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Journal of Petrology | Volume 45 | Number 5 | Pages 1013-1043 | 2004
Journal of Petrology 45(5) © Oxford University Press 2004; all rights reserved.
The Relationship between Tectono-metamorphic Evolution and Argon Isotope Records in White Mica: Constraints from in situ 40Ar–39Ar Laser Analysis of the Variscan Basement of Sardinia
1 ISTITUTO DI GEOSCIENZE E GEORISORSE–CNR, VIA MORUZZI 1, I-56124 PISA, ITALY
2 DIPARTIMENTO DI SCIENZE DELLA TERRA, UNIVERSITÀ DI PISA, VIA S. MARIA 53, I-56126 PISA, ITALY
3 MUSEO NAZIONALE DELL'ANTARTIDE, UNIVERSITÀ DI SIENA, VIA LATERINA 8, I-53100 SIENA, ITALY
RECEIVED JUNE 7, 2003; ACCEPTED OCTOBER 30, 2003
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL BACKGROUNDEXPERIMENTAL PROCEDURESRESULTSDISCUSSIONREFERENCES |
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The basement of Sardinia represents a nearly complete section of a segment of the Variscan belt that experienced a polyphase tectono-metamorphic evolution and Barrovian metamorphism. This basement is well suited to investigate the relationship between tectono-metamorphic evolution and argon isotope records in white mica. Micaschists from the garnet zone (maximum T of up to 520–560°C) contain two texturally and chemically resolvable generations of white mica: (1) deformed celadonite-rich flakes, defining a relict S1 foliation preserved within the main S2 foliation or enclosed in rotated albite porphyroblasts; (2) celadonite-poor white micas aligned along the main S2 foliation. The S1 foliation developed earlier and at a deeper crustal level with respect to that at which the thermal peak was reached. From the staurolite zone (T of up to 590–625°C) to the sillimanite + K-feldspar zone, white mica is nearly uniform in composition (muscovite) and is predominantly aligned along the S2 foliation or is of later crystallization. In situ 40Ar–39Ar laser analyses of white mica yielded ages of
340–315 Ma in the garnet zone, and 320–300 Ma in the staurolite and sillimanite + K-feldspar zones. Results highlight a close link between argon isotope records in white mica and both textures and structure-forming major elements. The oldest ages were detected in samples where the earlier syn-D1 white mica generation did not texturally and chemically re-equilibrate at upper-crustal levels. This study suggests that the white micas that escaped recrystallization retain argon isotope records of an earlier metamorphic stage that survived a later event characterized by temperatures higher than 500°C. KEY WORDS: K–Ar systematics; 40Ar–39Ar laserprobe dating; white mica; Barrovian metamorphism; Variscan belt
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDEXPERIMENTAL PROCEDURESRESULTSDISCUSSIONREFERENCES |
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In the last decade, K–Ar mineral ages have widely been interpreted to record cooling below a critical temperature defined as the closure temperature (Tc; Dodson, 1973
). Minerals that developed and/or re-equilibrated at temperatures above their Tc yield K–Ar ages that represent cooling ages. The use of the closure temperature concept, however, implies that temperature was the only parameter controlling the rate of isotope transport during tectono-metamorphic events and that argon diffusion parameters are known, so that for a given grain size and cooling history the appropriate Tc can be calculated. Nevertheless, metamorphic rocks commonly exhibit disequilibrium textures, and each mineral often comprises different generations developed under different P–T conditions and/or different deformational regimes and fluid activities. Regarding white mica, numerous studies have shown that variations in argon isotope age records can be closely linked to microtextural and microchemical variations (e.g. Chopin & Maluski, 1980; Hammerschimdt & Frank, 1991; Hames & Cheney, 1997; Di Vincenzo et al., 2001), thus assigning a major role in the control of argon transport to chemical reactivity, strain and circulation of fluids. These studies have also shown that, in certain circumstances, the argon diffusion rate in white mica may be significantly slower than previously believed. However, most of those studies dealt with rock samples that had experienced very extreme physical conditions and/or were based on bulk 40Ar–39Ar analyses. Other studies highlighted age gradients in micas independent of crystal-chemical variations and compatible with volume diffusion processes (e.g. Hodges et al., 1994). It is therefore important to study the relationship between tectono-metamorphic evolution and argon isotope records in white mica from carefully selected, natural samples with well-defined P–T–deformation paths. The basement of Sardinia represents a nearly complete section of a segment of the Variscan belt comprising low- to high-grade metamorphic rocks. This is a suitable setting in which to study the relationship between tectono-metamorphic evolution and argon isotope records in white micas. This paper reports the results of a detailed microtextural and microchemical study and in situ 40Ar–39Ar laser experiments on metamorphic samples from the garnet zone to the sillimanite + K-feldspar zone of northern Sardinia. The research aimed to examine the relationship between argon age records and both superimposed fabric-forming episodes and chemical variations in white mica.
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GEOLOGICAL BACKGROUND |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDEXPERIMENTAL PROCEDURESRESULTSDISCUSSIONREFERENCES |
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Sardinia and Corsica represent a segment of the Southern European Variscan belt extending from the Axial Zone in Corsica and northern Sardinia to the External Zone in southwestern Sardinia (Carmignani et al., 1994
). The crystalline basement of the Sardinia–Corsica block formed in the Early Carboniferous as a result of the collision between the northern Armorica and southern Gondwana continents, which produced intense deformation and synkinematic regional metamorphism accompanied by widespread igneous activity.
In the metamorphic basement the metamorphic grade increases from very low grade in SW Sardinia to medium and high grade in the northern sector of the island. Polyphase deformation and metamorphism were recorded during the collisional and post-collisional stages (Carmignani et al., 1994, and references therein). The study area (Fig. 1) represents a portion of the medium- to high-grade basement belonging to the Internal Nappes and High-Grade Metamorphic Complex. The southern part of the study area consists of micaschists, porphyroblastic paragneisses, and Ordovician granodioritic orthogneisses and granitic augen gneisses. The Posada–Asinara suture zone (Fig. 1, Posada–Asinara Line, Cappelli et al., 1992) separates the High Grade Metamorphic Complex from the Internal Nappes (Fig. 1). Migmatites and migmatitic gneisses with lenses and bodies of amphibolites, granulites and retrogressed eclogites crop out north of the Posada–Asinara Line (Franceschelli et al., 1982). The progression from low-grade to high-grade metamorphism is considered a prograde evolution caused by collision-related Barrovian metamorphism (Franceschelli et al., 1989).
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Metamorphic basement
Structural data indicate the presence of four main phases of ductile deformation. South of the study area, the D1 deformation phase is characterized by overturned, SW-facing folds and syn- to late-D1 ductile to brittle shear zones with a top-to-SW sense of shear (Franceschelli et al., 1982
; Carmignani et al., 1994). An S1 penetrative axial plane foliation developed during D1 deformation. A heterogeneous mylonitic D2 deformation, related to a transpressional phase (Carosi & Palmeri, 2002), overprinted the D1 structures on a regional scale. The D2 strain increases northward until the main pervasive foliation becomes the S2 foliation. The S2 foliation strikes WNW–ESE and strongly dips SSW. The D2 deformation accompanied much of the exhumation of the low- to medium-grade metamorphic rocks (Carosi & Palmeri, 2002). The main foliation and stretching lineation in the migmatitic complex has the same attitude as that of the mylonitic foliation in the micaschists and gneisses. The D2 deformation is overprinted by two main systems of open to tight folds [F3 and F4 according to Carosi & Palmeri (2002)]. Franceschelli et al. (1982) mapped seven metamorphic zones (from south to north): (1) chlorite; (2) biotite; (3) garnet; (4) staurolite + biotite; (5) kyanite + biotite; (6) sillimanite + muscovite; (7) sillimanite + K-feldspar. Franceschelli et al. (1989) highlighted the diachroneity of mineral growth. The index minerals in the low-grade rocks are aligned along S1. Starting from the garnet zone, the index minerals are predominantly post-D1 and pre- to syn-D2 porphyroblasts. The index minerals (white mica, sillimanite and K-feldspar) in the high-grade zones are aligned along the S2 foliation.
Variscan plutonic rocks
The Variscan Sardinia–Corsica Batholith is one of the largest batholiths in SW Europe. Widespread plutonic activity consisting of two orogenic associations took place during and after the Variscan tectono-metamorphic events (Rossi & Cocherie, 1991): (1) an earlier syntectonic Mg–K calc-alkaline association, cropping out only in northern Corsica and emplaced during peak amphibolite-facies metamorphism; (2) a late- to post-tectonic, high-K calc-alkaline association that postdates the D2 structures. In Sardinia, the high-K calc-alkaline association also includes strongly peraluminous granitoids, which were emplaced at the same time as the high-K calc-alkaline sensu stricto granitoids (Di Vincenzo et al., 1994a, and references therein).
Previous geochronology of the Variscan cycle in the Sardinia–Corsica basement
The cumulative probability distribution diagrams in Fig. 2 display previous geochronological data for the metamorphic basement, for high-K calc-alkaline granitoids from both Sardinia and Corsica, and for Mg–K calc-alkaline granitoids from Corsica. Published data indicate that the emplacement of late- to post-tectonic, high-K, calc-alkaline granitoids in both Corsica and Sardinia occurred from 310 to 280 Myr ago (Fig. 2a). The youngest ages (280–290 Ma) refer to the post-tectonic leucogranites. The earliest syntectonic Mg–K association, instead, was emplaced at 330–345 Ma (Fig. 2a). Rb–Sr data on biotite and muscovite from Sardinian calc-alkaline granitoids yielded ages of 275–295 Ma and 290–305 Ma for biotite and muscovite, respectively (Fig. 2b). The emplacement of the peraluminous Sos Canales pluton, cropping out in the study area (Fig. 1b), was dated to 298 ± 4 Ma by whole-rock Rb–Sr data; biotite and muscovite yielded Rb–Sr ages (mineral– whole-rock pairs) of 285 ± 4 to 299 ± 4 Ma (Di Vincenzo et al., 1994b).
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Figure 2c shows that available muscovite and biotite Rb–Sr and K–Ar ages from the metamorphic basement largely overlap at 300–320 Ma. Most of the biotite ages are younger (280–290 Ma) and overlap with the youngest biotite ages of the high-K calc-alkaline granitoids. Older ages were reported by Del Moro et al. (1991)
for muscovite– whole-rock pairs (ages of 336 ± 8 and 350 ± 16 Ma) and for actinolite 40Ar–39Ar data (age of 345 ± 4 Ma), from micaschists and a metagabbro sampled in the Internal Nappes of NW Sardinia. Figure 2c also reports the Rb–Sr age of 344 ± 7 Ma obtained using six bands from a stromatic migmatite with trondhjemitic leucosomes (Ferrara et al., 1978), which was considered to approximate the Variscan metamorphic climax. |
EXPERIMENTAL PROCEDURES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDEXPERIMENTAL PROCEDURESRESULTSDISCUSSIONREFERENCES |
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Petrographical observations on polished thin sections (derived from the opposite face of the rock chip used for 40Ar–39Ar dating) were carried out using a light microscope and by scanning electron microscopy (SEM) using a Philips XL30 instrument operating at 20 kV and equipped with an X-ray energy dispersive system (EDS). Electron microprobe (EMP) analyses were carried out using a JEOL JX 8600 electron microprobe fitted with four wavelength-dispersive spectrometers, using an accelerating voltage of 15 kV and sample current of 10 nA. Natural standards were used for calibration.
Rock chips 8·5 mm in diameter were drilled from polished sections (300 µm thick) for in situ 40Ar–39Ar laserprobe analysis. They were irradiated for 30 h in the TRIGA reactor at ENEA-Casaccia (Rome). The neutron flux was monitored using the standard MMhb-1 (age 520·4 Ma; Samson & Alexander, 1987). In situ laser analyses were carried out along the length of the mica grains (i.e. parallel to the basal cleavage) using a CW diode-pumped Nd–YAG infrared (IR) laser (mainly for biotite) connected to an external computer-controlled shutter and a pulsed Nd–YAG ultraviolet (UV) laser (frequency quadrupled and Q-switched). Although this orientation differs from that commonly used in diffusion studies (i.e. perpendicular to the cleavage plane), it allows the direct comparison between argon data and microstructural information. Single short pulses (5–7 ms) of the IR laser (operating at 12 W) were focused to produce circular melt pits 100 µm in diameter. The UV laser, operating at 20 Hz and 0·5–1 mJ per pulse, was focused to 10 µm and repeatedly rastered by a computer-controlled x–y stage over areas of 0·0036–0·010 mm2, to produce pits 50–100 µm deep. The sample was observed by a CCD camera coaxial with the laser beam. After cleanup (10–15 min for IR and UV analyses, respectively), extracted gases were equilibrated via automated valves with an MAP215-50 mass spectrometer fitted with a Balzers SEV217 secondary electron multiplier. Blanks were analysed every two to four analyses. The interference factors, measured on phases of Ca and K, were: (36Ar/37Ar)Ca = 0·00029; (39Ar/37Ar)Ca = 0·00069; (40Ar/39Ar)K = 0·0076; (38Ar/39Ar)K = 0·013. The Ca/K and Cl/K ratios were calculated following the procedure given by Di Vincenzo & Palmeri (2001). After 40Ar–39Ar analyses, the location of the analysed pits in the irradiated rock chips of the finer-grained samples (micaschists) was checked by SEM.
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RESULTS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDEXPERIMENTAL PROCEDURESRESULTSDISCUSSIONREFERENCES |
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Studied samples
Samples (location in Fig. 1b) are metasedimentary rocks and orthogneisses from the garnet zone (C3, C15, GFS280), the staurolite + biotite zone (C6), and the sillimanite + K-feldspar zone (RS2BA, RS5A, GFS294).
Sample C15 is a porphyroblastic micaschist collected at the beginning of the garnet zone. It consists of millimetre-sized albite porphyroblasts showing evidence of post-D1 and pre-D2 growth and including an internal foliation (S1) defined by inclusion trails of white mica, quartz, garnet and minor biotite. The external foliation envelops porphyroblasts and is an advanced, discrete, S2 crenulation cleavage dominated by white mica and minor chlorite and biotite. Garnet, oligoclase and opaque minerals are also found along S2. Microlithon relics of the S1 foliation and F1 fold hinges are preserved within S2. Sample C3 is a micaschist collected 5 km north of sample C15. Although its petrographic and structural features are comparable with those of sample C15, it is characterized by a better-developed D2 deformation. GFS280 is an augen gneiss collected 2 km from the contact with the Sos Canales Intrusion (Fig. 1b). The orthogneiss consists of K-feldspar, quartz and plagioclase with a crenulated foliation (S2) defined by biotite and white mica; zircon, apatite and opaque minerals are accessory phases. Chlorite is a common secondary phase after biotite.
Sample C6 is a micaschist with pre-D2 millimetre-sized garnet and staurolite porphyroblasts. It was collected at the start of the staurolite zone (Fig. 1b), where D2 strain was high. It is characterized by a well-developed D2 shear-band cleavage, a significant coarsening of grain size in the matrix and by the disappearance of primary chlorite. White mica is more abundant than biotite. Biotite locally shows partial retrogression to chlorite, staurolite to white mica, and garnet to a fine-grained association of micas.
Samples RS5A and GFS294 are orthogneisses consisting of K-feldspar, quartz and plagioclase, with biotite and white mica aligned along the main foliation (S2); apatite, zircon and opaque minerals are accessory phases. Biotite is locally chloritized in RS5A, whereas it seems pristine in GFS294. Sample RS2BA is a metasedimentary stromatic migmatite; mesosomes consist of biotite, quartz, plagioclase, K-feldspar and white mica, with accessory zircon, apatite and opaque minerals, whereas leucosomes consist of K-feldspar, plagioclase and quartz, minor biotite, small garnet grains and apatite. Mineral phases in the mesosomes grew along the main foliation (S2). Biotite is partially replaced by chlorite.
Microtextural and microchemical data
Samples C3 and C15 contain different texturally resolvable generations of potassic white mica: (1) relics of the S1 foliation; (2) slightly crenulated, fine-grained flakes along the S2 foliation and enveloping porphyroblasts (Fig. 3a–c). Rare, undeformed, later flakes cross-cut the main foliation or occur with biotite as degradation products of garnet. Syn-D1 white micas occur: (1a) within albite porphyroblasts (Fig. 3a); (1b) in microlithons where white micas along with biotite–chlorite constitute the hinges of microfolds enveloped by syn-D2 micas (Fig. 3b); (1c) ‘coarse-grained’ relics enveloped by mica sheaves aligned along S2 and consisting of finer-grained white micas with biotite–chlorite interlayers of variable thickness (Figs 3b and c, and 4a). Type-1 white micas are invariably celadonite-rich and paragonite-poor, and show no significant chemical zoning (Table 1, Fig. 4). Low-celadonite and high-paragonite compositions were detected along with subordinate high-celadonite and low-paragonite contents in white micas on the S2 foliation (Table 1). X-ray maps of the microtextural type-2 white micas (Fig. 4c) show that the celadonite content is locally heterogeneously distributed at the microscopic scale. White micas in both samples exhibit nearly continuous variations, from earlier high-celadonite and low-paragonite to later low-celadonite and high-paragonite compositions (Fig. 5). White micas along S2 in sample C3, when compared with those of C15, are characterized by lower celadonite and higher paragonite contents (Fig. 5b). Microtextural and mineral features in these samples suggest that white mica was involved in different continuous metamorphic reactions. The AKFM reaction (Thompson, 1976) chlorite + K-white mica + quartz = garnet + biotite + H2O (R1) can be regarded as responsible for the observed paragenesis in both syn-D1 and syn-D2 assemblage. In addition, type-1 and type-2 white micas are thought to be the result of a reaction in which celadonite-rich white mica is a reactant and celadonite-poor white mica a product (Thompson, 1976): cel-rich K-white mica + Fe-rich chlorite = cel-poor K-white mica + Mg-rich chlorite + biotite (R2). The presence of oligoclase in the syn-D2 assemblage after the albite porphyroblasts and the decrease of the grossular content from core to rim of syn-D2 garnets (Carosi & Palmeri, 2002) suggest the involvement of the reaction (Crawford, 1966) Ca-rich garnet + albite + K-white mica + chlorite + quartz = Ca-poor garnet + oligoclase + biotite (R3). Although microtextural variations in biotite are similar to those in white mica, biotite shows no significant chemical variations (Fig. 6). A few data points for C3 biotite show higher AlIV and XFe, and lower K, and are attributed to the possible occurrence of interlayered chlorite. White mica in sample GFS280 occurs as: (1) rare, deformed, small flakes enveloped by the main foliation, or within biotitic glomeroblastic associations suggesting a previous crystallization with respect to the micas along the main foliation; (2) millimetre-sized flakes along the main foliation; (3) rare, later flakes showing no preferred orientation. Celadonite-rich compositions were rarely detected within the glomeroblastic associations, whereas lower celadonite contents characterize the other microtextural occurrences. White mica data for sample GFS280 define a nearly flat trend in Fig. 5b, with paragonitic contents significantly lower than those of micaschists C3 and C15. Ti contents (Fig. 5c) increase with decreasing celadonite contents and are, at intermediate celadonite values, higher than those of white micas from micaschists C3 and C15. Again, the biotite composition is uniform (Fig. 6).
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Micaschist C6 (staurolite + biotite zone) also shows white mica characterized by different microtextural features: (1) rare and tiny earlier flakes within porphyroblasts; (2) sheaves aligned along the main foliation; (3) later flakes cross-cutting the main foliation and/or replacing staurolite and (with biotite) garnet. In spite of the different microtextural occurrences, EMP analyses reveal only the presence of celadonite-poor white micas in this metamorphic zone (Fig. 5). Celadonite contents are comparable with those of white micas growing along the main S2 foliation of the lower-grade zone. As shown in Fig. 5b, however, the paragonite content of white micas in sample C6 (XNa up to 0·30) is higher than that of white micas in micaschists C3 and C15 (XNa up to 0·25). In this metamorphic zone, the growth of staurolite coexisting with biotite and idioblastic garnet and the absence of primary chlorite suggest that white mica was involved in the discontinuous reaction of the KFMASH system (Hoschek, 1969
): chlorite + K-white mica = staurolite + biotite + quartz + H2O (R4). Moreover, the decrease of the grossular content from core to rim of porphyroblastic garnets and the occurrence of only celadonite-poor white micas suggest that reactions (R3) and (R2) were still active. Biotite microtextural variations are comparable with those of white mica. Biotite locally occurs as very tiny flakes within white mica sheaves (Fig. 3d). Lower XFe compositions at nearly constant AlIV characterize some biotite flakes enclosed in staurolite.
White mica and biotite in samples from the sillimanite + K-feldspar zone (RS2, RS5, GFS294) mostly grow along the S2 foliation (Fig. 3e and f). Minor flakes cross-cut the main foliation or replace feldspars and sillimanite (white mica) or garnet (white mica + biotite). EMP analyses reveal that white micas are celadonite- and paragonite-poor (<10% paragonite) but contain more Ti (>0·025 up to 0·07 a.p.f.u.; Fig. 5c) than lower-grade samples. It should be noted that chemical data are much more homogeneous for sample GFS294 than for RS2BA or RS5B (Fig. 5). Slight but systematic variations in Ti contents were observed within single grains of RS5A white mica, with TiO2 ranging from 0·8 to 1·1 wt % from core to rim (Table 1). The sillimanite + K-feldspar zone is marked by the disappearance of prograde white mica and by the generation of stromatic migmatites. White mica was therefore involved in the dehydration reaction of the KFMASH system (Le Breton & Thompson, 1988): K-white mica + albite + quartz = K-feldspar + sillimanite + melt. This reaction accounts for both the disappearance of white mica during the prograde evolution and the growth of muscovite during cooling. Biotite compositions in all samples are nearly homogeneous (Fig. 6).
Pressure and temperature estimates and P–T path
Different calibrations of the garnet–biotite (Fe–Mg exchange) thermometer were used to constrain the syn-D1 and the syn-D2 thermal conditions of the garnet zone (Table 2). Temperatures were evaluated using different garnet–biotite pairs enclosed in albite porphyroblasts, with the assumption of equilibrium between garnet and biotite inclusions, or growing along the main S2 foliation. The calibrations based on the experimental dataset (i.e. Hodges & Spear, 1982; Perchuck & Lavrent'eva, 1983) gave comparable results, whereas calibrations (i.e. Ganguly & Saxena, 1984; Kleeman & Reinhardt, 1994) based on a thermodynamic dataset yielded slightly lower and higher values than those obtained using the former (Table 2). Results yielded temperatures of 497–544°C for the syn-D1 couples and of 521–560°C for the syn-D2 phases, and suggest a slight increment of temperature from the D1 to D2 phase. Temperature estimations are remarkably consistent with previous determinations on the same metamorphic zone by Franceschelli et al. (1989). A minimum pressure of 1·0 GPa can be deduced for the D1 stage based on the celadonite content in white mica relics (Massonne & Schreyer, 1987). The GFS280 orthogneiss, which belongs to the same metamorphic zone and contains a limiting assemblage consisting of phengite, K-feldspar, biotite and quartz, yielded a pressure of 1·1 GPa for the rare syn-D1 relics. Pressures of 0·7–0·9 GPa were calculated for the syn-D2 peak using different calibrations of the garnet–biotite–muscovite– oligoclase barometer (Table 2). The staurolite + biotite zone (Table 2) is characterized by peak temperatures of 588–624°C (garnet rim, matrix biotite) and pressures of 0·6–0·9 GPa (garnet rim, matrix biotite, muscovite, oligoclase).
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Pressure and temperature estimates, along with petrographical constraints, were used to reconstruct the P–T paths of the studied rock samples from the garnet and staurolite + biotite zones (Fig. 7). The reconstructed trajectories in Fig. 7 are also compared with those derived for the higher-grade zones by Franceschelli et al. (1989)
. After the D1 deformation, the path of the lower-grade samples describes a stage with rising temperatures and decreasing pressures, followed by a stage with falling temperatures and pressures. Samples from the staurolite zone to sillimanite zone record only the retrogressive path after the thermal climax.
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40Ar–39Ar dating
40Ar–39Ar investigations concentrated on white mica and, when resolvable, on biotite. The IR laser was mainly used to analyse biotite, whereas the UV laser was used to obtain most of the data on white mica because of the possible poor absorption of IR radiation in white mica. Unfortunately, biotite was not identified within all the analysed rock chips of samples from the lower-grade zones. Biotite in these samples is a minor phase commonly interlayered with chlorite. In situ IR spot-fusion data and UV laser-ablation analyses are listed in Table 3 and presented in the frequency diagrams of Figs 8 and 9 (white mica and biotite, respectively). In addition, data for selected areas are reported in Fig. 10 to illustrate the intra-sample distribution of argon in analysed micaschists. Table 3 also lists the Ca/K and Cl/K ratios from neutron-produced 39ArK, 38ArCl and 37ArCa. Because of the low Cl and Ca contents of micas and the small lasered sample volume, 38ArCl and 37ArCa intensities were in most instances close to the detection limits. Ca/K and Cl/K ratios derived from argon isotopes are therefore affected by large uncertainties.
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and include the uncertainty in the J value.
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40Ar–39Ar apparent ages for white mica of sample C15 (garnet zone) range widely, mainly from
320 Ma to 338 Ma (Fig. 8). The oldest ages were detected in syn-D1 white mica within an albite porphyroblast and in microlithons (Fig. 10a). Analyses along the main S2 foliation show a nearly continuous variation from 320·2 ± 2·4 Ma to 333·5 ± 1·8 Ma (Fig. 10 and Table 3). White micas in sample C3 display a comparable continuous age variation, but with the lower limit shifted toward slightly younger apparent ages (315·1 ± 2·6 to 336·3 ± 3·0 Ma). Ages younger than 315 Ma come from areas where white mica is contaminated by biotite–chlorite interlayering. Younger ages, not yet well-defined, were also detected for biotite–chlorite interlayering along late shear bands (Fig. 10b). Two IR spot-analyses on biotite yielded ages of 319·4 ± 9·2 Ma and 299 ± 11 Ma. The oldest white mica ages were detected in syn-D1 white mica enclosed in albite porphyroblasts (Fig. 10b and Table 3) or the inner portions of mica sheaves (Fig. 10b and c). In contrast, orthogneiss GFS280, collected within the same metamorphic zone but not far from the contact with an 300 Ma pluton, yielded significantly younger apparent ages mainly ranging from 290 to 310 Ma (Fig. 8). It is worth noting that in this sample most of the white mica ages cluster at 300 Ma. One analysis carried out within a glomeroblastic association yielded an older age (323·8 ± 4·2 Ma). Biotite apparent ages range from 276·2 ± 4·1 Ma to 292·3 ± 5·5 Ma (Fig. 9), with a weak positive correlation between age and radiogenic argon content (Table 3).
The white mica in sample C6 (staurolite + biotite zone) yielded argon ages mainly ranging from 310 to 320 Ma (Fig. 8). Figure 10 shows that a millimetre-sized syn-D2 white mica flake exhibits a homogeneous intragrain argon distribution; the weighted average of 10 UV analyses from a single grain yielded an age of 317·4 ± 1·4 Ma and an MSWD of 0·43. The homogeneous spatial pattern at the single-grain scale argues against the importance of argon loss by solid-state diffusion. Younger ages (300–310 Ma) were detected where white mica is replacing staurolite or, intergrown with biotite, garnet (Fig. 10d).
Samples RS2BA and RS5A (sillimanite + K-feldspar zone) yielded comparable white mica ages of 300 to 320 Ma (Fig. 8). In sample RS2BA there is no clear core–rim relationship within single flakes, whereas in RS5A the youngest ages were observed along the rims and the oldest in the inner portions (Table 3). Biotite ages in both samples range widely from 240 Ma to 305–310 Ma (Fig. 9). Ages younger than 290 Ma are the least radiogenic (Table 3) and come from areas where biotite exhibits pronounced parting along the basal cleavage and/or appears variably chloritized. Sample GFS294 yields a narrow range of white mica apparent ages (300–310 Ma; Fig. 8). Eleven 40Ar–39Ar analyses yield a weighted mean of 304·6 ± 1·6 Ma and an MSWD of 2·3, which attests to only a slight excess of scatter. The age interval of the GFS294 biotite (290–310 Ma) is narrower than that of RS2BA and RS5A biotite (Fig. 9). Most biotite ages overlap with those of white mica.
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDEXPERIMENTAL PROCEDURESRESULTSDISCUSSIONREFERENCES |
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Tectono-metamorphic evolution and microstructural–microchemical variations in white mica
It is widely acknowledged that white micas from metamorphic rocks of greenschist to lower amphibolite facies may not always be in equilibrium with each other or with other constituent minerals of the rocks. White mica may preserve chemical zoning, and/or single rock samples may contain diachronous generations of white mica with distinct chemical compositions. The preservation of multiple generations of white mica is well documented within mafic and felsic single rock samples of high-pressure (HP) and ultra-high-pressure (UHP) metamorphic units. During recent decades, numerous studies have shown that phengites, relics of the HP or UHP stages, may survive subsequent retrogression at upper-crustal levels at temperatures of 450–650°C (e.g. Heinrich, 1982
; Saliot & Velde, 1982; Klemd et al., 1991; Nowlan et al., 2000). Other studies have similarly shown that pre-metamorphic igneous relics (Frey et al., 1976) and detrital features (Dietrich, 1983) in white micas survived metamorphism at 450–550°C. HP and UHP metamorphic rocks are peculiar systems, which experienced very extreme physical conditions characterized by limited mass transfer and exceedingly sluggish reaction kinetics. However, multiple generations of white micas have also been preserved in metamorphic rocks that experienced less extreme conditions. In a detailed investigation of white micas in the metasediments of the Dalradian block of the SE Scottish Highlands, Dempster (1992) reported a large compositional range in micas within single samples. Compositional variations were strongly controlled by the structural age of white mica. The studied micas, much like our samples, showed a marked reduction in chemical variability starting from the staurolite zone. Dempster (1992) concluded that diffusion-controlled exchange reactions in the studied white micas had a relatively high ‘closure temperature’ (higher than 500°C) below which the exchange of structure-forming major elements was negligible; most of the chemical variations preserved in white mica may reflect equilibria controlled by deformation recrystallization.
Microtextural and microchemical data from our study indicate that samples from the garnet zone contain diachronous white mica generations that refer to different P–T–deformation stages of tectono-metamorphic evolution. The preservation of different compositional features in white micas with distinct microstructural ages indicates disequilibrium and suggests that the later syn-D2 generation of white micas developed under physical conditions that did not allow major-element exchange reactions. The preservation of high-celadonite and low-paragonite compositions, and the lack of significant chemical zoning in white mica relics enclosed in syn-D2 mica sheaves suggest that chemical exchange between white micas, even those in direct contact, was very limited. Estimated P–T conditions suggest a slight rise in temperature and a decrease in pressure from the D1 to D2 deformational events. This P–T evolution, typical of thermal relaxation after a homogeneous thickening stage (England & Thompson, 1984), agrees with the following petrological constraints: (1) both syn-D1 garnets (within albite porphyroblasts) and syn-D2 garnets in micaschists display growth zoning; (2) the grossular content in syn-D2 garnets decreases from core to rim (Carosi & Palmeri, 2002) and plagioclase compositions change from albite in the porphyroblasts to oligoclase in the syn-D2 matrix (i.e. the grossular/anorthite ratio decreases; Spear et al., 1991). Besides the absolute rise in temperature from the D1 to D2 deformational phases, it is important to note that rock samples from the garnet zone, at least starting from the baric peak, evolved entirely in the stability field of almandine-rich garnet coexisting with chlorite and biotite. In summary, petrological data suggest that the early formed white micas (syn-D1) escaped re-equilibration during the thermal peak (syn-D2). This implies that the effective critical temperature for the celadonite and paragonite exchange in studied white micas exceeded 500°C, in agreement with estimates for other areas with comparable P–T regimes. The exchange of major elements, however, does not simply depend on a fixed temperature, but probably results from the complex interplay between extrinsic (e.g. strain, circulation of fluids, thermal history) and intrinsic (e.g. rock and mineral composition, grain size) variables, which may vary from sample to sample. In this respect it is important to note that orthogneiss GFS280, from the garnet zone, preserves only rare white mica relics compatible with the medium-pressure D1 event. White mica in this sample, in contrast to that in micaschists, mainly re-equilibrated at upper-crustal levels. Additionally, lower paragonite and higher Ti contents differentiate the main low-Si population observed in this sample from the white micas along the main S2 foliation in micaschists C3 and C15. These findings suggest a lithological control (i.e. quartz–feldspathic vs pelitic compositions) or, taking into account the diffuse alteration of this sample, the intervention of additional processes that gave rise to a somewhat different T–fluid regime after the D2 deformational phase. The closeness of the orthogneiss to a 300 Ma granite may be an important factor.
In the staurolite + biotite zone, microtextural features such as the increase in quartz inclusions from the core to rim of garnet and the growth zoning of almandine (increase in pyrope and decrease in both spessartine and grossular from core to rim; Carosi & Palmeri, 2002) suggest that syn-D2 garnet grew during a rise in temperature and decrease in pressure (Crawford, 1977; Tracy, 1982). The presence of staurolite coexisting with biotite suggests temperatures above 550°C at 0·8 GPa (Fig. 7), and the preservation of zoning in garnet indicates a peak temperature below 650°C (Tracy, 1982). This temperature interval agrees with our estimates of 590–625°C. Under these metamorphic conditions white micas do not retain any chemical record of the earlier D1 stage. The effective critical temperature for cation exchange was therefore overstepped in this metamorphic zone. Based on the above arguments, the critical temperature for the exchange of structure-forming major elements in the studied white mica was >500°C to <570°C, assuming that the remaining extrinsic variables (strain and fluid activity) were nearly constant from the garnet zone to the staurolite + biotite zone.
Samples from the sillimanite + K-feldspar zone were affected by higher temperatures (T up to 750°C; Palmeri, 1993) at pressures of 0·6–0·7 GPa, under which anatexis processes developed. The high-grade conditions were sufficient to totally recrystallize earlier white micas. White micas from this metamorphic zone are chemically distinguishable from those of the lower-grade zone by their rather uniformly low paragonite and celadonite compositions, and high Ti contents irrespective of the rock type.
Time constraints from geological and previous geochronological data
Geological and stratigraphical data indicate that the Variscan continental collision in Sardinia was preceded by a passive margin stage (Carmignani et al., 1994). The maximum age for the beginning of the collisional stage can be constrained by the upper age of the sediments deposited onto the passive margin and then deformed by the Variscan tectono-metamorphic events. According to Carmignani et al. (1994), the presence of Upper Devonian and Lower Carboniferous (Tournaisian) sediments in the External Nappes of southern Sardinia dates the collisional stage to 345 Ma. However, taking into account that only the Upper Devonian (Middle Famennian) is palaeontologically documented in the marbles of the Internal Nappes (Pili & Saba, 1975; G. Bagnoli, personal communication, 2003), the Variscan continental collision and concomitant Barrovian metamorphism cannot have commenced in north–central Sardinia earlier than 360 Ma. Furthermore, taking into account the time required to change from a passive margin setting to a collisional one and the theoretical time lag (20 Myr) before elevation of a thickened crust commences (Thompson et al., 1997), the above geological constraint agrees with the actinolite 40Ar–39Ar age of 345 ± 4 Ma (Del Moro et al., 1991) from a metagabbro and with the Rb–Sr whole-rock age of 344 ± 7 Ma (Ferrara et al., 1978) obtained from a banded migmatite with trondhjemitic leucosomes. In interpreting the meaning of the latter 345 Ma age, one should take into account that the composition of the analysed migmatite indicates that it underwent metamorphic differentiation (Carmignani et al., 1992). This would imply that the metamorphic stage recorded by this sample preceded the melting reaction involving the breakdown of white mica. Alternatively, assuming that trondhjemitic leucosomes formed through water-fluxed melting, one should take into account that they represent the earliest appearance of melt during a prograde metamorphism of a thickening orogenic belt (Patiño Douce & Harris, 1998). The 345 Ma age could therefore be close to the collisional stage or represent the beginning of exhumation. This interpretation agrees with the conclusion of Pin & Paquette (2002), who inferred that the Early Carboniferous was the oldest limit for the beginning of the Variscan continental collision in the Massif Central (France). In summary, geological constraints and available geochronological data suggest the following temporal evolution for the study area: (1) the maximum age for the beginning of the collisional stage in the Internal Nappes is <360 Ma (345 Ma, age of the thickening stage or beginning of exhumation in the High-Grade Metamorphic Complex); (2) the second intrusive cycle of late- and post-tectonic granitoids took place at 310–290 Ma and plutonism ended at 280 Ma; (3) the closure ages of muscovite and biotite from intrusive rocks are 305–290 Ma and 295–275 Ma, respectively. These data can be further used to temporally constrain the development of deformational events: both the D1 and D2 phases took place later than 360 Ma but earlier than the emplacement of the widespread late-tectonic plutonism, which postdates the D2 structures, i.e. earlier than 310 Ma.
Interpretation of 40Ar–39Ar data
The key issue involved in the interpretation of 40Ar–39Ar mineral data is whether apparent ages (1) reflect crystallization ages or (2) reflect cooling below the Tc, or (3) are meaningless as a result of incorporation of excess argon. Regarding points (1) and (2), there is still no general consensus on the argon-retentive properties of white mica and, more in general, on the effective role of thermal activated solid-state diffusion or recrystallization processes in controlling argon transport. Studies on eclogite-facies metamorphic rocks have shown that celadonite-rich white micas, relics of the HP stage, may yield 40Ar–39Ar ages that overlap with those obtained by more retentive dating systems (e.g. Schmädicke et al., 1995; Di Vincenzo et al., 2001; Jahn et al., 2001; compilation by Glodny et al., 2002). This evidence has been indifferently used to either affirm high argon-retentive properties of white micas [following hypothesis (1)] or infer a very rapid cooling rate [following hypothesis (2)]. Although evidence of argon transport controlled by volume diffusion processes, namely, argon isotope gradients independent of crystal-chemical variations and compatible with Fickian behaviour, has been inferred for white micas from various areas (e.g. Hodges et al., 1994; Reddy et al., 1996), other studies have highlighted a close link between the mobility of radiogenic argon and that of structure-forming major elements. Temperature was not considered the main factor controlling the rate of argon transport in these natural samples; argon mobility was instead attributed to mineral reactivity (Chopin & Maluski, 1980; Hammerschimdt & Frank, 1991; Hames & Cheney, 1997; Di Vincenzo et al., 2001). Most importantly, those studies documented incomplete age resetting for white micas, which experienced metamorphic overprinting at temperatures of >425–450°C and <600°C; that is, above the commonly quoted Tc (350–400°C).
Bulk theoretical Tc for a given grain size, mineral geometry and cooling rate may be derived using Dodson's formula (Dodson, 1973), provided that the appropriate diffusion parameters are known. However, experimental diffusion data on white micas are scarce and the possible compositional effect, inferred from crystal-chemical data (Dahl 1996), has not yet been experimentally investigated. It is important to recall that the closure temperature concept assumes that: (1) migration of an isotope (solute atom) occurs by thermally activated volume diffusion through a homogeneous crystal structure (i.e. isotope transport obeys Fick's second law); (2) diffusing isotopes are efficiently removed at the grain boundary (i.e. the concentration of that isotope at the grain boundary is always zero). Complications in the application of the closure temperature concept to natural samples may arise from intrinsic variables such as mineralogical complexities, even at the scale of a single grain (i.e. coexistence of diachronous multiple generations), which produce a temporally and spatially inhomogeneous crystal structure. Additionally, externally controlled variables may also represent potential pitfalls, namely, complicated thermal histories deviating from linear cooling, heterogeneous repartitioning of strain, and fluid-absent or fluid-present conditions, which may all influence intrinsic variables. Interpretation of 40Ar–39Ar mineral data should therefore be sustained by detailed microscale investigations of structure, texture and chemistry.
40Ar–39Ar laser ages of white micas from this study show large inter-sample and intra-sample variations that correlate with metamorphic grade and microtextural and microchemical features. In particular, micaschists from the garnet zone, which preserve relict syn-D1 celadonite-rich white micas, retain the oldest 40Ar–39Ar ages. Ages as old as 335–340 Ma fall within the age interval defined by different dating techniques and are compatible with independent geological constraints. This argues against the importance of excess argon in the studied rocks. Furthermore, ages of 335–340 Ma are geologically compatible with the time of maximum thickening in the Internal Nappes and may closely approximate the white mica crystallization age during the D1 phase. Conversely, only ages 
320 Ma were detected in samples that lack syn-D1 white mica relics. It is important to note that the age clustering at 315–320 Ma for most syn-D2 white mica in micaschists from the garnet zone and staurolite zone is older than the emplacement age of calc-alkaline granitoids and the closure ages of both muscovite and biotite in these granitoids (Fig. 2). This finding, along with the observation that syn-D1 white mica began to retain radiogenic argon much earlier than 320 Ma and the inter-sample consistency of 40Ar–39Ar ages, suggests that the 315–320 Ma interval most probably represents the end of the chemical re-equilibration (including neo-crystallization) of white mica at upper-crustal levels during the D2 phase. A thorough comparison of EMP analyses with argon data for all the investigated samples reveals that 40Ar–39Ar laser ages are remarkably consistent with microtextural and microchemical data: (1) micaschists C3 and C15 from the garnet zone, which preserve distinct, early-formed, high-celadonite populations (Fig. 11), also retain old ages; (2) when compared with C15, sample C3 shows the mode of the low-celadonite population at lower Si contents (Fig. 11) and yields younger ages (Fig. 8); (3) white micas in orthogneiss GFS280 (garnet zone) contain only rare phengite relics and mainly yield argon ages younger than 320 Ma; (4) in the sillimanite + K-feldspar zone, white micas from samples RS2BA and RS5A display a larger chemical variability and wider age interval than sample GFS294 (Figs 8 and 11). There are some intra-sample inconsistencies between chemical data and argon ages derived from samples in the garnet zone. The bimodal distribution of celadonite content in micaschists C15 and C3 (Fig. 11) is not reflected in the age frequency, which conversely exhibits a roughly single mode with a pronounced spread at 320–330 Ma (Fig. 8). Orthogneiss GFS280 preserved rare celadonite-rich relics but did not retain ages as old as 335–340 Ma. Such discrepancies may arise from the significant difference between the spatial resolution of EMP data and that of 40Ar–39Ar laser analyses (2 x 10–5 mm2 and a few micrometres depth vs 
3·6 x 10–3 mm2 and several tens of micrometres depth). Complex small-scale mineralogical heterogeneities, such as white micas aligned along the S2 foliation with heterogeneous celadonite content or biotite– chlorite interlayers, may have been averaged by the argon laserprobe.
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Argon data from biotite also yielded a large time span (>
20 Ma), which is not paralleled by significant compositional variations. In all investigated samples, biotite ages are generally younger than those of white mica from the same sample. This observation agrees with the widely acknowledged lower retentive properties of biotite with respect to those of coexisting white mica, sustained both by experimental data and on theoretical grounds (Dahl, 1996). However, as stressed by Dahl (1996), this pattern may arise irrespective of isotope loss mechanism (diffusion, recrystallization, alteration, etc.). Although it is difficult to draw straightforward conclusions because of the lack of biotite argon data along the whole north–south transect, biotite ages seem to be very sensitive to secondary alteration processes. This interpretation is supported by the observation that younger ages mainly come from areas where biotite is visibly chloritized or characterized by pronounced parting along the basal cleavage, and by the negative correlation between apparent ages and atmospheric argon contents (Fig. 12). This finding agrees with the outcome of other studies (e.g. Roberts et al., 2001; Di Vincenzo et al., 2003), which demonstrated that even incipiently altered biotite may yield younger apparent ages with correspondingly higher atmospheric argon contents.
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The most important finding of this study is that microtexturally and microchemically earlier generations of white micas in micaschists of the garnet zone preserve older ages that can be referred to the D1 deformational event. This is surprising, as the whole tectono-metamorphic evolution, starting from the baric peak, developed at temperatures under which white mica is generally considered open to argon isotope exchange. This means that both texturally and chemically resolvable syn-D1 and syn-D2 white micas should have recorded comparable ages. It is important to note that old ages were detected not only in white mica enclosed in albite porphyroblasts, which could be considered armoured relics, but also within the mica sheaves aligned along the S2 foliation where petrographic investigation revealed the presence of celadonite-rich relics.
This is not the first report of white mica that experienced metamorphic overprinting and escaped argon isotope resetting even when reheating exceeded the commonly quoted Tc. Previous studies, however, dealt with rock samples that experienced very extreme physical conditions (i.e. HP or UHP conditions, Chopin & Maluski, 1980; Giorgis et al., 2000; Di Vincenzo et al., 2001), or metamorphic overprinting under lower-grade conditions (425–450°C, Wijbrans & McDougall, 1986; 425–525°C, Hames & Cheney, 1997), and/or were based on bulk sample analyses (e.g. Wijbrans & McDougall, 1986; Andriessen, 1991; Hammerschmidt & Frank, 1991). Possible explanations for the survival of argon isotope relics in early-formed white mica include: (1) episodic thermal overprinting of short duration (Wijbrans & McDougall, 1986); (2) the grain boundary network was not free of argon (i.e. argon concentration at the grain boundary was >0) because of external buffering (fluid flow advecting argon in solution; Wheeler, 1996) or internal buffering (limited mass transfer and sluggish reaction kinetics owing to extreme physical conditions; Foland, 1979; Kelley & Wartho, 2000); (3) generally limited argon mobility in the absence of chemical re-equilibration processes, i.e. mobility of radiogenic argon was strictly coupled with that of structure-forming major elements, with the implication that petrographic relics ensure isotopic inheritance (Villa, 1998). Whereas short-lived thermal overprinting is clearly not compatible with the tectono-metamorphic evolution of the studied rocks [hypothesis (1)], the latter two hypotheses should be carefully addressed. However, a straightforward discussion of these issues is complicated by the lack of detailed information on argon partitioning between fluids and K-rich minerals and, more in general, on the behaviour of Ar in the crust. The role of additional and concomitant variables, such as composition and pressure, which may have enhanced argon retention (Dahl, 1996) in the studied white micas, appears of secondary importance. Celadonite-rich syn-D1 white mica and celadonite-poor syn-D2 white mica represent a disequilibrium sub-assemblage; texturally and chemically resolvable white micas grew at different times, and argon ages conform to white mica structural ages. We stress that the study area is a natural setting in which petrographic and argon isotope relics are preserved in rock samples that did not experience very extreme physical conditions, but metamorphism of intermediate thermal gradient (Barrovian-type), evolved under lower amphibolite-facies conditions and experienced a protracted tectono-metamorphic evolution. As such, results from the present study may have a broader significance. We conclude that: (1) mineral reactivity (recrystallization) was the most important factor controlling the rate of argon transport in the studied white micas; (2) the assumption that, in the absence of recrystallization, white mica in the studied metamorphic rocks retained radiogenic argon only below the commonly quoted Tc, would have led to a significant underestimation of the actual argon-retentive properties of white mica. Argon retention in syn-D1 white mica from the garnet zone was high enough to survive the syn-D2 thermal climax, which occurred at a temperature above 500°C.
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ACKNOWLEDGEMENTS |
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We are grateful to G. Giorgetti (Dipartimento di Scienze della Terra, Siena) and F. Olmi (IGG-CNR, Firenze) for support during SEM investigation and EMP analyses, respectively. The journal reviews of T. Dempster, K. Hammerschmidt and I. M. Villa, and the editorial handling of K. Bucher are acknowledged. The 40Ar–39Ar laserprobe facility was funded by ‘Programma Nazionale di Ricerche in Antartide’ (PNRA). Research was financially supported by ‘Consiglio Nazionale delle Ricerche’ (CNR).
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FOOTNOTES |
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* Corresponding author. Telephone: +39 050 3152270. Fax: +39 050 3152360. E-mail: g.divincenzo{at}igg.cnr.it
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REFERENCES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDEXPERIMENTAL PROCEDURESRESULTSDISCUSSIONREFERENCES |
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