The PT path of the ultra-high pressure Lago Di Cignana and adjoining high-pressure meta-ophiolitic...

The P–T path of the ultra-high pressure Lago Di Cignana andadjoining high-pressure meta-ophiolitic units: insights into theevolution of the subducting Tethyan slab

C. GROPPO, M. BELTRANDO AND R. COMPAGNONIDepartment of Mineralogical and Petrological Sciences, University of Torino, via Valperga Caluso 35, 10125 Torino, Italy([email protected])

ABSTRACT The Lago di Cignana ultra-high-pressure unit (LCU), which consists of coesite–eclogite faciesmetabasics and metasediments, preserves the most deeply subducted oceanic rocks worldwide. Newconstraints on the prograde and early retrograde evolution of this ultra-high pressure unit and adjoiningunits provide important insights into the evolution of the Piemontese–Ligurian palaeo-subduction zone,active in Paleocene–Eocene times. In the LCU, a first prograde metamorphic assemblage, consisting ofomphacite + Ca-amphibole + epidote + rare biotite + ilmenite, formed during burial at estimatedP < 1.7 GPa and 350 < T < 480 �C. Similar metamorphic conditions of 400 < T < 650 �C and1.0 < P < 1.7 GPa have been estimated for the meta-ophiolitic rocks juxtaposed to the LCU. Theprograde assemblage is partially re-equilibrated into the peak assemblage garnet + omphacite +Na-amphibole + lawsonite + coesite + rutile, whose conditions were estimated at 590 < T <605 �C and P > 3.2 GPa. The prograde path was characterized by a gradual decrease in the thermalgradient from �9–10 to �5–6 �C km)1. This variation is interpreted as the evidence of an increase in therate of subduction of the Piemonte–Ligurian oceanic slab in the Eocene. Accretion of the Piemonteseoceanic rocks to the Alpine orogen and thermal relaxation were probably related to the arrival of morebuoyant continental crust at the subduction zone. Subsequent deformation of the orogenic wedge isresponsible for the present position of the LCU, sandwiched between two tectonic slices of meta-ophiolites, named the Lower and Upper Units, which experienced peak pressures of 2.7–2.8 and<2.4 GPa respectively.

Key words: pseudosection analysis; P–T evolution; subduction zone; ultra-high pressure Lago diCignana Unit; Western Alps.

INTRODUCTION

The study of the P–T evolution of deeply buriedophiolitic units that have been accreted to orogenicbelts may provide important insights into the struc-ture and evolution of palaeo-subduction zones. Sev-eral studies of heat flux, seismic activity or seismictomography carried out on present-day subductionzones have shown that subducting slabs display awide range of geometries and thermal structures. Thegeometry of a subducting plate is the result of theinterplay of several factors, including the viscositystructure of the surrounding mantle (Royden &Husson, 2006), slab width (Schellart et al., 2007),absolute motion of the overriding plate (Royden &Husson, 2006), presence of oceanic plateau, sea-mounts or slices of continental crust (Martinod et al.,2005). The thermal structure of a subduction zonemay also vary widely, particularly in relation to (i)changes in the subduction rate, and (ii) the thermalstructure of the subducting lithosphere, which in turndepends on its age and on the sediment thickness(Peacock & Wang, 1999). Low thermal gradients are

characteristic of fast-subducting slabs, as conductiveheating of the sinking lithosphere is hindered by itshigh vertical velocity (Peacock & Wang, 1999). On theother hand, slow-subducting slabs display lessthermal disequilibrium with the surrounding mantleand, consequently, have higher thermal gradients.Changes in subduction rates are widely reported fromcurrently active subduction zones, where they areinterpreted to result from the complex interplaybetween density and geometry of the slab, and theviscosity structure of the surrounding mantle (Royden& Husson, 2006).

Thanks to the knowledge acquired on the connec-tion between thermal structure and geodynamicbehaviour of present-day subduction zones, attemptscan be made to use the P–T evolution of ophiolites toconstrain the evolution of palaeo-subduction zones.The Lago di Cignana Unit (Western Alps), which wassubducted to depths in excess of 90 km during theEocene (Reinecke, 1991; Rubatto et al., 1998; Lapenet al., 2003; Gouzu et al., 2006), is the oceanic unit thatexperienced the deepest burial among the remnants ofthe Tethys cropping out along the Alpine–Himalayan

J. metamorphic Geol., 2009, 27, 207–231 doi:10.1111/j.1525-1314.2009.00814.x

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belt. A detailed petrological study by means ofpseudosection analysis and conventional thermoba-rometry has been undertaken on metabasics from boththe ultra-high pressure (UHP) Lago di Cignana Unitand its adjoining units, to constrain their P–T pathsand to give new information about the thermal evo-lution of the Piemontese–Ligurian subduction zone.On the basis of these results, combined with consi-derations on middle Eocene plate kinematics in theWestern Tethys, it is proposed that the metamorphicevolution of the UHP Lago di Cignana Unit and itsadjoining units record an increase in the rate of sub-duction, followed by the accretion of oceanic units tothe Alpine chain, probably as a consequence of thearrival of more buoyant continental crust at the sub-duction zone.

GEOLOGICAL SETTING

The study area is located in the Piemonte Zone ofcalcschists with meta-ophiolites, which is commonlyinterpreted as a remnant of the Piemonte Ocean thatopened in the Late Jurassic between the Europeancontinent to the NW and the Apulia plate to the SE(Dal Piaz, 1974; Dewey et al., 1989; Polino et al., 1990;Lombardo et al., 2002 and references therein). In theupper Valtournenche (Val d�Aosta), the PiemonteZone consists of a pile of tectonic slivers includingboth epidote–blueschist to lawsonite–blueschist facies(�Combin Zone�) and eclogite facies (�Zermatt–SaasZone�) Alpine metamorphic rocks (Bearth, 1967).These thin units of meta-ophiolites and metasedimentsare sandwiched between the overlying AustroalpineDent Blanche unit and the underlying Penninic MonteRosa unit, both consisting of continental crust(Fig. 1a).

In the southern side of the lower Conca di Cignana,upper Valtournenche, the discovery of coesite in the1990s (Reinecke, 1991, 1998) led to the identification ofan UHP metamorphic unit, named Lago di CignanaUnit (LCU). Detailed geological mapping of the arearevealed that the LCU is very thin (less than �200 m)and is dismembered into three main flat-lying lensesmeasuring �1400, 350 and 300 m along their N–S axes(Forster et al., 2004) (Fig. 1b). The LCU consists ofeclogite facies metabasics and metasedimentary rocks,which are thought to represent a coherent segment offormer oceanic crust (Compagnoni & Rolfo, 2003).Coesite-bearing eclogites are variably retrogressed togreenschist facies prasinite (van der Klauw et al.,1997).

The LCU is sandwiched between two tectonic slicesof meta-ophiolites, which hereafter will be referred toas Upper (UU) and Lower Unit (LU) respectively(Fig. 1b,c). While the LU is generally considered to bepart of the Zermatt–Saas Zone, the UU has beenalternatively attributed to the Combin Zone (Forsteret al., 2004) and to the Zermatt–Saas Zone (Pleugeret al., 2007). In this paper, we provide P–T data sup-

porting this latter interpretation, which was originallybased on structural arguments concerning the positionof the contact between the Combin and Zermatt–Saasunits. Pleuger et al. (2007) placed this contact at thebottom of the Cime Bianche slice, which overlies theUU, and named it �Combin thrust� (Fig. 1b).The prograde and retrograde P–T path of the LCU

has been described in detail by several authors. Theprograde history has been reconstructed using growthzoning and mineral inclusions of garnet in botheclogites (van der Klauw et al., 1997) and metasedi-ments (Reinecke, 1991). King et al. (2004) showed thatgarnet from eclogites of the LCU preserves trace ele-ment evidence of discontinuous reactions (such as thebreakdown of clinozoisite + titanite to grossu-lar + rutile + quartz ⁄ coesite + H2O) that occurredduring its prograde growth. Peak metamorphic con-ditions have been estimated at 615 ± 15 �C and2.8 ± 1.0 GPa (Reinecke, 1991, 1998; Reinecke et al.,1994) from metasediment parageneses. Similar peakconditions have been obtained from eclogites (Rei-necke et al., 1994), and the exhumation history of theunit has been reconstructed on the basis of micro-structural analysis of metabasics (van der Klauw et al.,1997). Lu–Hf and Sm–Nd garnet geochronology sug-gest a c. 50 and c. 40 Ma for the age of the progradeand peak metamorphic events respectively (Amatoet al., 1999; Lapen et al., 2003). Older ages of c. 44 Mahave been estimated for the UHP metamorphism bothwith in situ U ⁄Pb dating of metamorphic zircon(Rubatto et al., 1998) and with 40Ar ⁄ 39Ar high-resolution laser dating of phengitic mica (Gouzu et al.,2006).As to the Zermatt–Saas Zone, peak metamorphic

conditions for the eclogite facies rocks have beenoriginally estimated with conventional thermobaro-metry at 450–700 �C and 1.0–2.2 GPa (Ernst & DalPiaz, 1978; Oberhansli, 1980, 1982; Meyer, 1983a,b;Barnicoat & Fry, 1986; Cartwright & Barnicoat, 2002),i.e. at pressures significantly lower than those esti-mated for the LCU. However, more recent data basedon pseudosection analysis (Bucher et al., 2005) indicatepeak conditions of 550–600 �C and 2.5–3.0 GPa, justat the boundary between quartz–eclogite and coesite–eclogite facies.The metamorphic evolution of the Combin Zone has

not been well constrained so far. The sequence ofcalcschist with subordinate meta-ophiolite is charac-terized by a pervasive greenschist facies overprint (DalPiaz & Ernst, 1978). However, the presence ofglaucophane included in garnet and epidote from ametabasite (Dal Piaz, 1974; Le Goff, 1986; Wagner-Zweigel, 1993) is indicative of HP metamorphism priorto the greenschist facies overprint. Flattened rhombsconsisting of chlorite + carbonate, most probablypseudomorphous after former lawsonite, have beenreported from metabasic rocks, for which peak meta-morphic conditions of 450 �C and 1.2 GPa (i.e. P–Tconditions considerably lower than those estimated for

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the Zermatt–Saas Zone) have been estimated (Cart-wright & Barnicoat, 2002).

METABASALTS

In this work, three metabasics sampled from the LCUand the LU and UU respectively have been studied indetail (Fig. 1b,c). In the field, metabasics from all theunits mainly occur as deformed lenses and boudins,several metres long and up to few metres across,bounded by a flat-lying schistosity. Eclogite OF2512(LCU) has been collected from the UHP LCU, justsouth of the Lago di Cignana dam; eclogite OF2595(LU) comes from an outcrop localized to the SE ofPalud, in the upper part of the quartz–eclogite faciesLU; and metabasite OF2933 (UU) from the UU, onthe southern slope of Mt. Pancherot (Fig. 1b). Thethree outcrops are located within a distance of <1 km.

Samples OF2512 and OF2595, coming from theLCU and LU respectively, have been selected out ofseveral tens of fresh eclogites, all studied in thin section,and are similar to eclogites that have been previouslydescribed by several authors (e.g. van der Klauw et al.,1997; Compagnoni & Rolfo, 2003; King et al., 2004)thus confirming that the chosen samples are represen-tative of the eclogites commonly occurring in the LCUand in the adjacent units. They are foliated and lineatedeclogites, and the main foliation in the matrix is definedby the preferred orientation of the omphacite nemato-blasts, developed during a UHP ⁄HP deformation event(Muller & Compagnoni, 2008). In contrast, in the UU,all the eclogite boudins are considerably retrogressed togreenschist facies mineral assemblages and no pre-served eclogites have been discovered, resulting in itsattribution to the Combin Zone (Forster et al., 2004).Our OF2933 (UU) sample, however, shows fresh garnet

(a) (b)

(c)

Fig. 1. (a) Simplified tectonic sketch-map of the Italian Western Alps. Helvetic Domain: Mont Blanc-Aiguilles Rouges (MB); PenninicDomain: Grand St Bernard Zone (SB), and Monte Rosa (MR), Gran Paradiso (GP), Dora-Maira (DM) and Valosio (V) InternalCrystalline Massifs; the Piemonte Zone of Calcschists with meta-ophiolites is shown in light (calcschists) and dark (meta-ophiolites)grey respectively; Austroalpine Domain: Dent Blanche nappe (DB), Mont Emilius klippe (ME) and Sesia–Lanzo Zone (SZ); SouthernAlps (SA); Embrunais-Ubaye Flysch nappe (EU); Canavese line (CL); Sestri-Voltaggio line (SVL). The Lago di Cignana region isshown in the box (see the white arrow). (b) Geological map of the Lago di Cignana region, modified after Compagnoni et al. (2000),Forster et al. (2004) and Pleuger et al. (2007). Austroalpine Domain (Dent Blanche Nappe): (1) Valpelline Series; (2) Roisan Zone; (3)Arolla Series. Pennine Domain (Piemonte Zone): (4) Combin Zone; (5) Pancherot – Cime Bianche – Bettaforca Unit; (6) Zermatt–SaasZone serpentinite; (7) Lago di Cignana Unit (LCU); (8) Zermatt–Saas Zone eclogite and metagabbro. The white stars show loca-tions of the studied samples. (c) Simplified geological cross-section through the UHP LCU, modified from Forster et al. (2004).

P – T PA TH O F UH P L A G O D I C I G NA N A AN D A D J O I N I N G H P U N I T S 20 9


porphyroblasts already visible in hand specimen, andthe P–T estimates presented below support the attri-bution of the UU to the Zermatt–Saas unit as proposedby Pleuger et al. (2007).

Bulk-rock compositions of the studied samples wereacquired by two different methods: (i) as average of 10SEM-EDS analyses on 4.70 · 3.20 mm areas, and (ii)ICP-MS analysis by ALS Chemex, Vancouver, Can-ada. The results, reported in Table 1 and Fig. 2, andcompared with that of a typical Mid Ocean RidgeBasalts (MORB) from the Mid-Atlantic Ridge(Carmichael et al., 1974, p. 376, col. 1), show that:

1 The two methods, area-scan at SEM-EDS andICP-MS, give very similar results suggesting that thearea-scan method is a reliable alternative approach toobtain bulk-rock chemical compositions suitable forpseudosection modelling, provided the sample ishomogeneous and fine-grained;

2 Major element compositions of eclogites OF2512(LCU) and OF2595 (LU) are comparable with that ofa typical MORB, thus suggesting a mid-ocean ridgetholeiitic protolith for both samples (see, e.g. Bearth &Stern, 1971, 1979; Pfeifer et al., 1989). The slightenrichment in Na2O is consistent with an incipientlow-temperature albitization of primary plagioclaseduring ocean-floor metamorphism (Pfeifer et al.,1989); however, the studied samples do not showcompositions in the spilite field as defined by Mullen(1983) (Na2O ⁄CaO > 0.66). Sample OF2933 (UU) isenriched in Ca with respect to typical MORB (highmodal amounts of epidote and carbonate) and is morehydrated, showing affinities with the interpillowmaterial described by Bearth & Stern (1979) andBarnicoat & Fry (1986).3 Alteration-discriminant element abundances, withNb ⁄Y (0.13–0.16), Zr ⁄ (P2O5 · 104) (0.05–0.07) andP2O5 (0.30–0.39 wt%) confirm the oceanic tholeiiticparentage for all the studied samples (Floyd & Win-chester, 1975). Ba ⁄Zr (0.04–0.07) and Zr ⁄Y (4.56–5.08)ratios are intermediate between the values ofnormal-MORB (N-MORB) and transitional-MORB(T-MORB) (Saunders & Tarney, 1984).

PETROGRAPHY AND MINERAL CHEMISTRY

Garnet, omphacite, amphibole, phengite, epidote andbiotite from the three metabasite samples were analy-sed with a Cambridge Stereoscan 360 SEM equippedwith an EDS Energy 200 and a Pentafet detector(Oxford Instruments) at Department of Mineralogicaland Petrological Sciences, University of Torino (Italy).The operating conditions were: 50 s counting time and15 kV accelerating voltage. SEM-EDS quantitativedata (spot size = 2 lm) were acquired and managedusing the Microanalysis Suite Issue 12, INCA Suiteversion 4.01; the raw data were calibrated on naturalmineral standards and the FqZ correction (Pouchou &Pichoir, 1988) was applied. Representative SEM-EDSanalyses of these minerals are reported in Tables 2–5.The Fe3+ ⁄Fetot ratio in garnet, omphacite andamphibole has been estimated by stoichiometry fromthe SEM-EDS analyses.Mineral abbreviations are from Bucher & Frey

(2002) and Fettes & Desmons (2007).

Table 1. Bulk compositions of the studied metabasites.

Analyses Sample Bulk composition (wt%) Trace elements (ppm)

SiO2 TiO2 Al2O3 FeOtot MnO MgO CaO Na2O K2O P2O5 LOI Total Ba Cr Cu Dy Ga La Nb Ni Pb Sr V Y Yb Zn Zr

SEM-EDS OF2512 50.77 2.22 17.46 9.17 – 5.48 9.77 4.80 0.34 – – 100.0

ICP-MS OF2512 50.20 2.24 15.35 9.77 0.16 5.82 8.88 4.62 0.26 0.39 0.57 98.3 10.1 100 20 9.38 23.8 10.0 7.9 61 <5 137 264 50.2 5.47 118 255

SEM-EDS OF2595 49.60 1.65 17.15 10.10 – 4.67 11.52 5.30 – – – 100.0

ICP-MS OF2595 48.90 1.73 16.20 9.72 0.17 5.07 11.25 5.23 0.11 0.30 0.85 99.5 13.5 200 72 7.93 20.3 7.7 5.5 77 <5 174 218 42.1 4.52 96 192

SEM-EDS OF2933 49.81 1.96 16.03 8.62 – 5.32 14.29 3.43 0.56 – – 100.0

ICP-MS OF2933 48.00 1.91 14.50 7.36 0.15 4.74 13.65 2.60 0.29 0.36 5.73 99.3 12.9 120 31 7.90 23.3 8.3 5.9 67 <5 372 254 42.0 4.67 82 197

MORBa 49.20 2.03 16.10 10.22 0.18 6.44 10.50 3.01 0.14 97.8 5 220 87 1 190 280 160

aMid-Atlantic Ridge; Carmichael et al. (1974) , p.376, col. 1.

Fig. 2. Projections of measured bulk-rock compositions (arealanalyses at SEM-EDS and ICP-MS data) and determined min-eral compositions (omphacite, garnet and glaucophane in thepeak assemblages) onto the Na2O–Al2O3–Ca(Mg,Fe)Si2O6

plane, projected from water and quartz. Bulk composition of atypical MORB from the Mid-Oceanic Ridge (Carmichael et al.,1974) is shown for comparison. Note that measured composi-tions are compatible with the observed mineral assemblages.

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Eclogite OF2512 (LCU)

Eclogite OF2512 (LCU) is a fine-grained foliated rockmainly consisting of omphacite, garnet, blue amphi-bole, minor phengite and accessory rutile. The matrixfoliation (Sm) is mainly defined by the alignment ofelongated omphacite crystals and curves around

porphyroblastic garnet up to 5 mm in diameter(Fig. 3a).

Garnet

A complex growth history can be inferred for gar-net porphyroblasts on the basis of their strong

Table 2. Representative analyses of garnet, omphacite, phengite, biotite and epidote from eclogite OF2512.

Garnet Omphacite Phengite Biotite Epidote

In matrix In Grt

Anal. 3.1a

rim

3.3

mantle

3.5 3.7 3.9 3.11

core

1.20a

core

1.23

rim

1.24

outer rim

2.29 Grt

core

8.31 Grt

rim

6.1a 6.3 4.1

after Lw

2.39

Bt I

3.42

Bt II

3.35

Ep I

2.55

Ep II

SiO2 38.03 37.03 36.77 36.48 36.80 37.09 SiO2 56.84 56.13 56.16 56.14 55.14 SiO2 54.05 54.00 51.48 36.38 35.31 SiO2 37.73 38.01

Al2O3 20.98 20.70 20.25 20.46 20.48 20.28 Al2O3 12.65 11.38 10.96 10.61 8.83 TiO2 0.00 0.00 0.00 1.11 0.00 Al2O3 24.66 27.77

FeO 28.28 30.96 30.23 29.13 27.37 24.70 FeO 3.97 4.97 5.04 7.33 8.08 Al2O3 24.63 24.63 27.83 16.61 19.24 FeO 11.16 6.89

MnO 0.00 0.00 1.02 1.59 3.17 5.53 MgO 6.86 7.36 7.81 7.53 7.13 FeO 2.08 1.87 2.70 20.09 21.02 MnO 0.00 0.00

MgO 6.57 3.41 2.68 2.14 1.73 1.29 CaO 11.22 12.60 14.03 12.45 13.18 MnO 0.00 0.00 0.00 0.00 0.00 MgO 0.00 0.00

CaO 5.91 7.52 8.99 9.91 10.86 11.44 Na2O 8.26 7.58 7.09 7.31 6.92 MgO 4.46 4.67 3.22 11.37 11.15 CaO 24.75 25.46

Total 99.77 99.62 99.95 99.71 100.41 100.33 Total 99.80 100.02 101.09 101.37 99.28 Na2O 0.00 0.32 0.82 0.57 0.61 Total 98.31 98.14

K2O 10.97 10.86 10.98 8.96 8.13

Si 2.97 2.95 2.93 2.92 2.92 2.95 Si 2.01 1.99 1.98 2.00 1.99 Total 96.19 96.35 97.03 95.09 95.46 Si 2.98 2.97

Al 1.93 1.94 1.90 1.93 1.92 1.90 Al 0.53 0.48 0.45 0.45 0.38 Al 2.30 2.55

Fe3+ 0.14 0.16 0.24 0.24 0.24 0.19 Fe3+ 0.03 0.07 0.08 0.06 0.13 Si 3.56 3.55 3.39 2.78 2.69 Fe3+ 0.66 0.40

Fe2+ 1.71 1.90 1.77 1.70 1.58 1.45 Fe2+ 0.09 0.08 0.07 0.12 0.13 AlIV 0.44 0.45 0.61 1.22 1.31 Fe2+ 0.00 0.00

Mn 0.00 0.00 0.07 0.11 0.21 0.37 Mg 0.36 0.39 0.41 0.40 0.38 AlVI 1.47 1.46 1.55 0.28 0.41 Mn 0.00 0.00

Mg 0.76 0.40 0.32 0.25 0.21 0.15 Ca 0.42 0.48 0.53 0.48 0.51 Ti 0.00 0.00 0.00 0.06 0.00 Mg 0.00 0.00

Ca 0.49 0.64 0.77 0.85 0.92 0.98 Na 0.57 0.52 0.48 0.50 0.48 Fe2+ 0.11 0.10 0.15 1.28 1.34 Ca 2.10 2.13

Mn 0.00 0.00 0.00 0.00 0.00

XCa 0.16 0.21 0.25 0.28 0.31 0.35 XNa 0.57 0.52 0.48 0.52 0.49 Mg 0.44 0.46 0.32 1.30 1.26 XZo 0.28 0.52

XFe 0.59 0.66 0.65 0.64 0.62 0.59 XFe 0.25 0.27 0.27 0.31 0.40 Na 0.00 0.04 0.10 0.08 0.09 XEp 0.66 0.41

XMg 0.25 0.13 0.10 0.08 0.06 0.05 XAeg 0.03 0.07 0.08 0.06 0.13 K 0.92 0.91 0.92 0.87 0.79

XMn 0.00 0.00 0.02 0.03 0.07 0.12 XJd 0.53 0.45 0.41 0.44 0.35

aCompositions of garnet rim and omphacite and phengite cores used for the thermobarometric estimates of the peak metamorphic event.

Table 3. Representative analyses of amphibole from eclogites OF2512 and OF2595 and metabasite OF2933.

OF2512 OF2595 OF2933

In Grt core In Grt mantle Blue amphibole Symplectites In Grt core Blue amphibole Sympl Around Grt In Grt

Anal. 2.33 2.41 3.3 3.33 1.13

core

4.5 4.4

outer rim

1.8

after Gln

2.26

after Omp

10.25 9.25 7.3

core

7.6

rim

7.7

outer rim

1.18

after Omp

10.33 2.41 2.42

SiO2 42.44 41.37 43.03 45.31 59.78 59.08 47.51 52.28 49.42 40.57 41.80 57.18 56.96 42.16 44.56 45.58 55.50 53.74

TiO2 0.84 0.00 0.47 0.00 0.00 0.00 0.43 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Al2O3 10.73 14.27 13.89 13.42 11.68 13.24 11.70 5.02 6.30 13.15 12.77 10.41 10.60 15.42 10.38 11.47 3.50 2.73

FeO 23.33 20.32 22.83 21.09 6.36 6.16 15.82 9.59 19.51 21.61 20.35 8.83 11.65 19.09 17.49 17.92 9.78 11.18

MnO 0.56 0.51 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.84 0.74 0.01 0.01 0.01 0.01 0.01 0.00 0.00

MgO 6.88 7.64 5.10 6.59 12.74 12.28 10.27 16.79 10.50 7.22 7.22 11.37 9.37 6.90 10.44 8.65 17.68 16.91

CaO 9.43 9.62 7.37 7.25 0.58 0.47 8.77 12.91 9.50 7.37 7.54 0.91 0.89 7.75 9.67 6.81 11.14 11.44

Na2O 3.87 4.04 4.95 5.00 7.17 6.62 4.00 1.56 3.05 5.47 5.15 6.98 6.89 4.93 3.73 5.00 1.38 1.20

K2O 0.30 0.43 0.00 0.00 0.00 0.00 0.22 0.00 0.00 0.01 0.01 0.01 0.01 0.51 0.27 0.30 0.00 0.00

Total 98.38 98.20 97.64 98.66 98.31 97.85 98.71 98.15 98.28 98.04 96.89 96.01 96.62 97.31 97.12 96.80 98.98 97.20

Si 6.42 6.18 6.47 6.65 7.94 7.83 6.85 7.45 7.23 6.12 6.35 7.92 7.95 6.33 6.66 6.78 7.66 7.63

Ti 0.10 0.00 0.05 0.00 0.00 0.00 0.05 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Al 1.91 2.51 2.46 2.32 1.83 2.07 1.99 0.84 1.09 2.34 2.29 1.70 1.74 2.73 1.83 2.01 0.57 0.46

Fe3+ 0.81 0.79 0.68 0.68 0.28 0.45 0.35 0.00 0.60 1.27 0.94 0.30 0.23 0.59 0.62 0.64 0.69 0.86

Fe2+ 2.14 1.75 2.20 1.91 0.43 0.24 1.56 1.25 1.79 1.46 1.64 0.72 1.13 1.81 1.56 1.59 0.44 0.46

Mn 0.07 0.07 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.11 0.10 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Mg 1.55 1.70 1.14 1.44 2.52 2.42 2.21 3.57 2.29 1.62 1.63 2.35 1.95 1.54 2.33 1.92 3.64 3.58

Ca 1.53 1.54 1.19 1.14 0.08 0.07 1.35 1.97 1.49 1.19 1.23 0.14 0.13 1.25 1.55 1.09 1.65 1.74

Na 1.13 1.17 1.44 1.42 1.85 1.70 1.12 0.43 0.87 1.60 1.52 1.88 1.86 1.43 1.08 1.44 0.37 0.33

K 0.06 0.08 0.00 0.00 0.00 0.00 0.04 0.00 0.00 0.00 0.00 0.00 0.00 0.10 0.05 0.06 0.00 0.00

Na(M4) 0.47 0.46 0.81 0.86 1.92 1.93 0.65 0.03 0.51 0.81 0.77 1.87 1.87 0.75 0.45 0.91 0.35 0.26

Na(A) 0.66 0.71 0.63 0.56 0.00 0.00 0.47 0.40 0.36 0.79 0.74 0.01 0.00 0.68 0.63 0.53 0.02 0.07

AlVI 0.33 0.69 0.93 0.97 1.77 1.89 0.84 0.29 0.32 0.46 0.63 1.63 1.69 1.06 0.49 0.80 0.23 0.09

AlIV 1.58 1.82 1.53 1.35 0.06 0.17 1.15 0.55 0.77 1.88 1.65 0.08 0.05 1.67 1.34 1.22 0.34 0.37

XMg 0.34 0.40 0.28 0.36 0.78 0.78 0.54 0.74 0.49 0.37 0.39 0.70 0.59 0.39 0.52 0.46 0.76 0.73



compositional zoning and the distribution of inclu-sions. Three main zones have been recognized(Fig. 3a): a garnet core crowded with very smallinclusions of omphacite, Ca-amphibole, epidote,ilmenite and very rare biotite, locally defining aninternal foliation (Sm)1); a garnet mantle character-ized by coarser-grained, randomly oriented inclu-sions, concentrated at the core–mantle interface,of NaCa-amphibole, omphacite, quartz, rutile andlozenge-shaped epidote + paragonite ± phengiteaggregates after former lawsonite; a garnet rim, very

thin (50–100 lm) compared to the garnet coreand mantle, free of inclusions. It is worth notingthat quartz is absent in the inclusion-rich core,whereas it is present in the mantle. Neither quartznor coesite inclusions have been observed in thegarnet rim.The garnet core, mantle and rim, as defined above,

display also different compositions. A compositionalprofile through a garnet porphyroblast is reported inFig. 4(a). The Ca and Mn contents progressivelydecrease from core to rim (XCa: 35–24 in the core,

Table 4. Representative analyses of garnet, omphacite and epidote from eclogite OF2595.

Garnet Omphacite Epidote

In matrix In garnet

Anal. 10.1a

rim

10.3 10.5 10.7 10.9

core

2.1

core

2.2a

mantle

2.5

rim

2.6

outer rim

9.20

Grt core

9.22

Grt rim

9.35

Ep I

10.29

Ep II

SiO2 37.90 37.63 37.29 37.76 37.38 SiO2 56.26 56.55 54.58 56.06 55.35 56.63 SiO2 39.37 38.07

Al2O3 20.75 20.49 20.03 20.32 20.34 Al2O3 9.52 10.18 10.37 9.10 5.86 10.19 Al2O3 29.19 24.55

FeO 29.33 28.49 28.81 25.26 24.61 FeO 9.00 8.23 9.16 9.58 8.55 7.03 FeO 5.91 12.34

MnO 1.00 1.32 2.73 4.40 6.95 MgO 6.28 6.13 7.29 6.64 9.46 6.99 MnO 0.00 0.00

MgO 3.48 2.96 2.12 2.70 2.03 CaO 10.76 10.04 11.65 11.83 15.41 11.16 MgO 0.00 0.00

CaO 8.14 9.36 8.78 9.34 8.90 Na2O 8.10 8.52 7.02 7.55 5.44 7.95 CaO 23.96 23.12

Total 100.60 100.27 99.76 99.79 100.21 Total 99.92 99.65 100.07 100.76 100.07 99.95 Total 98.43 98.07

Si 2.98 2.98 2.99 3.00 2.98 Si 2.02 2.01 2.02 2.02 2.00 2.02 Si 3.03 3.01

Al 1.93 1.91 1.89 1.90 1.91 Al 0.40 0.43 0.45 0.39 0.25 0.43 Al 2.64 2.29

Fe3+ 0.11 0.14 0.13 0.09 0.13 Fe3+ 0.13 0.14 0.00 0.11 0.13 0.09 Fe3+ 0.34 0.73

Fe2+ 1.82 1.74 1.79 1.59 1.51 Fe2+ 0.14 0.13 0.15 0.15 0.13 0.12 Fe2+ 0.00 0.00

Mn 0.07 0.09 0.19 0.30 0.47 Mg 0.34 0.32 0.40 0.36 0.51 0.37 Mn 0.00 0.00

Mg 0.41 0.35 0.25 0.32 0.24 Ca 0.41 0.38 0.46 0.46 0.60 0.43 Mg 0.00 0.00

Ca 0.69 0.79 0.75 0.80 0.76 Na 0.56 0.59 0.50 0.53 0.38 0.55 Ca 1.97 1.96

XCa 0.24 0.27 0.27 0.29 0.30 XNa 0.58 0.61 0.52 0.54 0.39 0.56 XZo 0.64 0.29

XFe 0.62 0.60 0.64 0.59 0.60 XFe 0.44 0.45 0.28 0.42 0.34 0.36 XEp 0.34 0.72

XMg 0.14 0.12 0.09 0.12 0.10 XAeg 0.13 0.14 0.00 0.11 0.13 0.09

XMn 0.02 0.03 0.06 0.10 0.16 XJd 0.40 0.43 0.45 0.39 0.25 0.43

aCompositions of garnet rim and omphacite mantle used for the thermobarometric estimates of the peak metamorphic event.

Table 5. Representative analyses of garnet, phengite and epidote from metabasite OF2933.

Garnet Phengite Epidote

Ep I Ep II Ep III

Anal. 2.1a

rim

2.4 2.7 2.11 2.15 2.18

core

3.12a

core

3.20 3.14

rim

1.21

In Grt

2.45

In Grt

2.45

In Grt

1.28

core

1.27

rim

5.12

rim

SiO2 37.18 36.84 36.84 36.87 36.80 36.71 SiO2 50.02 48.97 48.53 SiO2 37.19 37.25 37.72 37.76 37.10 36.87

Al2O3 20.86 21.11 20.98 21.06 21.09 20.68 TiO2 0.00 0.00 0.35 Al2O3 24.77 25.69 28.12 28.77 23.64 23.68

FeO 27.01 26.88 28.39 27.07 27.51 25.83 Al2O3 26.45 27.11 28.70 FeO 11.79 10.29 6.99 6.43 13.26 12.49

MnO 0.60 0.00 0.55 1.11 1.95 2.66 FeO 2.24 1.80 2.12 MnO 0.00 0.00 0.00 0.00 0.00 0.00

MgO 2.02 2.04 2.19 1.83 1.80 1.31 MnO 0.00 0.00 0.00 MgO 0.00 0.00 0.00 0.00 0.00 0.00

CaO 11.71 11.92 10.33 12.16 10.81 12.01 MgO 3.90 3.57 2.71 CaO 24.32 24.05 24.85 24.31 23.65 23.80

Total 99.38 98.79 99.28 100.10 99.96 99.20 Na2O 0.50 0.46 0.64 Total 98.06 97.28 97.69 97.72 97.65 96.84

K2O 10.35 10.14 10.11

Si 2.96 2.95 2.94 2.92 2.93 2.94 Total 93.46 92.05 93.16 Si 2.95 2.96 2.95 2.94 2.97 2.97

Al 1.96 1.99 1.98 1.97 1.98 1.95 Al 2.32 2.41 2.59 2.64 2.23 2.25

Fe3+ 0.12 0.12 0.14 0.20 0.17 0.16 Si 3.41 3.38 3.31 Fe3+ 0.70 0.62 0.41 0.38 0.80 0.76

Fe2+ 1.68 1.68 1.76 1.60 1.66 1.57 AlIV 0.59 0.62 0.69 Fe2+ 0.00 0.00 0.00 0.00 0.00 0.00

Mn 0.04 0.00 0.04 0.07 0.13 0.18 AlVI 1.53 1.58 1.62 Mn 0.00 0.00 0.00 0.00 0.00 0.00

Mg 0.24 0.24 0.26 0.22 0.21 0.16 Ti 0.00 0.00 0.02 Mg 0.00 0.00 0.00 0.00 0.00 0.00

Ca 1.00 1.02 0.88 1.03 0.92 1.03 Fe2+ 0.13 0.10 0.12 Ca 2.07 2.05 2.08 2.03 2.03 2.05

Mn 0.00 0.00 0.00

XCa 0.32 0.33 0.29 0.33 0.30 0.33 Mg 0.40 0.37 0.28 XZo 0.27 0.37 0.55 0.61 0.20 0.22

XFe 0.58 0.59 0.62 0.58 0.59 0.56 Na 0.07 0.06 0.08 XEp 0.70 0.62 0.41 0.38 0.80 0.76

XMg 0.08 0.08 0.08 0.07 0.07 0.05 K 0.90 0.89 0.88

XMn 0.01 0.00 0.01 0.02 0.04 0.06

aCompositions of garnet rim and phengite core used for the thermobarometric estimates of the peak metamorphic event.

2 12 C . G R O P P O E T A L .


26–21 in the mantle and 19–16 in the rim; XMn: 15–2 inthe core, 2–0 in the mantle and 0 in the rim), whereasthe Mg content gradually increases outward (XMg: 5–10 in the core, 11–17 in the mantle and 17–25 in therim). The Fe content is more variable: it increases fromcore to mantle (XFe = 56–65 in the core, 65–68 in themantle) and decreases from mantle to rim (XFe = 65–59 in the rim) (Table 2), resulting in a Fe ⁄Mgratio decreasing from 11 in the core to 2 in therim [XCa = Ca ⁄ (Ca + Mg + Fe)·100; XMn = Mn ⁄(Mn + Ca + Fe + Mg) · 100; XMg = Mg ⁄ (Mg +Fe + Ca) · 100; XFe = Fe ⁄ (Fe + Mg + Ca) · 100].The contact between garnet core and mantle is marked

by an increase in XFe, whereas the contact betweenmantle and rim is marked by a sharp increase in XMg.Chlorite, which replaces garnet at the rim, is in turnpartially replaced by biotite just at the contact withgarnet (Fig. 3a).

Omphacite

Omphacite occurs as elongated crystals both in thematrix and in the garnet (Fig. 3a,d). Omphacite inthe matrix includes rutile and is locally replaced at themargins by a symplectite of Ca-amphibole + albite(Fig. 3d). Omphacite is strongly zoned, with XNa

(a)

(b) (c)

(d)

(e)

Fig. 3. Representative microstructures of eclogite OF2512 (LCU) as seen under optical microscope and SEM. (a) Strongly zonedporphyroblastic garnet (Grt) set in an omphacite (Omp) + glaucophane (Gln) matrix. In garnet, three main zones are evident: a core,characterized by an internal foliation (Sm)1) mainly defined by very fine-grained inclusions of omphacite + amphibole; a mantle, withcoarser-grained randomly oriented inclusions of omphacite, amphibole (Am), former lawsonite, quartz (Qtz) and rutile (Rt); a verythin, inclusion-free rim. Garnet is partially replaced by biotite (Bt). Plane Polarized Light (PPL). (b) Coarse-grained glaucophane (Gln)in the matrix, partially replaced by a symplectite of Ca-amphibole + albite (Sympl) (PPL). (c) Glaucophane (Gln) rimmed with green-blue NaCa-amphibole (Am) and a phengite (Phg) flake in the matrix (PPL). (d) Back-Scattered Electron (BSE) image of zonedomphacite (Omp) from the matrix. (e) BSE image of a zoned glaucophane (Gln) rimmed with blue–green amphibole (Am) and partiallyreplaced by a green amphibole + albite symplectite (Sympl). Rutile (Rt) inclusions occur in the glaucophane core.



[XNa = Na ⁄ (Na + Ca) · 100] ranging from 57 in thecore (XJd = 48–53, XAeg = 1–5) to 51 in the rim(XJd = 45–47, XAeg = 5–7) (Fig. 5a & Table 2).Locally, a thin rim, enriched in the aegirine component(XNa = 45–48; XJd = 41–45, XAeg = 6–8), discon-tinuously overgrows omphacite.

Omphacite included in both garnet cores andmantles is relatively homogeneous as to the XNa

(XNa = 48–52); higher XFe is the characteristic ofomphacite included in garnet mantles (XFe = 31–33 ingarnet cores; XFe = 36–46 in garnet mantles) [XFe =Fe2+ ⁄ (Fe2+ + Mg) · 100] (Fig. 5a & Table 2).

Amphibole

Different amphibole generations have been recog-nized on the basis of microstructural relationshipsand mineral chemistry: (i) a green and (ii) a bluish–green amphibole included in garnet cores and man-tles respectively (Fig. 3a); iii) a blue amphibole in thematrix (Fig. 3b,c,e); and iv) a fine-grained green

amphibole in the symplectites after blue amphiboleand omphacite (Fig. 3d,e). Fine- to coarse-grainedgreen to bluish–green amphibole included in garnetcores and mantles occur as both single and multipleinclusions in association with omphacite. Amphiboleincluded in garnet cores has a lower NaM4 con-tent (NaM4 = 0.32–0.62) with respect to that ingarnet mantles (NaM4 = 0.67–0.97), whereas AlIV isrelatively constant (AlIV = 1.07–1.86) (Fig. 6b &Table 3). Ca-amphibole included in garnet cores isa tschermakite–magnesiohornblende, whereas NaCa-amphibole in garnet mantle is a magnesiotaramite–magnesiokatophorite (Leake et al., 1997).A coarse-grained Na-amphibole is found in the

matrix, where it defines together with omphacite themain foliation. It is strongly zoned, with a pale blueglaucophane core and a discontinuous bluish–greenbarroisite rim, and often includes rutile (Fig. 3b,c).The NaM4 and AlIV contents in glaucophane rangefrom 1.71 to 1.95 and from 0 to 0.21 respectively(Fig. 6b & Table 3).

(a) (b)

(d)(c)

Fig. 4. (a) Zoning profile of a porphyroblas-tic garnet from eclogite OF2512 (LCU). Thecore-mantle contact is marked by an increasein XFe, whereas the mantle-rim contact by asharp increase in XMg [XCa = Ca ⁄ (Ca +Fe + Mg + Mn) · 100; XFe = Fe ⁄ (Ca +Fe + Mg + Mn) · 100; XMg = Mg ⁄ (CaFe + Mg + Mn) · 100; XMn = Mn ⁄ (Ca+ Fe + Mg + Mn) · 100]. (b) Simulatedgarnet zoning for eclogite OF2512 (LCU)along a straight line from 500 �C, 2.0 GPa to600 �C, 3.3 GPa. (c) Zoning profile across aporphyroblastic garnet from eclogite OF2595(LU). The core–rim contact is marked by asmallXMg increase, and by the sharp decreasein XMn. (d) Simulated garnet zoning foreclogite OF2595 (LU) along a straight linefrom 550 �C, 2.0 GPa to 600 �C, 2.8 GPa.

2 14 C . G R O P P O E T A L .


Symplectites, developed at the expense of bothglaucophane and omphacite, consist of green ferro-hornblende + albite (Fig. 3d,e), finer grained at thecontact with the amphibole ⁄ omphacite and progres-sively coarsening outward.

Phengite

Phengite occurs both in the matrix, where it shows aweak preferred orientation parallel to the matrix folia-tion (Sm) (Fig. 3c), and as aggregates with paragoniteand epidote, which are pseudomorphs after formerlawsonite. Thematrix phengite is rare andweakly zoned,with Si content ranging from3.56 a.p.f.u. (on the basis of11 oxygen) in the core to 3.54 a.p.f.u. in the rim. In thepseudomorphs after lawsonite, phengite has lower Si

contents (3.35–3.41 a.p.f.u.) (Fig. 6a and Table 2).Phengite has never been observed included in garnet.

Biotite, epidote, chlorite and paragonite

Two generations of biotite and epidote occur. Rare andvery fine-grained biotite I and epidote I are present inthe inner garnet cores; epidote II occurs in the pseud-omorphs after lawsonite, whereas biotite II locallyreplaces chlorite just at the contact with garnet(Fig. 3a). Biotite I has little Ti, whereas biotite II is Ti-free (Table 2). Epidote I is Al-poorer than epidote II(XZo = 28–30 in EpI, 45–52 in EpII; Table 2). Chloritemainly develops at the expense of garnet, at its rim andalong fractures. Paragonite is only present in thepseudomorphs after lawsonite.

Eclogite OF2595 (LU)

Eclogite OF2595 (LU) is very similar to OF2512(LCU). It mainly consists of omphacite, porphyrob-lastic garnet, blue amphibole, epidote and paragonitepseudomorphs after former lawsonite, and accessoryrutile and titanite. The weak foliation defined by thealignment of omphacite and glaucophane curvesaround garnet porphyroblasts. Phengite and biotite arelacking.

Garnet

Garnet porphyroblasts are zoned but show a growthhistory simpler than that observed in OF2512 (LCU).Two main zones have been recognized (Fig. 7a): (i) agarnet core, crowded with small inclusions of ompha-cite, bluish–green amphibole, epidote + paragonitepseudomorphs after former lawsonite, and ilmeniterimmed by rutile, which define an internal foliation(Sm)1) (Fig. 7c); (ii) a garnet rim that contains only fewinclusions of omphacite and rutile. Quartz is neverfound included in garnet.

Garnet porphyroblasts are chemically zoned(Fig. 4c), with element profiles similar in shape tothose of garnet from OF2512 (LCU): Ca and Mn aredecreasing from core to rim (XCa = 32–26 in the core,27–24 in the rim; XMn = 16–4 in the core, 4–1 in therim), and Mg and Fe increasing (XMg = 8–12 in thecore, 11–14 in the rim; XFe = 58–62 in the core, 62–64in the rim) respectively (Table 4). The Fe ⁄Mg ratioprogressively decreases from core to rim (Fe ⁄Mg =7–4). Although the transition between garnet core andrim, as defined on the basis of mineral inclusions, issharp, chemical zoning is gradual. At the rim, garnet ispartially replaced by green magnesiokatophorite(Fig. 7a).

Omphacite

Matrix omphacite occurs as elongated, strongly zonedcrystals, with a zoning pattern more complicated

(a)

(b)

Fig. 5. Omphacite composition from eclogites OF2512 (LCU)(a) and OF2595 (LU) (b), plotted in the Morimoto (1988) dia-gram. Both omphacite in the matrix and omphacite included ingarnet porphyroblasts are shown.



than that of OF2512 (LCU). Three main zones maybe recognized: a relict core with XNa = 52–58(XJd = 38–40, XAeg = 6–13), a wide mantle with XNa

= 57–61 (XJd = 43–48, XAeg = 8–15) and a rim withXNa = 50–54 (XJd = 43–46, XAeg = 0–6) (Fig. 5b &Table 4). Locally, a very thin and discontinuousrim enriched in the aegirine component hasbeen observed (XNa = 54–57; XJd = 39–43, XAeg =10–11). Omphacite is replaced along grain margins bya symplectite of albite + edenite–pargasite.

Omphacite included in garnet cores shows a variablecomposition, with XNa ranging from 42 to 53(XJd = 25–40, XAeg = 6–17). Omphacite included ingarnet rims has the same composition as that of themantle of matrix omphacite (XNa = 56–63; XJd = 43–50, XAeg = 8–11) (Fig. 5b & Table 4).

Amphibole

Amphibole occurs in different microstructural sites: abluish–green amphibole in garnet cores, a blue amphi-bole in the matrix, a deep green amphibole replacinggarnet and a fine-grained green amphibole in the sym-plectite after omphacite. The bluish–green amphiboleincluded in garnet cores is generally taramite withNaM4 = 0.77–0.91 a.p.f.u. and AlIV = 1.65–1.88a.p.f.u. (Fig. 6b & Table 3), whereas the late deep greenamphibole replacing garnet is a Mg-katophorite. Thecoarse-grained, sharply zoned, blue amphibole in thematrix is glaucophane (Fig. 7b & Table 3), character-ized by an enrichment in Fe from core to rim; it is rim-med by taramite toMg-taramite. The green amphibole inthe symplectite after omphacite plots in the edenite–pargasite fields according to Leake et al. (1997).

Epidote and paragonite

Two different epidote generations have been observed:epidote I, richer in zoisite component (XZo = 51–65),is included in garnet cores and locally includes

omphacite; epidote II (XZo = 27–35) occurs, in asso-ciation with paragonite, in the pseudomorphs afterlawsonite, which are found both in the matrix and inthe garnet porphyroblasts.

Metabasite OF2933 (UU)

Metabasite OF2933 (UU) is a foliated rock mainlyconsisting of chlorite, epidote, albite, garnet, phengite,minor quartz and amphibole, accessory titanite, andpatches enriched in carbonate. The matrix foliation(Sm), defined by millimetre-thick layers of chlo-rite + epidote alternating with albite ± epidotedomains, curves around porphyroblastic garnet up to1 cm in diameter (Fig. 8a) and relict phengite flakespartially replaced by chlorite.

Garnet

Garnet porphyroblasts are crowded with small inclu-sions of amphibole, epidote and titanite, which define afolded internal foliation (Sm)1) (Fig. 8a); coarser-grained inclusions of quartz are locally abundant in theoutermost garnet rim. XMn gradually decreases fromcore to rim (XMn = 6–0), whereas XMg slightlyincreases outward (from 5 to 8). XFe is lower in thecore (XFe = 56), it increases in the mantle (up toXFe = 62) and it finally decreases again in the rim(XFe = 57); XCa ranges from 29 to 34 and shows anantipathetic trend with respect to XFe (Table 5).Garnet is partially replaced by biotite at the rim.

Epidote

Epidote is one of the main components of the rock.Different generations may be recognized on micro-structural and chemical grounds. Epidote included ingarnet is gradually zoned, with a Fe-richer core (EpI,XZo = 27–40) and an Al-richer rim (EpII, XZo = 48–58). Epidote in the matrix is sharply zoned: an Al-rich

(a) (b)

Fig. 6. (a) Phengite compositions from eclogite OF2512 (LCU) and metabasite OF2933 (UU) plotted in the Si v. (Fe + Mg) diagram.The line represents the ideal celadonitic substitution. Si is given as atoms per formula unit (a.p.f.u.) on the basis of 11 oxygen.(b) Compositions of amphibole from samples OF2512 (LCU) and OF2595 (LU) plotted in the Al(IV) v. Na(M4) diagram.

2 16 C . G R O P P O E T A L .


core (EpII, XZo = 45–61) is partially resorbed andovergrown by a Fe-richer rim (EpIII, XZo = 20–22)(Fig. 8b).

Phengite

The matrix foliation Sm curves around relict phengiteflakes, partially replaced by chlorite at the rim. Theirceladonite contents decrease from core to rim(Si = 3.41–3.31 a.p.f.u.); spot analyses in the Si vs.(Fe + Mg) diagram plot above the ideal celadoniticsubstitution line (Fig. 6a), suggesting that some Fe3+

is present. No phengite flakes have been observed asinclusions in garnet porphyroblasts.

Amphibole

A green idioblastic and homogeneous actinolite occursas fine- to medium-grained crystals, which define an Si

(Sm)1) in garnet.

METAMORPHIC EVOLUTION

On the basis of microstructural relationships andmineral chemistry, several mineral assemblages havebeen recognized in each of the studied metabasics. Inthe following (and cf. Fig. 9), mineral assemblages thatrecorded a progressive increase in P–T will be groupedin the �prograde assemblages� section (A1), as opposedto the �peak assemblages� (A2) and to the �retrogradeassemblages� (A3).

Eclogite OF2512 (LCU)

Prograde assemblages (A1)

These assemblages are preserved in the core andmantle of porphyroblastic garnet and consist ofomphacite, Ca- (tschermakite) and NaCa-amphibole(Mg-katophorite), epidote I, lawsonite (now pseudo-morphically replaced by paragonite + epidoteII ± phengite), rare biotite I, quartz, ilmenite andrutile. On the basis of garnet zoning (Fig. 9), threemain stages may be recognized in the progradeevolution:Stage A1-I: corresponds to the Sm)1 foliation pre-served in the garnet cores and is defined by omphacite,Ca-amphibole, epidote I, rare biotite I, and ilmenite.Stage A1-II: corresponds to the growth of the garnetcore.Stage A1-III: corresponds to the growth of the garnetmantle, which was in equilibrium with omphacite,NaCa-amphibole, lawsonite, quartz, and rutile.

This sequence of stages indicates that eclogiteOF2512 (LCU) evolved from the epidote to thelawsonite stability field and from the quartz-absent tothe quartz-present field (Fig. 9).

Peak assemblage (A2)

This assemblage consisted of matrix omphacite, garnetrim, Si-rich phengite, glaucophane and lawsonite (nowpseudomorphically replaced by paragonite + epidote

(c)(a)

(b)

Fig. 7. Representative microstructures from eclogite OF2595 (LU) as seen under optical microscope. (a) Strongly zoned porphy-roblastic garnet (Grt) set in an omphacite (Omp) + glaucophane (Gln) matrix. Garnet is partially replaced at its rim byMg-katophorite (Am) (PPL). (b) Coarse-grained glaucophane (Gln) in the matrix, rimmed with a green amphibole (Am). Coarse-grained titanite (Ttn) is also present in the matrix (PPL). (c) Detail of the garnet core, with the internal foliation (Sm)1) defined by thepreferred orientation of the omphacite (Omp) and rutile (Rt) inclusions (PPL).



II ± phengite aggregates). No relics of coesite havebeen found in either garnet rims or omphacite.

Retrograde assemblage A3

During early post-peak evolution lawsonite wasreplaced by epidote II + paragonite ± phengite. Atstill lower pressures, within the plagioclase stabilityfield, both omphacite and glaucophane were partiallyreplaced by albite + ferrohornblende symplectites.Garnet was partially replaced by chlorite along frac-tures and at the rim; a discontinuous rim of biotite IIlocally developed between chlorite and garnet.

Eclogite OF2595 (LU)

Prograde assemblages A1

This assemblage, preserved in the core of porphy-roblastic garnet, consisted of omphacite, NaCa-amphibole (taramite), epidote I and lawsonite (nowpseudomorphically replaced by paragonite + epidoteII). The garnet core also belongs to this progradeparagenesis. Two main stages may be recognized(Fig. 9):

Stage A1-I: corresponds to the development of theSm)1 foliation and consists of omphacite, NaCa-amphibole, epidote I, ilmenite and titanite.Stage A1-II: corresponds to the growth of the garnetcore in the lawsonite stability field, probably in equi-librium with the core of matrix omphacite.

Peak assemblage A2

This assemblage consists of omphacite (the mantle ofthe matrix omphacite and the omphacite included inthe garnet rim), garnet rim, glaucophane, lawsoniteand rutile.


During the following evolution, an aegirine-rich rimovergrew omphacite, and taramite overgrew glauco-phane. Lawsonite was replaced by epidote II + par-agonite pseudomorphs both in the matrix and garnetporphyroblasts, and garnet was replaced by NaCa-amphibole. At still lower pressures, within the plagio-clase stability field, omphacite was partially replacedby an edenite–pargasite + albite symplectite and thematrix rutile by titanite.

(a) (b)

Fig. 8. Representative microstructures frommetabasite OF2933 (UU) as seen underoptical microscope and SEM. (a) Porphy-roblastic garnet (Grt) showing a deformedSi (Sm)1) mainly defined by the alignment offine-grained epidote (EpI), actinolite (Am)and titanite. (b) BSE image of zoned epidotefrom the matrix, characterized by a zoisite-richer core (EpII) and an epidote-richer rim(EpIII).

2 18 C . G R O P P O E T A L .



Prograde assemblage A1

It is preserved as inclusions in garnet porphyroblastsand consists of actinolite, epidote (EpI) and quartz(Fig. 9).

Peak assemblage A2

It consists of garnet, phengite and epidote (EpII).Neither omphacite relics nor its retrograde symplectitehave been observed, possibly because the rock matrixis pervasively recrystallized under greenschist faciesconditions or, more probably, because omphacite hasnever been stable in such an oxidized system. In con-trast, a Na-amphibole was probably stable at peakP–T conditions, now replaced by the albite and chlo-rite present in the matrix.


During the later evolution, both phengite and epidotere-equilibrated and developed a strong zoning. Albiteand chlorite extensively crystallized in the matrix(probably at the expense of former Na-amphibole) andbiotite partially replaced the garnet rim.

PSEUDOSECTION MODELLING

Model system and bulk-rock compositions

To model prograde, peak and retrograde conditions ofeclogites OF2512 (LCU) and OF2595 (LU), P–Tpseudosections were calculated using bulk-rock com-positions obtained by the area-scan method at SEM-EDS (Table 1), in the model N(K)CFMASH system.For OF2595 (LU), a correction has been applied to themeasured CaO considering that titanite is a veryabundant phase in the matrix: a CaO content equal toTiO2 was subtracted from the bulk-rock composition,in the plausible assumption that titanite was aprograde mineral (Table 6).

Fe2O3 was neglected because epidote was mainlyobserved in the retrograde mineral assemblage A3 (inthe pseudomorphs after lawsonite) and the calcu-lated amount of Fe3þ in both omphacite and garnet islow. Because garnet contains low spessartine compo-nent, MnO was neglected in the calculation. This is aquite delicate problem in doing pseudosection calcu-lations because even a small amount of MnO in thebulk-rock composition may expand the garnet stabilityfield towards lower-T (e.g. Spear & Cheney, 1989;Symmes & Ferry, 1992; Mahar et al., 1997; Tinkhamet al., 2001; Wei et al., 2004; Matsumoto et al., 2005).However, the studied rocks contain very little MnO(MnO = 0.16–017 wt%, determined by ICP-MS) andthe only Mn-bearing phase is garnet. As a conse-quence, small errors in the estimate of the MnO con-tents could cause very large errors in pseudosectiontopology. In fact, as garnet is the only Mn-hostingphase, if MnO is overestimated, this would result in thestabilization of a larger volume of garnet, with a con-sequent depletion in some elements used to stabilizeother minerals such as omphacite.

The pseudosection approach has not been applied forsample OF2933 (UU) because: (i) it is not homogeneous

Fig. 9. Metamorphic evolution of the studied metabasics.Sm = matrix foliation; Sm)1 = relict foliation, earlier than Sm,preserved in microlithons or as internal foliation (Si) in garnetporphyroblasts.



as to the grain size (large garnet porphyroblastsenclosed in a fine-grained matrix); (ii) it is stronglyoxidized (epidote is themost abundant phase in the rockmatrix) and it is difficult to reliably estimate the oxygenfugacity conditions at each step of metamorphic his-tory; and (iii) it is locally enriched in carbonates, thussuggesting the presence of a mixed H2O–CO2 fluid atleast during some stages of its evolution.

Fractionation effects on the bulk-rock composition

As zoned garnet porphyroblasts with relict progradecore and mantle are present in both eclogites OF2512(LCU) and OF2595 (LU), the rocks must haveundergone a progressive chemical fractionation as aconsequence of porphyroblast growth (Marmo et al.,2002). This means that the bulk-rock composition maynot be representative of the effective equilibriumcompositions during each stage of the P–T history.Therefore, following the method proposed by Marmoet al. (2002), the fractionation effects on the bulkcomposition of eclogite OF2512 (LCU) have beencalculated. Microstructurally, the garnet zoning maybe inferred from the inclusion distribution: the garnetcore is crowded with very small inclusions defining aweak foliation (Sm)1), whereas the garnet mantlecontains coarser and randomly oriented inclusions.Image analysis has been used to calculate the modalabundances of garnet cores and mantles at 3 and 6vol.% respectively. Average compositions of garnetcore and mantle have been calculated and three dif-ferent equilibrium compositions have been used: (i) theentire bulk-rock composition, which corresponds tothe equilibrium composition during the progradegrowth of the garnet core (OF2512a); (ii) the bulk-rockcomposition minus the garnet core composition(OF2512b), which gives the equilibrium compositionduring the growth of prograde garnet mantle; and (iii)the bulk-rock composition minus the garnet core andmantle compositions (OF2512c), which gives theequilibrium composition during the growth of garnetrim at peak conditions (Table 6). Three pseudosectionshave been modelled for the three equilibriumcompositions.

Pseudosection calculation

The pseudosections were calculated following theapproach of Connolly (1990) and using the internally

consistent thermodynamic data set and equation ofstate for H2O of Holland & Powell (1998, revised2002). The phases considered in the calculation were:garnet, omphacite, lawsonite, zoisite ⁄ clinozoisite,amphibole, phengite, biotite, paragonite, quartz ⁄ coe-site, plagioclase and chlorite. The following solid-solution models have been used: garnet, phengite(Holland & Powell, 1998), omphacite (Holland &Powell, 1996), amphibole (Dale et al., 2005), biotite(Powell & Holland, 1999), plagioclase (Newton et al.,1981) and chlorite (Holland et al., 1998). The follow-ing assumptions were made: (i) zoisite ⁄ clinozoisite wasused as a proxy for epidote; (ii) H2O was assumed to bepresent in excess (aH2O = 1): this assumption is sup-ported in part by the abundance of the hydrous phasesphengite, amphibole, lawsonite, epidote and parago-nite.Amphibole has been modelled using the Dale et al.

(2005) solid-solution model, which takes into accountthe known solvi between Ca- and Na-amphiboles.Roughly, the amphibole modelled at high pressure issodic, whereas the amphibole modelled at low pressureis mainly calcic. Nevertheless, the amphibole compo-sitional isopleths modelled in the pseudosection are notstrictly in agreement with the observed compositions.In particular, in the modelling, the amphibole incor-porates too much Na, especially at low pressures,which results in a reduction of the paragonite stabilityfield. However, as a detailed investigation of thepseudosection topology at low pressures is beyond theaim of this work, we have considered the obtainedresults as a fair approximation of the observed retro-grade parageneses.The assemblages on the high temperature side of the

pseudosections may be metastable with respect to melt-bearing assemblages (Stıpska & Powell, 2005). Thesehigh-temperature assemblages are not relevant to themodelling of the P–T paths for the studied rocks and,consequently, this problem has been neglected. In anycase, for the studied bulk-rock compositions, the pro-portion of melt in this P–T range is small (Stıpska &Powell, 2005), and the topology of phase relationshipsremains approximately the same in both melt-absentand melt-bearing systems. Miscibility gaps wereobserved in omphacite, plagioclase and amphibole inthe calculated phase diagrams, typically at low P–T(Stıpska et al., 2006).

RESULTS

Pseudosection analysis of eclogite OF2512 (LCU)

Pseudosection (a): unfractionated bulk composition

The bulk-rock composition OF2512a (Table 6) wasused to model the first step of the prograde evolutionof eclogite OF2512 (LCU), which corresponds to thedevelopment of Sm)1 (stage A1-I) and to the growthof garnet cores (stage A1-II). The pseudosection,

Table 6. Bulk compositions used for pseudosection calculation.

wt% Na2O MgO Al2O3 SiO2 K2O CaO FeO

OF2512

Entire rock (a) 4.80 5.48 17.46 50.77 0.34 9.77 9.17

Grt core isolated (b) 4.94 5.56 17.34 51.12 0.35 9.75 8.66

Grt mantle isolated (c) 5.25 5.68 17.10 51.92 0.37 9.89 7.37

OF2595

After Ttn correction 5.45 4.80 17.64 51.01 – 10.72 10.39

2 20 C . G R O P P O E T A L .


calculated in the P–T range 0.5–3.5 GPa and 300–900 �C, consists of di-, tri-, quadri- and pentavariantfields, which are represented in Fig. 10(a) in white,light-, medium- and dark-grey respectively.

The pentavariant Bt + Am + Omp + Zo field,stable at 350 < T < 480 �C and P < 1.7 GPa, welldepicts the prograde assemblage (stage A1-I) preservedin the garnet cores. Modelled isomodes show thatbiotite is modally scarce (<3 vol.%) in agreement withmicroscopic observations.

The XCa and XFe isopleths of garnet modelled in thepseudosection suggest that garnet cores (XCa = 35–24;XFe = 56–65) mainly grew in the lawsonite stabilityfield at 480 < T < 600 �C and 2.4 < P < 2.8 GPa(stage A1-II; Fig. 10b,c). Prograde omphacite includedin garnet cores has a XNa = 48–52. The XNa isoplethsmodelled in the pseudosection (Fig. 10d) show thatomphacite formed at 450 < T < 600 �C atP > 1.9 GPa.

Pseudosection (b): unfractionated bulk composition minusthe garnet cores

Pseudosection (b), calculated using the bulk composi-tion obtained by subtracting the garnet cores(OF2512b, Table 6), models the prograde portion ofthe eclogite evolution, which corresponds to thegrowth of the garnet mantle in equilibrium withomphacite, lawsonite, NaCa-amphibole and quartz(stage A1-III). The topology of both pseudosections(a) and (b) is the same. Only the garnet-in curve isslightly shifted by �0.05 GPa towards higherpressures, but the position of garnet and omphacitecompositional isopleths does not change (Fig. 10f).This means that the bulk composition barely changesas a consequence of fractionation during the growth ofthe garnet core. Consequently, either pseudosection (a)or (b) can be used to model the prograde evolution ofeclogite OF2512 (LCU); for this reason the details ofpseudosection (b) are not reported in Fig. 10.

The quartz-bearing assemblage, only observed in thegarnet mantle, is represented by the trivariant fieldPhg + Am + Grt + Omp + Lws + Qtz ⁄Coe. TheXCa and XFe isopleths of garnet modelled in thepseudosection (Fig. 10b,c) suggest that the garnetmantle (XCa = 26–21; XFe = 68–65) mainly grew at520 < T < 600 �C and 2.7 < P < 3.2 GPa. Pro-grade omphacite included in the garnet mantle has aXNa = 48–52: the XNa isopleths modelled in thepseudosection (Fig. 10d) show that omphacite formedat 450 < T < 600 �C at P > 1.9 GPa.

Combining the information given from pseudosec-tions (a) and (b) with the garnet and omphacite com-positional isopleths, the three prograde stagesdescribed in the metamorphic evolution (Fig. 9) maybe constrained at the following conditions:Stage A1-I (Sm)1 foliation): 350 < T < 480 �C andP < 1.7 GPa, in the epidote stability field; the penta-variant Bt + Am + Omp + Zo field predicts the

stability of omphacite at relatively low P–T, undertypical blueschist facies conditions. Similar occur-rences of prograde omphacite stable at < 500 �C havebeen reported, for example, from the Pouebo andDiahot terranes of the Pam Peninsula, NE NewCaledonia (Carson et al., 1997; Clarke et al., 1997,2006), from the Sivrihisar Massif, Turkey (Davis &Whitney, 2006), and from the Sifnos island, Cyclades,Greece (Groppo et al., 2009).Stage A1-II (garnet core): 480 < T < 600 �C and2.4 < P < 2.8 GPa, in the lawsonite stability field.Stage A1-III (garnet mantle): 520 < T < 600 �C and2.7 < P < 3.2 GPa. The stability field of garnetmantle modelled in the pseudosection lies in the Coe-bearing trivariant field Phg + Am + Grt +Omp + Lws + Coe at >575 �C, whereas at <575 �C(for which pressures would be compatible with thepresence of quartz) the system is undersaturated in SiO2

and quartz is unstable. The modelled pseudosectionthus suggests that the first SiO2-phase, formed inequilibrium with the garnet mantle, should have beencoesite; however, quartz included in the garnet mantledoes not seem to derive from coesite inversion. Theformation of quartz instead of coesite may be due to aslightly lower water activity (i.e. aH2O < 1) than thathypothesized. Lowering the water activity, the topo-logy of the pseudosection does not change, but all theunivariant reactions defining the boundaries among themultivariant fields shift towards lower-T (�30 �C lowerfor aH2O = 0.75). The isopleth position does notchange (indicating that the stability field modelled forthe garnet mantle remains unchanged), but at <575 �Cthe garnet mantle could have grown in equilibrium withquartz. The sluggish kinetics of the quartz to coesiteinversion reaction probably prevented the growth ofcoesite after quartz.

Pseudosection (c): unfractionated bulk composition minusgarnet cores and mantles

Pseudosection (c), calculated using the whole bulk-rock composition minus garnet cores and mantles(OF2512c; Table 6), models the peak and retrogrademetamorphic conditions. Also in this case, the topol-ogy of pseudosections (a) and (c) and the position ofgarnet and omphacite compositional isopleths are verysimilar to the high-pressure portion (i.e. in the lawso-nite stability field): this means that the bulk-rockcomposition does not change as a consequence offractionation during the growth of both garnet coreand mantle. As for pseudosection (b), the garnet-incurve is shifted by �0.1 GPa towards higher pressures(Fig. 10f). The topology of pseudosection (c) slightlydiffers from pseudosection (a) at lower pressureswithin the epidote stability field.

The UHP peak assemblage A2, consisting of thegarnet rim, Si-rich phengite, glaucophane and law-sonite (+possibly coesite), is modelled by the triva-riant field Phg + Am + Grt + Omp + Lws + Coe,



occurring at 550 < T < 630 �C. The XCa and XFe

isopleths of garnet modelled in the pseudosectionwell constrain the garnet rim (XCa = 19–16;

XFe = 65–59) growth at 590 < T < 605 �Cand>3.2 GPa (Fig. 10b,c). These P–T conditions arein good agreement with those suggested by the XNa

(a)

(e)(d) (f)

(c)

(b)

Fig. 10. (a) P–T pseudosection for eclogite OF2512 (LCU) calculated in the NKCFMASH system at a(H2O) = 1, using theunfractionated bulk composition (OF2512a) of Table 6. White, light-, medium- and dark-grey fields are di- tri-, quadri- and penta-variant fields respectively. (b–e) Compositional isopleths for garnet (b: XCa; c: XFe), omphacite (d: XNa) and phengite (e: Si a.p.f.u.).Dotted ellipses constrain P–T conditions at which assemblages A1-I, A1-II, A1-III and A2 (cf. Fig. 9) developed. (f) XCa isopleths forgarnet calculated using the unfractionated bulk composition (whole rock: OF2512a in Table 6) and the bulk compositions from whichgarnet cores and mantles have been subtracted (OF2512b and OF2512c in Table 6). The Grt-in curve is shifted towards higherpressure, but no significant variation in the isopleth position is observed.

2 22 C . G R O P P O E T A L .


isopleths of the core of matrix omphacite (XNa = 55–57; Fig. 10d). Isopleths of the maximum celadonitecontent in the phengite cores (Si = 3.56 a.p.f.u.)(Fig. 10e) suggest a peak pressure slightly lower thanthat obtained from the garnet and omphacite isopleths(P � 3.1 GPa at 600 �C). However, as phengite is rare,it is possible that the analysed relict cores do not cor-respond to their maximum celadonite content (up toSi = 3.60 a.p.f.u.).

The retrograde assemblage is less constrained thanprograde and peak assemblages. Both phengite andomphacite in the matrix are zoned. Si(Phg) isoplethsmodelled in the pseudosection are mainly pressuredependent: the lower celadonite contents of phengitegrown in the pseudomorphs after lawsonite (Si = 3.35a.p.f.u.) are compatible with a decompressional tra-jectory, but no information about the temperature maybe obtained from these isopleths. Omphacite in thematrix is zoned, with XNa ranging from 57 in the coreto 51 in the rim. XNa isopleths modelled in thepseudosection (Fig. 10d) are mainly T-dependent andsuggest that the decompressional P–T path was asso-ciated with a slight cooling from 605 �C at a peakpressure to �540 �C at �2.3 GPa.

Another constraint on the decompressional evolu-tion of eclogite OF2512 (LCU) derives from theobservation that, although most of the prograde P–Tpath lies in the lawsonite stability field, no lawsoniterelics have been observed both in the studied sampleand in the other eclogites from the UHP Cignana Unit(Compagnoni & Rolfo, 2003 and references therein).However, lozenge-shaped aggregates of epidote andparagonite ⁄phengite are common in the eclogites fromthis unit and may be interpreted as pseudomorphsafter lawsonite (Fry & Fyfe, 1969; Pognante, 1989;Evans, 1990; Will et al., 1998; Ballevre et al., 2003;Schmadicke & Will, 2003). The preservation of law-sonite mainly depends on the P–T path. Zack et al.(2004) suggested that lawsonite may be preserved onlyif the retrograde P–T path occurs on the low-T side ofthe reaction Lws = Ky + Zo + Qtz ⁄Coe + H2O(Poli & Schmidt, 1998), which marks the upper sta-bility of lawsonite (Fig. 12). The absence of relictlawsonite thus means that the UHP Cignana Unitcrossed this reaction during the decompressional evo-lution. Such a P–T path is typical of most Alpine-typeeclogites (Ernst, 1988), which initially formed withinthe lawsonite stability field, but during further sub-duction and ⁄or uplift entered into the zoisite stabilityfield (Zack et al., 2004).

Simulation of the garnet growth along the progradeP–T path

In the pseudosection (a), the garnet zoning resultingfrom the prograde P–T path has been simulated(excluding XMn) along a straight line between500 �C ⁄ 2.0 GPa and 600 �C ⁄ 3.3 GPa. The simulatedgarnet zoning (Fig. 4b) is very similar to the actual

garnet composition. In particular, the XFe increase inthe mantle has been well modelled, together with thesteep increase in XMg towards the rim. Furthermore,not only the shape of the chemical profiles, but also thenumerical values of X(Ca,Mg,Fe) are very close to themeasured values. The slightly higher XCa simulated forgarnet core is due to the absence of Mn in the modelledpseudosection. The results of the garnet-zoning simu-lation confirm the prograde P–T path inferred fromthe pseudosection analysis.

Pseudosection analysis of eclogite OF2595 (LU)

The effects of chemical fractionation have beenconsidered negligible as discussed for OF2512(LCU), and then the bulk-rock chemical compositionhas been considered as representative for the wholeevolution of the sample. The pseudosection, calcu-lated in the P–T range 0.5–3.5 GPa and 400–800 �C,consists of di-, tri-, quadri- and pentavariant fields,which are shown in Fig. 11(a) in white, light-, med-ium- and dark-grey respectively. The progradeassemblage A1-I, preserved as inclusions in garnetcores and consisting of omphacite + NaCa-amphi-bole + epidote I (+titanite + rare ilmenite), corre-sponds to the pentavariant field Am + Omp + Zolocated in the range 400–650 �C and 1.0–1.7 GPa.The prograde assemblage A1-II, which consists ofgarnet core + omphacite (core of the matrix om-phacite) + lawsonite ± NaCa-amphibole (+rutile),corresponds to the quadrivariant field Am + Grt +Omp + Lws of Fig. 11(a). The epidote to lawsonitetransition is marked by a narrow trivariant field. Theprograde garnet core has XCa = 32–26 andXFe = 58–62, and cores of matrix omphacite haveXNa = 56–58. The compositional isopleths modelledin the pseudosection (Fig. 11b–d) constrain thegrowth of garnet cores and omphacite cores at530–600 �C and 2.4–2.7 GPa respectively, atapproximately the same pressures as stage A1-II ofeclogite OF2512 (LCU), but at slightly highertemperatures.

The two prograde stages recognized in eclogiteOF2595 (LU) may be, then, constrained at thefollowing P–T conditions:Stage A1-I (Sm)1 foliation): 400 < T < 650 �C and1.0 < P < 1.7 GPa, in the epidote stability field.Stage A1-II (garnet core): 530 < T < 600 �C and2.4 < P < 2.7 GPa, in the lawsonite stability field.

The peak assemblage (A2), consisting of omphacite(mantle of the matrix omphacite) + garnet rim +glaucophane + lawsonite (+rutile), mainly differsfrom the prograde A1-II assemblage as to the garnetand omphacite compositions. XCa and XFe of thegarnet rim are in the range 27–24 and 62–64 respec-tively, and XNa in omphacite in the range 58–61. Themodelled compositional isopleths for garnet(Fig. 11b,c) and omphacite (Fig. 11d) suggest thatgarnet rims and omphacite mantles grew at 570–



605 �C and 2.7–2.8 GPa. These P–T conditions mainlyplot in the trivariant field Am + Grt + Omp +Lws + Qtz ⁄Coe, very close to the quartz ⁄coesite inversion curve. Neither quartz nor coesitehave been observed in the studied eclogite, but theirmodal amounts (isomodes) modelled in thepseudosection suggest that the quartz ⁄ coesite modalcontent at the estimated P–T conditions is lowerthan 3 vol.%. Peak P–T conditions of eclogiteOF2595 (LU) occur at approximately thesame temperature estimated for the UHP eclogiteOF2512 (LCU), but at pressures �0.4–0.5 GPalower.

As for eclogite OF2512 (LCU), the decompressionalevolution of eclogite OF2595 (LU) is much less con-strained than the prograde and peak stages. Thebreakdown of lawsonite to epidote + paragonitesuggests that the lawsonite to epidote reaction curvewas crossed during decompression; rim composition ofomphacite (XNa = 50–54) points to <1.75 GPa, butno information about the temperature may beobtained from the pseudosection.

Simulation of the garnet growth along the progradeP–T path

As for eclogite OF2512 (LCU), the garnet zoningresulting from the prograde P–T path has been simu-lated along a straight line from 550 �C ⁄ 2.0 GPa to600 �C ⁄ 2.8 GPa respectively. The resulting simulatedgarnet zoning is reported in Fig. 4(d), and comparedwith the actual garnet composition. The increase inboth XFe and XMg from core to rim has been wellmodelled, thus confirming the prograde P–T evolutioninferred from the pseudosection analysis on the basisof the garnet and omphacite compositional isopleths.

THERMOBAROMETRIC ESTIMATES

Eclogites OF2512 (LCU) and OF2595 (LU)

Peak metamorphic conditions for the studied sampleswere also calculated using THERMOCALCTHERMOCALC v. 3.25 �Aver-age PT� method (Powell & Holland, 1988, 1994) andthe latest version of the Holland and Powell data set.

(a)(d)

(b) (c)

Fig. 11. P–T pseudosection for eclogite OF2595 (LU) calculated in the NCFMASH system at a(H2O) = 1. White, light-, medium-and dark-grey fields are di- tri-, quadri- and pentavariant fields respectively. (b–d) Compositional isopleths for garnet (b: XCa; c: XFe)and omphacite (d: XNa). Dotted ellipses constrain P–T conditions at which assemblages A1-I, A1-II and A2 (cf. Fig. 9) developed.

2 24 C . G R O P P O E T A L .


The �Average PT� method (also called �optimal ther-mobarometry�, Powell & Holland, 1994) estimates P–Tconditions using an independent set of reactions rep-resenting all the equilibria among the end-members ofthe equilibrium assemblage. It shows a number ofadvantages over conventional thermobarometry,including: (i) its consistency with other mineral equi-libria methods (like pseudosections); (ii) the possibilityof realistic estimates of uncertainties; and (iii) a sta-tistical analysis assessing the reliability of thermo-barometric information that can be extracted from aspecific mineral assemblage. Mineral compositionsused for thermobarometric calculations [garnet rim,omphacite core and phengite core, corresponding tothe highest XMg(Grt), XNa(Omp) and Si(Phg)] arereported in Tables 2–5. Peak mineral assemblages ofsamples OF2512 (LCU) and OF2595 (LU) allowed aset of seven and six reactions to be defined respectively.P–T conditions of 573 ± 34 �C, 2.58 ± 0.30 GPa,and 542 ± 34 �C, 2.12 ± 0.72 GPa were obtained forsamples OF2512 (LCU) and OF2595 (LU) respectively(Table 7 & Fig. 12). The considerably large uncer-

tainties in the pressure estimate for eclogite OF2595(LU) is due to the lack of phengite in the equilibriumassemblage.

The Krogh Ravna & Terry (2004) calibration forkyanite-free eclogites (for which mean standard devi-ations of ±45 �C in temperature and ± 0.32 GPa inpressure were estimated) was also applied to thephengite-bearing peak assemblage of eclogite OF2512(LCU), resulting in P–T conditions of 587 �C and2.7 GPa (Table 7 and Fig. 12).

Pressure estimates obtained using Average PT andconventional thermobarometry are then lower thanthose inferred from the pseudosection analysis,whereas temperatures are approximately the same.This difference may be explained in different ways: (i)for sample OF2512 (LCU), the highest celadoniticcontent measured in the phengite core may not corre-spond to the maximum one present in the sample dueto the limited number of analyses, as previously sug-gested on the basis of the Si(Phg) isopleth trend; (ii) forsample OF2595 (LU), pressure conditions derivedfrom the pseudosection analysis plot within the very

Fig. 12. P–T paths for the UHP LCU and adjacent Zermatt–Saas LU and UU as inferred from eclogites OF2512 (LCU) and OF2595(LU) and metabasite OF2933 (UU). Continuous and dashed ellipses show prograde (A1-I, A1-II and A1-III), and peak (A2) con-ditions for eclogite OF2512 (LCU) and OF2595 (LU) respectively, as derived from the compositional isopleths modelled in thepseudosections (Figs 10 & 11). Dotted ellipse shows the peak conditions for metabasite OF2933 (UU) as derived from conventionalthermobarometry. The lawsonite stability field, as inferred from the pseudosections (Figs 10 & 11), is reported in pale grey. TheLws = Ky + Zo + Qtz ⁄Coe reaction is from Poli & Schmidt (1998). In the inset, the P–T constraints derived from conventionalthermobarometry and Average PT method (Powell & Holland, 1988) are reported and compared with the P–T paths reconstructed onthe base of pseudosection analysis. Numbers in italics refer to the Si(Phg) isopleths modelled for eclogite OF2512 (LCU) (Fig. 10e).KR&T04, Krogh Ravna & Terry (2004); K&R78, Krogh & Raheim (1978).



large uncertainties of THERMOCALCTHERMOCALC pressure estimates(±0.72 GPa).


For metabasite OF2933 (UU), the peak assemblagedid not define enough reactions for Average PT towork. The peak-T conditions have been estimatedusing the garnet–phengite thermometer of Krogh &Raheim (1978) and Green & Hellman (1982) due to theabsence of relict omphacite. Compositions of garnetrim and phengite core, corresponding to the highestXMg(Grt) and Si(Phg) values respectively were used(Tables 2–5). The garnet–phengite thermometer gives�600 �C at 2.0 GPa (Table 7 and Fig. 12), similar tothe peak temperature for eclogite OF2512 (LCU). Asin the Green & Hellman (1982) experiments, total Fe inphengite was treated as Fe2+, hence the equilibrationtemperature of 600 �C obtained for a nominal pressureof 2.0 GPa should be considered as a maximum value.

The Si content in phengite has been used to con-strain peak-P to a first order, by comparison with theSi(Phg) isopleths modelled for eclogite OF2512 (LCU)(Fig. 10e). The highest Si content (Si = 3.41 a.p.f.u.)measured in the phengite cores suggests �2.4 GPa at600 �C; as phengite in metabasite OF2933 (UU)probably contains some Fe3+ (see Fig. 6a), thispressure should be considered as a maximum value.Conventional thermobarometry thus constrains peakconditions for metabasite OF2933 (UU) at maximumP–T conditions of 2.4 GPa and 600 �C (Fig. 12). TheseP–T conditions plot just at the boundary betweenlawsonite and epidote stability fields, at approximatelythe same temperature and �0.8 GPa below the UHPpeak estimated for eclogite OF2512 (LCU).

At these P–T conditions, omphacite is generallystable in metabasic rocks: however, neither omphaciterelics nor pseudomorphs after omphacite have beenobserved in sample OF2933 (UU). The lack ofomphacite in the peak paragenesis may be related to:(i) the strongly oxidized bulk composition, and (ii)higher values of lCO2 (and lower values of lH2O)than in eclogites from LCU and LU, as suggested bythe presence of carbonate-enriched zones. Both fac-tors, which can be related to an ocean-floor alteration,enlarge the stability field of blueschist facies assem-blages at the expense of eclogite facies assemblages.Similar cofacial blueschist and eclogitic assemblageshave been reported from pillow lavas of the Zermatt–Saas Zone (Bearth, 1959, 1967, 1973) and have beeninterpreted as due to variation in bulk and fluid com-

position (Oberhansli, 1982; Barnicoat, 1988). A highX(CO2) in the fluid phase may also explain the absenceof lawsonite and its pseudomorphs in metabasiteOF2933 (UU), though peak P–T conditions lie insidethe lawsonite stability field.

DISCUSSION

P–T pseudosection analysis and conventional ther-mobarometry was used to reconstruct the P–T pathsfor metabasalts belonging to the LCU and adjoiningmeta-ophiolites. These three units experienced similarpeak temperatures (�600 �C) at considerably differentpeak pressures (>3.2, 2.7–2.8 and <2.4 GPa respec-tively) (Fig. 12). The resulting P–T path of UHPsample OF2512 (LCU) is similar to that estimated byReinecke (1998), being only slightly shifted towardslower temperatures and higher pressures.From these P–T paths, it is evident that the geother-

mal gradients changed significantly through time dur-ing subduction. More specifically, initial gradients of�9–10 �C km)1 were followed by lower gradients of�5–6 �C km)1, close to the �forbidden zone�, peaking inthe lawsonite eclogite facies field. Exhumation wascharacterized by a renewed increase in thermal gradient,coeval to the juxtaposition of the Lago di CignanaUHPunit with the other units of the Piemonte Zone. Hottergeotherms resulted in the retrogression of lawsonite, afeature commonly observed in orogenic belts.

Prograde evolution

The thermal gradient decrease in the prograde pathmay provide important insights on the evolution of thePiemonte subduction zone, which accommodated theconvergence between the African and European platesin the area that later became part of the Western Alps.Existing geochronological data indicate that this sub-duction zone has been active approximately from c.65 Ma, when the Sesia–Lanzo continental unit wasaccreted to the evolving orogen, to c. 44–40 Ma, whenthe more external Brianconnais continental unitreached the subduction zone (Rosenbaum & Lister,2005). Plate kinematic reconstructions support thisview, revealing that the convergence between Africaand Europe slowed down dramatically between c. 65and 55 Ma, presumably as a result of the collisionbetween the Sesia–Lanzo microplate and the evolvingorogen (Rosenbaum et al., 2002). Subsequently, sub-duction re-started, and the convergence rate underwenta progressive increase up to �10 cm year)1.

Table 7. Results of conventional thermo-barometry and Average PT calculations

Sample Method Assemblage Temperature (�C) Pressure (GPa)

OF2512 Average PT Grt–Omp–Gln–Phg–Lws–Coe–Ru–H2O 573 ± 34 2.58 ± 0.30

Krogh Ravna & Terry (2004) Grt–Omp–Phg 587 ± 45 2.70 ± 0.32

OF2595 Average PT Grt–Omp–Gln–Lws–Qtz–Ru–H2O 542 ± 34 2.12 ± 0.72

OF2933 Krogh & Raheim (1978) Grt–Phg 600 (at 2.0 GPa)

All uncertainties are given at the 2r level.

2 26 C . G R O P P O E T A L .


Lu–Hf garnet–omphacite–whole-rock isochron agesof 48.8 ± 2.1 Ma have been interpreted by Lapenet al. (2003) to date the growth of garnet cores similarto those described in our sample OF2512 (LCU),indicating that the LCU was already at depths inexcess of 30–35 km by that time (Fig. 13a). Peakpressure conditions have been estimated to have oc-curred at 44 ± 1 Ma with U ⁄Pb dating of zircon rims(Rubatto et al., 1998) and at 43.2 ± 1.1 and44.4 ± 1.5 Ma with 40Ar ⁄ 39Ar in situ dating ofphengitic mica included in garnet (Gouzu et al., 2006).A Sm–Nd garnet age of 40.6 ± 2.6 Ma has also beenproposed for the pressure peak (Amato et al., 1999),but this interpretation seems questionable in the lightof 40Ar ⁄ 39Ar and Rb–Sr dating of greenschist faciesfabrics from the Piemonte units from that area, whichyielded ages of 42–40 Ma (Reddy et al., 1999, 2003;Cartwright & Barnicoat, 2002).

Therefore, subduction of the LCU from �30–35 kmto 100–110 km depth took place in 3–7 Ma, betweenc. 48.8 ± 2.1 and c. 44 Ma. During this time span, thethermal gradient at the subduction zone changed from�9–10 to 5–6 �C km)1 (Fig. 13a,b). Studies carried outon present-day subduction zones have shown that thethermal structure of a subduction zone varies as aconsequence of: (i) changes in the subduction rate and(ii) thermal structure of the subducting lithosphere(Peacock & Wang, 1999). More specifically, low ther-mal gradients are widely reported from fast-subductingslabs, as conductive heating of the sinking slab ishindered by its high vertical velocity (Peacock &Wang, 1999). On the other hand, slowly subductingslabs are characterized by a less-pronounced thermaldisequilibrium with the surrounding mantle and, con-sequently, by higher thermal gradients. Changes in thevertical velocity of subducting slabs as subduction

(a)

(b)

(c)

Fig. 13. Simplified evolution of the Piemonte subduction zone in the 49- to 40-Ma interval: (a) early subduction of oceanic crust andsediments is recorded by metamorphic re-equilibration at gradients of �9–10 �C km)1. (b) As more and more oceanic slab sinks intothe upper mantle, the vertical velocity of the subducting plate increases. As a result, metamorphic mineral assemblages in the LCUrecord lower thermal gradients. (c) At c. 44–40 Ma, the arrival of more buoyant Brianconnais continental basement rocks locks thesubduction zone, stopping the retreat of the subducting plate and allowing the accretion of oceanic units to the Alpine belt. Starindicates the approximate position of the Lago di Cignana unit during the different stages.



progresses have been related to the presence of agreater length of oceanic lithosphere in the uppermantle, acting as an engine for subduction (Royden &Husson, 2006). This effect is due to the fact that slabsare undersupported by viscous stresses at depths of�150 km (Royden & Husson, 2006).

Therefore, we propose that the change in thermalgradient recorded by the LCU between 48.8 ± 2.1 andc. 44 Ma may have been related to an increase in thevertical velocity of the subducting Piemonte slab, ofwhich the LCU was part (Fig. 13a,b). The existence ofa time-lag between the initiation of subduction of thePiemonte slab, which in the Western Alps post-dates c.65 Ma but could be as young as 55 Ma (Rosenbaumet al., 2002), and the increase in its vertical velocity are,therefore, consistent with models and observations ofpresent-day subduction zones, where such an increasepost-dates the initiation of subduction by a few millionyears (Royden & Husson, 2006).

A corollary of this change in the sinking rate of thePiemonte slab is that the subduction zone hinge wouldhave undergone a progressive retreat with respect tothe overlying plate, a feature observed in most cur-rently active subduction zones (Schellart et al., 2007).This process provides an explanation for the observa-tion that, while over 200–250 km of Piemonte oceanlithosphere were subducted beneath Adria (Stampfliet al., 1998), only a very limited amount of this oceaniclithosphere and crust is presently exposed in the Wes-tern Alps, indicating that most of it escaped accretion(Manatschal & Muntener, 2008). Tomographic imagesof the Western Alps area show the presence of abun-dant lithospheric material underneath Adria (Lippitschet al., 2003), which may represent the remnants of thePiemonte slab.

Several lines of evidence indicate that the Brian-connais continental basement reached the subductionzone at 44–40 Ma (e.g. Stampfli et al., 1998; Rosen-baum & Lister, 2005). The arrival of buoyant crust atsubduction zones results in a decrease in subductionrate, eventually stopping the retrograde motion of thesubduction zone hinge (Martinod et al., 2005;Fig. 13c). In the case of the Western Alps, the buoyantBrianconnais block is the best candidate for the lock-ing of the subduction zone, which resulted in theaccretion of the Piemonte oceanic units to the evolvingorogen. Locking of a subduction zone is an efficientmechanism to induce the accretion of oceanic units: (i)by slowing down the vertical motion of the sinking slaband (ii) because ongoing convergence is accommo-dated through the nucleation of a new subduction zonein a more external position. As a result, a new thrustfault forms underneath the more internal units, therebypotentially causing accretion (Fig. 13c).

Retrograde evolution

The retrograde evolution preserved in the studiedsamples allows constraints to be placed only on the

early part of the exhumation of the Lago di Cignanaand surrounding units. After metamorphic re-equili-bration under UHP conditions, the Lago di Cignanaunit was exhumed to 2.3 GPa and 540 �C (Fig. 12). Asthe exhumation path of the LCU overlaps with thehighest P–T conditions recorded by the LU and UU, itseems plausible to suggest that the three units werejuxtaposed during the early stages of exhumation atthe footwall of the extensional Gressoney shear zone,whose activity has been dated at c. 42–38 Ma (Ballevre& Merle, 1993; Reddy et al., 1999, 2003; Wheeleret al., 2001; Cartwright & Barnicoat, 2002; Forsteret al., 2004; Beltrando et al., 2008 for a discussion).The current position of the UU above the LCU is

probably related to the late-stage top-to-the-westthrusting that partially reactivated the top of theeclogitic units under greenschist facies conditions at c.39–37 Ma (Reddy et al., 1999; see discussion in Beltr-ando et al., 2008). This thrusting event is particularlywell expressed at the base of the Pancherot-CimeBianche unit (Pleuger et al., 2007).

CONCLUSIONS

The detailed petrological study of multistage meta-morphism in meta-ophiolitic units can provideimportant insights into the evolution of palaeo-subduction zones and into the early post-accretionstages. Halving of the thermal gradients in the Pie-monte slab has been interpreted as the evidence of aprogressive increase in the subduction rate betweenc. 49 and c. 44 Ma. Such increase in the verticalvelocity, which post-dates the initiation of subductionof the Piemonte slab in the Western Alps by 6–15 Ma,is consistent with observations and models that requirea significant length of slab to be present in the uppermantle to act as an engine for subduction. This processlikely resulted in the retreat of the subduction zonehinge. Therefore, estimates of the behaviour of palaeo-subduction zones of the kind presented here are provedto be extremely important for understanding the evo-lution of orogens located in their hangingwall.

ACKNOWLEDGEMENTS

The authors are grateful to B. Lombardo, D. Castelli,I. Gabudianu and T. Zack for useful discussions andsuggestions. K. Bucher and T. Hirajima are gratefullyacknowledged for careful and constructive reviews of aformer version of the manuscript. The detailed reviewsby N. Froitzheim and an anonymous referee aregreatly appreciated.

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Received 27 August 2008; revision accepted 27 January 2009.



The PT path of the ultra-high pressure Lago Di Cignana and adjoining high-pressure meta-ophiolitic...

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