Mineralogy and Petrology (1998) 62:29-60 Mineralogy Rn(1

Petrology © Springer-Vertag 1998 Printed in Austria

Metamorphogenic barite-pyrite (Pb-Zn-Ag) veins at Pollone, Apuane Alps, Tuscany: vein geometry, geothermobarometry, fluid inclusions and geochemistry

P. Costagliola 1, M. Benvenuti 2, P. Lattanzi 2,3, and G. Tanelli 2

1 Museo di Mineralogia e Litologia dell'UniversitY, Firenze, Italy 2 Dipartimento di Scienze della Terra, Firenze, Italy 3 Dipartimento di Scienze della Terra dell'UniversitY, Cagliari, Italy

With 13 Figures

Received December 12, 1996; revised version accepted December 19, 1997

Summary

The barite-pyrite-(Pb-Zn-Ag) deposit of Pollone is located in the southernmost tip of the Apuane Alps metamorphic core complex, and is hosted by a siliciclastic formation of pre-Norian age. The southern sector of the deposit mainly consists of stratiform, supposedly syngenetic, barite-pyrite orebodies, whereas the northern area is characterized by a barite-pyrite-(Pb-Zn-Ag) vein system. Vein geometry in the northern area is controlled by a shear zone, developed during the greenschist facies metamorphism which affected the Apuane Alps core complex between 27 and 8 Ma, that was responsible for fluid focusing and vein emplacement. At Pollone, arsenopyrite and chlorite geothermometers show broadly comparable results, and suggest local metamorphic peak temperatures between 320 and 350°C. Phengite geobarometry indicates minimum pressures of about 3.5k bar. Fluid inclusion data and mineral equilibria suggest that the mineralizing fluids were initially hotter than the country rocks (about 450°C at 3.5-4.0kbar). Rocks in direct contact with the orebodies are depleted in Rb and enriched in Sr in comparison to similar rocks elsewhere in the area. This is attributed to the presence of Rb-poor muscovite and Sr-rich barite. Rb-depleted muscovites suggest mineral-fluid interaction in a rock reservoir characterized by a different (modal) mineralogical composition than the Pollone host rocks. The progressive decrease of Sr in barite with increasing distance from the orebodies may be explained with a temperature decrease along the infiltration paths of mineralizing fluids (i.e., from the vein into the wall rocks). The similar O-isotope composition of quartz from veins and host rocks is explained with the overall homogeneous O-isotope

30 P. Costagliola et al.

composition of the Alpi Apuane basement rocks. This indicates a limited interaction between mineralizing fluids and the rocks exposed at Pollone. Remobilization of syngenetic orebodies was conceivably of minor importance in the production of metamorphogenic veins. Fluid cooling along a major tectonic lineament is thought to be responsible for barite deposition.

Zusammenfassung

Die metamorphogenen Baryt-Pyrit (Pb-Zn-Ag) Giinge von Pollone, Apuanische Alpen, Toskana: Geometrie der Giinge, Geothermobarometrie, Fliissigkeitseinschliisse und Geochemie

Die Baryt-Pyrit (Pb-Zn-Ag) Lagerstfitte von Pollone liegt im s~idlichsten Ende des metamorphen Kern-Komplexes der Apuanischen Alpen, und sitzt in einer siliziklas- tischen Formation pr~i-Norischen Alters auf. Der stidliche Sektor der Lagerst~itte besteht haupts~ichlich aus stratiformen, wahrscheinlich syngenetischen Baryt-Pyrit-Erzk6rpern, w~ihrend der n6rdliche Teil des Gebietes durch ein Baryt-Pyrit (Pb-Zn-Ag) Gangsystem charakterisiert wird. Die Geometrie der G~nge im Nordteil wird dutch eine Scherzone kontrolliert, die w~ihrend einer grtinschieferfaziellen Metamorphose entstanden ist, die den Kemkomplex der Apuanischen Alpen zwischen 27 und 8 Ma betroffen hat. Diese Scherzone war auch ftir die Zufuhr der Fluide und die Platznahme der G~inge verantwortlich. In Pollone zeigen Arsenopyrit- und Chlorit-Geothermometrie weithin vergleichbare Ergebnisse und weisen auf lokale Maximaltemperaturen der Metamor- phose zwischen 320 und 350 °Chin. Phengit-Geobarometrie lfiBt Minimal-Drucke von ungef~ihr 3,5 kbar erkennen. FluidfltissigkeitseinschluB-Daten und Mineral-Gleichge- wichte zeigen, dab die erzbringenden Fluide ursprtinglich heiger als die Wirtsgesteine waren (ca. 450°C ftir P von 3,5 bis 4kbar). Gesteine, die im direkten Kontakt rnit den Erzk6rpern sind, zeigen eine Abreicherung an Rb und eine Anreicherung an Sr, im Vergleich mit fihnlichen Gesteinen, die im Gebiet anzutreffen sind. Dies wird auf das Vorkommen von Rb-armen Muscovit und Sr-reichen Baryt zurtickgeftihrt. An Rb- abgereicherte Muscovite legen Mineral-Fluid-Reaktionen nahe, die in einem Gesteins- reservoir abliefen, das durch eine andere mineralogische Zusammensetzung als die Wirtsgesteine von Pollone charakterisiert war. Der zunehmende Verlust von Sr im Baryt mit zunehmender Entfernung von den Erzk6rpem, kann durch einen Temperaturabfall entlang der Infitrations-Pfade der erzftihrenden L6sungen erkRirt werden (d.h. yon Gang in die Nebengeseine). Die fihnliche Sauerstoff-Isotopen-Zusammensetzung fiir Quarz aus den G~ingen und den Nebengesteinen RiBt sich auf die allgemein homogene Sauerstoffisotopen-Signatur des Basements der Apuanischen Alpen zuriickftihren. Dies weist auf beschr~nkte Wechselwirkung zwischen erzftihrenden L6sungen und den in Pollone anstehenden Gesteinen hin. Die Remobilisation von syngenetischen Erzk6rpern in Pollone war nut von geringer Bedeutung ftir die Entstehung der metamorphogenen G~inge. Abktihlung der Fluide an einem wichtigen tektonischen Lineament gilt als Ursache ftir den Absatz von Baryt.

Introduction

The Apuane Alps (northwestern Tuscany) have been interpreted as a metamorphic core complex composed of several tectono-metamorphic units subjected to Oligocene-Miocene greenschist-facies metamorphism and multiple episodes of deformation (Carmignani et al., 1987; Carmignani and Kligfield, 1990). Pressure

Metamorphogenic barite-pyrite (Pb-Zn-Ag) veins 31

(P)-temperature (7) metamorphic conditions of the central zone of the complex are fairly well known (P = 3.5 kbar; T= 350-450 °C) but are rather poorly constrained in the southern region, which hosts most of the mineral deposits (see Lattanzi et al., 1994). In this region, the Pollone, Monte Arsiccio and Buca della Vena deposits constitute a barite district, exploited up to 1987, that has been extensively studied (see references in Lattanzi et al., 1994). Most of these deposits show a marked stratabound character, a stratiform near-conformable lens-shaped morphology ("banchi"), banded textures of the orebodies, a distinct mineral zoning that reflects changes of host rock lithology, and a very simple main mineralogy (barite, pyrite and Fe-oxides). These features, as well as ore textures and stable isotope geochemistry, lead to the conclusion that these deposits had a sedimentary- diagenetic origin, whereas minor barite veins were thought to be the results of metamorphic hydrothermal remobilization (Cortecci et al., 1989). As a peculiar feature, the northern area of Pollone is characterized by the widespread presence of economic barite-pyrite-(Pb-Zn-Ag) veins; they are discordant to subconcordant with respect to host rock foliation ("filoni") and suggest that in this area (metamorphic?) hydrothermal circulation was important. Cortecci et al. (1989), essentially on the basis of stable isotope data, suggested a "closed system" circulation at Pollone, whereby the vein bodies were interpreted as the products of metamorphic Ba mobilization from the stratiform ores.

In this paper we present field, textural, fluid inclusion, mineral and rock chemistry data in order (1) to define the P-T conditions of regional metamorphism and of mineral deposition at Pollone, (2) to investigate the relationships between deformation, hydrothermal circulation and ore deposition and (3) to test the reliability of a "closed system" model.

Geological setting The lowermost metamorphic unit of the Alpi Apuane core complex is the so-called "Nucleo Metamorfico Apuano" (NMA), that is tectonically overlain, in the southern Apuane Alps, by the "Fomovolasco-Panie Unit" (FPU; see Ciarapica and Passeri, 1982, for a detailed geological description). According to Carmignani and Kligfield (1990), two major deformation phases affected the metamorphic units: a compressional phase (D1) followed by an extensional phase (D2), dated at 27 and 12-8 Ma, respectively (Kligfield et al., 1986). Carmignani and Kligfield (1990), following the model of Platt (1986), proposed that during the D2 phase several shear zones developed in response to the gravitational collapse of the Apuane Alps core complex. Rocks belonging to different structural levels were juxtaposed along the shear zones during this phase.

The polymetamorphic basements of the NMA and the FPU show many similarities, and are predominantly composed of siliciclastic and volcanic lithologies, probably spanning from Cambrian to Triassic in age (Conti et al., 1991). The Pollone deposit crops out in the S. Anna tectonic window, in the southernmost zone of the core complex (Fig. 1). Similar to the other barite deposits of the district, it is hosted by a dominantly siliciclastic metamorphic formation belonging to the "Fomovolasco-Panie" Unit, usually referred to as "Scisti di Fornovolasco" (Ciarapica and Passeri, 1982). The nature and age of this formation

32 E Costagliola et al.

Valdicastello

,RSICCIO

Valdicastello

10' 1 g'

MINERALIZATION

Ba, Fe Cu, Pb, Ag sulphides

~ SOUTHERN NAPPE

~ GREZZONI ] FORNOVOLASCO PANIE ~ SCISTI DI JUNIT

FORNOVOLASCO

] FORMATIONS METAMORFICO ~ GREZZONI APUANO

m. 500

SW

Pollone Shear Zone

/ NE

Fig. l. Geological map of the S. Anna tectonic window showing the most important mineral deposits. An idealized schematic structural SW-NE cross section, along the trace indicated by the arrows in the map, is given at the bottom; hatched symbol indicates the shear zones (modified after Carmignani and Kligfield, 1990)

Metamorphogenic barite-pyrite (Pb-Zn-Ag) veins 33

are still open to discussion (cf. Cortecci et al., 1985). A Triassic age was suggested by the occurrence of dolomitic levels, lithologically similar to the overlying Norian "Grezzoni" formation (Fig. 1), that would indicate a gradual transition from siliciclastic to carbonatic sedimentation (Ciarapica and Passeri, 1982). On the other hand, Carmignani et al. (1976), on the basis of strong lithological similarities with formations of the Apuane Alps basement, ascribed the Scisti di Fornovolasco to the Palaeozoic.

Carmignani et al. (1975) proposed that the most prominent foliation in the S. Anna Window ($2) is related to the second regional deformational event (D2). D2 is also responsible for the development of a prominent shear-zone that, in particular, affected the northern area of the deposit ("Pollone shear zone"; Fig. 1).

Petrography of the Scisti di Fornovolasco formation

In the S. Anna tectonic window, the Scisti di Fornovolasco formation is composed mainly of siliciclastic rocks containing quartz and muscovite as the main minerals. Quartz is mostly of detrital origin and generally shows rounded shape, subgrain division and undulose extinction. This detrital quartz typically occurs in narrow horizons, whereas neomorphic quartz grains are disseminated more pervasively. Newly-formed quartz is particularly abundant over short distances (generally less than 1 meter) around vein bodies, indicating that silicification took place during the mineralizing event. No other prominent hydrothermal alteration has been observed, apart from moderate replacement of rare plagioclase and K-feldspar by white mica. Muscovite, generally less abundant than quartz, is frequently localized along $1 and, less commonly, $2 surfaces. Pyrite, quite often in euhedral crystals, is the most common accessory mineral. Tourmaline, albite, K-feldspar and, in minor amounts, chlorite are present as isolated crystals in the quartz-muscovite matrix. However, chlorite and tourmaline are locally abundant. Tourmaline associated with quartz locally forms 1-5 cm thick levels of a rock locally described as "tormalinolite" (cf. Benvenuti et al., 1989). Rare euhedral arsenopyrite crystals 1-5ram large were found in the northern Pollone area. Arsenopyrite and pyrite are in textural equilibrium; they form crystal aggregates elongated along the $2 surfaces, suggesting a synkinematic (syn-D2) growth. In the upper part of the Scisti di Fornovolasco, some dolomitic layers, 1-2 to 10-15cm thick, are present (Ciarapica and Zaninetti, 1983). In addition to dolomite and calcite, accessory pyrite, quartz and, locally, tourmaline, are present. The rare occurrence of talc has been reported by Orberger (1985) within these carbonatic layers.

In the study area, the Scisti di Fornovolasco are associated with some rhyodacitic and rhyolitic bodies (Orberger, 1985). Rhyodacites are characterized by abundant chlorite replacing a preexisting mafic phase, probably biotite. Chlorite is frequently associated with euhedral and fine-grained hematite. Quartz is present both as phenocrysts and as microcrystals in the groundmass. Alkali-feldspars occur as minor phases. Locally, rhyodacites may contain minor amounts of tourmaline, pyrite, iron oxides and (rare) calcite. Massive rhyolitic bodies ("porfiroidi"; cf. Carmignani et al., 1976) mainly consist of porphyritic and microcrystalline quartz, and of muscovite. Calcite, tourmaline, K-feldspar and pyrite are present in minor amounts. A lithotype with quite similar mineralogy, "scisti porfirici"

34 R Costagliola et al.

( = porphyritic schists; Carmignani et al., 1976) crops out in the S. Anna tectonic window, and is considered to be the product of subaerial reworking of the rhyolitic bodies. For descriptive purposes, we have distinguished three types of rocks in the S. Anna window, according to their distance from ore bodies: (a) "wall rocks", those in direct contact with orebodies; (b) "host rocks", which crop out in the mining area and (b) "country rocks", located away (>_300 m) from the mineralized area. Wall and host rocks lithologically belong to the Scisti di Fornovolasco formation sensu stricto. Country rocks, in addition to the Scisti di Fornovolasco formation, also include volcanic lithologies.

Vein geometry, mineralogy and textures

The northern sector of the Pollone mine is characterized by veins at an angle with $2 varying between 15 ° (low angle ore bodies: Fig. 2) and 80 ° (high angle orebodies: Fig. 3, 4 and 5), Both low and high angle veins contain mainly barite, quartz, and pyrite, with variable amounts of base-metal sulfides, fluorite and calcite. Barite forms white microcrystals (50 and 200~tm). Pyrite and quartz crystals have similar dimensions and are euhedral to subeuhedral.

Ore bodies, particularly the low angle veins (Fig. 2), may exhibit a banded texture due to the local abundance of sulfide layers (mainly pyrite) which extend along the main direction of the ore bodies. Wall rock alteration is moderate (see above). Quartz is the first phase in the paragenetic sequence. It is followed by pyrite, barite, a second generation of quartz (quartz II) and, locally, fluorite. Pyrite, barite, quartz II and fluorite are in textural equilibrium. However, some exceptions have been noticed for barite and fluorite, which may be replaced by quartz. Where present, calcite is in equilibrium with the other gangue minerals with very limited replacement of fluorite after carbonates. Locally, quartz, barite and pyrite may be

Fig. 2. Low angle ore body composed of barite (white). $2 schistosity is approximately horizontal

Metamorphogenic barite-pyrite (Pb-Zn-Ag) veins 35

Fig. 3. Weakly deformed high angle orebody. The vein has a sigmoidal shape and is folded along the subhorizontal $2 direction

Fig. 4. Strongly deformed high angle orebody. The vein is folded along the $2 direction that is approximately horizontal. The width of the photograph is about 1.5 m

36 R Costagliola et al.

High angle strongly deformed orebodies

//:2___2--: . . . .

Low angle orebodies

Fig. 5. Sketch of vein geometry and foliation orientation at Pollone (northern sector)

brecciated and cemented by a later generation of the same minerals. Sphalerite galena and sulphosalts are in textural equilibrium with pyrite; exceptionally, they may partially replace the iron sulfide.

In the southernmost occurrences (Fig. 1), represented by stratiform ore bodies, barite and pyrite are the prevailing minerals. Northward, the ore bodies become remarkably enriched in Pb, Zn and Ag sulfides and sulfosalts; however, barite- pyrite orebodies also occur at the contact between the Scisti di Fornovolasco and the Grezzoni formations. For a complete list of minerals found at Pollone, the reader is referred to Brizzi and Olmi (1989).

Low angle ore bodies are weakly deformed although, locally, boudin-like structures occur or barite may be fractured and cemented by sulfides. More frequently, low angle ore bodies are boudinaged. High angle ore bodies show a variable degree of deformation: veins may be practically undeformed, or weakly deformed resulting in sigmoidal-shaped bodies (Fig. 3). Increasing deformation leads to tight folding of the ore bodies along an axial plane coincident with $2 direction (Fig. 4). Cavities filled by quartz and barite are sometimes present in high angle veins (Fig. 5). High and low angle veins never crosscut each other, at least in the presently accessible exposures.

At the microscale, the presence of barite veins in the Scisti di Fornovolasco is rare. The most distinctive occurrence is in brittle structures in the "toI~alinolite" layers (such as "tension gashes" and "bookshelf sliding" surfaces - Ramsay and Huber, 1987), that are occupied by quartz-barite-(pyrite) microveins (Fig. 6). Mineral growth was syntaxial; quartz invariably occupies the walls, whereas barite and very minor amounts of pyrite are restricted to the center of the vein (Fig. 6).

No spatial relationships between veins in the northern and southern sector are observed. This is due to two principal reasons: 1) in the southern section veins are

Metamorphogenic barite-pyrite (Pb-Zn-Ag) veins 37

m

w

• _ _ i J ~

Fig. 6. Top: Schematic representation of a "bookshelf sliding" structure in a "tormali- nolite" layer (modifed after Ramsay and Huber, 1987). Bottom: Thin section of a quartz- barite vein hosted by the "tormalinolite" which is composed by an alternation of thin layers of tourmaline (dark gray) and quartz (light gray). The fracture is filled with quartz (transparent on the wall rocks) and barite (light gray in the center). Microvein width is about 1 mm

less developed; 2) from internal reports of EDEM - the company formerly exploiting the deposits - as well as from direct field observations, it appears that probably the most significant example of a contact between a high angle and a stratiform orebody was encountered in underground mining workings. However, this contact was destroyed during exploitation, and is no more accessible today.

Geothermometry and geobarometry

Most estimates of metamorphic pressures and temperatures in the Apuane Alps region refer to the NMA, and fall in the range of 3-4 kbar and 350-450 °C (Di Pisa et al., 1985; Preite Martinez et al., 1978; Carmignani et al., 1978, 1987; Kligfield, 1979). Slightly higher pressures have been proposed by Franceschelli et al. (1992).

38 R Costagliola et al.

Recently, temperatures as high as 470 °C have been calculated from the isotopic fractionation of oxygen in metamorphic terranes belonging to the NMA (Cortecci et al., 1994).

In the UFR the only available P-T data were provided by Orberger et al. (1986), who estimated metamorphic pressures at 2-3 kbar, based on the Si content of muscovite. The same authors report the presence of green biotite, from which they suggest temperatures of 425-460 °C, in the Scisti di Fornovolasco formation. Green biotite, however, was not observed in the course of the present study. In the UFR the general lack of a silicate mineral assemblage that can be used for geothermometry precludes the application of the so-called "phengite geobarom- eter", which actually shows an appreciable temperature dependence (Massonne and Schreyer, 1987). However, empirical calibrations of chlorite composition as a function of temperature (Cathelineau and Nieva, 1985; Cathelineau, 1988) may be used for temperature determinations. The chlorite geothermometer was calibrated in geothermal wells, and de Caritat et al. (1993) have warned against its indiscriminate use (see also Essene and Peacor, 1995; Lu et al., 1996). However, applications in greenschist facies terranes have provided estimates in good agreement with independent petrologic evidence (Brown, 1993). Moreover, at Pollone most chlorite occur in (meta) volcanic rocks having bulk compositions similar to those for which the geothermometer was calibrated.

The arsenopyrite geothermometer is based on variations of the arsenopyrite composition, that evolves toward As-richer terms with increasing temperature and pressure, provided that f(Sa) is buffered by coexisting phases (Sharp et al., 1985). This geothermometer has been frequently applied in metamorphic environments. The results are controversial, but in general appear reliable at low to medium metamorphic grades (Sharp et al., 1985), provided that textural analysis can relate arsenopyrite formation to a specific metamorphic stage. Arsenopyrite composition is weakly influenced by pressure values lower than 5 kbar (Sharp et al., 1985). The arsenopyrite-pyrite assemblage does not allow univocal temperature determination, since the system is divariant; therefore only a maximum temperature is calculated.

Analytical methods

Chlorite, muscovite ~tnd arsenopyrite have been analyzed using an ARL-SEMQ microprobe at the Centro CNR di Studi Geominerari e Mineralurgici, Cagliari. The following analytical conditions have been adopted: accelerating voltage 20kV; sample current 20nA; beam diameter, 5-10 ~tm. The following standards were used: rhodonite for Mn, microcline for A1, Si and K, albite for Na, olivine for Mg, anorthite for Ca, ilmenite for Ti, chromite for Cr and fayalite for Fe. Raw data were reduced by the correction program MAGIC IV. The Asp200 standard of Kretschmar and Scott (1976) was used for As, S and Fe. Pure elements were employed for Co, Ni, Cu and Sb. Microthermometric measurements of fluid inclusions were performed with a U.S.G.S. type gas flow heating-freezing stage. The thermocouple of the stage was calibrated using synthetic fluid inclusions. Measurements were made on double polished sections having a thickness of about 0.3ram. Temperature of homogenization and ice melting were generally reproducible within -+-2 and ±0.3 °C, respectively.

Metamorphogenic barite-pyrite (Pb-Zn-Ag) veins 39

Chlorite and arsenopyrite geothermometers

To estimate metamorphic temperatures, chlorite and arsenopyrite were selected from samples of country and host rocks, i.e. far away from mineralized veins (from some tens to some hundreds of meters). In this way, thermal perturbation induced by infiltrating fluids at higher temperatures (cf. Chamberlain and Rumble, 1988) is avoided. Arsenopyrite, in textural equilibrium with pyrite, crystallized synkine- matically along $2 structures (see above), indicating a syn-metamorphic formation. Chlorite has been sampled from massive rhyolitic and rhyodacitic bodies, where it is mainly present as a pseudomorphic phase, probably after biotite, closely associated with hematite. Chlorite and arsenopyrite compositions are shown in Tables 1 and 2, respectively. In the literature several empirical composition (AlXV)- temperature curves for chlorite (Cathlineau and Nieva, 1985; Cathelineau, 1988) have been proposed, but all do not take into account the presence of Fe 3+ in the structural formula. This may represent an acceptable approximation in some instances, since Fe 3+ contents in chlorite are generally very low (Cathelineau and Nieva, 1985). However, in the assemblage studied here, chlorite is frequently associated with hematite, requiring a correction for Fe 3+. For our samples, both Fe 3+ content and temperatures have been calculated using the program "Clorita" compiled by E Tornos, that is based on the algorithms of Walshe (1986) and Cathelineau (1988). The calculated temperatures fall in the range 300-350 °C, with a frequency peak between 310 and 330 °C (Fig. 3). Slightly higher estimates have been obtained by arsenopyrite geothermometry, that gives (maximum) formation temperatures ranging between 300 and 400 °C (Fig. 7), with most values falling in

Table 1. Electron microprobe analyses of chlorite (expressed as oxides, in wt.%) and calculated cation occupancies on the basis of 14 oxygens

sample sarl SiO2 25.14 24.74

AI203 20.48 20.79

TiO2 0.04 0.1E FeO 24.72 24.34 MnO 0,18 0.27 MgO 16.09 15.75 CaO 0.02 0,01 Na20 0.00 0.00

K20 0.00 0.00

sat6 25.26 25.35 23.79 24.85 24.86 23.87 25.45 25.51

20.95 20.73 19.36 21.09 19.11 21.44 20.45 20.53 0.05 0.14 0.13 0.34 0.07 0.05 0.06 0.08

26.33 25.93 26.12 25.77 25.37 25.96 26.03 25.90 0.21 0.22 0.27 0.22 0.21 0.24 0.20 0.20

14.99 15,15 14,83 15.52 15.71 15.80 14.65 15.72 0.03 0.00 0.01 0.03 0.02 0.02 0.00 0.00 0.00 0.00 0.00 0.00 0,00 0.00 0.00 0.00

0.03 0.04 0.00 0.02 0.05 0.02 0.03 0.06

E23 25.17 24.63 24.50 20.84 20.79 20.28

0.08 0.09 0.05 23.12 23.89 23.63

0.12 0.13 0.17 16.03 16.53 15.04 0.00 0.02 0.00 0.00 0.00 0.00 0.20 0.00 0.02

isar2 24.89 25.03 25.14 23.77 23.74 23.46 20.37 20.68 20.79 22.37 22.27 21.47

0.06 0.07 0.07 0.03 0.06 0.05 25.18 25.36 25.17 27.20 27.26 27.34

0.24 0.22 0.17 0.12 0.12 0.10 15.83 15.24 15.47 13.41 13.48 13.36 0.00 0.00 0.00 0.00 0.01 0.02 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.00 0.02

Tot. 86.66 86.05 87.85 87.56 84.51 87.83 85.40 87.40 86.87 87.99 85.56 86.26 63.68 86.57 86.61 86.83 86.89 86.92 85.83

sar6 E23 sar2 sample sarl Si (=V) 2.68 2.65 AI tot 2.57 2.63

Al(~V) 1.32 1.35

AI (v~) 1.25 1.28 Ti 0.00 0.01 Fe 2.20 2.18 Mn 0.02 0.02 M 9 2.56 2.52

2,67 2.69 2.64 2.63 2.71 2.55 2.72 2.69 2.61 2.59 2.53 2.63 2.45 2.70 2.58 2.55

1.33 1.31 1.36 1.37 1.29 1.45 1.28 1.31

1.29 1.28 1.17 1.26 1.16 1.24 1.30 1.24 0.00 0.01 0.01 0.03 0.01 0.00 0.01 0.01 2.33 2.30 2.42 2.28 2.31 2.32 2.33 2.28 0.02 0.02 0.02 0.02 0.02 0.02 0,02 0.02! 2.37 2.39 2.45 2.45 2.55 2.51 2.33 2.47

2.69 2.65 2.69 2.63 2.61 2.63

1.31 1.35 1.31

1.32 1.26 1.32 0.01 0 .01 0.00 2.07 2.13 2.17 0.01 0.01 0.02 2.56 2.63 2.47

2.66 2.68 2.68 2.56 2.56 2.57 2.57 2.61 2.61 2.84 2.83 2.77

1.34 1.32 1.32 1.44 1.44 1.43

1.24 1.29 1.29 1.40 1.39 1.34 0.00 0.01 0 .01 0,00 0,00 0.00 2.25 2.27 2.24 2.45 2.46 2.50 0.02 0.02 0.02 0.01 0.01 0.01 2.53 2.43 2.46 2.15 2.16 2.18

40 R Costagliola et al.

Table 2. Representative electron microprobe analyses of arsenopyrite (in wt. %), calculated composi- tion in at. % and statistic parameters of As content (at. %)

sample AMA1 S 21.95 22.11 21.79 As 42.26 42.16 43.06 iFe 35.13 35.29 35,15 Co 0.19 0.03 0.05 Ni 0.25 0.01 0.04 Tot, 99.78 99.60 100.09

A1 22.42 22.01 41.78 42.86 35.76 35.42

0.04 0.05 0.01 0.00

100.01 100.34

A2 21.76 22.58 21.86 21.89 22.15 42.54 41.83 42.55 42.13 42.15 35.29 35.87 35.73 35.96 35.91

0.24 0.04 0,09 0.07 0,05 0.03 0.06 0.08 0.11 0.02

99.86 100.38 100.31 100.16 100.28

~,M1 21.66 21.89 21.88 21.15 42.02 42.07 41.86 43.05 36.19 35.77 36.15 35,86 0.07 0.07 0.06 0.13 0.05 0,06 0.03 0,17

99.99 99.86 99.98 100.36

I I A' sample AMA1"

IS t% 136,46 36.60 36,08 36.86 iAs 13004 29.87 30.511 29.39 Fe I 33.50 33.53 33.41[ 33.75

I A2 JAM1 1

36.27 36.13 36.97 36.09 36.15 36.43~ 35.85 36.23 36.14 35.16 30.221 30.23 29.31 30.06 29.77 29.67~ 29.76 29.79 29.59 30.62[ 33.511 33.64 33.72 33.86 34.09 33.90 / 34.39 33.98 34.27 34,221

As (at. %) statistics Mean St. Dev. Min. Max. Counts 29.90 0.37 29.31 30.66 31

the 320-370 °C range. The lack of significant compositional zoning in arsenopyrite suggests that temperature and f(S2) either were relatively constant, or covariated allowing arsenopyrite to maintain a constant composition. Kligfield et al. (1986) stated that, during the D2 deformation phase, metamorphic pressure and temperature in the Apuane Alps were close to peak values. Textural evidence indicates that arsenopyrite formation was synchronous with D2. Hence we suggest that arsenopyrite geothermometry reflects the metamorphic temperature peak in the UFP. On the other hand, textural studies indicate that chlorite is a retrograde phase and its relationships to metamorphic structures are unclear. Chlorite geothermo- metry yields temperatures slightly lower than those obtained by arsenopyrite. This difference may be either an artifact arising from uncertainties in the application of the geothermometers or, if it is real, may suggest that biotite replacement by chlorite took place under conditions slightly below the metamorphic peak, e.g. during an early retrograde stage.

Phengite geobarometer

Phengite is frequent at Pollone and occurs in $2 foliation planes in host and country rocks, and in microveins hosted by the "tormalinolite". These veins show mineralogical, textural and structural characteristics similar to the macroveins present in the mining area, which opened during the D2 stage of deformation (see the following discussion). The analyzed phengites do not coexist with a limiting assemblage (e.g., K-feldspar, phlogopite and quartz; Massonne and Schreyer, 1987). Therefore, their compositions (Table 3) is not uniquely a function of pressure at a given temperature. However, Massonne and Schreyer (1987) suggest

Metamorphogenic barite-pyrite (Pb-Zn-Ag) veins 41

(/) ¢-

E (1)

4--- O

O

Z

.

6-

4 -

2-

0 280 300

12- 10.

8

6

4

2 0

J ~=~ Chlorite I

I

320 340 360 Temperature (°C)

(a)

I I

380 400

: x~×xxxxx=

I

Arsenopyrite [

I

(b)

NNN I

280 300 320 340 360 380 400 Temperature (°C)

12 10

8

6

4

2

0

~ Rock Microveins

0

()

\ \ \ \ \ \ \ \ \ \ " ~\\\\\\\'~ 222222/ (((<<((((~ C , / / # # # / / # / ,

/ / / / / / / / / / ,

/ / / / / / / / / / ! I I !

1 2 3 4 5

Pressure (Kbar) Fig.

Fig. 7. a Histogram of temperatures obtained by applying the A1TM chlorite geothermometer of Cathelineau (1988). b Histogram of temperatures obtained by applying Kretschmar and Scott's (1976) geothermometer for arsenopyrite in textural equilibrium with pyrite, e Histogram of pressures obtained by applying the phengite geobarometer (Massonne and Schreyer, 1987) to Pollone host- and country-rocks for a reference temperatures of 350 °C, based on arsenopyrite geothermometry

that phengite composition, in the presence of a Mg-Fe silicate, can provide a minimum pressure estimate. The analyzed phengite was sampled from a mineral assemblage that, beside tourmaline, includes chlorite. Due to the temperature dependence of this barometer, application of the phengite "geobarometer" requires an independent temperature estimate. Textural evidence indicates formation of the analyzed phengite during the D2 stage, for which arsenopyrite geothermometry gives a mean value of about 350 °C. In the case of phengite hosted by microveins, it was assumed that the ratio between the fluid volume and the wall rock area was

42 E Costagliola et aL

Table 3. Representative electron microprobe analyses of muscovite (expressed as oxides, in wt. %). Cation occupancies calculated on the basis of 24 oxygens and statistic parameters of Si occupancy

S,,ample E16C EgA 3.8189 E13B AI203 31,11 33.72 32.19 32.66 3231 33 .04 32.38 31.04 SiO2 47.13 48.38 48.64 45,64 46.85 47.05 45.55 46.85 Na20 0.52 0.52 0.51 0.69 0.55 0.62 0.77 0,54 MgO 2.00 1.32 1.39 1.46 1,37 1.39 1,39 2.21 K20 10.02 9.90 9.59 10.03 10.39 10,48 9.75 10.01 CaO 0.00 0.06 0,06 0.03 0.00 0.01 0.13 0,02 TiO2 0,24 0.27 0.28 0.28 0.31 0.25 0.32 0.21 0r203 0.01 0.01 0.00 0.03 0.01 0.02 0.02 0.01 MnO 0.01 0.00 0.01 0.00 0.01 0.02 0,04 0.02 FeO 1.53' 1.86 1.85 1,98 1.88 1.92 2.23 1.54 TOT 92.57 96.05 94.53 92.80 93,87 94.81 92.68 92.44

Sample E-3 AI20~ 29.24 SiO2 46.53 46.68 Na20 0.12 0.12 MgO 1.64 1.67 K20 10.92 10,85 CaO 0.01 0.01 Ti02 0.93 0.81 P-,r20~ 0.04 0.18 MnO 0.02 0.01 FeO 3,13 3.15 'TOT 92.58

CDM4 29.3 29.40 29.98 30.62 . . . .

46.40 0.13 1.62

10.97 0.00 0.77 0.02 0.00 3.11

29.71 29.84 29,32 45.31 44.82 44,99 44.12 45.79

0.16 0.14 0.20 0.20 0.20 1.81 1.77 1.71 1.79 1,85

10.19 10.22 10.37 10.46 10.44 0.02 0.01 0.00 0.03 0.03 0.91 0.87 0.68 0.98 0.95 0.01 0.00 0.00 0.00 0.00 0.06 0.05 0.06 0.08 0.02 3.14 3.63 3.49 3.59 3.61

. . . . . . . , , , , , ,

92.83 92.42 91.32 91.35 91.47 91.85 92.21

s.ample E16c AI 5.47 Si 7.03 Na 0.15 Mg 0.44 K 1.91 Ti 0.03 Fe 0.19

E9a 3.8.89 IEt3B I

5,70 5.52 5,76 5.66 5.70 5.73i 5,47 6.94 7.08 6.83 6.92 6.89 6.831 7.01 0.15 0.14 0.20 0.16 0.18 0.231 0.16 0.28 0.30 0.32 0.30 0.30 0.311 0.49~ 1.81 1.78 1,91 1.96 1.96 1.871 1.91 0.03 0.03 0.03 0.03 0.03 0.041 0.02 0,22 0.23 0,25 0.23 0.24 0.281 0J9

sample E- 3 AI 5.22 Si 7.05 7.05 7.04 Na 0.03 0.04 0,04 Mg 0.37 0,38 0.37 K 2.11 2.09 2.12 Ti 0.11 0,09 0,09 Fe 0.40 0.40 0.40 ~

CDM4 5.23 5.26 5.r S.27 5.37 5.41 5,43

6,94 6.89 6.91 6.77 6.98 0.05 0.04 0.06 0.06 0.06! 0.41 0.40 0.39 0.41 0.42 1.99 2.00 2.03 2.05 2.03 0.10 0.10 0.08 0.11 0 0.40 0.47 0.45 0.46 0

Si content statistics ]

J Mean St, Dev. Min. Max. Count 6.94 0.10 6.60 7.08 41

Metamorphogenic barite-pyrite (Pb-Zn-Ag) veins 43

quite low since the low fluid flux through the vein section; consequently, the temperature of the fluids was controlled by wall rock temperature. By applying the phengite geobarometer for a temperature of 350 °C, a frequency peak at about 3.5 kbar was obtained (Fig. 7). Note that both rock and microvein phengites have similar compositions, leading to the conclusion that the mineralizing episode was characterized by Pfluid e.o Pload-

Fluid inclusions

Fluid inclusions were studied in barite, fluorite and quartz I and II hosted by syn-D2 orebodies (both low and high angle veins). Only fluid inclusions considered as primary have been considered. Fluid inclusions in quartz occur both along planes and as isolated individual inclusions, and may show negative crystal or irregular shapes. Necking-down is relatively common; in order to avoid errors due to post- entrapment density (composition) changes, inclusions showing necking-down were not considered. Fluid inclusions in barite occur both isolated and along planes. Irregularly-shaped inclusions are predominant. Fluid inclusions in fluorite commonly show a negative crystal morphology. They are frequently located along the growth planes of the mineral.

Inclusion fluids belong either to the HzO-NaC1 or to the H20-NaC1-CO2 system. Barite and fluorite host mainly CO2-bearing fluid inclusions which, by contrast are uncommon in quartz. Melting temperatures of CO2 are generally close to -56.6 °C, with values down to -58.3 °C indicating traces of other gases (CH4 and N27). CO2 densities, calculated from the CO2 homogenization temperature, are in the range 0.6-0.7 (g/cm 3) and 0.6-0.9 (g/cm 3) in fluorite and barite, respectively. XCO2, was calculated considering the CO2 density, the aqueous fluid composition and the relative volumes of the aqueous and carbonic phases. XCO2 generally varies between 0.08-0.15. All the inclusions homogenize to the liquid phase. In barite samples, total homogenization was seldom observed, due to widespread decrepitation (Fig. 8).

Salinity varies in a relatively narrow range between 5-13 eq. wt.% NaC1 (Fig. 8). In CO2-bearing inclusions, salinity was calculated using the clathrate melting point. Where traces of CH4, (inferred from CO2, melting temperatures) are present this lead to a slight underestimation of true salinity, because CH4 raises the melting point of clathrate. For example, the barite-hosted fluid inclusions showing an apparent salinity lower than 7-8 eq. wt.% NaC1 (Fig. 8) may have a true salinity closer to the bulk of the data, which falls in the range 8-11 eq. wt.% NaC1. Homogenization temperature (Th) in quartz ranges from about 110 to 285 °C, with most of the values plotting between 200 and 250 °C (Fig. 8). No systematic differences have been observed between fluid inclusions hosted by the first and second generation of quartz. Th is generally reached without decrepitation, with a few exceptions. On the contrary, barite-hosted fluid inclusions show a tendency to stretch and decrepitate before total homogenization. Therefore, barite Td (temperature of decrepitation) values shown in figure 8 have been plotted only to show the salinity of the inclusions and are not meaningful for calculating fluid PTV properties. Fluorite-hosted fluid inclusions homogenize in the 220-230 °C range (Fig. 8) without undergoing significant stretching. In fact, CO2 density

4-4 R Costagliola et al.

300

¢-,

0

"" 250 N_ c ~ O

o ~ 200

"1-

o 150

100 Q.

E I---

I o

0 A

quartz barite (Th) oO barite (Td) 0 0 fluorite <> <> ~ .~i~

<> <>'~. <> <>

<> ^ <><> 8 <>2 ° 0

Salinity (wt % NaCI eq.)

Fig. 8. Homogenization temperature (Th) vs. salinity (in wt.% NaC1 eq.) diagram of fluid inclusions hosted in ore minerals at Pollone. Th homogenized fluid inclusions in barite. Td decrepitated fluid inclusions in barite

~" 3

2

0 I , I , I , I , I , I

250 300 350 400 450 500

F~

Temperature °C

Fig. 9. Trapping temperatures estimated by pressure correction applied to mean isochores calculated from microthermometric data in fluid inclusions hosted by barite, quartz and fluorite. See comments in the text about reliability of temperature estimates

Metamorphogenic barite-pyrite (Pb-Zn-Ag) veins 45

measurements performed before and after total homogenization do not show appreciable differences. Isochores for quartz hosted fluid inclusions have been calculated for H20-NaC1 fluids from Th and salinity values using Zhang and Frantz's (1987) equation. The bulk density ofCO2-bearing inclusions (barite and fluorite) was calculated by taking into account the volume fractions (visual estimates) and the densities of the aqueous and carbonic phases; the corresponding isochores have been calculated using Bowers and Helgeson's (1983) equation. The mean isochores for barite, quartz and fluorite are shown in Fig. 9.

During the uplift of metamorphic terranes, fluid inclusions trapped close to P-T peak metamorphic conditions may be subjected to large pressure differentials (both negative and positive) between the internal pressure and the confining pressure (~Pload; Sterner and Bodnar, 1989; ~tyk and Bodnar, 1995). This mechanism could lead to volume changes or fluid inclusion decrepitation, especially for "soft" phases such as fluorite and barite (Bodnar and Bethke, 1984; Ulrich and Bodnar, 1988). Quartz is considered a more resistant host mineral, at least for pressure differentials up to 1-2kbar (Sterner and Bodnar, 1989; Vityk and Bodnar, 1995). As a consequence, if under- or over-pressures were produced during uplift, fluid inclusions in fluorite and barite are more likely to have undergone volume changes than those hosted by quartz. At Pollone, Costagliola et al. (1994) have suggested that the mean barite isochore is nearly coincident with the regional uplift path; however, it seems more conservative to assume that inclusions in quartz provide the most reliable estimate of original fluid density, and thus they should give more reliable temperature estimates. By assuming a pressure correction of 3.5 kbar, isochores for fluid inclusions in quartz give temperatures of about 450 °C (Fig. 9).

5 6 ] XC02 = O. l /

~, 4 ~, ~-of ~t¢ b~¢°lx ~ ° c h ° ~ ' ' ' ' ~

3 o×

~ x~¢, P =4"1KC~ 2 = 60 °

1 , I w i I i i I , i I =

350 400 450 500

Temperature (°C)

Fig. 10. P-T estimate obtained by the intersection of the mean isochore for fluorite-hosted fluid inclusions with the "talc in" isograd, calculated according to Powell and Holland (1988) for XCO2 = 0.1

46 R Costagliola et al.

The rare occurrence of talc has been reported by Orberger (1985) in the northern area of Pollone, probably reflecting the prograde reaction:

3 Dolomite ÷ 4 Quartz + H20 = Talc + 3 Calcite + 3 CO2

in dolomitic layers of the Scisti di Fornovolasco. Orberger (1985) does not specify the textural relationships of talc, but it is likely that this mineral was formed under peak metamorphic conditions (cf. Winkler, 1976). The dolomitic layers of the Scisti di Fornovolasco host syn-metamorphic (syn-D2) mineralized veins. CO2-bearing fluids, possibly reflecting the above decarbonation reaction, have been observed in fluid inclusions in fluorite from these veins. The °'talc in" isograd, calculated for the XCO2, measured in fluid inclusions, intersects the mean fluorite isochore at P-T conditions of about 4 kbar and 450 °C (Fig. 10).

Geochemistry

Analytical methods

Major clement analyses were performed in the Dipartimento di Scienze della Terra, Universit?t di Firenze, using a combination of AAS (Na, K) and XRF (Mg, Na, K, A1, Si, P, Ca, Ti, Mn, Fe) techniques. Trace element contents (Sr, Ba, Y, Rb) were also determined by XRE The XRF analyses were corrected for matrix effects using the method of Franzini et al. (1972). A specific calibration was necessary for samples containing large amounts of Ba (0.2-20wt.%) to account for matrix effects. Standard curves were obtained by adding known amounts of Ba to the samples. For contents below 2000ppm, the Ba contribution to the absorption coefficient was considered negligible. Quartz samples for O-isotope analyses were purified by heavy liquids to eliminate barite and pyrite and, if necessary, by fusion with KHSO4 to eliminate micas. Repeated treatments yielded samples with >> 95% quartz, as confirmed by binocular microscope observations and quantitative X-ray diffractometry. Powdered samples were analyzed for O-isotopes at the ETH in Zurich. Oxygen was extracted by reaction with BrFs following the method of Clayton and Mayeda (1963), and its isotopic composition was determined on a VG Isogas 903 mass spectrometer. The results are presented in the standard 6 notation relative to SMOW.

Major and trace elements

Chemical analyses for wall, host and country rocks are listed in Table 4, and presented in Fig. 11 and 12. Generally, compositions of wall and host rocks do not show wide variations. This reflects the overall quite simple mineralogical composition of these lithotypes: quartz and muscovite are the main phases, often in similar proportions, in all samples. Relatively large deviations in chemical composition correspond to local enrichment of accessory phases such as plagio- clase, K-feldspar, chlorite, calcite, dolomite, and tourmaline. This is the case, for example, for the samples taken in the upper part of the Scisti di Fornovolasco formation. They display a high LOI (>15 wt.%) and high CaO due to intercalation of some carbonate-rich horizons. Generally, country rocks may exhibit a relatively

Metamorphogenic barite-pyrite (Pb-Zn-Ag) veins 47

Table 4. Major (expressed as oxides, in wt. %) and trace (ppm) element analyses of the Scisti di Fornovolasco formation. (1) siliciclastic facies; (2) rhyolites and rhyodacites. W wall rock; H host rock; C country rock

MAJOR ELEMENTS TRACE ELEM. Sample Rock SlO2 AI203 Fe203 CaO Na20 K20 MgO TIO2 P205 MnO BaO LOI

V25EA C(1) 63.66 13.61 11,46 0 .28 0.01 4.94 0,82 0,55 0.1 0 .08 0 .23 4.27 lm3589/1 C(1) 42.29 15,63 11,11 11.19 0.64 1.7 0.59 0.96 0.22 0.32 0.18 15.17 CDM88/2 C(1) 67.84 15.75 3,76 0 ,24 1.58 4.24 1,77 0.7 0.13 0.04 0 . 5 3 3.42 CDM5 SF1 SF2 SF3C SF4 E1 El l E12

C(1) 68.01 14.06 6.63 0 .28 0.07 5.4 0.7 0.87 0.18 0.16 0 .18 3.47 C(1) 62.34 10.91 3.98 5 .56 0.13 3.56 3.71 0.41 0.09 0.08 0.05 9.2 C(1) 71.6 1 4 . 9 2 . 4 6 0 .39 2.72 3.33 1.68 0.68 0.26 0,01 0 . 0 6 1.92 C(1) 66.71 16.51 4.42 0 .23 3.45 3.17 2.24 0.66 0.13 0.02 0 .09 2.38 C(1) 49.14 27.72 4.29 0.3 0.58 8.64 1.63 1.37 0.05 0.01 0 . 9 3 5.34 C(1) 67.55 17.78 2.66 0 ,22 1.79 5,46 0.7 0.76 0.1 0.13 0 .33 2.49 C(1) 57.33 18.68 3.85 4 .35 0.21 4,47 2.08 0,84 0.18 0.12 0.6 7,3 C(1) 65.61 21.14 2.34 0 .06 0.42 6,01 1.24 0.29 0.09 0.01 0.51 2.29

E4 E5 MA4/1 MA89/1 MA89/2 MA89/3 MA89/4 MA89/5 MA89/7 MA89/8

H(1) 65,93 18.62 2 . 5 9 0 .19 2.18 5,04 0.69 0.81 0.13 0.13 1.24 2.46 H(1) 86.92 17.34 2.06 1.44 0.34 5,29 1.24 0.41 0.11 0 .08 1 .23 3.53 H(1) 63.21 8.72 3.3 0 .33 0.19 0.39 0.93 0.5 0.01 0.01 1.31 21.1 H (1) 73.61 16.38 0.8 0.01 0.21 4.68 0.58 0.7 0.01 0.01 1.15 1.88 H(1) 76.2 13.39 1.9 0 .02 0.04 3.73 0.59 0.47 0,01 0.01 0 . 8 5 2.81 H(1) 64.82 20.75 2.36 0 .16 0.28 6.25 0.82 0.85 0.04 0.01 0.51 3.16 H(1) 66.39 20.99 0.97 0 .18 0.21 5.97 0.69 0.77 0.11 0.01 1.03 2.68 H(1) 57.48 25.63 1,51 0 .25 0:28 7.83 0.87 1.11 0.14 0.01 1.19 3.7 H(1) 68.8 1 5 , 8 3 . 8 7 0.61 0.31 3.86 1.43 0,58 0.05 0.01 1.25 3.42 H(1) 66.83 16.86 3 . 0 2 0.16 0.2 4.65 0.49 0.62 0.1 0.01 2 . 1 8 4.89

MA5S11 H (1) 70.72 3.21 4.84 3.47 E24 E27 E29 E13 E15 E16 E17 E16 E19

0 1.27 0.61 0.12 0,03 0 12,68 3.05 H(1) 63,22 22.21 1.3 0 .28 0.28 7.08 0,86 0.83 0.09 0.01 1,03 2.83 H(1) 66.7 20.66 1.31 0 .02 0.21 6.12 0.53 0,84 0 0.01 0 . 6 8 2.93 H(1) 76.45 13.15 1.83 0.1 0.04 3.65 0.35 0.55 0 0 1,23 2,75 H(1) 63.36 11.64 3.25 6 .19 0.26 3.32 3.08 0.45 0.25 0.15 0,81 7.24 H(1) 73.13 15.35 1.11 0 .72 0.35 4.53 1.41 0.75 0.13 0.02 0 . 5 5 1.95 H(1) 54,43 18.05 2.55 6.6 0.37 5,55 2.45 0.86 0,11 0.15 0,91 7.98 H(1) 50.62 22.03 3.35 4.01 0.28 6.76 2.27 0.86 0.14 0.13 1 .17 8.38 H(1) 71.29 14.96 3.41 0.21 0.1 4.04 0,53 0.55 0.15 0.01 0.7 4.07 H(1) 72.15 16,28 1.24 0 ,19 0.21 5.05 0.69 0.62 0.15 0.01 1 2.42

E20 W(1) 72.3 10.86 0.67 1.97 0.06 3.38 0.21 0.34 0.06 0.01 7 . 9 8 2.15 E28A W(1) 77.29 8.74 2.33 0 ,04 0.01 2.49 0.17 0,28 0 0 5 . 4 3 3.22 E30 W (1) 69.34 9.35 0.5 0.4 0.03 2.8 0.13 0.3 0.01 0 16.58 0.92 VBA62 W(1) 86.92 5.61 0.39 0 ,09 0,39 1.64 0.55 0.22 0,05 0 3.3 0.85 )3589/1 C(2) 70,04 17.69 2,48 0 ,18 0.28 5.02 0.7 0.36 0.11 0.1 0 . 5 4 2.52 3589/2 (3(2) 26.68 18.53 5.95 12.92 0.37 4.38 6.11 0,72 0.04 0.47 0.86 22.99 S3589/3 C(2) 67,73 18.51 1.38 1.6 0.31 5.12 0.82 0.26 0.11 0.05 0 .16 3.94 SAR1 SAR2 SAR3 SAR4 SAR5 SAR6 SAR7 E6 E8 E21 E23 CDM88/1 C(2) 69.67 18.41

C(2) 53,97 17.64 10.62 0 ,39 2.96 2.16 6.2 1.76 0.27 0.1 0.04 3.9 C(2) 57 19.76 7.82 0.19 2.3 4.21 3.97 0.73 0.15 0.05 0.1 3.74 C(2) 67.49 16.96 3,21 0 .99 3.61 3.22 1.52 0.38 0.11 0.04 0 . 0 6 2.42 C(2) 60,41 18.89 5.97 0.41 1.19 5.3 2.98 0.85 0.1 0 .04 0 . 7 2 3.14 C(2) 63.04 18.44 5.76 0 .26 2,34 4.14 2.3 0.69 0.15 0.06 0 . 0 7 2.77 C(2) 68.39 16.15 3.4 0 .56 1,92 4.54 1.59 0.42 0.12 0.4 0.08 2.7 (3(2) 67.79 15.94 3.2 1 .27 3.55 3.08 1.59 0.4 0.12 0.09 0 ,05 2.92 C(2) 28.54 10.16 7.1 18.19 0.1 2.52 7.11 0.21 0.04 0.49 0.48 24.99 C(2) 66,2 18.63 1.37 1.98 0,24 5.17 1.23 0.26 0.11 0.06 0 .17 4.59 C(2) 62.97 16.79 3,82 2 .16 2.78 3.71 1.98 0.53 0.13 0.08 0 .69 4.36 C(2) 64.77 17.64 4.58 0 .65 2.45 4.42 1.48 0.77 0.18 0.02 0 .65 2.39

1.13 0 .98 0.24 4.96 0.88 0.24 0.12 0.03 0 .35 2.99

Rb Sr Y 172 42 16 70 331 36 175 88 22 222 55 17 127 53 19 131 75 37 130 81 20 337 122 26 192 76 14 223 179 24 205 50 11 178 130 21 196 67 17 19 462 19 185 51 11 144 28 4 229 74 17 233 77 25 310 116 20 111 303 15 156 456 13 10 2355 8

241 34 22 197 25 14 106 143 3 117 187 16 138 27 22 189 162 25 256 88 22 121 30 4 176 39 16 64 1580 7 57 1036 6 10 3620 12 36 805 1 206 57 13 180 165 20 189 67 13 82 51 31 151 40 31 115 103 17 2O8 50 20 146 55 27 166 51 19 117 103 18 117 271 23 190 71 10 163 80 16 177 53 16 181 181 11

48 R Costagliola et al.

o ~ ('4

O O4

m

0.6 ,~

0.5 []

0.4

0 " 3 t ~ " 0.2

0.1 t ~ 0 .0 ' ,

0

30

25-

2O- 15: []

¢o

0 10-" O4

5- 0

20 ' 3'0

Country Rock (1) Host Rock (1) Wall Rock (1) Country Rock (2)

A • A

g 1'0 1'~ ...... 2o BaO (wt. %)

• %

mP A • A A

A

' 4 ' 0 ' 5 ' 0 ' 6 ' 0 ' 7 ' 0 8'0'9 '0

SiO 2 (wt.%)

Fig. l l. a K20/A1203 vs. BaO diagram. The K20/A1203 ratio is constant around 0.3, i.e., the stoichiometric value of mus- covite, in both host and wall rocks, b A1203 vs. SiO2 diagram. (1) siliciclastic facies; (2) rhyolites and rhyodacites

16J 14 12

o ~ lo ¢-1

~" 6 4 2 ¸ 0

3 , 5 '

3.0 ~

2.5 ~

ca 2.0 ~

g'15~ d

0.5 1.0

[ [] Country Rock (1) • Host Rock (1) " Wall Rock (1) ,1 Country Rock (2)

~ ' , . . .

6 ; 1'o 1'5 BaO (wt. %)

[] rl []

m~ mm

°~m B

1;s 22o 22s

A

31o

A

Log (Sr)

20

B

4.0

Fig. 12. a Rb/A1203 vs. BaO dia- gram showing the progressive Rb decrease with increasing Ba content in the mineralized area. b LogBa vs. LogSr diagram. Sr content decrea- ses with increasing distance from orebodies. (1) siliciclastic facies; (2) rhyolites and rhyodacites

Metamorphogenic barite-pyrite (Pb-Zn-Ag) veins 49

broader variation in chemical composition compared to both host and wall rocks, in response to a relatively more variable mineralogical assemblage. However, the main chemical differences between host and country rocks mostly concerns Ba and Si.

Figure 11 indicates that samples show an inverse correlation between A1203 and Sit2. High SiO2 contents reflect wall rock silicification. Samples with high silica contents are also Ba-enriched, indicating that silicification was accompanied by barite deposition, as could be expected from the close association of quartz and barite in the mineralized veins. Elements like A1, Y and Ti, usually considered immobile in low-grade metamorphic environments (Etheridge and Cooper, 1981; Wedepohl, 1978; Kerrich, 1988), are strongly depleted with increasing SiO2 contents, because of passive dilution due to increase in rock volume as the veins formed (cf. Nabelek, 1987). Immobility of A1 and Ti is further suggested by the constant A1203/TiO2 ratio in both wall rocks and country rocks. K20/A1203 in wall and host rocks (around 0.3; Fig. 11) is close to the stoichiometric value in muscovite, which is the only K-A1 phase present in most samples. In some instances, country rocks have K20/A1203 values ranging from about 0.1 to 0.4; the lowest values are indicative for the presence of Al-bearing and K-poor minerals such as chlorite and tourmaline, whereas the highest values reflect the presence of K-feldspar. From Figure 11 it can be observed that the constancy of K~O/AI~O3 ratio is maintained irrespective of SiO2 (and St: not shown) contents, whereas the Rb/A1203 ratio decreases with increasing Ba concentrations in host and wall rocks (Fig. 12). Major chemical variations induced by mineralizing fluid circulation are evidenced by the logSr vs. logBa plot, in which two main groups of samples are clearly distinguishable (Fig. 12). A positive correlation between these two elements in the first group is not observed in the second one.

Rb geochemistry

KzO/A1203 and K20/Rb values are constant in most wall and host rocks, whereas Ba (and Sr) enriched wall rocks have variable and higher K20/Rb values. The latter may be explained by an absolute decrease of Rb, rather than by an increase in K20. At constant values for K20/A1203 Rb/A1203 decreases. Low values for Rb/A1203 are also observed in a few samples of country and host rocks, reflecting the presence of Al-bearing (Rb-poor) minerals such as chlorite and tourmaline. Assuming that the geochemical behavior of Rb is controlled primarily by the substitution of Rb for K in silicates, the low Rb contents of wall rocks suggest the presence of Rb- depleted muscovite with respect to the rocks cropping out in the area. This implies a low-Rb content in the mineralizing fluids as well. Under metamorphic conditions, the Rb content of fluids is mainly controlled by partitioning of this element between Rb-poor (K-feldspar) and Rb-rich (muscovite and biotite) mineral phases (Barbey and Cuney, 1982; Kerrich, 1988). K/Rb increases in (meta)siliciclastic rocks as long as K-feldspar forms at the expenses of Rb-rich phases such as biotite and muscovite, At intermediate metamorphic grades, biotite typically buffers the Rb rock content (Barbey and Cuney, 1982). In granulitic facies, the disappearance of biotite causes a dramatic drop in the Rb content of the rock, and, consequently, a significant enrichment in the concentration of Rb in metamorphic fluids in equilibrium with the biotite-consuming reactions (Kerrich, 1988). It

50 R Costagliola et al.

follows, therefore, that Rb-depleted fluids could form from the breakdown of K- feldspar (ksp) to produce biotite (bi) and/or muscovite (mu), following a reaction like:

ksp + fluid = bi (mu) + Rb - poor fluid

K-feldspar is an accessory phase in most lithotypes of the Scisti di Fornovolasco formation; relatively high amounts are observed in rhyodacites and rhyolitic bodies. However, the absolute quantities are always fairly small. Therefore, the breakdown of K-feldspar at Pollone is not the major reason for the depletion of Rb in the mineralizing fluids. Hence, the mineralizing fluids must have equilibrated with a rock originally containing a greater amount of K-feldspar (i.e. a rock with a different modal composition) than the Scisti di Fornovolasco in the Pollone area. This suggests an external source for the mineralizing fluids and, probably, also for Ba and the other elements present in the orebodies. The Rb depletion of the wall rocks is accompanied by an alteration which generally affects a relatively narrow band (1-20 cm) along the vein. Moderate silicification, plagioclase and K-feldspar breakdown to form white mica are the main alteration processes observed. The data suggest that the chemical components of mineralizing fluids, except Rb, were close to equilibrium with the main wall rock minerals (quartz and muscovite). In most igneous rocks, Rb and K behave in a coherent way because of the ready substitution of Rb for K due to the similar ionic radii. As a consequence the K/Rb ratio is remarkably constant in many magmatic rocks. Howevei, Rb has a larger ionic radius than K, and therefore it is strongly concentrated in late stage magmatic fluids. Geochemical studies on fresh and altered andesites related to porphyry copper deposits indicate that Rb is strongly enriched (with consequently low K/Rb ratios) in altered rocks that experienced intense magmatic hydrothermal circulation. These Rb anomalies were interpreted as the result of hydrothermal alteration by a Rb-enriched aqueous phase separated from a cooling magma. The distribution of the K/Rb ratio in the Pollone wall rocks is quite different from that characteristic of magmatic-related ore deposits or pegmatites, which typically show low K/Rb ratios (see e.g. Armbrust et al., 1971; C~rny, 1982). This is a further support to the concept that syn-tectonic ore deposits of the Apuane Alps are not necessarily related to a supposed, but so far unrecognized, deep-seated plutonic source, but may be ascribed to fluids generated in a metamorphic environment (cf. Benvenuti et al., 1989).

Sr geochemistry

Further support for K-feldspar breakdown as a possible source of Rb-poor fluids comes from 87Sr/86Sr isotopic data. Three barite samples from different stratiform (syngenetic?) barite bodies (Pollone, Monte Arsiccio and Buca della Vena; see Lattanzi et al., 1994, for deposit location) show very similar 87Sr/86Sr ratios of about 0.7102, suggesting a common Sr source (Barbieri et al., 1982). A range of 0.71017-0.71122 was given by Orberger et al. (1986) in samples from the southern sector of Pollone. On the other hand, samples from the studied vein bodies in the northern sector of Pollone appear isotopically inhomogeneous, with 87Sr/86Sr values generally higher (range 0.71098-0.71581) with respect to stratiform

Metamorphogenic barite-pyrite (Pb-Zn-Ag) veins 5l

orebodies (Costagliola et al., 1997). This would imply that the source of Sr for metamorphogenic barite was not the stratiform barite alone. The contribution of a more 87Sr-rich source, probably from the breakdown of K-feldspar (cf. Etheridge and Cooper, 1981; Chaudhuri and Clauer, 1993), is therefore suggested (Costagliola et al., 1997). The Sr and Ba distribution in the studied rocks is controlled by the presence of barite, iThe only other mineral phases that may contain appreciable amounts of Ba and Sr are carbonates (calcite, dolomite), which are, in general, quite scarce in the study area. Most barites from Pollone have a SrO content of about 0.5-1.0 wt.% (Barbieri et al., 1982; Checcacci, 1983; Orberger, 1985). Consequently, a positive correlation between Ba and Sr should be expected. As shown in Fig. 12, a good positive correlation exists for samples taken in close proximity to the orebodies. As distance from the orebodies increases, the Sr contents decrease towards values lower than the mean composition of the bulk continental crust (260 ppm; Taylor and Mc Lennan, 1985), whereas the Ba contents are constantly higher than mean crustal abundances. This indicates that the Sr content in barite tends to decrease as the distance from the ore bodies increases. The partition of Ba and Sr between barite and fluid can be modeled assuming that the two elements (having similar ionic radii) are essentially equivalent in the barite structure. Hence, assuming activity coefficients close to unity, the amount of Sr substituting for Ba in the barite structure at any temperature is dependent on the solubility product of the two sulfates (Church, 1979). The ideal distribution coefficient of Sr between fluid and barite (KSrT) can therefore be evaluated at any temperature by the following equation:

KSrT -- Kd[BaSO4] _ [Ba2+][SrSO4] Kd[SrSO,] [Sr2+][BaSO4]

where Kd are solubility constants at a given temperature. The solubility data of barite and celestite are scarce and controversial (Blount, 1977; Gundlach et al., 1972; Strubel, 1966). However, KSrT values, calculated by the available data, generally display a positive correlation with temperature in NaC1 bearing fluids (Gundlach et al., 1972); for values higher than 200°C, KSrT shows a marked increase with temperature (Orberger, 1985). These evidences indicate that barite with progressively lesser amounts of Sr may precipitate from cooling fluids, provided that the Sr2+/Ba 2+ ratio in the fluid remains constant. The linear relation between Sr and Ba contents in mineralized wall rocks (Fig. 12) suggests that this latter requirement seems to have been attained at Pollone (alternatively, SrZ+/Ba 2+ and temperature covariated to produce the observed linear array). Consequently, the variation of Sr content in barite from vein bodies towards the host rocks may be explained by a progressive cooling of mineralizing fluids during lateral infiltration into the host rocks. This interpretation is in accordance with geothermometric data, indicating that the temperature of the mineralizing fluids was higher than the host rock temperature.

Oxygen isotopes

The 5180 values for quartz fall within a relatively narrow range (14.4-19.0 %O(smow) Table 5). No significant difference among veins, stratiform mineralizations and

52 E Costagliola et al.

Table 5. O-isotope composition of quartz (expressed as 61SO%o(smow)). HAO High Angle Orebodies; LAO Low Angle Orebo- dies; H host rocks; C country rocks

Sample VBA 3 VBA 2

E30 Ang PRZ

E 28 QZ 2.3.90/4 2.3.90/1

El4 El9

MA 89 5 SF2 SFI

CDM 88.1

Setting] O-Isotope composition HAO ' 17.2 HAO HAO HAO HAO HAO t LAO LAO LAO

H H C C C

m

16.2 17.6 19.0 17.1 17.7 16.2 17.0 16.9 16.0 14.4 16.8 17.4 15.3

r-

E

Q

"6

_

3- !

.

HAO LAO

13 14 15 16 17 18

160 (SMOW) °/oo

19

Fig. 13. O-isotopic composition of quartz at Pollone. HAO High Angle Orebodies; LAO Low Angle Orebodies; see text for explanations of terms reported in the legend

Metamorphogenic barite-pyrite (Pb-Zn-Ag) veins 53

wall, host and country rocks has been observed (Fig. 13). Quartz from country rocks, lacking any evidence of fluid circulation, is indeed isotopically indis- tinguishable from vein and wall rock quartz.

At the estimated deposition temperature (450°C), mineralizing fluids in equilibrium with quartz would have had an 6180 composition of about 10-13%O(smow), within the so-called "metamorphic fluids field" (e.g., Taylor, 1974). Temperatures of homogenization of quartz-hosted fluid inclusions also fall in a relatively narrow range suggesting that quartz deposition took place under broadly isothermal conditions at Pollone. This evidence indicates that the observed O-isotope compositions may reflect an isotopically homogeneous mineralizing fluid reservoir. The 6180 values of quartz from other deposits in the Apuane Alps, such as Bottino (Benvenuti et al., 1991) and Frigido (Benvenuti et al., 1995), as well as from basement rocks (Pochini, 1979; Cortecci et al., 1994), range between 14-18 %O(smow). Therefore, the homogeneous O-isotope composition of the Pollone samples, that could be an indication of an isotopically closed system, could alternatively be explained by an external provenance of fluids (open system), possibly from different structural levels of the Apuane Alps basement. Following the latter hypothesis, fluids originating in a lower level of the basement, once injected into upper levels, could have precipitated quartz in apparent isotopic equilibrium with host rocks, notwithstanding the observed lack of thermal and chemical equilibrium between parent fluids and infiltrated rocks.

Vein emplacement and fluid-rock interaction

Carmignani and Kligfield (1990) demonstrated that the present geological setting of the Apuane Alps is largely controlled by shear zones and other structures developed during the second stage of regional deformation, D2. Vein geometry at Pollone shows a clear link to this deformation, notably to a shear zone cutting across the S. Anna tectonic window. Directions of both high and low angle veins are compatible with an interpretation as Riedel's shear (Fig. 5) and with the vein setting in many ore deposits emplaced along shear zones (Robert and Brown, 1986a; Roberts, 1990 and reference therein).

The variable degree of deformation in the high angle orebodies at Pollone may be the consequence of strain partitioning. However, as these veins may form at any stage of shearing (Roberts, 1990), the variable amount of folding of the high angle orebodies (Fig. 3 and 4) suggests a dynamic development of veins during deformation. Veins formed in the course of the early stages of D2 were subjected to a more pronounced strain than late stage veins which are thus only weakly deformed. The lack of deformation observed in low angle orebodies reflects their attitude with respect to the shear direction. As they lie approximately in the plane of shearing, they cannot be folded but only extended, leading to the observed boudinage. Multiple generations of barite, quartz and sulfides, and the layered textures observed in the low angle orebodies, imply that deformation progress led to progressive growth of preexisting veins through repeated opening and filling episodes, as typically occurs in shear zones (Ramsay, 1980; Moritz and Crocket, 1990). In conclusion mineralized veins at Pollone were emplaced in consequence

54 R Costagliola et al.

of syn-deformation and syn-metamorphic fluid circulation, which was focused by a shear zone affecting the northern sector of the deposit.

The P-T conditions during deformation and vein growth are typically those of the brittle-ductile transition corresponding to greenschist facies metamorphism. At Pollone, metamorphic P-T estimates (about 350 °C and 3.5 kbar) are close to the lowest values that have been proposed for the underlying NMA (Di Pisa et al., 1985; Preite Martinez et al., 1978; Carmignani et al., 1978, 1987; Kligfield, 1979; Cortecci et al., 1994). This evidence suggests that UFP and the NMA belonged to slightly different structural (metamorphic) levels. At a more detailed scale, P-T estimates at Pollone indicate that host and country rocks record metamorphic temperatures of about 350 °C, whereas appreciably higher values (around 450 °C) were found for the mineralizing fluids, as also supported by the Sr geochemistry of the mineralized rocks. This temperature difference could reflect a thermal disequilibrium between mineralizing fluids and rocks, possibly indicating that the fluids were not produced locally, but came from a higher temperature reservoir (deeper level?) than the Pollone rocks.

Fluid inclusion data do not show any evidence of mixing of fluids having different salinifies or temperatures. Departures from the lithostatic toward the hydrostatic pressure regime may be ruled out, taking into account the lack of fluid unmixing evidences in fluid inclusions (cf. Robert and Kelly, 1987) possibly suggesting that, during mineralization, fluid pressure maintained in equilibrium with load pressures. A possible source of Ba for the metamorphogenic vein bodies at Pollone is constituted by the syngenetic orebodies. Pervasive fluid circulation could have had scavenged elements from these bodies reprecipitating them in veins. However, the possible thermal disequilibrium between mineralizing fluids and host rocks may imply also a chemical disequilibrium due to a rapid flux throughout the Pollone structural level. The latter hypothesis is supported by the Rb geochemistry, which suggests that the mineralizing fluid reservoir cannot be recognized in the Scisti di Fornovolasco. On the other hand, shear-related fractures at Pollone represent highly permeable conduits that could have minimized the extent of interaction between an externally derived mineralizing fluid and syngenetic anomalies in the southern sector of the deposit because the unrestricted flow into major structural channelways likely kept the fluid out of contact with its immediate wall rocks. A similar process has been already recognized for fluids moving along major tectonic structures, i.e., along high permeability conduits generated during deformation (Fyfe et al.. 1978; see Oliver et al., 1993 and literature therein; Yardley, 1983; Robert and Brown, 1986a, b). The insulation between mineralizing fluids and wall rocks may have been enhanced by the early precipitation of quartz along the walls of the epigenetic veins at Pollone (cf. Robert and Brown, 1986b; Cathles, 1991). A moderate amount of fluid-rock interaction occurred during the early stage of vein opening, as testified by the silicification of vein walls. However, the amount of alteration is limited, indicating that disequilibrium between mineralizing fluids and wall rocks was minor. This may suggest that the fluid rock reservoir had a similar mineralogical composition with respect to the Pollone host rock. Under such conditions, large amounts of external fluids may affect metamorphic rocks without leaving any detectable geochemical trace (Thompson and Connolly, I992). The lack of O-isotope zonation fl~om

Metamorphogenic barite-pyrite (Pb-Zn-Ag) veins 55

country rocks towards the veins may indicate a closed system fluid circulation, or that the (external) fluid reservoir had a broadly similar O-isotope composition with respect to Pollone host rocks.

In agreement with the geological evidence, the high K/Rb ratio of wall rocks rules out a possible magmatic source of fluids. A metamorphic fluid equilibrated with a rock where K-feldspar reacted to form mica may satisfy the requirement for the observed K, Rb and Sr geochemistry. Unfortunately, these results do not constrain precisely a possible reservoir, but only indicate that mineralizing fluids were (a) not magmatic in origin, (b) out of thermal and, at least in part, chemical equilibrium with the country rocks. In this context, it seems likely that remobilization of Ba from the syngenetic orebodies of the southern sector of Pollone was of minor importance, as for this case clearer equilibrium evidence between the country rocks and the fluids should be expected. Possibly, mineralizing fluids scavenged Ba, and other elements such as Zn, Pb and Ag, which are fairly scarce in the syngenetic orebodies, from other sources. Therefore, the occurrence of epigenetic orebodies in the northern sector at a relatively short distance from the syngenetic ones of the southern sector is more a coincidence rather than a causal relation.

The exact nature and the location of the source(s) of elements in the mineralizing fluids remains undefined. An ideal source of Ba for the Pollone deposit could be represented by a barite and/or Ba-feldspar (pyrite) bearing rock, that is not known in the studied area. This mineral association is relatively uncommon in metamorphic terranes (Clark et al., 1990). However, barite-rich, feldspathic (celsian), pyritiferous layers in amphibolite facies terranes were recently described in the Archean sequences from the Hemlo deposit (Pan and Fleet, 1991). This barite-feldspar rock was considered the element source for extensive syn-metamorphic Ba metasomatism.

Summary and conclusions

In the Pollone area, Tertiary syn-ldnematic metamorphism took place at T and P of about 350 °C and 3.5 kbar, respectively. Fluid circulation was focused mainly along shear zones leading to barite-pyrite-(Pb-Zn-Ag sulfides) precipitation. Barite deposition, silicification, slight Rb depletion and Sr enrichment indicate that the fluids were hotter than and in moderate chemical disequilibrium with the country rocks; these evidences, as well as the differences in Sr-isotope composition between vein and syngenetic barite, suggest an external derivation of the mineralizing fluids, and an overall scarce remobilization of Ba from syngenetic deposits in the area. Mineralizing fluids probably equilibrated at some deeper structural level with siliciclastic basement rocks, scavenging Ba and other elements from a Ba-rich source such as pre-existing ores and/or Ba-feldspars.

The widespread precipitation of quartz and barite, both having a prograde solubility in the investigated P-T-X conditions (Blount, 1977; Fournier, 1985), the lack of evidence for fluid mixing and of extensive alteration in the vein wall rocks suggest that mineral deposition was mainly controlled by a sudden change of the geothermal gradient in this area, i.e., by physical (cooling) rather than chemical factors. Cooling and rapid changes of thermal gradients can be an effective ore

56 P. Costagliola et al.

forming mechanism (Phillips and Powell, 1993), especially for minerals that have low solubility such as barite (Blount, 1977). In metamorphic complexes sudden changes of the thermal gradient are often controlled by tectonics (cf. Connolly and Thompson, 1989). The syn-D2 tectonic collapse of the Apuane Alps core complex caused the juxtaposition along shear zones of rocks belonging to different crustal levels, which conceivably were at different temperatures. According to this interpretation, the gravitational collapse of the Apuane Alps was not only responsible for the vein opening at the Pollone shear zone, but also built up a thermal trap for mineralizing fluids.

Acknowledgments

The research was supported by the Ministero dell'Universit~ e Ricerca Scientifica e Tecnologica (40% and 60% grants to P. Lattanzi and G. Tanelli), and by the Consiglio Nazionale delle Ricerche through the Centro Studi Minerogenesi e Geochimica Applicata. Prof. C. Garbarino assisted with electron microprobe analysis. We thank E Tornos of the Instituto Geominero of Madrid, Spain, for providing the program "Clorita" used for calculation of chlorite temperatures and compositions. The authors are grateful to G. FruehGreen of the ETH Zurich for the help provided with the O-isotope analyses. We thank several referees for their constructive and detailed criticism of various versions of this paper, and Prof. Stumpfl for his painstaking editing.

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Authors' addresses: P. Costagliola, Museo di Mineralogia e Litologia dell'Universith, via La Pira 4, Firenze, Italy; M. Benvenuti, G. Tanelli, Dipartimento di Scienze della Terra, via La Pira 4, Firenze, Italy; P. Lattanzi, Dipartimento di Scienze della Terra, via Trentino 51, Cagliari, Italy.