreement with slip systems active at high temperature and low strain rates with the primary Burger vector [100] (e.g. Nicolas & Poirier, 1976). These LPOs have been considered to be inherited mantle fabrics, although the rocks may have suffered structural and chemical changes during tectonic and metamorphic activity. In the following, textures of the type (010)[100] will be referred to as mantle LPO.

Numerous occurrences of peridotites are known in the Alpine orogen. They contain different structural and metamorphic records due to their different palaeogeographic settings and thus different geodynamic histories. Many of these rocks in the Western, Central and Ligurian Alps are not ophiolite sensu stricto, because they are lherzolites of subcontinental origin (Trommsdorff et al., 1993). Petrological and structural work was carried out on several of these occurrences with LPO data available (e.g. Boudier, 1978; Hoogerduijn Strating, 1991). Some of the investigated peridotites match olivine “mantle LPOs”, as they are ex-mantle fragments without significant Alpine overprint. Other peridotites in the Alps underwent subduction and a metamorphic overprint to high pressure and temperature. They now form part of the Alpine nappe system. These peridotites with an Alpine metamorphic signature display LPOs different from the mantle LPO outlined above (Frese et al., 2003).

The perhaps best known peridotite with an Alpine HP metamorphism is the garnet peridotite of Alpe Arami. Its olivine LPO is known since over 30 years (Möckel, 1969) and is exceptional, with olivine a axes, i.e. [100], oriented subnormal to foliation and c axes, [001], parallel to lineation within the foliation plane. This fact has led Dobrzhinetskaya et al. (1996) to conclude that the Arami LPO is yet another characteristic of extraordinary high pressure as had been inferred from the occurrence of ilmenite rods in olivine.

The LPOs of olivine from the unaltered garnet lherzolite at CdG and AA, however, show consistent patterns (Fig. 4a-c). The olivine [100] axes are distributed along a partial girdle normal to the lineation, with the highest concentrations subnormal to the foliation. The olivine [010] pole figures have a maximum near the foliation plane away from the lineation, and olivine [001] are concentrated in a point maximum close to the lineation. The first eigenvector for [100] plots nearly normal to the foliation plane and the first eigenvector for [001] plots nearly parallel to the lineation. The first eigenvalues for [100], [010] and [001] are generally twice as high as the second and third, which are about equal, indicating point-like distributions with a major random component.

Fig. 4.

Lattice preferred orientation for (a) olivine porphyroclasts from the AA garnet lherzolite, (b, c) olivine in two samples from the CdG garnet lherzolite, and (d)enstatite porphyroblasts from the CdG garnet lherzolite. Lower hemisphere equal area projection. Normal to foliation plane is vertical, and lineation is horizontal. All measurements done by EBSD. The texture index pfJ, the minimum and maximum density is in the contoured pole figures, the eigenvalues (E1, E2, E3) and eigenvectors (▲,■,●) of the corresponding orientation ellipsoid, and the number of data points per sample are given. After Frese et al. (2003).

Fig. 4.

Lattice preferred orientation for (a) olivine porphyroclasts from the AA garnet lherzolite, (b, c) olivine in two samples from the CdG garnet lherzolite, and (d)enstatite porphyroblasts from the CdG garnet lherzolite. Lower hemisphere equal area projection. Normal to foliation plane is vertical, and lineation is horizontal. All measurements done by EBSD. The texture index pfJ, the minimum and maximum density is in the contoured pole figures, the eigenvalues (E1, E2, E3) and eigenvectors (▲,■,●) of the corresponding orientation ellipsoid, and the number of data points per sample are given. After Frese et al. (2003).

The orientation distribution of enstatite (Fig. 4d) at CdG (Mg160Tro) shows strong maxima of [100] normal to the foliation, of [010] normal to lineation within the foliation plane, and of [001] parallel to lineation. The first eigenvector for enstatite [100] plots nearly normal to the foliation plane, and the one for [001] lies parallel to lineation. The first eigenvalues for all three pole figures are three times higher than the second and third, which are similar. All this indicates strong point maximum distributions, and a (100)[001] preferred orientation for enstatite as well as for olivine. The enstatite LPO is very sharp whereas the LPO of olivine is less distinct, but all distributions possess approximately orthorhombic sample symmetry.

The LPO of the first olivine generation from Alpe Arami (Fig. 4a) determined in this study confirms the results of Möouml;ckel (1969), with preferred orientation of olivine [100] subnormal to the foliation. The significant new result is that CdG olivine I grains and AA olivine porphyroclasts have a similar LPO.

Because the garnet peridotite at CdG formed its LPO at pressures around 3 GPa, the LPO at AA cannot be taken as argument for higher pressures.

Criteria determining the deformation conditions

The LPO of olivine in mantle peridotites is typically characterised by maxima of the [010] and [001] axes perpendicular to foliation, and of the [100] axes parallel to lineation. These fabrics originate from glide and dynamic recrystallisation with the dominant slip systems (010)[100] and (0kl)[100]. The (010)[100] system was experimentally determined for dry deformation at high temperatures (> 1200 °C; Nicolas et al., 1973).

Recently Jung & Karato (2001) produced several new olivine fabrics in experiments under wet conditions in nearly simple shear deformation (ca. 2 GPa/1200 °C). Their “type C” fabric is characterised by an oblique pattern with concentrations of [001] 20° off the shear direction, and of the (100) plane 20° off the shear plane. This fabric was produced at high H2O activity and modest shear stress (~ 250 MPa).

None of the investigated fresh garnet lherzolites of the Central Alps from CdG and AA show LPOs for olivine with [100] maxima parallel to the lineation, but mostly with [100] subnormal to foliation and [001] towards lineation. These LPOs are similar to the experimental “type C” fabric presented by Jung & Karato (2001). Their LPO suggests “easy slip” on the (100)[001] slip system in simple shear deformation. This system was observed to be dominant under high stress at temperatures below 900 °C in dry experiments (Carter & Ave Lallemant, 1970; Phakey et al., 1972). Jung & Karato (2001) argued that slip along [001] may be similarly enhanced by water at higher temperature as by high stress under dry conditions. A qualitative comparison of the LPOs measured at CdG and AA with the experimentally produced LPOs suggests deformation for AA and CdG with similar mechanisms being activated.

The LPO of enstatite in the CdG peridotite has similar maxima and is much stronger than that of olivine. This is in good agreement with LPO work done on the chlorite-enstatite-olivine schists from Val Cama, Central Alps (Trommsdorff & Evans, 1969). This LPO pattern is consistent with crystal plastic slip on (100)[001], the dominant slip system of enstatite (Muegge, 1898; Raleigh, 1965; Mackwell, 1991). The LPO is dominated by an “easy slip” orientation for this slip system with respect to foliation and lineation in a simple shear regime.

Based upon the distribution of the ilmenite rods within olivine I at CdG the following may also be considered: ilmenite palisades occur within olivine (001) planes (Risold et al. 2001), and the preferred orientation of their poles is therefore identical to the LPO of [001] of olivine. [001] of olivine is oriented preferentially parallel to the lineation within the foliation plane, with the rods aligned within a plane normal to the lineation. The olivine grains hosting the ilmenite palisades display shape and lattice preferred orientation along [001], with a mineral elongation which marks a macroscopically visible lineation. Normal to lineation, late extensional cracks normal to [001] of olivine are observed. These cracks have given access to fluid migration and to the formation of fluid inclusion trails parallel to the (001) plane of olivine. These extensional cracks are interpreted as having formed from the release of residual stress parallel to the lineation. The mineral growth must have occurred under glide active on (100)[001] of olivine, which is considered the reason for the observed fabric.

Eastern Alpine domain

The Ulten zone (Andreatta, 1935; 1948; Hoinkes & Thöni, 1993) is a high-grade basement unit of the Italian Eastern Alps which is composed of migmatites, Grt + Ky ± Sil (fibrolitic) gneisses, lenses of amphibolitised eclogites and slices of Spl/Grt peridotites (mineral abbreviations after Kretz, 1983). It has been classically attributed to the Upper Austroalpine system of the Central Eastern Alps (Fig. 5), a nappe pile of Cretaceous age (Thöni, 1981).

Fig. 5.

Geological sketch and main tectonic lineaments of the Nonsberg area (after Obata & Morten, 1987, and Nimis & Morten, 2000). (1) Bt-Ms paragneiss; (2) Grt-Ky paragneiss; (3) migmatite and orthogneiss; (4) siltstone, sandstone and conglomerate of the Insubric Flysch (Upper Cretaceous); (5) Triassic dolomite; (6) ultramafic rocks; (7) amphibolite; (8) tonalite; (9) marble.

Fig. 5.

Geological sketch and main tectonic lineaments of the Nonsberg area (after Obata & Morten, 1987, and Nimis & Morten, 2000). (1) Bt-Ms paragneiss; (2) Grt-Ky paragneiss; (3) migmatite and orthogneiss; (4) siltstone, sandstone and conglomerate of the Insubric Flysch (Upper Cretaceous); (5) Triassic dolomite; (6) ultramafic rocks; (7) amphibolite; (8) tonalite; (9) marble.

The Upper Austroalpine system consists of metasedimentary cover and upper-to-lower crust slices derived from the Mesozoic passive margin of the Adria microplate (Dal Piaz, 1993). According to Flügel (1990) and Neubauer & von Raumer (1993), the Austroalpine system of the Eastern Alps is a piece of the Variscan belt. Its hypothetical evolution history, as pointed out by various authors, can be briefly summarised, according to Godard et al. (1996), as follows. (i) Ocean closure before Devonian and collision of continental and arc elements during the Early Carboniferous (Neubauer & von Raumer, 1993) produced the Central European part of the Variscan fold belt (Neugebauer, 1990), possibly enclosing minor elements of pre-Variscan age (Ziegler, 1984, 1993; Maggetti & Flisch, 1993; von Raumer & Neubauer, 1993; Neubauer & von Raumer, 1993). (ii) This belt was dismembered and uplifted during Late Carboniferous/Early Permian times along transtensive fault systems that transected the orogen (Ziegler, 1984). The Variscan basement was further unroofed and dismembered by Permian and Mesozoic lithospheric thinning and rifting along normal faults (Eberli, 1988; Dal Piaz, 1993). (iii) The Late Triassic-Late Jurassic rifting and drifting stages created the Adria microplate (Africa), in relation to the opening of the Tethys basin between Africa and Europe. The inversion of these normal faults during the Cretaceous (Eo-Alpine) orogeny led to the thrusting of the Austroalpine system (i.e., the Adria microplate) northward and northwestward, over the colliding Penninic domain of the European plate margin (Froitzheim & Eberli, 1990).

In northwestern Trentino, the Austroalpine system is separated from the Southern Alps domain by the major Tertiary Periadriatic (Insubric) lineament, which is composed of the strike-slip to transpressive Tonale and Giudicarie faults (Fig. 5). North of these faults, the Austroalpine system is divided into the Tonale and Ortler nappes (Thöni, 1981) by the Peio line (Andreatta, 1948), a Cretaceous thrust which was folded and reactivated during the Tertiary (Martin et al., 1991). The northern cover-bearing Ortler nappe is overlain along the Peio thrust by the cover-free Tonale nappe, which includes the Tonale and Ulten zones (“series”, Andreatta, 1948).

The pre-Alpine basement of the Ortler nappe consists largely of Grt + St ± Sil micaschists with minor amphibolites of tholeiitic composition, calc-alkaline metagranitoid, marbles and serpentinite slivers. The Grt + Sil paragneisses in the Tonale zone of the Tonale nappe include similar marbles and mafic-ultramafic intercalations, although they display a higher-grade pre-Alpine metamorphic overprint. This lithological assemblage partly represents a disrupted ophiolitic suite within the Variscan basement (Martin & Prosser, 1993). A back-arc basin environment may be envisaged in comparison to the Speik Complex (Eastern Austroalpine system; Neubauer, 1988; Neubauer & Frisch, 1993). The basement was retrogressed during the late- or post-Variscan exhumation and was locally deformed under greenschist facies conditions during the Alpine orogeny (Thöni, 1981).

By contrast, the Ulten zone of the Tonale nappe is mainly composed of kyanite-bearing micaschists, migmatites and gneisses with well-preserved pre-Alpine high-grade metamorphic signatures, weakly overprinted by the Alpine metamorphism. It is bordered by two belts of sillimanite-bearing paragneisses, which represent the eastern continuation of the Tonale zone (Fig. 5). The Ulten zone is divided from the Bt - Ms ± Sil metasedimentary rocks of the Tonale zone by the Rumo line (Morten et al., 1976); the higher-grade rocks of the Ulten zone are located north of the Rumo line (Nonsberg area) and consist of migmatites, Grt-Ky gneisses, with boudins of amphibolitised eclogites and Spl/Grt peridotites (Fig. 5).

Nonsberg ultramafic rocks

The Nonsberg ultramafic rocks mainly consist of barrel-shaped bodies tens to hundreds of metres in size, which are generally located along the boundary between the underlying strongly-foliated gneisses and the overlying migmatites and orthogneisses (Fig. 5). They are predominantly peridotites of harzburgitic to lherzolitic composition (Bondi et al., 1992). Compositional banding, primarily due to modal variation of pyroxene is conspicuous in some outcrops. Layers or bands of pyroxenite, some of which are complexly folded, also occur (Morten & Obata, 1983).

On a textural and grain size basis two petrographic types have been distinguished (Obata & Morten, 1987): coarse type and fine type. The coarse type rocks are coarsegrained (up to few centimetres in grain size) and are relatively undeformed. The fine type rocks are of much finer grain size (0.2-1 mm) and show various metamorphic textures. In the Nonsberg area, the fine type prevails in abundance over the coarse type. In some ultramafic bodies, the coarse type sporadically occurs in the more typical fine-grained matrix. The field and petrographic observations of many typical and transitional rock types have indicated that the fine types have been derived from the coarse types by syntectonic recrystallisation and that the Grt-bearing rocks derived from previous Spl-bearing peridotites (Obata & Morten, 1987). Various hypotheses have been put forth to explain the transformation from spinel peridotites to garnet peridotites and to depict the evolution of the Nonsberg ultramafic rocks. Proposed scenarios include: (i) upwelling of deep mantle material and partial melting in the spinel lherzolite field followed by isobaric cooling and then tectonic emplacement into the crust (Herzberg et al., 1977); (ii) high-pressure metamorphism after emplacement of spinel lherzolite into the crust (Rost & Brenneis, 1978); (iii) tectonic emplacement of spinel lherzolite into the crust followed by “burial” into deeper levels (Rost et al., 1979); (iv) emplacement of high-temperature spinel lherzolite into the lower crust followed by conductive cooling of the body essentially at the same depth (Obata & Morten, 1987); (v) emplacement of spinel lherzolite bodies into the crust and then sinking through the less dense, ductile, sialic lower crust (Obata & Morten, 1985); (vi) emplacement into the crust of sliced fragments of a mantle wedge overlying a subducting slab in a collisional environment (Godard et al., 1996). More recently, on the basis of P-T estimates, the inferred P-T path and physical models for subduction zones, Nimis & Morten (2000) proposed the following scenario. High temperatures, exceeding 1400 °C, were attained in the innermost portion of a mantle wedge overlying a subducting continental slab. Hydrous melts were produced within these hot mantle portions and rose into the overlying, lower-T spinel lherzolites (~ 1200 °C, 1.31-1.58 GPa), where the rising melts produced high-T (> 1400 °C) pyroxenitic segregates. Convection induced by the movement of the subducting plate caused wedge peridotites to cool while flowing towards the slab at essentially constant depth (isobaric cooling path). The mantle material was then driven to greater depth by the downward flow near the wedge-slab interface (T decrease and P increase path) and was eventually incorporated in the crust as tectonic slices or as sinking blobs. Mantle flow and entrainment in the crust caused the peridotites to cool down to ~ 850 °C before, or while, being subducted together with the slab to a depth of about 90 km. The latest subduction stage, which led to the formation of garnet, is represented by a nearly isothermal path. The transformation from spinel peridotite to garnet peridotite (from coarse type to transitional and fine type) is accompanied by significant input of crustal metasomatic agents (Morten & Obata, 1990) as indicated by the crystallisation of abundant, up to 23 vol%, amphibole with crustal geochemical signatures (Rampone & Morten, 2001). During the subsequent exhumation history, peridotites and continental crust shared the same decompressional and cooling path, which is testified by mineral zoning patterns in ultramafic rocks and is best recorded by mafic and pelitic country rocks, as well as by kelyphite minerals after peridotitic garnets (Godard et al., 1996).

Textures of the Nonsberg peridotites

Coarse type

Most of the coarse type peridotites are spinel lherzolite with protogranular texture (Fig. 6a,b) similar to that observed in many ultramafic xenoliths in alkali basalts (Mercier & Nicolas, 1975). The constituent minerals are olivine (up to 5 mm grain size), enstatite (2-4 mm), diopside (0.7-2 mm), and Cr-Al spinel (0.6-1 mm). In some samples Ca-amphibole (actinolitic) occurs along pyroxene grain boundaries and locally in finegrained recrystallised parts. Large grains of olivine and enstatite show undulating extinction and kink band textures. Enstatite contains exsolution lamellae of diopside and spinel, and diopside contains lamellae of enstatite and spinel. Spinel occurs as large independent grains, as well as exsolutions in pyroxenes, and they are typically intergrown with enstatite (Fig. 6c).

Fig. 6. (a)

Protogranular texture of coarse type Spl-lherzolite. Both pyroxenes exsolve lamellae of Ca-poor and Ca-rich pyroxene, respectively. Crossed polarisers. The field of view is 3 mm. (b) Protogranular texture of coarse type Spl-lherzolite. The clinopyroxene (left bottom corner) carries exsolution lamellae of Ca-poor pyroxene and spinel. Crossed polarisers. Width of view is 2 mm. (c) Coarse type Spl-lherzolite. Spinel intergrowths with enstatite. Plane polarised light. Field of view is 1.5 mm.

Fig. 6. (a)

Protogranular texture of coarse type Spl-lherzolite. Both pyroxenes exsolve lamellae of Ca-poor and Ca-rich pyroxene, respectively. Crossed polarisers. The field of view is 3 mm. (b) Protogranular texture of coarse type Spl-lherzolite. The clinopyroxene (left bottom corner) carries exsolution lamellae of Ca-poor pyroxene and spinel. Crossed polarisers. Width of view is 2 mm. (c) Coarse type Spl-lherzolite. Spinel intergrowths with enstatite. Plane polarised light. Field of view is 1.5 mm.

In a few coarse type samples, garnet occurs in addition to spinel. The texture is transitional from protogranular to porphyroclastic, where large grains of olivine and pyroxene appear to be more strained than those in normal spinel lherzolites and the smaller, unstrained grains (neoblasts, approximately 0.3 mm grain size) are more abundant. Garnet occurs surrounding spinel (Fig. 7a) and also forms 10-50 μm thick exsolution lamellae, parallel to (100) of the host (Fig. 7b), in both enstatite and diopside.

Fig. 7.(a)

Coarse type Spl-Grt peridotite. Garnet porphyroblast with an inclusion of corroded Cr-spinel. Partly crossed polarisers. Field of view is 2 mm. (b) Coarse type Spl-Grt peridotite. Both clinopyroxene (left) and orthopyroxene (right) contain exsolution lamellae of garnet. Crossed polarisers. Field of view is 2 mm. (c) Coarse type Spl-Grt peridotite. Clinopyroxene with exsolution lamellae of orthopyroxene, garnet and amphibole. Crossed polarisers. Field of view is 2 mm.

Fig. 7.(a)

Coarse type Spl-Grt peridotite. Garnet porphyroblast with an inclusion of corroded Cr-spinel. Partly crossed polarisers. Field of view is 2 mm. (b) Coarse type Spl-Grt peridotite. Both clinopyroxene (left) and orthopyroxene (right) contain exsolution lamellae of garnet. Crossed polarisers. Field of view is 2 mm. (c) Coarse type Spl-Grt peridotite. Clinopyroxene with exsolution lamellae of orthopyroxene, garnet and amphibole. Crossed polarisers. Field of view is 2 mm.

In addition to the garnet and pyroxene lamellae, the presence of very thin (less than 1 μm thick) lamellae of Cr-spinel was confirmed in enstatite parallel to (100) of the host by transmission electron microscopy (Obata & Morten, 1987). Some large diopside grains contain Ca-amphibole lamellae, whose composition is identical to the neoblasts in the same sample (Fig. 7c).

The large spinel grains are always mantled by garnet; they are brown to reddish brown and amoeboid to holly leaf shaped, with a maximum grain size of 2 mm. The smaller the size, the darker the color and more chromian the composition. In finegrained, recrystallised parts of the thin section, the garnet lherzolite assemblage, i.e., olivine + enstatite + diopside + garnet, occurs together with hornblende, but without spinel. It is evident from petrographic observations reported above that the spinel-garnet lherzolite was originally a protogranular spinel lherzolite that has been partly transformed to garnet lherzolite.

Fine type

As mentioned above, the fine type is more abundant in this region. It is finer grained and the texture is porphyroclastic to tabular or mosaic equigranular. It contains more varieties of mineral assemblages than the coarse type. Observed mineral assemblages are summarised as follows (Obata & Morten, 1987; Morten, 1993):

 

formula

Pyrope-rich garnet is the stable phase in the first two groups. Clinopyroxene is minor or absent. Ca-amphibole, containing up to 1 wt% K2O (Obata & Morten, 1987; Rampone & Morten, 2001) is an important hydrous, Ca-bearing phase. It was considered (Obata & Morten, 1987) to grow according to the clinopyroxene and garnet consuming reactions (Obata & Thompson, 1981): Cpx + Opx + Grt +H2O ↔ Ca-amph + Ol and, under even higher partial pressure of H2O, Grt + Ol + H2O ↔ Ca-amph + Opx + Spl.

No recent data exist of LPO of olivine from the Nonsberg peridotites. However Andreatta (1934) measured with a U-stage the LPO of olivine from both granular-textured spinel and garnet peridotites. In both cases the LPO resulted in [010] normal to foliation, [001] and [100] parallel to foliation, i.e., the mantle LPO (see above).

In porphyroclastic rocks, porphyroclasts of olivine, enstatite, and diopside are up to 5 mm in size and are more or less elongated and show signs of deformation such as undulating extinction, kink band boundaries and subgrains (Fig. 8a). Kinked orthopyroxene porphyroclasts contain numerous exsolution lamellae of garnet and clinopyroxene (Fig. 8b and c), and the clinopyroxene porphyroclasts contain those of orthopyroxene. Olivine neoblasts show granular polygonal texture (Fig. 8a). Pyroxene neoblasts are unstrained and contain no exsolution lamellae. Spinel inclusions are commonly observed in olivine neoblasts, while they are very rare in the olivine porphyroclasts.

Fig. 8.(a)

Porphyroclastic texture of Grt-Spl peridotite. Large kinked olivine porphyroclast within neoblastic, granular polygonal matrix. Crossed nicols. The field of view is 2.5 mm. (b) Porphyroclastic texture of Spl-Grt peridotite. The highly strained orthopyroxene porphyroclast contains garnet lamellae. Crossed polarisers. The field of view is 3 mm. (c) Porphyroclastic texture of Spl-Grt peridotite. The kinked orthopyroxene porphyroclast contains garnet and clinopyroxene lamellae. Small clinopyroxene grains crystallised along the kink planes. Crossed polarisers. Field of view is 0.6 mm.

Fig. 8.(a)

Porphyroclastic texture of Grt-Spl peridotite. Large kinked olivine porphyroclast within neoblastic, granular polygonal matrix. Crossed nicols. The field of view is 2.5 mm. (b) Porphyroclastic texture of Spl-Grt peridotite. The highly strained orthopyroxene porphyroclast contains garnet lamellae. Crossed polarisers. The field of view is 3 mm. (c) Porphyroclastic texture of Spl-Grt peridotite. The kinked orthopyroxene porphyroclast contains garnet and clinopyroxene lamellae. Small clinopyroxene grains crystallised along the kink planes. Crossed polarisers. Field of view is 0.6 mm.

The amphibole is colorless or pale green to brownish green. It is typically zoned, being more tremolite-rich at the grain margins, where compositional change is so abrupt that sometimes Becke lines are visible within the grains. The amphibole grains typically include small, irregular shaped spinel (Fig. 9a) and, in few cases, thin lamellae of phlogopite occur in some amphibole grains.

Fig. 9.(a)

Fine type amphibole-bearing peridotite. Amphibole includes small, irregular shaped spinels. Parallel polarisers. Field of view is 2 mm. (b) Fine type Grt-Spl-Amph peridotite. Relict of coarse garnet porphyroblast with inclusion of corroded spinel set up in fine-grained matrix. Partly crossed polarisers. Field of view is 4 mm. (c) Fine type chlorite-bearing peridotite. Chlorite grains mark the rock foliation. Partly crossed polarisers. Field of view is 1.25 mm. (d) Granoblastic polygonal texture of a fine type peridotite. Dolomite grain surrounded by small apatite grains. Plane polarised light. Field of view is 1.3 mm.

Fig. 9.(a)

Fine type amphibole-bearing peridotite. Amphibole includes small, irregular shaped spinels. Parallel polarisers. Field of view is 2 mm. (b) Fine type Grt-Spl-Amph peridotite. Relict of coarse garnet porphyroblast with inclusion of corroded spinel set up in fine-grained matrix. Partly crossed polarisers. Field of view is 4 mm. (c) Fine type chlorite-bearing peridotite. Chlorite grains mark the rock foliation. Partly crossed polarisers. Field of view is 1.25 mm. (d) Granoblastic polygonal texture of a fine type peridotite. Dolomite grain surrounded by small apatite grains. Plane polarised light. Field of view is 1.3 mm.

Garnet occurs as large porphyroclastic crystals (up to 2-3 mm size) and also as small grains which are in textural equilibrium with olivine, pyroxenes, amphibole, and perhaps with spinel. Large garnet crystals typically include vermicular grains of Cr-spinel, which are partly armoured by olivine (Fig. 9b). Spinel inclusions are less common in small garnet but olivine, pyroxene and, sometimes, carbonate (dolomite) inclusions are commonly observed.

Chrome spinel is a common accessory phase and it is dispersed as small (less than 0.2 mm in diameter) translucent brown grains. Although the spinel included in the garnet may be a relic of earlier assemblages, these small grains dispersed among the olivine and pyroxene matrix are considered to be in equilibrium with garnet (Obata & Morten, 1987).

In the chlorite-bearing assemblage, chlorite appears to be in textural equilibrium with olivine, orthopyroxene and amphibole, while spinel is nearly opaque and rare: this indicates that in this assemblage chlorite, instead of spinel, is the stable aluminous phase (Fig. 9c).

Other phases such as phlogopite, dolomite, and apatite were recognised in some samples. In a garnet-bearing sample, several apatite grains occur around a large grain of dolomite (Fig. 9d). Locally, the peridotites have undergone retrogression to an olivine + tremolite + cummingtonite + chlorite + talc assemblage.

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Acknowledgements

The authors would like to acknowledge scientific support by Katrine Frese and Anne Chantal Risold, technical support by Roberto Braga and financial support to L. M. by MIUR and CNR. A careful review by G. B. Piccardo helped to improve the manuscript and is also acknowledged.

Figures & Tables

Fig. 1.

Eocene high pressure metamorphism in the Adula–Cima Lunga Nappe (modified after Frese et al., 2003). Quartz eclogites (shaded dashed isograds and values in squares) indicate an increase in P–T from north to south (Heinrich, 1982). Eclogites and garnet peridotites show values around 800 °C and 3 GPa; some values for garnet peridotites (dots) are given after Nimis & Trommsdorff (2001). The isograds of the later amphibolite facies regional metamorphism (Oligocene decompression) crosscut the nappe boundaries.

Fig. 1.

Eocene high pressure metamorphism in the Adula–Cima Lunga Nappe (modified after Frese et al., 2003). Quartz eclogites (shaded dashed isograds and values in squares) indicate an increase in P–T from north to south (Heinrich, 1982). Eclogites and garnet peridotites show values around 800 °C and 3 GPa; some values for garnet peridotites (dots) are given after Nimis & Trommsdorff (2001). The isograds of the later amphibolite facies regional metamorphism (Oligocene decompression) crosscut the nappe boundaries.

Fig. 2.

Photomicrographs of HP metaperidotites from the Central Alps, after Frese et al. (2003). Lineations and trace foliation are parallel to the long edge of the images. Crossed polarisers. All mineral abbreviations after Kretz (1983). (a) Cima di Gagnone garnet peridotite (sample Mg160Tro). The foliation is defined by elongated and slightly flattened grains of olivine, enstatite and garnet. (b) Prograde poikiloblastic pyrope-rich garnet (dark shading) with inclusions of enstatite, olivine and amphibole from Cima di Gagnone (sample Mg160-4-8) overgrowing an older foliation. (c) To the left: oriented inclusions of ilmenite rods elongated parallel to [010] olivine and forming palisades parallel to (001) of olivine (Mg160-98-1). To the right: schematic representation of the crystallographic orientation of the rods and palisades after Risold et al. (2001). (d) Porphyroclastic texture with two generations of olivine from Alpe Arami. The older olivine I forms large, mm-sized grains, surrounded by a mortar textured matrix of recrystallised olivine II. (e) Poikiloblastic pyrope-rich garnet (dark shading) from Alpe Arami (sample AA4; courtesy of P. Nimis).

Fig. 2.

Photomicrographs of HP metaperidotites from the Central Alps, after Frese et al. (2003). Lineations and trace foliation are parallel to the long edge of the images. Crossed polarisers. All mineral abbreviations after Kretz (1983). (a) Cima di Gagnone garnet peridotite (sample Mg160Tro). The foliation is defined by elongated and slightly flattened grains of olivine, enstatite and garnet. (b) Prograde poikiloblastic pyrope-rich garnet (dark shading) with inclusions of enstatite, olivine and amphibole from Cima di Gagnone (sample Mg160-4-8) overgrowing an older foliation. (c) To the left: oriented inclusions of ilmenite rods elongated parallel to [010] olivine and forming palisades parallel to (001) of olivine (Mg160-98-1). To the right: schematic representation of the crystallographic orientation of the rods and palisades after Risold et al. (2001). (d) Porphyroclastic texture with two generations of olivine from Alpe Arami. The older olivine I forms large, mm-sized grains, surrounded by a mortar textured matrix of recrystallised olivine II. (e) Poikiloblastic pyrope-rich garnet (dark shading) from Alpe Arami (sample AA4; courtesy of P. Nimis).

Fig. 3.

Photomicrograph of garnet peridotite from CdG. Pokiloblastic garnet (grt) overgrows a preexisting microfold with diopside (dio), enstatite (en) and pargasitic amphibole (amph).

Fig. 3.

Photomicrograph of garnet peridotite from CdG. Pokiloblastic garnet (grt) overgrows a preexisting microfold with diopside (dio), enstatite (en) and pargasitic amphibole (amph).

Fig. 4.

Lattice preferred orientation for (a) olivine porphyroclasts from the AA garnet lherzolite, (b, c) olivine in two samples from the CdG garnet lherzolite, and (d)enstatite porphyroblasts from the CdG garnet lherzolite. Lower hemisphere equal area projection. Normal to foliation plane is vertical, and lineation is horizontal. All measurements done by EBSD. The texture index pfJ, the minimum and maximum density is in the contoured pole figures, the eigenvalues (E1, E2, E3) and eigenvectors (▲,■,●) of the corresponding orientation ellipsoid, and the number of data points per sample are given. After Frese et al. (2003).

Fig. 4.

Lattice preferred orientation for (a) olivine porphyroclasts from the AA garnet lherzolite, (b, c) olivine in two samples from the CdG garnet lherzolite, and (d)enstatite porphyroblasts from the CdG garnet lherzolite. Lower hemisphere equal area projection. Normal to foliation plane is vertical, and lineation is horizontal. All measurements done by EBSD. The texture index pfJ, the minimum and maximum density is in the contoured pole figures, the eigenvalues (E1, E2, E3) and eigenvectors (▲,■,●) of the corresponding orientation ellipsoid, and the number of data points per sample are given. After Frese et al. (2003).

Fig. 5.

Geological sketch and main tectonic lineaments of the Nonsberg area (after Obata & Morten, 1987, and Nimis & Morten, 2000). (1) Bt-Ms paragneiss; (2) Grt-Ky paragneiss; (3) migmatite and orthogneiss; (4) siltstone, sandstone and conglomerate of the Insubric Flysch (Upper Cretaceous); (5) Triassic dolomite; (6) ultramafic rocks; (7) amphibolite; (8) tonalite; (9) marble.

Fig. 5.

Geological sketch and main tectonic lineaments of the Nonsberg area (after Obata & Morten, 1987, and Nimis & Morten, 2000). (1) Bt-Ms paragneiss; (2) Grt-Ky paragneiss; (3) migmatite and orthogneiss; (4) siltstone, sandstone and conglomerate of the Insubric Flysch (Upper Cretaceous); (5) Triassic dolomite; (6) ultramafic rocks; (7) amphibolite; (8) tonalite; (9) marble.

Fig. 6. (a)

Protogranular texture of coarse type Spl-lherzolite. Both pyroxenes exsolve lamellae of Ca-poor and Ca-rich pyroxene, respectively. Crossed polarisers. The field of view is 3 mm. (b) Protogranular texture of coarse type Spl-lherzolite. The clinopyroxene (left bottom corner) carries exsolution lamellae of Ca-poor pyroxene and spinel. Crossed polarisers. Width of view is 2 mm. (c) Coarse type Spl-lherzolite. Spinel intergrowths with enstatite. Plane polarised light. Field of view is 1.5 mm.

Fig. 6. (a)

Protogranular texture of coarse type Spl-lherzolite. Both pyroxenes exsolve lamellae of Ca-poor and Ca-rich pyroxene, respectively. Crossed polarisers. The field of view is 3 mm. (b) Protogranular texture of coarse type Spl-lherzolite. The clinopyroxene (left bottom corner) carries exsolution lamellae of Ca-poor pyroxene and spinel. Crossed polarisers. Width of view is 2 mm. (c) Coarse type Spl-lherzolite. Spinel intergrowths with enstatite. Plane polarised light. Field of view is 1.5 mm.

Fig. 7.(a)

Coarse type Spl-Grt peridotite. Garnet porphyroblast with an inclusion of corroded Cr-spinel. Partly crossed polarisers. Field of view is 2 mm. (b) Coarse type Spl-Grt peridotite. Both clinopyroxene (left) and orthopyroxene (right) contain exsolution lamellae of garnet. Crossed polarisers. Field of view is 2 mm. (c) Coarse type Spl-Grt peridotite. Clinopyroxene with exsolution lamellae of orthopyroxene, garnet and amphibole. Crossed polarisers. Field of view is 2 mm.

Fig. 7.(a)

Coarse type Spl-Grt peridotite. Garnet porphyroblast with an inclusion of corroded Cr-spinel. Partly crossed polarisers. Field of view is 2 mm. (b) Coarse type Spl-Grt peridotite. Both clinopyroxene (left) and orthopyroxene (right) contain exsolution lamellae of garnet. Crossed polarisers. Field of view is 2 mm. (c) Coarse type Spl-Grt peridotite. Clinopyroxene with exsolution lamellae of orthopyroxene, garnet and amphibole. Crossed polarisers. Field of view is 2 mm.

Fig. 8.(a)

Porphyroclastic texture of Grt-Spl peridotite. Large kinked olivine porphyroclast within neoblastic, granular polygonal matrix. Crossed nicols. The field of view is 2.5 mm. (b) Porphyroclastic texture of Spl-Grt peridotite. The highly strained orthopyroxene porphyroclast contains garnet lamellae. Crossed polarisers. The field of view is 3 mm. (c) Porphyroclastic texture of Spl-Grt peridotite. The kinked orthopyroxene porphyroclast contains garnet and clinopyroxene lamellae. Small clinopyroxene grains crystallised along the kink planes. Crossed polarisers. Field of view is 0.6 mm.

Fig. 8.(a)

Porphyroclastic texture of Grt-Spl peridotite. Large kinked olivine porphyroclast within neoblastic, granular polygonal matrix. Crossed nicols. The field of view is 2.5 mm. (b) Porphyroclastic texture of Spl-Grt peridotite. The highly strained orthopyroxene porphyroclast contains garnet lamellae. Crossed polarisers. The field of view is 3 mm. (c) Porphyroclastic texture of Spl-Grt peridotite. The kinked orthopyroxene porphyroclast contains garnet and clinopyroxene lamellae. Small clinopyroxene grains crystallised along the kink planes. Crossed polarisers. Field of view is 0.6 mm.

Fig. 9.(a)

Fine type amphibole-bearing peridotite. Amphibole includes small, irregular shaped spinels. Parallel polarisers. Field of view is 2 mm. (b) Fine type Grt-Spl-Amph peridotite. Relict of coarse garnet porphyroblast with inclusion of corroded spinel set up in fine-grained matrix. Partly crossed polarisers. Field of view is 4 mm. (c) Fine type chlorite-bearing peridotite. Chlorite grains mark the rock foliation. Partly crossed polarisers. Field of view is 1.25 mm. (d) Granoblastic polygonal texture of a fine type peridotite. Dolomite grain surrounded by small apatite grains. Plane polarised light. Field of view is 1.3 mm.

Fig. 9.(a)

Fine type amphibole-bearing peridotite. Amphibole includes small, irregular shaped spinels. Parallel polarisers. Field of view is 2 mm. (b) Fine type Grt-Spl-Amph peridotite. Relict of coarse garnet porphyroblast with inclusion of corroded spinel set up in fine-grained matrix. Partly crossed polarisers. Field of view is 4 mm. (c) Fine type chlorite-bearing peridotite. Chlorite grains mark the rock foliation. Partly crossed polarisers. Field of view is 1.25 mm. (d) Granoblastic polygonal texture of a fine type peridotite. Dolomite grain surrounded by small apatite grains. Plane polarised light. Field of view is 1.3 mm.

Contents

GeoRef

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