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Metamorphic evolution of preserved Hercynian crustal section in the SerreMassif (Calabria-Peloritani Orogen, southern Italy)

Gerolamo Angı̀, Rosolino Cirrincione, Eugenio Fazio, Patrizia Fian-nacca, Gaetano Ortolano, Antonino Pezzino

PII: S0024-4937(09)00496-4DOI: doi: 10.1016/j.lithos.2009.12.008Reference: LITHOS 2171

To appear in: LITHOS

Received date: 20 April 2009Revised date: 18 December 2009Accepted date: 19 December 2009

Please cite this article as: Ang̀ı, Gerolamo, Cirrincione, Rosolino, Fazio, Eugenio, Fi-annacca, Patrizia, Ortolano, Gaetano, Pezzino, Antonino, Metamorphic evolution ofpreserved Hercynian crustal section in the Serre Massif (Calabria-Peloritani Orogen,southern Italy), LITHOS (2009), doi: 10.1016/j.lithos.2009.12.008

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Metamorphic evolution of preserved Hercynian crustal section in the Serre Massif (Calabria-Peloritani Orogen, southern Italy) Gerolamo Angì,1 Rosolino Cirrincione,2 Eugenio Fazio,3 Patrizia Fiannacca,4

Gaetano Ortolano,5* Antonino Pezzino6

1TEL. – FAX – E-MAIL: 0957195786 – 0957195760 – g.angi@unict.it 2TEL. – FAX – E-MAIL: 0957195738 – 0957195760 – r.cirrincione@unict.it 3TEL. – FAX – E-MAIL: 0957195786 – 0957195760 – efazio@unict.it 4TEL. – FAX – E-MAIL: 0957195604 – 0957195760 – pfianna@unict.it 5*

TEL. – FAX – E-MAIL: 0957195786 – 0957195760 – ortolano@unict.it 6 TEL. – FAX – E-MAIL: 0957195746 – 0957195760 – pezzino@unict.it * Corresponding author: Gaetano Ortolano Authors’ affiliation Dipartimento di Scienze Geologiche, Università degli Studi di Catania, Corso Italia 57, 95129 Catania, Italy Abstract This paper presents and discusses the results of an integrated structural and petrological study, in order to entirely delineate the entire tectono metamorphic history of a still little known crystalline fragment of the southern Hercynian European Belt, currently framed within the central Mediterranean region after the superposition of the Alpine tectonics. These results were obtained by correlating P–T constraints yielded step by step with the sequence of the identified blasto-deformational relationships in an intermediate continental crustal level outcropping in the southern Serre Massif (Calabria). This allowed a detailed P–T evolution characterised by a multistage metamorphic history to be reconstructed. Structural investigations showed the presence of a pervasive mylonitic foliation, that obliterated most of the previous metamorphic textures. This fabric contains kinematic indicators consistent with an average top-to-ENE–NE sense of shear in the present-day geographic coordinates. In addition, the occurrence of late tectonic leucogranite rocks partly affected by subsolidus deformation, cut in turn by later undeformed ones, allowed the final stages of the shearing event to be bracketed at the same time as the Late Hercynian magmatic activity in the area. Microstructural investigation by quartz c-axis orientation pattern analysis allowed the temperature of shearing to be constrained as occurring under greenschist to amphibolite facies conditions. The latter are set in relation with the influence of the heat deriving from the intrusion of the Late Hercynian granitoids. Lastly, pressure temperature (P–T) pseudosection computations in the MnNaCaKFMASH system allowed a detailed P–T path to be reconstructed, consisting of an initial orogenic cycle characterised by a prograde lower amphibolite facies evolution, developing from P of 590 MPa at T of 500 °C to peak P–T conditions of 900 MPa at 530 °C. This stage was followed by retrograde quasi-adiabatic decompressional (P = 400 MPa; T = 500 °C), evolving towards an extensional deep-seated shearing, with P of 300 MPa at T of 470 °C. This last orogenic stage played a role in favouring the intrusion of granitoid bodies, which were indeed found to be partly affected by sub-solidus non-coaxial deformation. Progressive emplacement of large volumes of granitoid bodies gave rise to a gradually distributed thermal metamorphic overprint with thermal peak conditions at P of 300 MPa and T of 685 °C. This episode was finally followed by a low-pressure cooling path (P = 150 MPa; T = 500°C), consistent with the final unroofing stage of the former crystalline basement complex. A detailed reconstruction of the tectono-metamorphic evolution of this Hercynian continental crustal portion allowed a Late Palaeozoic geodynamic scenario to be envisaged, in which tectonic and magmatic processes mutually interacted to define, in a feedback-type evolution, the tectono-thermal regime operating during the gravitational collapse of a previously thickened Hercynian crust. Key words: Hercynian metamorphism; P–T pseudosection; quartz c-axis; Serre Massif, Calabria

1. Introduction

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The present-day distribution of European pre-Mesozoic basement blocks is mostly the result of

Palaeozoic orogenic processes, renewed by Alpine–Apennine large-scale nappe and strike-slip

tectonics, which locally produced a weak to pervasive metamorphic overprint (Fig. 1a).

In particular, the southern European Hercynian Belt was derived by accretion of the northern peri-

Gondwanan terranes to Laurussia and by subduction of small ocean basins in the Devonian–Early

Carboniferous (Stampfli and Borel, 2002; von Raumer et al., 2003), before final continental

collision in the Late Carboniferous. The subsequent slab rollback of the oceanic lithosphere was

responsible for the post-collisional extensional regime and the magmatic activity which affected the

European Hercynian belt during the Late Carboniferous–Early Permian (von Raumer et al., 2003,

and references therein).

The subsequent geodynamic evolution, leading to the formation of the present-day arcuate

Alpine–Apennine chains of the western Mediterranean area (Rosenbaum and Lister, 2004),

produced several sub-terranes, locally affected by early-to-late Alpine metamorphism (Bonardi et

al., 1987; Cheilletz et al., 1999; Pezzino et al., 2008).

The Calabria-Peloritani Orogen (CPO) (Fig. 1b) is an outstanding example of this complex

tectonic evolution. It is a composite segment of the western Mediterranean internal Alpine chain,

mostly comprising poly-orogenic multi-stage metamorphic rocks, presently merged with several

Hercynian (Pezzino, 1982; Atzori et al., 1984) or possibly older (Ferla, 2000; Micheletti et al.,

2007) sub-terranes. These rocks were locally overprinted during the different stages of the Alpine

metamorphic cycle, which also affected part of the Mesozoic oceanic-derived units and sedimentary

sequences (Liberi et al., 2006; Cirrincione et al., 2008; Fazio et al., 2008). Lastly, these basements

were definitively stacked by the Alpine–Apennine thin-skinned thrusting events in the central

Mediterranean area (Ortolano et al., 2005; Pezzino et al., 2008).

Within the CPO, the best-preserved relics of the southern European Hercynian Belt are

recognisable in the Sila and Serre Massifs, rather than in the Aspromonte Massif and Peloritani

Mountains, where a more intense Alpine reworking occurred (Pezzino et al., 1990; Cirrincione et al.,

1991, 2008; Atzori et al., 1994). In particular, the Serre Massif is one of the few places in the world

where it is possible to observe a nearly complete, tilted continental crustal section (Schenk, 1980,

1989, 1990) (Fig. 2a), such as is recognised only in other particular tectonic settings around the

world (e.g. Ivrea-Verbano zone in northern Italy, Dharwar craton in southern India, Fraser Range in

Western Australia, and Gold Butte block in the United States, among others). It therefore represents

a rare opportunity to study the relationship between different portions of the same crustal continental

section, emphasizing how different pressure (P) temperature (T) trajectories can be framed within

the same tectono-metamorphic evolution. Until now, the tectono-metamorphic evolution of the Serre

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Massif has been constrained by peak metamorphic estimates available only for the lower portions of

the crustal section, ranging from 750 MPa at 800 °C for the bottom levels to 550 MPa at 690 °C for

the top levels (Schenk, 1984, 1989). These data were later reviewed by Acquafredda et al. (2006,

2008), who suggested peak P–T values of 1100 MPa at 900 °C for the bottom of the lower crust, and

peak P–T values of 800 MPa at 700 °C for the top levels.

However, there are still no reliable P–T estimates available for the intermediate to upper levels

of the Serre Massif crustal section, characterised by two metamorphic complexes made up of

amphibolite to greenschist facies metamorphic rocks.

In this work, we provide the first integrated structural and petrological results aimed at

reconstructing the P–T evolution and deformation history of the amphibolite facies metapelites now

exposed at the bottom of the upper crustal levels of the southern Serre Massif, and also contributing

towards ascertaining the role played by the extensional (Del Moro et al 2000; Caggianelli et al.,

2000, 2007; Acquafredda et al., 2006) or compressional (Schenk, 1989) Late-Hercynian tectonics in

the exhumation of this sector of the Hercynian Belt.

In this view, after a detailed geological-structural investigation, we followed an integrated

approach consisting of petrographic-microstructural analysis (e.g. quartz c-axis orientation pattern

analysis) and thermodynamic modelling of the most informative identified metapelite samples by

means of P–T pseudosection computations in the MnNaCaKFMASH system.

2. Geological background

The Serre Massif represents the linkage between the southern (Aspromonte Massif and

Peloritani Mountains) and the northern (Sila and Catena Costiera) sectors of the CPO (Fig. 1b) and

can be briefly described as composed of three different complexes, as follows: a) the deepest

granulite facies metamorphic basement, made up of metagabbros, felsic granulites, metabasites, and

metapelitic migmatites (Maccarrone et al., 1983; Schenk, 1984, 1989; Fornelli et al., 2002, 2004;

Acquafredda et al., 2006, 2008); b) the middle crustal Late Hercynian batholith (Serre batholith)

composed of foliated tonalite with minor Qtz-diorite and gabbro, grading to more felsic and

peraluminous granitoid in upper crustal levels (D’Amico et al., 1982; Rottura et al., 1990; De Vivo

et al., 1992; Del Moro et al., 1994; Fornelli et al., 1994). At its south-western termination the Serre

batholith is intruded by the strongly peraluminous Cittanova granite (Atzori et al., 1977; Crisci et

al., 1979; D’Amico et al., 1982; Rottura et al., 1990; Graeßner et al., 2000) (Fig. 2a); c) the

intermediate to upper crustal portion, outcropping in the southern part of the Serre Massif,

composed of greenschist to amphibolite facies Palaeozoic metasedimentary and minor metavolcanic

successions (Colonna et al., 1973; Atzori et al., 1977; Bonardi et al., 1984; Acquafredda et al.,

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1987; Festa et al., 2003), locally intruded by discordant and concordant leucogranite dykes forming

an intricate network branching from the periphery of the main plutonic bodies (Colonna et al., 1973;

Borsi et al., 1976; Bonardi et al., 1984; Del Moro et al., 1994).

Early studies reported the Serre Massif as the result of the juxtaposition of several tectonic

slices, characterised by distinct tectono-metamorphic evolutions, that came into contact before the

intrusion of the Late Hercynian granitoids (Colonna et al., 1973; Amodio Morelli et al., 1976; Borsi

et al., 1976; Atzori et al., 1977; Gurrieri, 1980; Del Moro et al., 1986). In contrast, in more recent

studies (Schenk, 1980, 1984, 1989, 1990; Bonardi et al., 1984; Thomson 1994; Caggianelli et al.,

2000; Festa et al., 2003) it is regarded as a nearly complete continental crustal section that shows

shared tectonic evolution during most of the Hercynian and all of the Alpine orogenic cycles.

Studies of the deepest levels of the crustal section point to two different interpretations of the

tectono-metamorphic evolution of the Serre Massif. The first derives from the studies of Schenk

(1980, 1989, 1990), which indicate P–T evolution characterised by Early Hercynian granulite facies

metamorphism associated with an unusually high geothermal gradient, synchronous with the

development of penetrative deformational phases. This first orogenic metamorphic phase, for which

only poorly defined ages are given, exclusively for the initial prograde evolution (e.g. 450 ± 20 Ma;

Schenk, 1989), was followed by a static metamorphic event coinciding with extensive granitoid

magmatism at around 300 Ma (Rb-Sr micas ages and U-Pb zircon, monazite, and xenotime ages;

Schenk, 1980; Fornelli et al., 1994; Graeßner et al., 2000; Fiannacca et al., 2008).

By contrast, the second interpretation considers the bottom of the lower crustal section,

essentially comprising granulitic metagabbro, to be the result of multistage dehydration

decompression in relatively high pressure and high temperature conditions, passing from peak values

of 1100 MPa at 900 °C to 700–800 MPa at 650–700 °C for the retrograde P–T conditions

(Acquafredda et al., 2008). Conversely, the migmatitic paragneiss belonging to the top of the lower

Serre Massif crustal section would have been involved in P–T evolution consisting of medium-

pressure amphibolite facies prograde metamorphism, related to crustal thickening and heating,

followed by a multistage decompressional path, with an anatectic heating stage evolving to an

isothermal one, and by nearly isobaric cooling, associated with the final stages of exhumation

(Acquafredda et al., 2006). According to this view, Caggianelli et al. (2007) reviewed previous P–T

paths by means of thermobaric modelling, simulating the effects of extensional tectonics on the

geotherms during cooling of the mid-crustal granitoids. The results were considered by the authors

to be consistent with nearly isothermal decompression followed by isobaric cooling, supporting the

hypothesis that the magmatic and metamorphic evolution of the Calabria crust developed under

extensional tectonics, perhaps linked to a Late Hercynian slab break-off.

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The Serre Massif granitoids may be considered to belong to two different suites: one calc-

alkaline, metaluminous to weakly peraluminous, and one strongly peraluminous (Rottura et al.,

1990). Both are interpreted as late- to post-tectonic and were probably emplaced along ductile shear

zones in an extensional regime (Rottura et al., 1990; Caggianelli et al., 2007). In particular, the

strongly foliated calc-alkaline granitoids intruded earlier at somewhat deeper structural level,

whereas the unfoliated to weakly foliated strongly peraluminous and calc-alkaline types intruded

higher crustal domains.

Lastly, the intermediate to upper crustal section outcropping in the south-eastern Serre Massif

is characterised by the overlap of two complexes separated by a low-angle tectonic detachment,

dividing a lower metamorphic grade hanging wall complex (Stilo-Pazzano Complex) from a higher

metamorphic grade footwall metamorphic complex (Mammola Paragneiss Complex), both

overprinted by static metamorphism induced by the intrusion of the Late Hercynian granitoids

(Rottura et al., 1990; Fornelli et al., 1994) (Fig. 2b). The uppermost Stilo-Pazzano Complex (SPC)

includes low greenschist facies metapelite, marble, quartzite, and metavolcanic levels,

unconformably covered by a composite Mesozoic sedimentary succession (Festa et al., 2003). The

lowermost Mammola-Paragneiss Complex (MPC) comprises amphibolite facies paragneiss,

leucocratic gneiss, and amphibolite. The metamorphic rocks from both complexes are locally

intruded by Late- to post-Hercynian felsic to mafic dykes. To date, no geochronological data have

been produced to constrain the time of metamorphic evolution of the investigated crustal section,

although U-Pb monazite ages (Graeßner et al., 2000) for the upper crustal paragneisses of the

adjacent Aspromonte Massif indicate a metamorphic peak at 295–293 ± 4 Ma, coeval with the Serre

lower crust. Geochronological indications for early Hercynian events are given for metapelites of

southern Calabria in the Aspromonte Massif by the Rb/Sr biotite age of ca. 330 Ma (Bonardi et al.,

1987) and by a poorly constrained lower concordia intercept age of 377 ± 55 Ma (Schenk, 1990).

Our attention focussed on the reconstructing the tectono-metamorphic evolution of the MPC,

since it shows the best preserved evidence of the Hercynian multistage evolution, only locally

obliterated by static mineral and textural readjustments induced by the Late Hercynian thermal

overprint. This is clearly shown by the occurrence of: a) relic metamorphic assemblages

characterised by well-preserved porphyroblast zoning (e.g., zoned garnet), probably due to several

stages of prograde mineral growth; b) well-preserved syn-mylonitic textural and paragenetic

features, probably due to a later retrograde stage; c) local static mineral growth of cordierite and

biotite, clearly due to the last thermal overprinting stage.

3. Geological-structural features

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Field investigations (1:10000 scale) confirmed that the tectonic framework of the study area

consists of a middle to upper crustal metamorphic basement section, locally covered by an

unmetamorphosed Mesozoic sedimentary sequence. The basement rocks are juxtaposed along a

low-angle detachment, that subdivides: a) the lowermost Mammola-Paragneiss Complex,

comprising amphibolite facies paragneiss, leucocratic gneiss, and amphibolite; b) the uppermost

Stilo-Pazzano Complex, which includes low greenschist facies metapelites, marbles, quartzites, and

metavolcanics.

In agreement with Bonardi et al. (1984), we confirm that the two metamorphic complexes

share the same tectono-metamorphic evolution and are intruded by the same Late Hercynian plutonic

granitoids, represented here by the weakly peraluminous biotite ± amphibole granodiorite and

tonalite of the Serre batholith, as well as by several generations of widespread leucogranite dykes.

In particular, two main stages have been identified, consisting of a former prograde regional

metamorphic event followed by a retrograde regional one, both better preserved in the MPC,

accompanied by a thermal overprint due to the emplacement of Late Hercynian granitoid bodies.

Mineral growth associated to this latest phase outlasted deformation, as shown by randomly oriented

biotite and cordierite plates and centimetre-size andalusite spots in the country rock, most evident in

the phyllite of the SPC and, gradually, approaching the intrusive bodies.

Structural investigations showed that a pervasive syn-mylonitic texture, which developed

during the retrograde stages of the orogenic metamorphic cycle, defines foliation in the field.

Nevertheless, the older surfaces of previous metamorphic stages are preserved as relics within the

mylonitic foliation. The oldest one, related to the D1 deformational stage, is represented on outcrop

scale by relics of axial plane isoclinal folds (S1) (Fig. 3a), locally followed by a crenulation phase

(D2) with development of a crenulation cleavage (S2), more evident in micaceous-rich domains

(Fig. 3b) and in the SPC phyllites (Table 1).

The subsequent deformational event (D3) produced pervasive mylonitic foliation (S3) (Fig. 3c,

d) and a clearly defined stretching lineation (L3). The same event also locally involved some late

tectonic leucogranite dykes, in turn cut by later undeformed dykes (Fig. 3e).

S3, which partly obliterates the previous metamorphic structures, strikes NE–SW, with very

steep planes (strike-parallel stretching lineation) to the WNW–ESE, with average dip up to sub-

horizontal (dip-parallel stretching lineation) (Fig. 2c). L3, defined by elongate quartz and feldspar,

mainly runs ENE–SSW to NNW–SSE, plunging from 5° to 40°, with kinematic indicators showing

a constant top-to-ENE–NE sense of shear in the present-day geographic coordinates (Fig. 2). These

structural features indicate that a dextral transtensional shear zone developed along a non-planar

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surface (from sub-vertical to sub-horizontal), explaining the observed dispersion of the stretching

lineation (Fig. 2c).

The occurrence of weakly sheared late-tectonic leucocratic dykes, cut in turn by undeformed

ones (Fig. 3f), suggests that the final stages of the D3 phase occurred at the same time as Late

Hercynian magmatic activity in the area.

After the end of the Hercynian orogeny, the sedimentation of shallow-water Jurassic

limestones reveals a period of tectonic quiescence.

These results, now available for the middle to upper crustal levels, are consistent with existing

data from the middle-lower crust of the Serre Massif, which indicates shear zone activity affecting

metamorphic and granitoid rocks in high to low temperature conditions (Caggianelli et al., 2007, and

references therein).

Lastly, the sequential tectonic evolution consists of shallower ESE-verging metric

asymmetrical folding (D4) followed by a brittle thrusting stage (D5) (Fig. 3g). A subsequent NE–SW

brittle extensional fault system, accommodated by a NW–SE to N–S transtensional one, facilitated

the final stage of chain exhumation, within the framework of the eastward migration of the Apennine

southern orogen. According to Festa et al. (2003), this last evolutionary stage contributed to partly

disarticulating the previous Hercynian framework, playing a key role in the tilting of the present-day

Serre crustal section.

4. Reconstruction of blasto-deformational relationships

Petrographic studies defined the sequence of the porphyroblast growth-deformational

relationships of MPC metamorphic rocks. The observed relative timing relationships highlighted the

presence of multi-stage metamorphic evolution, consisting of an orogenic cycle partly overprinted

by a thermal one, both ascribable to the Hercynian orogenesis. WDS electron microprobe data were

obtained from thirty-two samples that show different stages of metamorphic evolution. Their mineral

compositions were used in conjunction with P-T pseudosections to estimate the P-T conditions of

the different stages of this evolution.

Mineral abbreviations, analytical conditions and representative analyses are reported in

Appendix 1.

4.1. Orogenic metamorphism

The orogenic cycle (i.e., M1–M3) consists of a prograde metamorphic stage defined by relic mineral

assemblages within zoned garnet, evolving towards retrograde stages, documented by garnet

resorption, followed by the development of greenschist facies mylonites. The kinematics and

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temperatures operating during this last orogenic metamorphic stage were also constrained by

analysis of quartz c-axis orientation patterns.

The earliest identified metamorphic stage (early-M1) is testified by relics of straight to

sigmoidal inclusion trails (S1), mainly composed of tiny zoisite grains within relatively high

spessartine garnet cores (Grt1 – Alm48–58Sps12–22Grs24–30Prp1–3) in assemblages with Pl1(An30–50),

Bt1(Ann35–45Phl25–35East10–15Sdph15–20), and Qtz (Fig. 4a; 5 – Table 1). This first stage evolves to a

second one defined by zoisite inclusion-free garnet (Grt2 – Alm62–70Sps7–10Grs16–26Prp6–10),

characterised by a smooth decrease in grossular and spessartine contents towards the outer core (Fig.

5), in equilibrium with Wmca1(Phg15–24) and Bt2(Ann32–40Phl18–25East12–18Sdph25–30). This evolution

suggests that garnet overgrowth (i.e., Grt2) originated at the expense of chlorite, white mica, quartz

and zoisite, according to the model reaction Chl + Wmca + Zo + Qtz = Grt + Bt + Pl + H2O (Menard

and Spear, 1993). Analysis of garnet zoning patterns shows the presence of a further overgrowth

stage (Grt3 – Alm75–78Sps1–4Grs4–14Prp12–18) in equilibrium with low anorthite plagioclase (Pl2 –

An37–3), probably linked to the orogenic peak metamorphic conditions (late-M1) (Figs. 4a, c and 5;

Table 1).

A subsequent crenulation event (D2) locally produced an S2 schistosity with syn-kinematic

growth of Qtz + Pl3 + Wmca2 ± Bt3 (M2 in Table 1) (Fig. 4b).

The orogenic metamorphic evolution proceeded towards a multi-stage retrograde history

consisting of an earlier stage, typified by garnet resorption (early-M3), and a late greenschist facies

mylonitic stage (late-M3).

This textural and mineralogical evolution is well preserved in the samples least affected by

thermal static effects. In such samples, resorption is clearly shown by embayed garnet rims that now

comprise aggregates of Wmca3(Phg3–10), Pl4(An30–35), Bt4(Ann36–44Phl34–38East7–12Sdph9–13) and

ripidolitic chlorite intergrowths with the observed spessartine-richest garnet (Grt4 – Alm52–55Sps20–

29Grs10–19Prp6–8) (Figs. 4c and 5; Table 1). This is indicative of a breakdown reaction, shown by the

inversion of the bell-shaped Mn zoning profile, which suggests resorption during a retrogressive

event (Spear, 1995).

The subsequent retrograde mylonitic deformational stage (D3) concludes the orogenic cycle,

producing a pervasive mylonitic fabric given by S/C texture, shear bands, oblique foliation

microstructure, boudinaged texture, and σ- and δ-type garnet and feldspar porphyroclasts (Fig. 4d, e;

Table 1). During this shearing event a strong quartz Lattice Preferred Orientation (LPO) developed,

due to the effects of a combined sub-grain rotation and grain boundary migration recrystallisation

regime (Passchier and Trouw, 1996).

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Textural analysis of kinematic indicators confirmed the top-to-ENE–NE sense of shear

already observed in the field. Syn-shearing crystal growth (late-M3) of Wmca4 (Phg6–10), ripidolitic

chlorite, epidote, and Pl5 (An24–26) (Fig. 5 – Table 1) documents shearing activity in greenschist

facies conditions. In addition, the widespread presence of boudinaged porphyroclasts suggests that

the shearing evolution operated under an extensional regime, which played a role in favouring

granitoid emplacement, as revealed by the presence of leucogranite dykes affected by syn-

emplacement shearing deformation (Figs. 3f and 4f).

Cross-cutting relationships in different outcrops are indeed highlighted, as early leucogranite

dykes are moderately foliated parallel to the field foliation, suggesting that they were involved in the

mylonitic event in near-solidus conditions. This evidence is supported by interfingered boundaries

between the host rock (i.e. mylonitic paragneiss) and the leucogranite dyke marked by a transitional

zone formed of both rock types sharing the same mylonitic foliation (Fig. 4 f).

Other microstructures emphasize a syn- to late-mylonitic recrystallisation regime, such as late

dynamic oligoclase overgrowth on porphyroclastic plagioclase, indicative of amphibolite facies

conditions (Figs. 5 and 6a, b). In this view, as the temperatures of the retrograde path of the orogenic

metamorphism were not above greenschist facies conditions, it is necessary to invoke an external

source of heat operating during the last stages of the shearing evolution (Kruhl and Vernon, 2005).

This conclusion is also supported by analysis of the quartz c-axis patterns of suitable quartz-rich

domains (Fig. 6a, c). Indeed analysis of quartz LPO, plotted on AVA diagrams (Sander, 1950), here

inferred by application of StereoNett 2.0 software (Stöckhert and Duyster, 1999; Appendix 2),

allowed both the kinematics (top-to-ENE sense of shear) and the temperature operating during the

shearing deformation to be constrained, as also amply demonstrated by Lister and Dornsiepen

(1982), Mainprice et al. (1986), Schmid and Casey (1986), and Heilbronner and Tullis (2006).

The inferred quartz c-axis patterns suggest that three different slip systems were activated

during shearing evolution. The AVA diagrams of Fig. 6c suggest, according to Heilbronner and

Tullis (2006), that a σ1 unfavourable slip system was replaced by a dominant basal <a> slip system,

consistent with the greenschist facies syn-kinematic mineral growth (Table 2) (Schmid and Casey,

1986). The activation of a prism <a> slip system may be imputed to the syn-tectonic increase in

temperature, consistent with late dynamic shearing deformational evolution developing in

amphibolite facies conditions (Schmid and Casey, 1986) (Table 2).

The occurrence of syn- to post-kinematic blastesis of feldspar and the activation of quartz slip

systems, both consistent with amphibolite facies conditions, may be related to the influence of heat

deriving from the intrusion of magmatic bodies, which took place during the last stages of the

greenschist facies shearing process.

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The emplacement of Late Hercynian plutons and minor dykes was followed by widespread

hypo-abyssal magmatism, leading to the generation of many undeformed leucogranite dykes which

crosscut the mylonitic foliation, representing a clear post-shear magmatic intrusion stage.

4.2. Thermal metamorphism

The thermal cycle (i.e., M4) linked to the emplacement of the main intrusive bodies, induced

mineralogical-textural re-equilibration in the host rocks, which partly obliterated the previous fabrics

by inducing annealing recrystallisation, without any evidence of partial melting effects.

The annealing process is revealed by randomly oriented blastesis, static mineral overgrowths,

and strain-free quartz aggregate levels with weak undulose extinction and straight grain boundaries,

as well as networks of triple junctions among grains of recrystallised quartz-feldspar phases (Table

1).

Spotted schist samples collected along transects perpendicular to and towards the contact with

the main plutonic bodies show a gradual increase in the metamorphic grade of the thermal event.

This is shown by the gradual change from the lowest-grade assemblages (Wmca + Bt + Crd + Pl +

Qtz ± Chl) to the thermal peak assemblage (Bt + And + Pl + Qtz + Crd ± St ± Wmca ± Sil ± Hc),

where andalusite crystals are commonly spatially associated with cordierite porphyroblasts and

occur as poikiloblasts enclosing randomly oriented biotite and opaque phases (Table 1). Staurolite

commonly occurs as a relic phase, replaced by And and/or Sil + Hc intergrowth or by growth of

Wmca + Crd. This last staurolite breakdown is explained by the hydration reaction of Pattison et al.

(1999): St + Bt + Qtz + H2O = Crd + Wmca.

Garnet-bearing paragneiss developed a static mineral assemblage given by tabular

porphyroblasts of biotite (Bt5 – Ann29Phl16East20Sdph35) in textural equilibrium with inclusion-free

almandine-rich garnet (Grt5 – Alm81–82Sps3–4Grs3–4Prp11–12), the latter forming sub-euhedral to

euhedral rims on the previous syn-tectonic garnet (Fig. 4g).

The peak of the thermal event was finally followed by a retrograde stage (late-M4), as

documented by the occurrence of retrograde andalusite in cordierite blasts and by partial to complete

sericitisation of andalusite crystal rims and pinitisation of cordierite blasts.

5. Thermo-barometric evolution

According to Zeh (2001), several samples which underwent the same tectono-metamorphic

evolution can record different steps of the same P–T evolution. This may be due to the different

bulk rock chemistry and/or different textural development associated with the single episodes of

polyphase metamorphic evolution. In this view, on the basis of the reconstructed blasto-

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deformational relationships, four suitable samples were selected to define the entire P–T evolution

of MPC metapelites.

For instance, well preserved assemblages of the relic prograde metamorphic stage were

observed in sample MA271 (Fig. 2; Table 3). This sample is characterised by the occurrence of

garnet with well-preserved growth zoning, probably due to a weak effect of the retrograde mylonitic

stage, as well as by a weak rehomogenisation effect due to the thermal metamorphic overprint.

Instead, the presence of highly resorbed porphyroblastic garnets highlights the fact that

sample AR221 may be considered the most suitable one to quantify the P-T conditions of the

earliest stages of the retrograde evolution. This is represented here by breakdown assemblages

which are well-preserved within garnet resorption embayments.

Although syn-mylonitic texture is well developed in many samples, sample GR164 was found

to be the most informative one, due to clearcut porphyroclastic plagioclase preserving weakly re-

equilibrated syn-mylonitic assemblages within its pressure shadows.

The effects of thermal metamorphism, due to an initial late-dynamic stage that evolved to a

static overprint, were observed in several samples characterised by both late-mylonitic annealed

textures and well-developed static mineral growth. The selected sample, GR166, best shows the

static effect of the thermal overprint, as testified by the observed foam texture and by randomly

oriented cordierite, biotite, and andalusite porphyroblasts.

For these four samples, thermodynamic modelling by the P–T pseudosection approach was

applied, to estimate the thermobarometric conditions of the identified mineralogical equilibria.

5.1. Methodological approach

Thermodynamic modelling of the observed phase equilibria was performed by a free energy

minimisation approach with Perplex software (Connolly, 1990; Connolly and Petrini, 2002).

The Perplex package consists of a suite of programs for calculating phase diagrams and

thermodynamic equilibria on the basis of variable solid solution compositions, approximated by a

series of arbitrarily defined components (i.e., pseudocompounds), dealing with solutions up to three

independent mixing sites and up to three species mixed on each site (Connolly, 1990). With this

approach, the obtained P–T constraints may be affected by two principal sources of uncertainty. The

first is due to the use of the experimentally derived thermodynamic datasets (e.g., Holland and

Powell, 1998), and can be estimated at about 20 MPa and 30 °C (Hetherington and Le Bayon,

2005). The second derives from the specific compositional spacing chosen among the end-members

of the solid solution models used (Cirrincione et al., 2008). This last source of error was recently

minimised in the latest version of Perplex by Connolly (2008), which uses a new computing

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approach called adaptive optimization strategy: it consists of the iterative redefinition of the

accuracy of the pseudocompound approximation by means of an increase (or decrease) in the

spacing of the solid solution models during the course of the calculation.

Taking into account the above limitations, the P–T pseudosections of the selected samples

were then constructed by using: a) the whole rock composition, as measured by XRF (Appendix 3);

b) the MnO, Na2O, CaO, K2O, FeO, MgO, Al2O3, SiO2, and H2O (MnNCKFMASH) oxides as a

suitable chemical system, assuming SiO2 as a saturated thermodynamic component and H2O as a

saturated phase component; c) the internally consistent thermodynamic database and the

compensated Redlich-Kwong fluid equation of state of Holland and Powell (1998), updated by the

same authors in 2002; d) solut_08.dat as the solid solution model database, reported in the most

recent version of Perplex (Appendix 3).

Using fixed bulk compositions in the pseudosection calculations implies that the chemical

systems in questions must be considered in overall equilibrium. This condition is reliable as long as

no chemical fractionation occurs as a consequence of the multi-stage growth of some minerals, such

as garnet and/or plagioclase, since this is potentially capable of modifying the effective bulk

composition (Stüwe, 1997) operating during metamorphic evolution.

The reliability of the inferred P–T constraints was thus verified step by step, considering the

XRF chemical data to be representative of the reacting rock volumes (i.e., assuming closed system

behaviour) only if: a) the fit between observed and calculated assemblages and their coexistence in

textural equilibria can be demonstrated (Connolly and Petrini, 2002; Cirrincione et al., 2008); b) the

intersection at least of three garnet compositional isopleths (e.g., almandine, grossular, and

spessartine) can define relatively small areas in the pseudosection P–T space (Evans, 2004).

5.2. P–T estimates

Textural and minero-chemical features of the selected rock samples of the MPC document

discrete segments of the entire identified multistage metamorphic evolution, consisting in an

orogenic cycle partly overprinted by a thermal metamorphism, both ascribable to the Hercynian

orogeny.

5.2.1. Sample MA271

The oldest identified mineralogical association was found to be well preserved in sample

MA271, characterised by a coarse-grained texture given by quartz-feldspar layers alternating with

subordinate lepidoblastic ones made of biotite that commonly show decussate texture. Both layers

also contain widespread porphyroblasts of zoned garnet (Fig. 7a) and diablastic biotite.

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The intersections of garnet inner core composition isopleths (Alm54Sps15Grs30Prp1) (field A –

Fig. 8), characterised by relatively high spessartine content (Fig. 7a) and zoisite-rich inclusion trails,

constrained the P–T conditions of the earliest identified metamorphic stage (early-M1) at a pressure

of 590 MPa and temperature of 500 °C (Table 4).

The equilibrium assemblage of field A in the pseudosection P–T space fits the observed

paragenetic equilibrium except for the absence of biotite, chlorite, and white mica of adequate

composition, because they were re-equilibrated during the following metamorphic stages (Table 4).

The subsequent metamorphic stage, marked by the breakdown reaction of the inner cores of

zoisite-rich garnet, developed a low-spessartine garnet overgrowth (Alm69Sps1-2Grs25-26Prp4) in

equilibrium with relatively anorthite-poor plagioclase (An20–22) (Table 4). The above assemblage,

characterised by the steep bell-shaped Mn garnet compositional profile (Fig. 7a), provided garnet

isopleths that intersected over a relatively large range of P-T conditions (Fig. 8). This intersection

was further restricted by anorthite isopleths, allowing average P–T values of 900 MPa at 530 °C to

be obtained (Fig. 8). The observed enlargement of the region defined by the intersection of garnet

isopleths in the pseudosection P–T space marks the beginning of the change in the effective bulk

composition of the system, probably due to fractionation processes during prograde metamorphism.

The following observed retrograde paragenesis, depicted by smooth spessartine garnet

enrichment towards the rim (Fig. 7a) (Alm74Sps4Grs4Prp22) in equilibrium with an anorthite-richer

plagioclase (An34–36), did not allow any useful intersection (Fig. 8). This suggests that most of the

earlier porphyroblasts did not take part in the reacting rock volume during retrograde evolution.

Lastly, no reliable intersections were observed for the equilibrium assemblage ascribable to the

thermal metamorphic reactions given by garnet rim (Alm82Sps3Grs3Prp12) in equilibrium with static

biotite (Fe2/(Fe2+Mg)61-64) (Table 4). This is clearly shown in the pseudosection of Fig. 8, where a

significant divergence of the mineral compositional isopleths occurs.

5.2.2. Sample AR221

The P–T estimates of the following prograde orogenic metamorphic stages up to relative peak

conditions, as well as the former stages of the retrograde evolution, were constrained by the

mineralogical assemblages of sample AR221. This sample has a fine-grained grano-lepidoblastic

matrix given by quartz-feldspar layers alternating with subordinate lepidoblastic layers of white

mica, biotite, and chlorite. These layers are interrupted by syn- to late-kinematic embayed garnet

porphyroblasts and by late- to post-kinematic staurolite and andalusite spots. Resorbed garnet locally

shows evidence of a new garnet rim.

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The oldest mineralogical assemblage identified is characterised by a garnet inner core of

intermediate spessartine content (Alm55Sps14Grs27Prp3) (Fig. 7b) in equilibrium with plagioclase

inclusions of intermediate anorthite content (An39-41).

This assemblage, interpreted as developing at the expense of the zoisite-rich garnet inner cores,

indicated the P–T trajectory of the prograde metamorphic evolution, yielding P-T estimates of 650

MPa and 520 °C (field A’, Fig. 9). These estimates are slightly higher then those constrained for the

earliest paragenesis identified in sample MA271 (Fig. 8; Table 4).

A further mineralogical assemblage, given by garnet overgrowth with a slight decrease in

grossular and spessartine content (Alm64Sps8Grs23Prp5) (Fig. 7b) in equilibrium with a low-

anorthite plagioclase (An35-37) and relatively high-phengite white mica (Phg14–20) yielded a

temperature of 550 °C and a pressure of 750 MPa (field B’, Fig. 9), constrained by garnet isopleth

intersections. The resulting P–T estimates are close to the peak metamorphic conditions observed in

sample MA271. These data suggest that no significant crystal fractionation occurred during

prograde metamorphism, and are confirmed by the good match between the chemical composition

of predicted and observed minerals (Table 4). The exception to this is the poor match between

predicted and observed compositions for biotite and chlorite, since they were entirely re-

equilibrated during the following evolutionary stages.

The observed marked increase in the spessartine content of garnet rim composition highlights

the irregular garnet zoning, and results in an inversion of the bell-shaped compositional profile (Fig.

7b). This feature is consistent with the effects of retrograde metamorphic evolution (Spear, 1995), as

confirmed by the close intersection of garnet rim isopleths (Alm52Sps29Grs10Prp9) in equilibrium

with intermediate phengite white mica (Phg10–7), which provided reliable pressure values of 470-375

MPa at slightly decreasing temperatures of 500–520 °C (field C’, Fig. 9) with respect to relative

peak conditions. The reliability of these data, supported by the good match between observed and

predicted phase equilibria (Table 4), suggests that no significant chemical fractionation occurred

during the first phases of retrograde metamorphic evolution. This was probably due to the

widespread development of fractures and embayments in the pre-kinematic garnet porphyroblasts,

which allowed the significant iso-chemical behaviour of the system to be maintained.

Lastly, the following observed quasi-static mineral assemblage of this sample, given by

staurolite and biotite crystal growth, sometimes replaced by static andalusite in textural equilibrium

with garnet rims, did not match any useful intersections in the pseudosection P–T space (Fig. 9).

This prevented the use of XRF data to estimate the P–T constraints of this specific assemblage

(Table 4). Nevertheless, the textural features are consistent with multi-stage thermal metamorphism

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consisting of a first late-dynamic stage, responsible for staurolite and biotite crystallisation, and a

later static stage, developing at lower pressure, marked by the growth of andalusite.

5.2.3. Sample GR164

The subsequent shearing stage, which concluded the evolution of the orogenic cycle, was

constrained by the phase relationships identified in sample GR164. This sample exhibits a strong

mylonitic texture given by widespread sigmoidal plagioclase and rare garnet porphyroclasts, mantled

by a syn-mylonitic assemblage of quartz, white mica, biotite, zoisite, and chlorite.

Rare and not very well-preserved mineral assemblages, given by scarce pre-kinematic garnets

(Alm54Sps18Grs24Prp4) in equilibrium with the inner cores of albite plagioclase and relatively high-

phengite white mica (Phg16-19) (Fig. 10a), did not allow any available P–T constraints to be applied

to the peak (late-M1) or the early retrograde metamorphic (early-M3) evolutionary stages (Fig. 11;

Table 4).

In contrast, the syn-mylonitic mineral assemblage, which is well-preserved in the pressure

shadows of pre-kinematic porphyroclasts (Fig. 10a), provided useful constraints on the retrograde

evolution of the orogenic metamorphic cycle. This assemblage, characterised by oligoclase reaction

rims (An26) in equilibrium with phengite-poor white mica (Phg6–9) and chlorite with Fe2/Fe2 + Mg

ratios from 56 to 59, did allow P–T constraints to be obtained (field A’’, Fig. 11; Table 4), at

pressure of 380-200 MPa and temperatures of 500-470 °C.

The reliability of the obtained P–T shearing estimates was confirmed both by the absence of

garnet in the P–T pseudosection stability field and by the good match between observed and

predicted mineral parageneses (Table 4). This last result was further strengthened by the significant

overlap between petrologically (Fig. 11) and microstructurally derived temperatures (Fig. 6; Table

2).

No reliable assemblages ascribable to the effects of thermal metamorphism were observed for

sample GR164.

5.2.4. Sample GR166

The P–T conditions of the thermal metamorphic cycle, linked to the emplacement of the main

intrusive bodies, were constrained by the mineral assemblages identified in sample GR166. This is a

garnet-free micaschist characterised by the quasi-foaming texture of quartz, cordierite, plagioclase

and andalusite in granoblastic layers and alternating subordinate layers of diablastic biotite with

minor white mica. Staurolite and sillimanite occur as relic minerals; chlorite is usually

pseudomorphic on cordierite blasts.

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The sequence of mineral parageneses identified here revealed a relic assemblage given by syn-

tectonic mineral growth of biotite (Fe2/(Fe2 + Mg)61–64), white mica (Phg10–5) and staurolite,

probably formed by the breakdown reaction of a pre-existing garnet + chlorite assemblage already

recognised in the other samples as the result of the initial increase in temperature during the latest

stages of the orogenic cycle.

This assemblage did not allow any P–T constraints to be made (Table 4), showing that the

chemical compositional system used is not representative of the specific effective bulk rock

chemistry (Fig. 12).

Conversely the following observed mineral paragenesis given by porphyroblastic biotite

plates (Fe2/(Fe2 + Mg)69–70) and widespread cordierite (Mg/(Fe2 + Mg)45–47) rarely accompanied by

sillimanite (Fig. 10b), provided useful constraints about the peak conditions of the thermal

metamorphic cycle by the available intersection between Fe2/Fe2 + Mg biotite and Mg/Fe2 + Mg

cordierite isopleths, taking into account the garnet-free cordierite field, (P = 300 MPa; T = 685 °C)

(field A’’’, Fig. 12), since garnet was not in equilibrium with the identified paragenesis (Table 4).

Peak temperature estimates are considered reliable in view of the good match between

observed and predicted mineral parageneses, except for the lack of plagioclase of adequate

composition (Table 4), probably not found because of its scantiness in our sample.

The last observed mineral paragenesis, given by andalusite in equilibrium with oligoclase

plagioclase (An24) and widespread pseudomorphic chlorite (Fe2/(Fe2 + Mg)61–63) on cordierite blasts,

allowed the P–T conditions of the retrograde stages to be constrained following the peak conditions

of the thermal cycle.

A pressure of 150 MPa at a temperature of 500 °C was obtained by means of Fe/Fe+Mg

chlorite and anorthite content isopleth intersections, which bracketed a potential intersection area

(Fig. 12c), further restricted to a field containing stable andalusite and without garnet (lacking in our

sample) (Fig. 12a). The final inferred P–T field (field B’’’, Fig. 12) is interpreted as the result of

cooling evolution due to the final exhumation stages, which followed the achievement of peak P–T

conditions in the thermal metamorphic cycle.

6. Derived P–T path and geodynamic implications

These results allow the tectono-metamorphic evolution of a representative sector of Hercynian

Calabrian crust exposed in the south-eastern Serre Massif to be constrained, correlating P–T

constraints yielded step by step with the sequence of the identified blasto-deformational

relationships identified in representative samples.

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Specifically, the tectono-metamorphic evolution of amphibolite facies garnet-bearing

paragneiss and micaschist of the Mammola Paragneissic Complex was investigated, since they were

the most suitable rock types revealing the entire Hercynian metamorphic history of the southern

Serre Massif basement rocks, a still little-known crustal portion belonging to the southern European

Hercynian Belt.

Results allowed a detailed P–T path to be constructed, and suggested that these crystalline

rocks underwent two Hercynian metamorphic cycles, the first consisting of a clockwise Barrovian-

type orogenic evolution, and the second defining late- to post-tectonic thermal episodes.

The P–T conditions associated with the orogenic cycle were constrained by P–T pseudosection

computations based on the XRF bulk rock chemistry of samples MA271, AR221 and GR164 (Figs.

8, 9, 11). The earliest metamorphic stage (early-M1) was identified by garnet isopleth intersections,

which yielded a pressure of 590 MPa and a temperature of 500 °C in the pseudosection P–T space of

sample MA271. This sample also yielded the orogenic peak conditions (late-M1) through isopleth

intersections of garnet outer core composition in equilibrium with observed medium-anorthite

plagioclase, defining a pressure of 900 MPa at a temperature of 530 °C (Fig. 13).

The garnet isopleth thermobarometry of sample AR221 allowed the following P–T stages of

prograde evolution to be constrained, yielding a pressure of 650 MPa at a temperature of 520 °C up

to (late-M1) conditions with corresponding 750 MPa and 550 °C (Fig. 13).

The inferred peak P–T conditions of samples MA271 and AR221 are consistent with relatively

high-pressure lower amphibolite facies metamorphism, which can be interpreted as due to the crustal

thickening stage of the Hercynian orogenic process.

Similar prograde P–T evolution (Fig. 14) was identified by Acquafredda et al. (2006) in the

migmatitic paragneiss of the uppermost part of the lower crustal section presently outcropping in the

northern part of the Serre Massif. This suggests that the migmatitic paragneiss and lower

amphibolite facies paragneiss of the Mammola Complex represent similar crustal levels at the end of

crustal thickening, as shown by both the shared early P–T values and comparable peak pressure

estimates (Fig. 14).

In addition, sample AR221 shows petrographic evidence of ragged edges in garnet

porphyroblasts as well as inversion of the bell-shaped Mn zoning-profile, suggesting a pervasive

resorption process during the first stage of the retrograde P–T trajectory (early-M3). A pressure of

400 MPa and a temperature of 500 °C were estimated by isopleth intersections of garnet rim

composition (Fig. 13). This P–T estimate, characterised by slightly decreasing temperature, depicts

a quasi-adiabatic decompression path from 900 MPa to 400 MPa, interpreted as due to relatively

fast crustal thinning. According to Escuder Viruete et al. (2000), such a decompression trajectory

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facilitated the preservation of the earlier prograde assemblages, as shown by the presence of well-

preserved zoned garnet in our samples. Conversely, according to Acquafredda et al. (2006), the

migmatitic paragneiss of the lower crustal portion underwent slow thermal re-equilibration, which

obliterated or attenuated the original prograde mineral growth zoning as a result of homogenisation

effects by intracrystalline diffusion. This process probably also causes an underestimate of the

pressure peak of Acquafredda et al. (2006) (850 MPa), which was indeed found to be slightly lower

than the peak pressure estimate reported in this paper (900 MPa). Similar or higher prograde peak

pressures have also been found for amphibolite facies rocks outcropping in the Hercynian segment

of northern Sardinia (Di Vincenzo et al., 2004; Franceschelli et al., 1989), giving a baric peak above

1000 MPa within a temperature range of 480—550 °C for garnet bearing rocks.

Subsequent P–T constraints show how the quasi-adiabatic decompression path, due to the

beginning of the exhumation chain, evolved towards a further uplifting stage, which developed along

a mylonitic shear zone at an average pressure of 300 MPa with temperatures of 500 °C to 470 °C

(Fig. 13). These P–T conditions were estimated by both phengite–chlorite equilibria and the

intersections between the stability fields of other syn-shearing parageneses observed in the

pseudosection P–T space of sample GR164. The inferred P–T values depicted the retrograde

trajectory of the orogenic path, which was interpreted as linked to a retrograde greenschist facies

shearing stage due to the activation of a regional-scale ductile shear zone. Shearing temperature

estimates were also confirmed by analysis of quartz c-axis orientation patterns, suggesting that the

main activated slip system is consistent with greenschist facies condition (Fig. 6; Table 2).

The mylonitic shearing stage gave rise to a pervasive field foliation, characterised by a well-

defined stretching lineation, which appears to disperse from WSW–ENE to SW–NE, locally passing

to N–S orientation, due to the variable orientation and inclination of the main shear surface.

However, kinematic indicators show an average top-to-ENE–NE sense of shear in the present-day

geographic coordinates.

Analogous retrograde P–T paths (Fig. 14), characterised by a first decompressional stage

followed by syn-shearing exhumation, are also known for the crystalline basement rocks of the

lower crust of the northern Serre Massif (Schenk, 1989; Acquafredda et al., 2006; Caggianelli et al.,

2007). Schenk (1989) considers this evolution to be related to the continental collision process

responsible for the first uplifting stage of the deep crust, whereas Acquafredda et al. (2006) and

Caggianelli et al. (2007) report that similar P–T evolution can be interpreted as due to Late

Hercynian extensional tectonics, perhaps linked to the orogenic collapse of the chain.

The latter interpretation fits better our hypothesis suggesting the presence of a dextral

transtensional shear zone, revealed by a stretching lineation sub-parallel to the direction of the main

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shear plane (Fig. 2) in rocks characterised by widespread pull-apart garnet porphyroclasts filled by

quartz and chlorite (Fig. 4e).

The activity of this extensional shear zone in the southern sector of the Serre Massif can be

correlated with the Late Carboniferous–Early Permian regional extensional regime described by

Ziegler (1993), and can be framed within a geodynamic scenario involving gravitational collapse of

the previously thickened Hercynian Belt, as reported for other peri-Mediterranean terranes (Concha

et al, 1992; Doblas et al., 1994 and refs. therein; Costa & Rey, 1995; Rossi et al., 2006; Giacomini

et al., 2006).

Late tectonic activity due to the extensional shearing process was accompanied by the early

intrusion of the Late Hercynian granitoids, in turn responsible for the syn- to post-kinematic

blastesis of feldspar over previously formed mylonitic porphyroclasts (Fig. 6a, b). This evidence is

also supported by the activation of quartz slip systems, consistent with amphibolite facies

conditions, partly replacing previously activated quartz slip systems related to the syn-mylonitic

greenschist facies process (Fig. 6c; Table 2).

In addition, the local cross-cutting relationships between mylonitic wall rocks and various

generations of Late- to post-Hercynian leucogranite dykes, and related textural features,

differentiate late-tectonic dykes from post-tectonic undeformed ones. The former, usually

discordant, are characterised by interfingered boundaries with the host rock, showing a moderate

internal foliation parallel to the mylonitic field foliation. They are considered as involved in the late

stages of the mylonitic process in near-solidus condition. The latter, which sharply cut the mylonitic

foliation or are at times para-concordant but never foliated, are interpreted as post-tectonic, since

they are totally devoid of deformation. In this view, the mylonitic shearing process can reliably be

constrained to the Late Hercynian, consistently with the already suggested late- to post-tectonic

emplacement ages of the granitoid bodies in the southern Serre Massif (Borsi et al., 1976; Rottura et

al., 1990; Del Moro et al., 1994; Caggianelli et al., 2000).

The emplacement of granitoids caused various mineralogical and textural adjustment to the

mylonitic texture. The P–T estimates for the consequent thermal metamorphic cycle are constrained

by the mineral assemblages identified in sample GR166 (Fig. 12). This sample was significantly

affected by thermal annealing and minero-chemical re-equilibration, providing assemblages which

constrained the peak and retrograde path of the thermal metamorphic cycle. Peak temperature

estimates of the thermal event were inferred by the available intersections between Fe2/(Fe2 + Mg)

biotite and Mg/(Fe2 + Mg) cordierite isopleths, yielding a pressure of 300 MPa at a temperature of

685 °C. Although relatively high temperature conditions were attained, the peak of the thermal

event did not trigger partial melting in the host rock, but only produced pervasive mineral static

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growth, weakly to strongly modifying the previous mylonitic microfabrics by an annealing

recrystallisation process. This was probably due to the fast exhumation from higher (900 MPa) to

lower pressure conditions (300 MPa).

At the same time, however, the migmatitic paragneiss belonging to the top of the lower Serre

Massif crustal section followed a decompressional path at a higher thermal regime (Fig. 14), which

caused pervasive anatexis (Acquafredda et al., 2006), highlighting the fact that Mammola

paragneiss and migmatitic paragneiss of the lower crustal section underwent different metamorphic

histories after prograde peak conditions had been attained.

This assumption may be explained by suggesting that the Mammola paragneiss and migmatitic

paragneiss shared the same crustal thickening evolution up to their pressure peaks. They then

followed independent P–T trajectories, marked by different retrograde evolution (Fig. 14),

characterised by: a) fast exhumation in the Mammola paragneiss, as shown by the observed quasi-

adiabatic decompression path; b) progressive heating in the migmatitic paragneiss, after the

attainment of the pressure peak shown by migmatite.

These discrepancies are explained by the different roles played by the extensional shear zone

in the two crustal sectors, which contributed to fast exhumation of the Mammola paragneiss,

whereas the migmatitic paragneiss remained at a deeper crustal level.

This reconstruction may also be viewed as consistent with new P–T data for the base of the

lower crustal section (Acquafredda et al., 2008), essentially composed of granulitic metagabbro with

peak P–T values of 1100 MPa at 900 °C (Fig. 14), much higher than the pressure peak estimates of

the previously described crustal sectors of the Serre Massif.

Lastly, the peak of thermal metamorphism was followed by retrograde evolution constrained

at 150 MPa and 500 °C by pseudomorphic chlorite on previous static cordierite in equilibrium with

retrograde plagioclase (Fig. 13). These inferred P–T estimates are interpreted as the result of

cooling due to the final exhumation stages, which followed the peak P–T conditions of the thermal

metamorphic cycle. This evidence is consistent with the analogous evolution identified by Schenk

(1989) who, for the lower crust of the Serre Massif, suggests a final exhumation stage along a

cooling trajectory throughout the tectonic quiescence of the Mesozoic (Festa et al., 2003).

7. Conclusion

This paper presents an integrated structural and petrological study of the lowermost part of the

upper crustal section of the Serre Massif. Results allow the Hercynian multi-stage evolution to be

constrained for this sector of the Calabrian Peloritani Orogen, and new geological and geodynamic

constraints are also provided.

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On an outcrop scale, the field foliation is given by a pervasive syn-mylonitic schistosity

surface. Older surfaces are preserved only as relics within the mylonitic foliation. Widespread

kinematic indicators show an average top-to-ENE–NE sense of shear in the present-day geographic

coordinates, suggesting a possible transcurrent dextral component along the main shear zone

surface. The local occurrence of late-tectonic weakly deformed leucogranite dykes, in turn cut by

later undeformed ones, also suggests that the final stages of the mylonitic event developed at the

same time as Late Hercynian magmatic activity in the area.

Petrographic studies allowed the blasto-deformational relationships of the investigated rock

types to be constrained, and these were useful in defining the P–T conditions of the assemblages

identified by the P–T pseudosection tool. Quartz c-axis orientation pattern analysis also indicated the

activation of quartz slip systems consistent with greenschist facies shearing, partly influenced by the

activation of other slip systems consistent with higher temperature conditions. This evidence is

interpreted as due to the temperature increase of the early stages of granitoid emplacement coeval

with the latest stage of the mylonitic phase.

Lastly, thermodynamic modelling allowed step-by-step definition of the single stages of the

tectono-metamorphic evolution on various selected samples, aiming at reconstructing a final P–T

path in which all these P–T estimates were summarised (Fig. 14). This reveals that the bottom levels

of the upper crustal portion of the Serre Massif underwent multi-stage metamorphism, consisting of

an orogenic cycle waning at the time of the first emplacement of late- to post-tectonic granitoids,

which were then responsible for a quasi-static thermal metamorphic overprint (rimming garnet and

plagioclase).

In detail, the orogenic cycle was characterised by: a) a Hercynian crustal thickening stage in

the prograde lower amphibolite facies, constrained by isopleth thermobarometry on bell-shaped

zoned garnet and plagioclase; b) a quasi-adiabatic decompression path, due to a first crustal thinning

episode, documented by the inversion of the garnet bell-shaped profile and by deep embayment on

previous garnet porphyroblasts; c) a retrograde greenschist facies mylonitic stage showing

prolongation of tectonic denudation, consistent with tectonic transport along a dominant extensional

shear zone. This last episode was associated with the emplacement of the Late Hercynian granitoids,

which caused late- to post-tectonic pervasive thermal metamorphism (blastesis of staurolite,

cordierite and sillimanite). This was followed by a final retrograde trajectory due to cooling and

exhumation, revealed by andalusite spots.

The results presented here are a new source of information, helping to clarify the Hercynian

metamorphic history of a complete exposed crustal section. They also delineate details of Palaeozoic

tectono-metamorphic evolution, consisting of an initial orogenic cycle characterised by prograde

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lower amphibolite facies evolution, followed by retrograde quasi-adiabatic decompression, evolving

towards deep-seated extensional shearing evolution, favouring granitoid intrusions in an extensional

setting.

Acknowledgment

We would like to thank Kevin Mahan, an anonymous reviewer, and the topic editor Ian Buick for

their useful comments and suggestions, which helped to improve the quality of this manuscript.

Financial support from MIUR (PRIN 2007 project: ‘‘Strain rate in mylonitic rocks and induced

changes in petrophysical properties across the shear zones’’) is gratefully acknowledged.

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APPENDIX A. Mineral abbreviations, microprobe equipment characteristics and representative analyses Used mineral abbreviations are after Kretz (1983), reviewed by Siivola and Schmid (2007). Mineral chemistry investigations were carried out by means of a CAMECA SX50 electron microprobe (EMP) equipped with four WDS spectrometers and one EDS spectrometer at the CNR-IGG, Unit of Padova. Operating conditions were set at 15 keV accelerating potential, 15 nA beam current and peak counting times of 15s. The PAP correction method was applied. Mineral formulae and ferric/ferrous iron ratios were calculated by MINPET 2.02 software (Richard, 1995), on the basis of 12 oxygens and 8 cations for garnet, 8 oxygens for feldspars, 24 oxygens for

white micas and biotite, and 36 oxygens for chlorite.

Representative mineral analyses of garnet Sample MA271 AR221 GR164 SiO2 37.63 37.87 37.81 37.58 37.49 37.43 37.50 37.68 36.81TiO2 0.29 0.00 0.02 0.00 0.01 0.17 0.14 0.05 0.09Al2O3 20.72 21.03 21.32 21.17 21.25 20.91 21.07 21.13 21.32Cr2O3 0.01 0.00 0.01 0.00 0.00 0.03 0.03 0.01 0.04FeOtot 25.00 35.29 37.15 36.84 36.19 25.19 26.15 28.67 28.23MnO 6.31 1.16 1.34 1.63 1.77 6.3 5.15 3.70 7.29MgO 0.32 2.99 3.00 2.84 2.7 0.77 0.76 1.17 1.63CaO 10.26 3.26 1.08 1.08 1.28 9.46 9.47 8.15 4.51Na2O 0.04 0.00 0.00 0.00 0.00 0.00 0.06 0.04 0.00Total 100.58 101.06 101.73 101.14 100.69 100.26 100.33 100.60 99.92

Si 3.009 2.997 3.002 3.004 3.009 2.998 2.998 3.006 2.973Al IV 0.00 0.003 0.00 0.00 0.00 0.002 0.002 0.00 0.027Sum_T 3.009 3.00 3.002 3.004 3.009 3.000 3.000 3.006 3.000AlVI 1.951 1.957 1.993 1.993 2.008 1.970 1.982 1.985 2.000Fe3+ 0.00 0.04 0.00 0.00 0.00 0.004 0.005 0.00 0.007Ti 0.017 0.00 0.001 0.00 0.001 0.010 0.008 0.003 0.005Cr 0.001 0.00 0.001 0.00 0.00 0.002 0.002 0.001 0.003Sum_A 1.969 1.997 1.995 1.993 2.009 1.986 1.997 1.989 2.016Fe2+ 1.672 2.296 2.466 2.462 2.429 1.683 1.743 1.913 1.899Mg 0.038 0.353 0.355 0.338 0.323 0.092 0.091 0.139 0.196Mn 0.427 0.078 0.09 0.11 0.12 0.427 0.349 0.25 0.499Ca 0.879 0.276 0.092 0.092 0.11 0.812 0.811 0.697 0.390Na 0.006 0.00 0.00 0.00 0.00 0.00 0.009 0.006 0.000Sum_B 3.022 3.003 3.003 3.004 2.982 3.014 3.003 3.005 2.984Sum_cat 8 8 8 8 8 8 8 8 8O 12 12 12 12 12 12 12 12 12

Alm 54.122 68.402 82.062 81.911 81.444 55.192 58.164 63.585 54.813And 0.00 2.66 0.00 0.00 0.00 0.219 0.255 0.00 0.466Gr 29.962 9.695 3.037 3.091 3.69 27.011 26.788 23.332 15.625Pyrope 1.302 15.767 11.859 11.31 10.831 3.095 3.030 4.667 8.172Spss 14.583 3.476 3.01 3.688 4.034 14.387 11.667 8.385 20.765Uvaro 0.032 0.00 0.031 0.00 0.00 0.096 0.095 0.032 0.160

Representative analyses of biotite and chlorite

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Mineral Biotite Chlorite Sample MA271 GR164 AR221 GR164 SiO2 35.525 35.126 34.91 40.178 23.914 24.948 24.776 TiO2 2.885 2.151 2.849 0.906 0.108 0.139 0.081 Al2O3 17.554 17.61 17.518 25.638 21.899 22.254 22.146 Cr2O3 0.063 0.064 0.017 0.051 0.080 0.016 0.011 FeOtot 23.216 21.877 21.335 14.403 31.30 25.495 27.424 MnO 0.096 0.032 0.044 0.037 1.271 0.455 0.271 MgO 7.468 8.411 7.651 4.961 8.372 14.424 12.694 CaO 0.012 0.011 0.000 0.034 0.059 0.015 0.000 Na2O 0.168 0.153 0.124 0.223 0.000 0.000 0.000 K2O 8.937 9.093 8.845 9.356 0.008 0.013 0.110 Total 95.924 94.517 93.293 95.787 87.011 87.759 87.513

Si 5.465 5.46 5.48 5.785 5.289 5.264 5.295 Al IV 2.535 2.54 2.52 2.215 2.711 2.736 2.705 AlVI 0.645 0.684 0.718 2.132 2.992 2.793 2.869 Ti 0.334 0.252 0.336 0.098 0.018 0.022 0.013 Fe2+ 2.987 2.844 2.801 1.734 5.789 4.499 4.902 Cr 0.008 0.008 0.002 0.006 0.014 0.003 0.002 Mn 0.013 0.004 0.006 0.005 0.238 0.081 0.049 Mg 1.713 1.949 1.790 1.065 2.760 4.537 4.045 Ca 0.002 0.002 0.000 0.005 0.014 0.003 0.000 Na 0.050 0.046 0.038 0.062 0.000 0.000 0.000 K 1.754 1.803 1.771 1.718 0.002 0.004 0.030 Cations 15.506 15.592 15.462 14.825 19.827 19.942 19.910 O 24 24 24 24 36 36 36

Fe/FeMg 0.64 0.59 0.61 0.62 0.68 0.50 0.55 Mg/FeMg 0.36 0.41 0.39 0.38 0.32 0.50 0.45

Representative analyses of white-mica Sample GR164 AR221 SiO2 46.741 47.657 45.55 45.067 47.338TiO2 0.346 0.326 0.67 0.600 0.471Al2O3 32.259 32.053 36.49 36.116 35.287Cr2O3 0.018 0.012 0.01 0.000 0.000FeOtot 2.850 2.737 0.87 0.856 1.118MnO 0.000 0.035 0.06 0.003 0.090MgO 1.382 1.607 0.37 0.527 0.772CaO 0.014 0.000 0.00 0.000 0.000Na2O 0.520 0.481 0.86 0.811 0.717K2O 10.128 10.209 9,96 10.198 10.003Total 94.258 95.117 94.84 94.178 95.795

Si 6.328 6.386 6.068 6.057 6.235Al IV 1.672 1.614 1.932 1.943 1.765AlVI 3.472 3.444 3.792 3.773 3.709Ti 0.035 0.033 0.067 0.061 0.047Fe2+ 0.323 0.307 0.097 0.096 0.123Cr 0.002 0.001 0.001 0.000 0.000Mn 0.000 0.004 0.007 0.000 0.01Mg 0.279 0.321 0.073 0.106 0.152Ca 0.002 0.000 0.000 0.000 0.000Na 0.136 0.125 0.223 0.211 0.183K 1.749 1.745 1.693 1.749 1.681Cations 13.998 13.98 13.953 13.996 13.905O 24 24 24 24 24

Fe/FeMg 0.54 0.49 0.57 0.48 0.45Mg/FeMg 0.46 0.51 0.43 0.52 0.55

Representative analyses of feldspar Sample MA271 AR221 GR164 SiO2 59.955 62.129 62.695 60.348 62.810 68.925 62.200TiO2 0.000 0.000 0.049 0.008 0.012 0.042 0.003Al2O3 25.739 24.257 24.015 25.073 23.679 19.682 23.720Cr2O3 0.000 0.000 0.000 0.000 0.000 0.000 0.000MnO 0.064 0.000 0.044 0.001 0.000 0.000 0.000MgO 0.000 0.031 0.000 0.000 0.000 0.000 0.007CaO 7.444 5.454 5.176 7.124 5.125 0.156 5.110Na2O 7.194 8.563 8.666 7.457 8.733 11.141 7.718K2O 0.085 0.127 0.106 0.080 0.157 0.092 1.161Total 100.48 100.561 100.751 100.09 100.51 99.408 99.919

Si 2.656 2.737 2.756 2.682 2.766 3.00 2.762Al 1.343 1.258 1.243 1.312 1.228 1.009 1.240Ti 0.000 0.00 0.002 0.000 0.000 0.001 0.000Mg 0.000 0.002 0.00 0.000 0.000 0.00 0.000Mn 0.002 0.00 0.002 0.000 0.000 0.00 0.000Ca 0.353 0.257 0.244 0.339 0.242 0.007 0.243Na 0.618 0.731 0.739 0.643 0.746 0.940 0.664K 0.005 0.07 0.006 0.005 0.009 0.005 0.066

Ab 63.3 73.5 74.7 65.10 74.80 98.7 68.20An 36.2 25.8 24.7 34.30 24.30 0.7 25.00Or 0.5 0.7 0.6 0.500 0.900 0.5 6.800Sum_cat 4.982 4.999 4.992 4.984 4.994 4.966 4.980Sum_oxy 8 8 8 8 8 8 8

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APPENDIX B. Quartz c-axis The crystallographic orientation of the quartz c-axis was computed with AVA diagrams (AVA= Achsenverteilungsanalyse; Sander, 1950) using StereoNett 2.0 software (Duyster, 1996). This software applies image-analysis techniques to calculate the azimuth and inclination of quartz c-axes for every position within the field of view by recording the changing birefringence colours while the microscope stage is rotated by 90° (AAVA method - Automated AchsenVerteilungs Analyse). A detailed description of the procedure is given in Appendix A of Stöckhert and Duyster, (1999) and on web site http://www.microtexture.de/StereoHTML/quarzava.htm.

APPENDIX C. Bulk rock analyses of samples used for thermodynamic modelling XRF chemical measurements on rock powder pellets were performed at the Department of Geological Sciences, University of Catania, on a Philips PW 2404 spectrometer equipped with a Rh anticathode; the matrix effect was corrected following Franzini et al. (1975). Calibration was carried out according to numerous international geo-standards; L.O.I. was determined by the gravimetric method and FeO by titration with KMnO4.

XRF-Bulk rock analyses of representative samples and simplified chemical system (MnNCKFMASH) Sample MA271 AR221 GR164 GR166

Lithotypes Biotite Paragneiss

Biotite Paragneiss

Garnet-Muscovite Schist

Garnet free micaschist

SiO2 57.89 55.03 60.73 74.56TiO2 1.57 0.79 1.18 0.65Al2O3 15.28 21.42 17.13 10.5Fe2O3 12.88 7.59 8.22 7.06MnO 0.19 0.18 0.08 0.06MgO 2.36 2.71 2.47 1.85CaO 2.84 3.78 1.98 0.55Na2O 2.74 3.31 3.57 1.03K2O 2.15 3.26 2.96 2.02P2O5 0.63 0.27 0.17 0.10L.O.I. 1.48 1.66 1.5 1.6TOT 100.01 100.00 99.99 100.00

MnNCKFMASH simplified system MnO 0.20 0.19 0.08 0.07Na2O 2.88 3.43 3.70 1.07CaO 2.98 3.91 2.05 0.57K2O 2.27 3.38 3.07 2.08FeOtot 12.26 7.11 7.73 6.55MgO 2.48 2.81 2.56 1.91Al2O3 16.07 22.18 17.77 10.83SiO2 60.86 56.99 63.01 76.91

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APPENDIX D. Mineral solid solution models

Summary tables of the solid solution models with brief commentary

Name and Phase solution model

End members Reference and brief commentary

Biotite - Bio (HP)

Mn-biotite - mnbi

KMn3AlSi3O10(OH)2 K[MgxFeyMn1–x–y]3–wAl 1+2wSi3–wO10(OH)2, x+y≤1 Speciation model, new parameters from THERMOCALC,

extended to cover Fe- and Mn-solution. Powell and Holland,

(1999)

Annite - ann

KFe3AlSi3O10(OH)2

Phlogopite - phl

KMg3AlSi3O10(OH)2

Eastonite - east

KMg2Al 3Si2O10(OH)2

mnts_i 1east+2/3mnbi-2/3phl

sdph_i 1east+2/3ann-2/3phl

Garnet - Gt (HP)

Almandine - alm

Fe3Al 2Si3O12 Fe3xCa3yMg3zMn3(1–x–y–z)Al 2Si3O12, x+y+z≤1 Quaternary garnet model (Holland and Powell, 1998). This model is characterised by a restricted subdivision range on Mn 0%<X<20%.

Grossular - gr

Ca3Al 2Si3O12

Spessartine - spss

Mn3Al 2Si3O12

Pyrope - py Mg3Al 2Si3O12

Chlorite - Chl (HP)

Mn-Chlorite - mnchl

Mn10Al 4Si6O20(OH)16 [MgxFewMn1–x–w]5–y+zAl 2(1+y–z)Si3–y+zO10(OH)8, x+w≤1 The application of this model was considered excluding in every computed pseudosections the afchl endmember because endmember has negligible contribution to the total energy of the solution (see fig 4 of Holland et al., 1998).

Daphnite - daph

Fe10Al 4Si6O20(OH)16

Amesite - ames

Mg8Al 8Si4O20(OH)16

Clinochlore - clin

Mg10Al 4Si6O20(OH)16

Whita Mica - Pheng (HP)

Celadonite - cel

KMgAlSi 4O10(OH)2 KxNa1–xMgyFezAl 3–2(y+z)Si3+y+zO10(OH)2

This model is entirely reported on Holland (2006) web-site: http://www.esc.cam.ac.uk/astaff/holland/ds5/muscovites/mu.html and assumes M2 (multiplicity 2) is split into 1 M2a site on which tri- and di-valent cations mix, and an M2b site occupied solely by Al.

Fe-Celadonite - fcel

KFeAlSi4O10(OH)2

Muscovite - mu

KAl 3Si3O10(OH)2

Feldspar

Anortite - an

CaAl2Si2O8 KyNaxCa1–x–yAl 2–x–ySi2+x+yO8, x+y≤1 Ternary-Feldspar Modeling and Thermometry. High structural state (Fuhrman and Lindsley, 1988).

K-feldspar Kfs

KAlSi 3O8

Albite - ab NaAlSi3O8

Staurolite – St(HP)

Staurolite St

Mg4Fe4Mn4Al 18Si7.5O48H4 Mg4xFe4yMn4(1-x-y)Al 18Si7.5O48H4, x+y≤1 After Holland and Powell (1998)

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FIGURE CAPTIONS

Fig. 1: (a) Distribution of pre-Alpine basement in Europe (after von Raumer et al., 2002). (b)

Distribution of Alpine and pre-Alpine (Hercynian and/or pre-Hercynian) basement rocks

in Calabrian-Peloritani Orogen and main tectonic alignments. Modified after Atzori and

Vezzani (1974), Amodio-Morelli et al. (1976), Schenk (1990), Bonardi et al. (2001),

Ortolano et al. (2005), Fazio et al. (2008).

Fig. 2: (a) Geological sketch-map of Serre Massif and location of study area (modified after

Graeßner et al., 2000). (b) Geological-structural map of study area and sample locations.

(c) Stereoplots (lower hemisphere) with location of structural field stations containing

contours of mylonitic foliation S3 and plunges of stretching lineation.

Fig. 3: Field evidence of D1–D5 deformational phases. (a) Relic S1 axial plane foliation (B1 axis)

embedded in mylonitic foliation S3; (b) S1 foliation folded by crenulation event (D2)

leading to local formation of S2 foliation; (c) Example of pervasive sub-vertical mylonitic

foliation (S3); (d) Asymmetric intrafoliar fold in quartz-feldspar level of mylonitic

paragneiss; (e) Field relationships between S3 foliation and late- to post-tectonic magmatic

dykes with detail of interfingered boundary between host rock and late-tectonic dyke; (f)

Post-tectonic undeformed leucogranite dyke discordantly cutting the mylonitic foliation

S3; (g) Thrust plane produced by brittle deformational stage D5, resulting from evolution

of shallow seated asymmetrical folding of D4 deformational stage.

Fig. 4: Representative photomicrographs of thin sections regarding sequence of the blasto-

deformational stages identified in Mammola Paragneissic Complex rock-types. (a) Early-

M1 assemblage given by tiny zoisite inclusion trails within garnet core (Grt1) in association

with Pl1, Bt1, and Qtz and prograde garnet overgrowth (Grt2) in equilibrium with Pl2 and

Wmca1 (crossed polars); (b) S2 schistosity defined by blastesis of Qtz + Wmca2 ± Bt3

developed during crenulation event D2 (parallel polars); (c) Prograde to peak assemblages

are represented by zoisite-free garnet outer core (Grt2) in equilibrium with Wmca1 and Bt2

and by later overgrowths of Grt3 with Pl2. Early-M3 retrograde stage is documented by

garnet embayments filled by intergrowths of Wmca3 + Pl3 + Bt4 + Chl + Ilm in equilibrium

with garnet rim overgrowth (Grt4) (parallel polars); (d, e) Non-coaxial syn-mylonitic

structures linked to late-M3 retrograde stage, producing Wmca4 + Chl + Ep + Pl5

assemblage. σ-type porphyroclast (d) provides a top-to-ENE–NE sense of shear (crossed

polars); boudinaged garnet porphyroclasts (e) testify to extensional characteristics of

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mylonitic deformational stage (parallel polars); (f) Syn-mylonitic deformational effects

exposed on late-tectonic leucogranite dyke. Microphotograph shows a transitional zone

developed at contact between late-tectonic dyke and mylonitic paragneiss (see Fig. 3f)

(high resolution scan: crossed polars); (g) Effects of thermal metamorphism, revealed by

static Grt5 growth forming limpid rims on previous syn-tectonic garnet, in equilibrium

with porphyroblastic plagioclase (Pl5) and biotite (Bt5) (parallel polars).

Fig. 5: Chemical compositional variations of garnet (inner core, outer core, rim), plagioclase,

white mica, biotite, and chlorite in relation to different stages of reconstructed blasto-

deformational history.

Fig. 6: Syn- to late-mylonitic textural features representative of greenschist facies up to

amphibolite facies conditions develop during mylonitic stage. (a) Thin section image of

mylonitic sample, with location of some representative syn- to late-mylonitic textural

domains (high resolution scan: crossed polars), (b) Evidence of late dynamic growth of

oligoclase rim over former plagioclase porphyroclast (crossed polars, λ plate inserted); (c)

Distribution of Lattice Preferred Orientation (LPO) pattern of two representative quartz-

rich domains plotted on AVA diagrams inferred via StereoNett 2.0 software (Duyster,

1996) by colour coding of reported look-up table. Left: colour-coded images of selected

quartz domains (see Appendix 2 for explanation). Centre: colour coding scheme of look-

up table with orientation of optical indicatrix axes (XYZ) and activation scheme of slip

systems (see also Table 2). Right: quartz c-axis contour plots.

Fig. 7: Representative SEM images, relative X-ray maps, and compositional profiles of garnet

porphyroblasts of sample MA271 (column a) and sample AR221 (column b).

Composition of some representative crystals is also shown.

Fig. 8: P–T pseudosection of sample MA271 in MnNCKFMASH system and P–T constraints.

H2O and quartz calculated as in excess. (a) P–T pseudosection with location of interpreted

P–T constraints; (b) Distribution of calculated compositional isopleths; (c) Potential

intersections of identified compositional assemblages.

Fig. 9: P–T pseudosection of sample AR221 in MnNCKFMASH system and P–T constraints.

H2O and quartz calculated as in excess. (a) P–T pseudosection with location of interpreted

P–T constraints; (b) Distribution of calculated compositional isopleths; (c) Potential

intersections of identified compositional assemblages.

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Fig. 10: Representative SEM images and relative compositional diagrams referring to mylonitic

stage as recorded by sample GR164 (a) and to thermal stage as recorded by sample GR166

(b). Symbols according to legend of Fig. 5.

Fig. 11: P–T pseudosection of sample GR164 in MnNCKFMASH system and P–T constraints.

H2O and quartz calculated as in excess. (a) P–T pseudosection with location of interpreted

P–T constraints; (b) Distribution of calculated compositional isopleths; (c) Potential

intersections of identified compositional assemblages.

Fig. 12: P–T pseudosection of sample GR166 in MnNCKFMASH system and P–T constraints.

H2O and quartz calculated as in excess. (a) P–T pseudosection with location of interpreted

P–T constraints; (b) Distribution of calculated compositional isopleths; (c) Potential

intersections of identified compositional assemblages.

Fig. 13: Integration of estimated P–T constraints and reconstruction of P–T path (thick black

arrows) of Mammola Paragneiss Complex, illustrated by multistage mineral growth

scheme representing discrete stages of Hercynian tectono-metamorphic history.

Fig. 14: Pressure–temperature trajectories reconstructed by various authors for parts of Serre

Massif crustal section: 1) P–T paths related to rocks belonging to uppermost and

lowermost parts of lower crust (after Schenk, 1989); 2) P–T path of uppermost part of

lower crustal segment (after Aquafredda et al., 2006); 3) P–T path reconstruction

considering thermobarometric estimates of this paper for bottom of upper crust; 4) P–T

path of lowermost part of lower crustal segment (after Acquafredda et al., 2008).

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Tables

Table 1. Relationship between deformational and crystallisation stages in various events of metamorphic

evolution

OROGENIC METAMORPHIC CYCLE

THERMAL

METAMORPHISM

Metamorphic

evolution prograde evolution retrograde evolution static events

Deformational

phases early D1 late D1 D2 D3 -

Field evidence -

Relics of isoclinalic folding surfaces S1 within mylonitic foliation

(S3)

Local transposition of S1 surface and formation

of a S2 crenulation cleavage

Pervasive mylonitic foliation S3

Randomly oriented biotite plates and cm-size andalusite and

cordierite spots

Metamorphic

events early-M1 late-M1 M2 early M3 late M3 M4 late M4

Petrographic

features

S1 defined by straight

to sigmoidal inclusion trails of zoisite in

garnet cores

Zoisite-free outer core garnet in

equilibrium with biotite,

plagioclase, white mica

S2 crenulation schistosity

Biotite, chlorite, white

mica, plagioclase

intergrowth in garnet rim

embayments

Mylonitic foliation S3 given by s- and d-type

porphyroclasts wrapped by chlorite, white mica, biotite, feldspar; S/C fabrics; shear bands; oblique

foliation

Foaming texture in

ribbon-like quartz

domains; randomly oriented

porphyroblasts; sub-

euhedral to euhedral

inclusion free garnet rim

Later retrograde static blastesis of

chlorite on previous

cordierite blasts and sericitisation

of andalusite rims

Crystallisation

events syn syn post syn post syn syn

Quartz ------------ ------------------ -------------- ----------- -------------- ---------------------------------

White mica ------------ -- High Phg content -- -------- ----------- -------------- ------------Low Phg content--------

Biotite --------- ------------------ -------------- ----------- -------------- --------------------------

Cordierite -----------------

Staurolite ---------

Chlorite ----------- -------------- ----------------

Plagioclase -Oligoclase-Andesine- ---Albite--- ---Oligoclase--- ----------Oligoclase-Andesine------

Garnet ------------ ------------------ -High Sps content- ----High Alm-Sps content-----

Clinozoisite ------------

Epidote ----------- --------------

Andalusite ------------------

Sillimanite -----

Tourmaline --------------

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Table 2. Relationship between active slip systems inferred by quartz c-axis orientation pattern analysis and approximate shearing temperature at specific metamorphic facies (see Fig.6c)

Metamorphic facies

Approx. T c-axis pattern active slip-system

lower greenschist facies

+ faster strain rates

400°- 450°C type 1

near Z maximum basal <a>

mid-greenschist facies

450°-500°C type 2 rhomb <a>

amphibolite facies 500°-550°C type 3

near Y maximum prism <a>

upper amphibolite facies

650°-700°C type 4

near X maximum prism <c>

low shear strain

-

type 5 c-axes close to s1

unfavorable slip

After Cirrincione et al. (2009) and reference therein

Table 3. Sample location* and brief minero-textural features of study samples

Sample N E Mineralogical assemblage

Texture

MA271 38°23’16’’ 16°13’12’’ Qtz, Grt, Bt, Pl, Chl, Ep,

(Ilm, Zr)

Coarse-grained. Granodiablastic texture, widespread porphyroblast of zoned garnets.

AR221 38°22’14’’ 16°20’13’’ Qtz, Grt, Bt, Wmca, And,

Pl, St, (Ilm, Zr)

Mostly fine-grained. Grano-lepidoblastic matrix interrupted by garnet, andalusite and staurolite porphyroblast.

GR164 38°22’21’’ 16°15’11’’ Qtz, Wmca, Bt, Ab, Grt,

Chl, Zo, (Rt, Ilm, Tur)

Fine-grained, non-coaxial texture, sigmoidal albite porphyroclasts with oligoclase reaction rims, accompanied by rare garnet ones. Mylonitic foliation

GR166 38°22’30’’ 16°15’02’’ Qtz, Crd, Pl, And, Bt,

Wmca, St, Sil, (Zr, Ap)

Fine-grained. Grano-lepidoblastic texture. Quasi-foaming granoblastic levels, alternating to lepidoblastic decussate aggregates

*GPS coordinates: World Geodetic System 84 (WGS84). Major minerals in decreasing order of abundance: bold style, minor minerals (<1 vol.%): italics, heavy and ore minerals: brackets.

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• Observed mineral assemblages have been recalculated: a) Garnet, considering all iron as FeO (i.e. on the basis of Alm-Sps-Grs-Prp

endmembers); b) White mica, considering the Sia.p.f.u., variable from 3.0 (i.e. Muscovite) to 4.0 (i.e. Celadonite), expressed as phengite

content (e.g. Sia.p.f.u. 3.30=Phe30); c) Chlorite and biotite expressed as Fe2/(Fe2+Mg) ratio and d) cordierite expressed as Mg/(Fe2+Mg).

Table 4. Comparison between observed and predicted mineral assemblages and P-T constraints on pseudosections

Sample MA271

Pseudosection: Fig. 8 Observed assemblages*

Constraining phases and P-T estimates

Computed assemblages

Oro

geni

c cy

cle

early

-M1

Grt(Alm54Grs30Sps15 Prp1) +Pl(An34)+Ep(Czo90)+Qtz

Garnet inner core isopleths: (Alm54Grs30Sps15Prp1)

(field A)

590 MPa 500°C

Grt(Alm54Grs30Sps15Prp1)+ Pl(An34)+Zo+Qtz+

Bt(Fe2/(Fe2+Mg)73)+ Chl(Fe2/(Fe2+Mg)70)+ Wmca(Phg14)

late

-M1

Grt(Alm69 Grs25-26Sps1-

2Prp4) +Pl(An20-22)+Qtz

Garnet outer core isopleths: (Alm69Grs25-26Sps1-2Prp4)+

Pl(An20-22) (field B)

900 MPa

530°C

Grt(Alm69Grs25-26Sps1-2Prp4)+ Pl(An20-22)+Qtz+

Bt(Fe2/(Fe2+Mg)64)+ Chl(Fe2/(Fe2+Mg)56)+ Wmca(Phg21)

early

-M3

Grt(Alm74Sps4Grs4Prp22) +Pl(An34-36)+Qtz

Not found in the pseudosection

The

rmal

m

etam

orph

ism

M4 Grt(Alm82Sps3Grs3Prp12)

+Bt(Fe2/(Fe2+Mg)61-64) Not found

in the pseudosection

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Table 4 - continued

Sample AR221

Pseudosection: Fig.9 Observed assemblages*

Constraining phases and P-T estimates

Computed assemblages

Oro

geni

c cy

cle

M1 Grt(Alm55Grs14Sps27

Prp3)+Pl(An39-41)

Garnet inner core isopleths: (Alm55 Grs14Sps27 Prp3)

(field A’)

650 MPa 520°C

Grt(Alm55Grs14Sps27Prp3)+ Pl(An38)+Zo+Qtz+

Bt(Fe2/(Fe2+Mg)57)+ Chl(Fe2/(Fe2+Mg)55)+ Wmca(Phg14)

late

-M1 Grt(Alm64Grs8Sps23

Prp5)+Pl(An35-37) +Wmca(Phg14-20)

Garnet outer core isopleths: (Alm64Grs8Sps23Prp5)

(field B’)

750 MPa 590°C

Grt(Alm64Grs8Sps23 Prp5)+ Pl(An35)+Zo+Qtz+

Bt(Fe2/(Fe2+Mg)45)+ Chl(Fe2/(Fe2+Mg)41)+ Wmca(Phg12)

early

-M3

Grt(Alm52Grs29Sps10

Pyr9)+Wmca(Phg10-7)

Garnet inner core isopleths: (Alm52 Grs29Sps10 Pyr9)

(field C’)

420 MPa 510°C

Grt(Alm52 Grs29Sps10 Pyr9)+ Pl(An38)+Qtz+

Bt(Fe2/(Fe2+Mg)60)+ Chl(Fe2/(Fe2+Mg)56)+ Wmca(Phg8)

The

rmal

met

amor

phis

m

M4

Grt(Alm52 Grs10Sps29

Prp8)+St+And

+Bt(Fe2/(Fe2+Mg)62)

Not found in the pseudosection

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Table 4 - continued

Sample GR164

Pseudosection:

Fig.11 Observed assemblages*

Constraining phases and P-T estimates

Computed assemblages

Oro

geni

c cy

cle Ear

ly-

stag

es

Grt(Alm54 Grs18Sps24 Prp4)+Ab+

Wmca(Phg16-19)

Not found in the pseudosection

Late

- re

trog

rade

Pl(An26)+Wmca(Phg6-9) +Chl(Fe2/(Fe2+Mg)56-

59)

Syn-mylonitic assemblage in porphyroclastic pressure shadow

domains: Pl(An26)+Wmca(Phg6-9) +Chl(Fe2/(Fe2+Mg)56-59)

(field A’’)

300 MPa 490°C

Pl(An26)+Wmca(Phg7) +Chl(Fe2/(Fe2+Mg)59) + Bt(Fe2/(Fe2+Mg)65)

Table 4 - continued

Sample GR166

Pseudosection: Fig.12

Observed assemblages* Constraining phases and P-T

estimates Computed assemblages

Oro

geni

c cy

cle

Ret

rogr

ade

stag

es

St+ Wmca(Phg10-5) +Bt(Fe2/(Fe2+Mg)61-64)

Not found in the pseudosection

The

rmal

met

amor

phis

m

Pea

k Bt(Fe2/(Fe2+Mg)61-

64)+Crd(Mg/(Fe2+Mg)45-

47)+Sil

Static porphyroblastic mineralogical growth:

Bt(Fe2/(Fe2+Mg)61-

64)+Crd(Mg/(Fe2+Mg)45-47) (field A’’’)

300 MPa

685°C

Bt(Fe2/(Fe2+Mg)65)+ Crd(Mg/(Fe2+Mg) 47)+Sil

Pl(An21)

Late

ret

rogr

ade

And+Pl(An24) +Chl(Fe2/(Fe2+Mg)56-59)

Retrograde pseudomorphic assemblage:

Pl(An24) +Chl(Fe2/(Fe2+Mg)56-59)

(field B’’’)

150 MPa 500°C

And+Pl(An24) +Chl(Fe2/(Fe2+Mg)57) +Bt(Fe2/(Fe2+Mg)69)+

Wmca(Phg3)

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