3.a. Metapelite

3.a.1 Migmatite gneiss

Migmatite gneiss covering ∼43 % of the VP at the map-scale (Fig. 2; Table 1) can be divided into four sub-types: (i) biotite + garnet + sillimanite-migmatite gneiss; (ii) biotite + garnet +  sillimanite ± cordierite-migmatite gneiss; (iii) biotite +  garnet +  sillimanite ± orthopyroxene-migmatite gneiss and (iv) fine-grained biotite-migmatite gneiss (Table 1).

Bt+Grt+Sil-migmatite gneiss is the most abundant sub-type, occurring as metric to decametric outcrops, interlayered with black amphibolite, sillimanite-gneiss, felsic granulite and marble. They show a gneissic structure made by a regular layering between millimetric- to pluricentimetric-thick brownish domains of biotite, garnet and sillimanite (i.e., melanosomes) and pluricentimetric-thick white quartz-feldspathic leucosomes (Fig. 4a and b). Melanosomes constitute a continuous to locally spaced foliation, defined by the elongation of biotite and sillimanite crystals making mm-thick layers. Leucosomes constitute plurimillimetric- to pluricentimetric-thick levels, locally appearing as lens-shaped decimetric pods wrapped by the melanosomes or cutting the main foliation. In some outcrops, they wrap decimetric fine-grained metabasic lenses. In detail, leucosomes are made of fine-grained quartz, plagioclase and garnet (blue star in Fig. 4a); locally centimetric K-feldspar occurs (red star in Fig. 4a). The modal abundance of sillimanite in melanosomes varies across the outcrops, with some nearly devoid or dominated by biotite. Conversely, in some locations, sillimanite modal abundance and grain size increase, reaching up to 2 mm in length. At the microscale, the gneissic structure of migmatite is defined by the isorientation of idiomorphic millimetric biotite and fibrous to locally prismatic sillimanite (Fig. 4b) with minor ilmenite. Biotite locally occurring as larger deformed relict grains shows an intense reddish-brown colour and very fine-grained inclusions of rutile. Garnet commonly displays rounded inclusions of biotite, quartz, sillimanite and minor plagioclase, K-feldspar and rutile, concentrated in the core. Garnet is wrapped by the biotite + sillimanite-rich main foliation. Relict deformed biotite and fibrous sillimanite, with quartz films in between, are present in microsites within skeletal garnet porphyroblasts. In some samples, garnet appears as anhedral millimetric blasts forming intergrowths with lobate quartz. The latter commonly shows a granoblastic structure with triple junctions in the less deformed or annealed leucosomes. A ‘chessboard’ structure is visible in the high-strain domains indicating deformation temperatures >700 °C (Stipp et al. Reference Stipp, Stünitz, Heilbron and Schmid2002; Law, Reference Law2014). Quartz also forms irregular sub-millimetric to millimetric films among grain boundaries, such as garnet or biotite. Plagioclase in leucosomes is idiomorphic and sub-millimetric to plurimillimetric, often showing a sharp undeformed twinning. It is in mutual contact with quartz and when present, K-feldspar.

Figure 4. (a) Outcrop of migmatite gneiss with the typical alternation between leucosomes and melanosomes marking the main foliation. The red star indicates the Kfs-bearing leucosome; blue star is the Pl-bearing one. (b) Microscale detail of Bt + Sil foliation in migmatite gneiss (PPL: plane-polarised light). (c) Migmatite gneiss with folded leucosome characterised by the occurrence of dark Crd porphyroblasts (Lac Mort). (d) Crd poprhyroblast in a leucosome of migmatite gneiss substituted by fine-grained Wm along fractures (XPL: crossed-polarised light). (e) Detail of migmatite gneiss with Opx grey blasts and Grt pink blasts (Dzovenno). (f) Microscale view of Opx and Grt porphyroblasts in a migmatite gneiss wrapped by the Bt-rich foliation (XPL). (g) Fine-grained Bt-migmatite gneiss levels cut by thin leucosomes. (h) Millimetric Grt blasts in a fine-grained Bt-migmatite gneiss (PPL). (i, j) Sillimanite-gneiss whose foliation is defined by prismatic sillimanite blasts wrapping around centimetric Grt porphyroblasts (j: XPL). (k) Felsic granulite characterised by widespread presence of plurimillimetric Grt in a white Pl-rich matrix. (l) Microscale detail of Grt porphyroblasts and Pl blasts overgrown by fine-grained white mica in felsic granulite. Rt crystals are also present (XPL).

At the Lac Mort area (N 45°55’8” E 7°29’8” WGS84; Fig. 2), the aspect of migmatite gneiss is slightly different: the structure becomes more complex with folded leucosomes forming a ptygmatic structure. Moreover, the number of metabasite lenses and aplitic or pegmatitic dykes here increases. Sillimanite, which together with biotite highlights the main foliation, is fine-grained (∼1 mm in length) with a fibrous habit.

The second sub-type of migmatite gneiss, Bt+Grt+Sil±Crd-migmatite gneiss, crops in a few spots along the valley (see yellow stars in Fig. 2). Cordierite appears as dark grey porphyroblasts of up to 3 cm within leucosomes or in the melanosomes wrapped by the biotite + sillimanite-rich foliation (Fig. 4c). At the microscale, cordierite is sub-idiomorphic or interstitial, millimetric to plurimillimetric in size, showing widespread fractures along which fine-grained aggregates of white mica and chlorite occur (Fig. 4d). It includes sub-millimetric rounded quartz, biotite and sillimanite. Moreover, cordierite (as well as garnet) is locally overgrown by late-stage biotite and sillimanite.

Bt+Grt+Sil±Opx-migmatite gneiss (Fig. 4e) crops out in one locality in the valley (N 45°51’55”; E 7°24’13”; red star in Fig. 2). These rocks show a gneissic structure defined by the alternance of discontinuous millimetric biotite-rich layers and plurimillimetric to pluricentimetric leucosomes. Plurimillimetric green to grey orthopyroxene porphyroblasts occur mostly in the leucosomes, together with garnet, and are wrapped by the biotite foliation (Fig. 4f). Sillimanite modal abundance is lower with respect to the other migmatite gneiss types. At the microscale, the foliation forms slightly discontinuous layers mainly made of biotite lamellae wrapping leucosomes. Orthopyroxene is characterised by the presence of rounded inclusions of quartz and biotite. Locally, biotite forms millimetric symplectitic intergrowths with quartz, wrapping or growing over orthopyroxene porphyroblasts.

Fine-grained Bt-migmatite gneiss occurs as metric to plurimetric-thick bodies (not visible at the map-scale), defining domains alternating with biotite + sillimanite-rich foliation and leucosomes or folded decimetric levels within Bt+Grt+Sil-migmatite gneiss (Fig. 4g). They are abundant at the Lac Mort and Dzovenno areas (Fig. 2). These rocks consist of sub-millimetric quartz, plagioclase, garnet and biotite, the latter defining a weak gneissic fabric whose orientation is parallel to the foliation of the surrounding migmatite gneiss. Locally, their main foliation is crossed by millimetric- to pluricentimetric leucosomes (Fig. 4g). At the microscale, this lithology shows a gneissic to granoblastic structure made of sub-millimetric quartz and plagioclase showing contacts with triple junctions. Garnet constitutes millimetric subhedral blasts with few rounded quartz inclusions. Biotite is mainly interstitial (Fig. 4h) but locally occurs as subhedral to idiomorphic millimetric lamellae defining a weak foliation.

3.b. Metabasite

3.b.1. Basic granulite

Basic granulite is mainly found two localities in the valley (N 45°55’8” E 7°29’8” and N 45°50’10”; E 7°21’47”) as plurimetric to pluricentimetric boudins (not mappable at the 1:10,000 map-scale; Table 1; Fig. 5a) and ellipsoidal lenses elongated parallel to the foliation pervasive in migmatite gneiss, or as close to isoclinally folded layers. These bodies are locally crosscut by a complex network of leucosomes and aplitic or pegmatitic dykes. In the rest of the unit, basic granulite crops out as smaller lenses (from 5 to 10 cm of thickness) within migmatite gneiss, sillimanite-gneiss and felsic granulite, locally separated by biotite + garnet-rich plurimillimetric to pluricentimetric portions (see zoom within Fig. 5a). The internal structure of basic granulite is granoblastic, locally gneissic, characterised by weak isorientation of fine-grained grey-greenish aggregates of pyroxenes and whitish plagioclase blasts. The orientation of their foliation is often discordant with respect to the pervasive foliation in migmatite gneiss. The contact between basic granulite and metapelite is marked by millimetric selvedges made of brown amphibole and biotite. Here, the modal abundance of amphibole gradually increases at the expense of pyroxene, up to becoming an amphibolite. Locally, a centimetric pale greyish amphibole occurs within the amphibole-rich layers, with no preferential orientation (Fig. 5b).

Figure 5. (a) Basic granulite level within migmatite gneiss at the Lac Mort area with a zoom on the selvedge between the level and migmatite on the bottom right. (b) Centimetric Oamp crystal in basic granulite at Thoules. (c) Detail of a basic granulite at the microscale made of granoblastic Opx, retrogressed Pl, Oamp and minor Qz and Bt (XPL). (d) Brown amphibolite boudins cut by Pl–Bt-pegmatite (red arrow) within migmatite gneiss at Oyace. (e) Microscale detail of brown amphibolite (PPL). (f) Black amphibolite outcrop crosscut by a pegmatite dyke. (g) Garnet megablast within black amphibolite surrounded by Pl + Qz-rich aggregates (red arrow). (h) Microscale view of black amphibolite made of Amp + Cpx-rich and Qz + Pl rich-levels (PPL). (i) Outcrop of an impure marble with pluricentimetric Qz-rich disrupted fold hinges. (j) Ol and Cpx blasts in a Cc-rich matrix in impure marble. The red arrow highlights a melt film (XPL). (k) Melt pocket (red arrow) moving from migmatite gneiss and cutting a basic granulite level. (l) Pegmatitic dyke cutting the migmatitic foliation and the basic granulite boudin at Lac Mort area.

At the microscale orthopyroxene and minor clinopyroxene, orthoamphibole, brown clinoamphibole and minor biotite and quartz are recognised (Fig. 5c): (i) orthopyroxene and clinopyroxene occur as anhedral sub-millimetric blasts; (ii) orthoamphibole is interstitial and colourless with twinning; (iii) plagioclase domains occur as grey to brownish fine-grained aggregates that consist of white mica and epidote (Fig. 5c) and (iv) brown amphibole forms sub-millimetric anhedral crystals growing after pyroxenes. Locally, plurimillimetric garnet porphyroblasts occur in plagioclase-rich domains with a subhedral and locally skeletal habitus including rounded sub-millimetric plagioclase, amphibole and minor quartz showing no preferential orientation.

3.c. Metacarbonates

This group comprises pure and impure marbles, and minor calcsilicates-bearing rocks (Fig. 5i; Table 1). Metacarbonates at the map-scale occur as decametric to hectometric bodies whose elongation is parallel to the NE–SW strike of the other lithologies (Fig. 2). At the outcrop scale, they occur both as interlayered with metapelites and as metric boudins within migmatite gneiss. Marbles are generally grey and locally yellowish, with a massive structure, which becomes locally gneissic due to the slight isorientation of calcite aggregates or quartz. Locally, greenish millimetric crystals of olivine can be recognised. Calcsilicate-bearing rocks occur mainly in the centre-northern part of the unit; they show a brownish-yellowish colour and are characterised by a foliated structure given by the SPO of calcite aggregates and thin quartz levels. In particular, the latter highlight disrupted fold hinges (Fig. 5i) or occur as up to 20 cm-clasts in the calcite-rich matrix.

At the microscale marble is made of sub-millimetric to millimetric calcite blasts, locally showing large, deformed twinning. Sub-millimetric anhedral olivine and clinopyroxene occur (Fig. 5j). Quartz and plagioclase are locally present together with calcite; plagioclase is overgrown by fine aggregates of white mica. In correspondence with some grain boundaries, thin films made of quartz occur (red arrow in Fig. 5j).

3.d. Aplite and pegmatite

Aplite and pegmatite are widespread throughout the unit and are found in all lithologies, mainly striking NNW–SSE and NW–SE. They form dykes up to 1 m-thick, parallel to the pervasive foliation at the map- and outcrop-scale or cutting the main foliation with complex patterns, made by different crosscutting generations of dykes. In particular, three main types of dykes have been distinguished based on their composition and structure: (i) Qz–Kfs-bearing aplites which occur in m-thick bodies in migmatite gneiss (Fig. 5k and l) and metacarbonate discordant to their main foliation; (ii) Pl–Bt-bearing pegmatite, occurring as 10 cm- to 1 m-thick dykes within brown and black amphibolites, or at the contact between these and migmatite gneiss and (iii) Amp-bearing pegmatites which occur in black amphibolite, cutting their main foliation. The first type forms 5 to 40 cm-thick dykes, with an isotropic structure made of plurimillimetric to centimetric sub-idiomorphic quartz, plagioclase and K-feldspar, plus minor biotite and muscovite. The second type shows an isotropic structure made of pluricentrimetric lamellae of biotite and whitish plagioclase crystals. Amp-bearing pegmatites have an isotropic structure made of dark green plurimillimetric to centimetric idiomorphic crystals in a whitish quartz and plagioclase matrix.

4.a. D₁ stage

The D₁ is well recognisable at the mesoscale in basic granulite and brown amphibolite as an S₁ foliation, which occurs within metre-sized boudins, having various orientations to the main foliation and elongation of main lithological boundaries (Fig. 2). The S₁ foliation has scattered orientations, with three main clusters: at Lac Mort, it dips towards the S, while at Oyace (N 45°51’1” E 7°22’54”) toward both W–SW and E (Fig. 2).

In basic granulite, the S₁ foliation is a discontinuous layering marked by the isorientation of pyroxene- and plagioclase-rich domains. At the outcrop scale, the S₁ is geometrically discordant to the long axis of the boudin and to the penetrative foliation that commonly wraps them (Fig. 6a). At the microscale, the S₁ is marked by the weak SPO of Opx I, Bt I (Fig. 7) and minor orthoamphibole. Pl I occurs as granoblastic domains, locally replaced by dark grey fine-grained aggregates.

Figure 6. (a) S₁ foliation in basic granulite boudin wrapped by S₂ foliation in migmatites. (b) S₁ folded foliation in brown amphibolite boudin. (c) Inclusions-rich Grt I core rimmed by the inclusions-poor Grt II. The latter includes rounded Bt I and is wrapped by the S₃ foliation (PPL). (d) Rounded Bt I blasts included in Crd in a migmatite gneiss (XPL). (e) S₂ foliation in migmatite gneiss at the Perquis area. f) S₂ foliation in migmatite gneiss at the microscale made by Bt II and Grt II and leucosome (XPL). (g) S₂ foliation in black amphibolite cut by Grt and Cpx-rich leucosomes. (h) S₂ foliation folded at the Lac Mort area. (i) S₃ foliation defined by Sil III + Bt III and locally Ilm in a sillimanite-gneiss (XPL). (j) Syn-kinematic Grt II porphyroblasts with an internal foliation in the rim concordant with the external S₃ (XPL). (k) Opx II blast in migmatite gneiss wrapped by the S₃ foliation made by Bt III + Qz II intergrowths (XPL). (l) Folded pegmatitic dyke intruding migmatite gneiss at the Lac Mort area.

Figure 7. Blastesis-deformation relationships diagrams for metapelite and metabasite relative to the HT stages.

In brown amphibolite, the S₁ foliation is defined by alternating brown amphibole-rich and plagioclase-rich mm-thick layers, where brown amphibole shows an SPO parallel to the layering (Fig. 6b). Biotite SPO may also contribute to the layering, while clinopyroxene occurs as isolated grains within amphibole-rich layers. At the microscale, Amp I is brown hornblende, and Cpx I occur as interstitial grains in amphibole-rich layers or are associated with Pl I microdomains.

In metapelites, a remnant of this phase is testified by Grt I inclusions-rich cores in bistadial garnets within sillimanite-gneiss. In particular, inclusions of Bt I, Sil I, Qz I and Pl I locally are isoriented marking a foliation discordant to the external one. However, the rounded Bt I, Qz I, Pl I and Sil I inclusions within Grt II and Crd are not always oriented (Fig. 6c and d).

4.b. D₂ stage

During the D_2, the main foliation (S₂) develops, and the main lithological boundaries are parallelised both at map- and mesoscale. The S₂ foliation, the lithological boundaries, and the alignment of basic granulite and brown amphibolite boudins tend to strike NE–SW and dip vertically toward NW or SE (Fig. 2). Sillimanite-gneiss, felsic granulite and black amphibolite bodies interlayered with migmatite gneiss are elongated parallel to the S₂ (Fig. 2). Metre-scale D₂ folds are recognised, folding the S₁ foliation and the lithological boundaries between amphibolite or basic granulite and migmatite gneiss (Fig. 6b). A₂ fold axes are scattered but with two main clusters dipping at a low angle toward the NE and SW. Axial planes (a.p.₂) mainly dip toward the S–SE with relatively high dip angles (Fig. 2).

The S₂ is the main foliation in migmatite gneiss and is represented by the alternation between leucosomes and melanosomes (Fig. 6e). In cordierite-bearing migmatite gneiss melanosomes are defined by idiomorphic Bt II, locally in apparent textural equilibrium with Grt II (Fig. 6f), and minor fibrous Sil II. In garnet- and orthopyroxene-migmatite gneiss the S₂ is made by melanosomes whereas biotite has a relict aspect and is deformed. Conversely, in sillimanite-gneiss, the S₂ is almost completely transposed or preserved within microlithons at the microscale. In felsic granulite, the S₂ foliation is marked by the weak isorientation of garnet-rich layers. At the microscale, the S₂ foliation is marked by the SPO of Bt II and minor Sil II in millimetric to plurimillimetric layers, alternating with the leucosomes (Fig. 6f). Grt II, Crd and Opx II porphyroblasts grow in metapelites (Fig. 7). Grt II occurs as inclusions-poor porphyroblasts within migmatite gneiss and sillimanite-gneiss (Fig. 6f), as the rim of inclusions-rich Grt I cores (Fig. 6c), or as blasts in equilibrium with cordierite or orthopyroxene (Opx II). In felsic granulite, the D₂ is testified by the gentle isorientation of Grt II + Pl II domains marking the S₂.

In black amphibolite, the S₂ is made by alternate dark amphibole- plus clinopyroxene-rich and quartz-feldspathic levels (Fig. 6g). The quartz-feldspatic portions of these amphibolites locally show irregular patterns, crosscutting the penetrative S₂ foliation; in these portions, up to 10 cm garnet or pluricentimetric clinopyroxene porphyroblasts occur. At the microscale, the S₂ is marked by the isorientation of the dark Amp II, Cpx II and Pl II + Qz II + Cc in the quartz-feldspathic plurimillimetric to pluricentimetric lithons. S₂ is also locally marked by Ttn blasts.

In marble, the S₂ is marked by the SPO of calcite and quartz-rich aggregates, with minor clinopyroxene or olivine. In calcsilicates, the S₂ foliation is locally mylonitic, wrapping around up to 30 cm quartz-rich porphyroclasts.

During the D₂ stage partial melting of metapelites and, in a minor amount, of black amphibolite occurs. In metapelites, this is mainly testified by the presence of leucosomes rich in quartz and feldspars marking the S₂ foliation. In black amphibolite, it could be testified by the presence of quartz and plagioclase-rich domains, which in some cases, crosscut the main foliation (Fig. 6g).

4.c. D₃ stage

The D₃ deforms the previous S₂ foliation producing tight to isoclinal folds at the meso- and map-scale, with rounded to acute hinges, locally disrupted, leading to the local transposition of the S₂ and the development of S₃ foliation. At the map and mesoscale, the S₃ is pervasive only in the sillimanite-gneiss and is locally parallel to the axial plane of the folds deforming the S₂. It dips mainly towards E–SE and NW with high dip angles. A₃ fold axes plunge mainly towards N–NE and E (Fig. 2).

At the mesoscale, in migmatite gneiss, the D₃ deformation phase is locally testified by close to isoclinal folds that deform the S₂ foliation (Fig. 6h), locally transposed into a S₃, the latter marked by the SPO of Bt III and Sil III. Leucosomes, likely produced during the D₂, occur also within syn-D₃ fold hinges or neck boudins (Fig. 6e). They are made of perfectly igneous granitic texture, but locally the grains show preferred orientation. In sillimanite-gneiss, the S₂ is completely transposed, and the pervasive S₃ is defined by the SPO of Sil III and minor Bt III (Fig. 6i) wrapping around the centimetric Grt I + Grt II porphyroblasts. However, some garnet porphyroblasts show a rim rich in fine sub-millimetric sillimanite inclusions, which outline a foliation concordant and parallel to the external S₃, suggesting that garnet is also syn-kinematic with respect to the D₃ phase (Figs. 6j and 7). At the Lac Mort area, asymmetric folds deform migmatite gneiss, also affecting cordierite blast aggregates. During this stage sillimanite and biotite locally grow after garnet or cordierite. In some cases, fine-grained intergrowths of Bt III + Qz II over Opx II occur (Figs. 6k and 7).

In black amphibolite, the D₃ deformation phase is only locally detectable as close to isoclinal folds deforming the main S₂ foliation.

Aplites and pegmatites intrude all the rock types locally along their main foliation, but in most cases, they crosscut the D₂ and D₃ structures. However, some of the dykes are also folded, probably those emplaced during the late-D₃ stage (Fig. 6l).

4.d. D₄ stage

This stage is localised in some hm-thick belts (Fig. 2) as an S₄ foliation made by white mica and chlorite or as ‘coronitic’ domains where the blastesis occurs without deformation, better recognisable at the microscale. In metapelites, fine-grained white mica aggregates grow over cordierite along its fractures, or over plagioclase and sillimanite. Chlorite partly pseudomorphs biotite and garnet along fractures. Fine-grained aggregates of epidote and white mica grow after plagioclase both in metabasite and metapelite. Locally, brown amphibole is characterised by a thin rim of green amphibole.

4.e. Inferred stable mineral associations

Based on the relationships between the mineral phases and the observed marking fabrics, it is possible to summarise the inferred mineral associations stable during each deformation phase (Fig. 7).

The D₁ in metapelite is defined by the stable association: Bt I + Sil I + Qz I + Pl I + Kfs I + Grt I ± Rt I.

In basic granulites is defined by Opx I + Cpx I + Qz I + Pl I + Oamp I ± Bt I ± Amp I and by Amp I + Cpx I +  Pl I + Qz I ± Bt I ± Oamp I in brown amphibolite.

The D₂ is defined by the stable associations: Grt II + Crd + Opx II + Pl II ± Rt II + melt (in metapelite) or Bt II + Grt II + Crd + Pl II + melt and Amp II + Pl II + Cpx II + Grt II ± Ttn ± melt (in black amphibolite).

The D₃ is characterised by the stable association: Sil III + Bt III ± Qz II ± Rt III ± Ilm ± Gf.

According to microstructural constraints, during the late D₃, the melt produced during the D₂ crystallises (Qz III + Pl III + Kfs II).

5.a. Metapelite

Metapelites show mineralogical differences confirmed by whole-rock chemical variations. In particular, SiO₂ in migmatite gneiss covers a relatively wide range of values from ∼42 to 75 wt%, while in sillimanite-gneiss this range is smaller, i.e., from ∼46 to 60 wt% (Fig. 8). Felsic granulite shows a SiO₂ content of ∼43 wt%, similar to sillimanite-gneiss (Fig. 8). Alkali content is similar among all analysed metapelites, between ∼3 and ∼6%, except for some modal-derived bulk compositions having relatively low values (<2%) or higher, up to ∼8% (Fig. 8a). Al₂O₃ wt% has highest values in correspondence of sillimanite-gneiss, while decreases in most of the migmatite gneiss samples (Fig. 8b). Pegmatite has a relatively low concentration of Al₂O₃ and a higher one of SiO₂ of ∼72 wt% (Fig. 8b). The iron content (Fe₂O₃ wt%) of migmatite gneiss varies between ∼3 and 20 wt%, depending on garnet modal abundance; in sillimanite-gneiss, it ranges from ∼7 to 14 wt% (Fig. 8c). Pegmatite has the lowest iron content (∼2%). The MgO and CaO (wt%) contents are constant through different samples: they range from ∼1 to 7 wt% (Fig. 8d) and ∼0.1% to ∼5% (Fig. 8e), respectively. A distinction between migmatite and sillimanite-gneiss is evident from the TiO₂ values: it is lower in migmatite gneiss in which it ranges between ∼0.1 and 2%, while is higher in sillimanite-gneiss (∼1.4 to 3 wt%; Fig. 8f).

5.b. Metabasite

Whole-rock analyses on metabasites confirm the differences among the described groups: black amphibolite shows a SiO₂ wt% content ranging between ∼40 and 47%, while brown amphibolite at 45 wt%; basic granulites show higher values, ranging from ∼45 to 55 wt% (Fig. 8). The alkali content varies between ∼1 and 6.8 wt%, with the lowest values in basic granulite samples from Nicot (Reference Nicot1977). Al₂O₃ ranges between ∼4 and 22% in all metabasites (Fig. 8b). Fe₂O₃ wt% content is higher in black amphibolite (∼13 – 20 wt%) and lower in brown amphibolite (∼12%) and basic granulite (∼7 – 15%; Fig. 8c). MgO content is between ∼4.8 and 7 wt% in most of the metabasites, with higher values (up to ∼17%; Fig. 8d) in brown amphibolite and two samples of basic granulite. The difference among metabasite is evident also looking at the CaO content: the highest values are in correspondence with black amphibolite and brown amphibolite (up to 11 wt%) and decrease in basic granulite (Fig. 8e). TiO₂ ranges between ∼0.4 and 6%, with the highest values in black amphibolite (Fig. 8f).

Figure 8. Whole-rock compositions of the main rock types of the VP with shaded areas indicating data from the literature of IVZ (Kunz et al. Reference Kunz, Johnson, White and Redler2014; Redler et al. Reference Redler, White and Johnson2013; Williams et al. Reference Williams, Kelsey and Rubatto2022). (a) SiO₂ vs. alkali content plot. (b) SiO₂ vs. Al₂O₃ wt% in metapelite. (c) SiO₂ vs. FeO + Fe₂O₃ wt% plot. (d) SiO₂ vs. MgO wt% plot. (e) SiO₂ vs. CaO wt% plot. (f) SiO₂ vs. TiO₂ wt% plot.

7.a. Grt–Bt exchange

The Grt–Bt exchange between garnet and biotite has been applied on three metapelite samples, on a total of 20 pairs. For these pairs have been used both analyses from garnet core and included biotite, or garnet rim and neighbouring biotite in the matrix or in garnet strain shadows. The Ferry and Spear (Reference Ferry and Spear1978) calibration has been chosen since garnet is mainly Fe- and Mg-rich. Sillimanite-gneiss shows temperatures ranging from 662 to 736 °C, from three pairs. Bt+Grt+Sil±Crd-migmatite gneiss has temperatures ranging from 556 to 730 °C while those from Bt+Grt+Sil±Opx-migmatite gneiss, taken from four pairs are between 680 and 770 °C (Fig. 11a; Table 2).

Figure 11. Summary of the T °C ranges of each analysed sample for different applied methods. (a) garnet-biotite thermometer (Ferry & Spear, Reference Ferry and Spear1978). (b) Zr-in-rutile thermometer of Tomkins et al. (Reference Tomkins, Powell and Ellis2007). (c) Ti-in-biotite thermometer of Henry et al. (Reference Henry, Guidotti and Thomson2005). (d) Ti-in-amphibole thermometer of Liao et al. (Reference Liao, Wei and Rehman2021).

Table 2. T °C estimated for different samples with the standard deviation reported.

^[1]Zr-in-Rt of Tomkins et al. (Reference Tomkins, Powell and Ellis2007);

^[2]Ti-in-Bt of Henry et al. (Reference Henry, Guidotti and Thomson2005);

^[3]Grt–Bt of Ferry and Spear (Reference Ferry and Spear1978);

^[4]Ti-in-Amp of Liao et al. (Reference Liao, Wei and Rehman2021). N are samples from Nicot, (Reference Nicot1977);

* is the value from Caso (Reference Caso2023).

7.b. Zr-in-rutile

The temperature of rutile crystallisation has been calculated based on its Zr content following the calibration of Tomkins et al. (Reference Tomkins, Powell and Ellis2007). Since the Zr-in-rutile thermometer is pressure-dependent, a value of 0.7 GPa has been used for the calculations. This pressure was estimated for the migmatitic events by Manzotti and Zucali (Reference Manzotti and Zucali2013) and Gardien et al. (Reference Gardien, Reusser and Marquer1994) based on conventional geobarometry. A total of 203 results have been obtained: 112 from sillimanite-gneiss, 43 from Bt+Grt+Sil-migmatite gneiss, and 48 from felsic granulite. Sillimanite-gneiss yielded temperatures ranging from 691 to 818 °C (see Table 2 for average T °C and errors), while Bt+Sil+Grt-migmatite gneiss temperatures vary between 683 and 773 °C (Figs. 11b and 12; Table 2). In felsic granulite, temperatures range between 667 and 843 °C (Fig. 11b).

Figure 12. Geological sketch map of the VP with the position of samples analysed for temperature estimations and zoom with the mean temperatures and errors reported.

7.c. Ti-in-biotite

The Ti-in-Bt thermometer (Henry et al. Reference Henry, Guidotti and Thomson2005) has been applied to biotites from three different microstructural domains (Bt I, Bt II and Bt III) on 18 samples (Fig. 11c). Four samples of sillimanite-gneiss collected at different localities have been analysed (139 biotite analyses used). Except for one sample (belonging to the west of Vessonaz stream) having temperatures between 649 and 727 °C, sillimanite-gneiss have values between 721 and 812 °C (see Table 2 for average T °C for each sample and error). Five samples of Bt+Grt+Sil-migmatite gneiss have been analysed (167 biotite analyses used) giving temperatures between 676 and 784 °C. Five Bt+Grt+Sil±Crd-migmatite gneiss samples (179 biotite analyses used) have temperatures ranging from 654 to 789 °C. Temperatures in Bt+Grt+Sil±Opx-migmatite gneiss (27 biotite analyses used) are quite higher: the two samples have temperatures ranging from 700 to 799 °C and from 760 to 812 °C pointing to higher average T °C (Table 2). Fine-grained Bt-migmatite gneiss (25 analyses used) yielded temperatures between 715 and 764 °C. Biotite from felsic granulite has been analysed only on one sample (from the Perquis area) due to its low modal abundance. Temperatures (13 biotite analyses used) are from 729 to 784 °C (Table 2). Since the chemical composition between different biotite generations is similar, the resulting temperatures do not show significant variations (Fig. 12).

7.d. Ti in-amphibole

The Ti-in-Amp calibration of Liao et al. (Reference Liao, Wei and Rehman2021) has been used on 34 analyses of black amphibolite (plus 18 analyses from Nicot, Reference Nicot1977) and 51 analyses of brown amphibolite (Fig. 11d). Black amphibolite has temperatures ranging from 746 to 880 °C with two main clusters at 767 ± 19 °C and 844 ± 21 °C (Figs. 11d and 12; Table 2). Brown amphibolite has relatively higher temperatures ranging from 858 to 916 °C (Figs. 11d and 12; Table 2). Among the Nicot (Reference Nicot1977) samples, D48 gave temperatures between 791 and 859 °C with a mean of 825 ± 48 °C; D16 gave temperatures between 695 and 893 °C with a mean of 783 ± 74 °C. B78 gives temperatures ranging between 750 and 901 °C with a mean value of 823 ± 68 °C (Figs. 11d and 12; Table 2).

8.a. Grt–Crd geobarometer

The Grt–Crd barometer (Bhattacharya, Reference Bhattacharya1986) has been used on the 33SC sample, corresponding to a Bt+Sil+Grt±Crd-migmatite gneiss. Three pairs of Grt–Crd (using cordierite rim and core, without significant P differences) have been used for the calculation, both with the minimum and maximum T °C estimated through the Ti-in-Bt thermometer on the same sample (i.e., 704 and 769 °C, respectively). Both temperatures used yielded similar values, ranging between 0.55 and 0.58 GPa, with an error of 0.01 GPa.

8.b. Pl–Amp geobarometer

The Pl–Amp geobarometer of Molina et al. (Reference Molina, Moreno, Castro, Rodríguez and Fershtater2015) has been used on a total of 17 pairs. Even for this sample, we used both the minimum and maximum T (i.e., 767 and 879 °C) for the P estimation, which was calculated through the Ti-in-Amp thermometer. Obtained pressure ranges between 0.8 ± 0.08 GPa (at 767 °C) and 1.1 ± 0.10 GPa (at 879 °C).

9.a. Pervasive fabrics associated with melt production and migmatisation at the unit scale

Differently to previous studies, which were merely lithological (i.e., Diehl et al. Reference Diehl, Masson and Stutz1952; the Monte Cervino Sheet n.70 of Geological Map of Italy; Dal Piaz et al. Reference Dal Piaz, Bistacchi, Gianotti, Monopoi, Passeri and Schivo2016) or restricted to more localised areas (Manzotti & Zucali, Reference Manzotti and Zucali2013), this work extends the description of rock types and associated structures at the whole unit scale.

This work detailed the distribution of partially melted metapelites with variable textures and %vol of leucosomes, with decreasing melt abundance: migmatite gneiss, sillimanite-gneiss and felsic granulite. Metabasites have also been thoroughly described and separated in terms of bulk rock chemistry and marking structures: (i) the metric-scale basic granulite and brown amphibolite lenses preserve the oldest recognised structures (i.e., S₁) while (ii) black amphibolites interlayered with metapelites mark the regional foliation (S₂).

Three deformation phases (D₁, D₂ and D₃) related to the HT evolution of the VP have been recognised. The oldest structures and stable assemblages (D₁ stage) preserved in basic granulite and brown amphibolite boudins as an S₁ foliation (green foliation traces in Fig. 13) are locally discordant to the S₂ foliation pervasive in metapelites (green traces in Fig. 13). The evidence of D₁ is detected just in a few localities in the valley: (i) in brown amphibolite metre-scale lenses at Oyace; (ii) in basic granulites and brown amphibolites decametric to metric-scale lenses and boudins in Thoules (as previously also described by Manzotti & Zucali, Reference Manzotti and Zucali2013 as S_pre-2) and (iii) at Lac Mort area within basic granulites and brown amphibolites lenses or folded levels. The S₁, marked by the stable association of Amp I + Opx I + Cpx I + Pl I, is not accompanied by an extensive production of melt, except for rare microscale Pl + Grt-rich domains, which may indicate localised melting.

Figure 13. Simplified foliation traces map and sketches of key outcrops of the VP.

The D₂ tectono-metamorphic stage is related to the development of the S₂ regional foliation pervasive at the map-scale in metapelites, black amphibolites and metacarbonatic rocks (blue traces in Fig. 13). The S₂ is marked in migmatite gneiss by Bt+Sil+Grt (i.e., melanosomes) and quartz-feldspathic domains. These domains, interpreted as leucosomes, are parallel to the S₂ regional foliation at the micro-, meso- and map-scales and testify to widespread melt production at different scales during the D₂ stage. Leucosomes, which mark the D₂-related fabrics, contain garnet, cordierite and orthopyroxene (Fig. 13). These three phases are interpreted as peritectic products related to melt production. Garnet occurs widely in the VP rocks, while cordierite and orthopyroxene are detected only in some localities (see yellow and red stars in Fig. 13). In detail, cordierite occurs with garnet; orthopyroxene is also together with garnet, but an Opx–Crd-bearing assemblage has not been detected.

The S₂ foliation is reactivated during the D₃ stage, leading to its local transposition and development of the S₃ axial plane foliation (Fig. 13). The S₃ is pervasive in sillimanite-gneiss and in correspondence of m-wide localised high-strain domains in the migmatite gneiss; it wraps garnet, cordierite and orthopyroxene. Gardien et al. (Reference Gardien, Reusser and Marquer1994) pointed out the presence of cordierite growing after regional foliation development. Conversely, our observations highlight that cordierite develops during the D₂ stage and is wrapped by the S₃ sillimanite-rich foliation (Caso, Reference Caso2023). Due to the local isoclinal geometry of the D₃ folds, the S₃ (black lines in Fig. 13) and S₂ are parallel, striking NE–SW.

9.b. Litho-structural relationships with melt distribution in the lower crust

Analysing the geometrical relationships between different fabric elements distributed in diverse rock types is fundamental for understanding the generation and mobilisation of melt throughout the crust. Taking the VP as an example, each rock type in the lower crust exhibits diagnostic features and structures revealing information about different stages of the continental crust evolution and partial melting processes (i.e., melt production, subsequent differentiation in melt-rich and restitic portions, melt extraction and melt consumption; Kreigsman & Hensen, Reference Kreigsman and Hensen1998; Waters, Reference Waters2001; White & Powell, Reference White and Powell2002; Kreigsman & Alvarez-Valero, Reference Kriegsman and Alvarez-Valero2010; Dyck et al. Reference Dyck, Waters, St-Onge and Searle2020; Yakymchuk, Reference Yakymchuk and Alderton2020).

As mentioned above, basic granulite (Fig. 14a) and brown amphibolite boudins (Fig. 14b and c), enclosed within metapelites and preserving the earlier mineral assemblages and structures, show little evidence of melt production. It is possible that the small amount of melt, eventually produced during the D₁ or D₂ stages, has been completely extracted thanks to the synchronous deformation, leaving a highly restitic rock made by amphibole, clino- and orthopyroxene and minor plagioclase.

Figure 14. Diagnostic microstructrures related to the migmatisation: (a) Opx I, Bt I and Amp I in basic granulite boudin (XPL). (b) Amp I and Pl I in brown amphibolite (PPL). (c) Grt + Pl-rich domain in brown amphibolite (XPL). (d) Grt II porphyroblast rich in rounded Qz I inclusions (PPL). (e) Relict deformed Bt I and fibrous Sil I reacting to produce melt and Grt II (PPL). (f) Myrmekites between Pl and Qz (XPL). (g) Sil III and Bt III S₃ foliation growing after Grt porphyroblast (PPL). (h) Bt III and Sil III growing over Crd (XPL). (i) Bt III and Qz II intergrowth after Crd (PPL).

The majority of metapelites in the studied area correspond to migmatite gneiss, characterised by a relatively high modal abundance of leucosomes (> 30%). Here, the boundary between leucosomes and fine-grained restitic lenses is frequently marked by biotite and sillimanite-rich selvedges, which may suggest limited melt extraction from these rocks and significant retrogression (Sawyer, Reference Sawyer2008). The rounded quartz and biotite inclusions within garnet (i.e., Bt I and Qz I; Fig. 14d) provide evidence for its growth during melt production. This is possible according to the reaction Bt + Sil + Qz = Grt + Kfs + melt (Fig. 14d; Kriegsman & Alvarez-Valero, Reference Kriegsman and Alvarez-Valero2010; White et al. Reference White, Powell, Holland, Johnson and Green2014). Relict and deformed Bt I and fibrous Sil I near or included in garnet porphyroblasts and separated by melt films (Fig. 14e) highlight the remnants of the biotite-dehydration melting reaction (Le Breton & Thompson, Reference Le Breton and Thompson1988). The same interpretation can be attributed to the few relict biotite and sillimanite grains re-oriented parallel to the S₂. The extensive partial melting during D₂ is also supported by the presence of myrmekites between quartz and plagioclase (Fig. 14f), diagnostic of melt production (e.g., Kreigsman & Alvarez-Valero, Reference Kriegsman and Alvarez-Valero2010). Felsic granulites made by garnet-rich domains, which mark the S₂ foliation, contain biotite or sillimanite in very low modal abundance. The few remnants of melt-related microstructures are recognisable only at the microscale as peritectic inclusions-rich garnet and plagioclase, and thin melt films along grain boundaries. They likely preserve the granulite-facies peak assemblage (i.e., Pl + Grt ± Rt), which may be due to significant melt-loss that prevented pervasive retrogression to lower temperature assemblages (e.g., White & Powell Reference White and Powell2002). According to this hypothesis, felsic granulite could represent melt extraction during the D₂ (Voshage et al. Reference Voshage, Hofmann, Mazzucchelli, Rivalenti, Sinigoi, Raczek and Demarchi1990; Harlov & Wirth, Reference Harlov and Wirth2000; White & Powell Reference White and Powell2002). On the other hand, it also could be possible these rocks derived from a granitic mica-poor protolith which may not have produced a considerable (< 10%) amount of melt (Kreigsman & Alvarez-Valero, Reference Kriegsman and Alvarez-Valero2010; Sawyer, Reference Sawyer2010).

Folded leucosomes within syn-D₃ fold hinges testify migration, during D₃, of the syn-D₂ melt towards the thickened hinges of these folds. Moreover, leucosomes are also localised within neck boudins. The transposition of the S₂ foliation is mainly localised in the sillimanite-gneiss, which is characterised by a pervasive continuous foliation (i.e., the S₃) marked by plurimillimetric prismatic sillimanite and minor biotite (Sil III and Bt III). This strong enrichment in sillimanite (Fig. 14g and h) that forms the pervasive foliation in sillimanite-gneiss could be produced in two ways: (i) destabilisation of white mica according to the reaction Ms + Qz = Sil + Kfs + melt or (ii) during retrogression in the presence of melt by back-crossing the reaction Bt + Sil + Qz + Pl = Grt + Crd + Kfs + melt (Waters, Reference Waters1988; Waters, Reference Waters2001; Kreigsman & Alvarez-Valero, Reference Kriegsman and Alvarez-Valero2010; White et al. Reference White, Powell, Holland, Johnson and Green2014). The S₃ sillimanite-rich foliation wraps peritectic mineral as garnet (Manzotti & Zucali, Reference Manzotti and Zucali2013) and cordierite (Caso, Reference Caso2023), likely represents the back-reaction with melt (Spear, Reference Spear1994; Kriegsman & Alvarez-Valero, Reference Kriegsman and Alvarez-Valero2010) involving garnet and cordierite. The intergrowths between quartz and biotite marking S₃ and wrapping cordierite or orthopyroxene (Fig. 14i) are also commonly interpreted as back-reaction products (Sawyer, Reference Sawyer1999; Waters, Reference Waters2001; Kreigsman & Alvarez-Valero, Reference Kriegsman and Alvarez-Valero2010). Therefore, S₃ may represent a preferential pathway for both melt migration and its interaction with the peritectic minerals.

The quartz-feldspathic layers or lenses, found in black amphibolite, contain clinopyroxene and/or garnet porphyroblasts, which could represent peritectic minerals as the product of partial melting through fluid-absent incongruent hornblende breakdown. Melt production in metabasic rocks needs sufficiently high temperatures to take place (> 800 °C; Rushmer, Reference Rushmer1991; Wyllie & Wolf, Reference Wyllie and Wolf1993; Kunz et al. Reference Kunz, Johnson, White and Redler2014) similar to those obtained through the Ti-in-amphibole thermometry. Since leucosomes of black amphibolite are parallel to S₂ foliation, it can be inferred that melting has occurred during the D₂ stage in amphibolites as well.

9.c. Thermal state during melt-production and migmatisation

Temperatures derived from different geothermometers differ depending on the methods used and vary across different localities. However, systematic variations exist among the different lithologies. Particularly evident is the temperature difference of ∼100 °C between metapelites (peak temperature ca. 760 °C) and amphibolites (peak temperature ca. 860 °C, locally up to 900 °C; Figs. 11 and 12). This marked difference may occur since biotite was no longer stable at the maximum reached temperature and therefore was not able to register peak temperatures like amphibole did. In fact, the biotite-dehydration melting (here testified by relicts Bt I) occurs at lower temperatures (Le Breton & Thompson, Reference Le Breton and Thompson1988) than amphibole-dehydration melting. Temperatures derived from biotite marking the S₃ foliation growing together with sillimanite after garnet, cordierite or orthopyroxene that reacted with melt during cooling after peak (Bt III; Stüwe, Reference Stüwe1997; Kreigsman & Hensen, Reference Kreigsman and Hensen1998; Brown, Reference Brown2002; Kreigsman & Alvarez Valero, Reference Kriegsman and Alvarez-Valero2010), are equal to those from Bt I and Bt II. Most of the temperatures derived from the garnet–biotite Fe–Mg exchange thermometer (Ferry & Spear, Reference Ferry and Spear1978) provide lower values (with a minimum of ∼560 °C). These values are related to late diffusional resetting after peak metamorphic conditions, and thus after D₂/D₃ stages (Pattison et al. Reference Pattison, Chacko, Farquhar and Mcfarlane2003 and references therein).

A temperature variation is visible among the different migmatite gneisses (i.e., Grt-, Crd- and Opx-bearing), with the highest temperatures calculated for Opx-bearing gneiss (up to ∼810 °C) and the lowest ones in Crd-bearing migmatites (750 °C; Fig. 15). However, these differences are relatively small, with values close to the error of each method, i.e., ±50 °C at > 700 °C (see Table 2 standard deviation and Henry et al. Reference Henry, Guidotti and Thomson2005).

Figure 15. P–T conditions of VP. Pl–Amp barometer calibration of Molina et al. (Reference Molina, Moreno, Castro, Rodríguez and Fershtater2015); Grt–Crd barometer of Bhattacharya (Reference Bhattacharya1986).

Partial Alpine retrogressions could also explain temperature variations in metapelites, where such retrogressions are localised along hectometre-wide Alpine-age deformation horizons (Fig. 2) of widespread greenschist-facies overprinting (Manzotti et al. Reference Manzotti, Ballèvre, Zucali, Robyr and Engi2014; Dal Piaz et al. Reference Dal Piaz, Bistacchi, Gianotti, Monopoi, Passeri and Schivo2016). These deformation zones could separate different domains within the VP that underwent different P–T conditions, explaining the temperature variation between Crd- and Oxp-bearing migmatites. According to the present map of rock distribution and Alpine horizons (Figs. 2 and 13), the reviewed temperature differences do not seem to be controlled by the Alpine horizons. Indeed, similar differences are recorded at the Lac Mort area, north of one km-scale Alpine horizons and at the Oyace area (Fig. 13).

9.d. Tectono-metamorphic evolution of the Valpelline series

The VP preserves a complex polyphasic evolution spanning from the pre-Alpine to Alpine age, under amphibolite-/granulite- to greenschist-facies conditions. The oldest remnants of the earlier tectono-metamorphic stage (D₁) have been detected in the basic granulite and amphibolite boudins within metapelites. An earlier subsolidus stage before partial melting occurring during the D₁ or early D₂ is preserved in garnet cores as relict biotite, sillimanite, plagioclase and rutile, included in porphyroblasts. Manzotti & Zucali (Reference Manzotti and Zucali2013) provided thermobarometric estimates of this event at T = 700 ± 50 °C and P = 0.57 ± 0.1 GPa (MZ1 in Fig. 15). Different values have been obtained by Gardien et al. (Reference Gardien, Reusser and Marquer1994), who fixed the P–T condition of their oldest stage at higher pressure (G1 in Fig. 15) from kyanite relicts within garnet cores (P = 0.9–0.1 GPa).

The D₂ stage during which partial melting occurs, is accompanied by garnet, cordierite or orthopyroxene growth in metapelites yielding different temperatures (considering the Ti-in-biotite thermometer). The temperature range of 650–780 °C obtained from cordierite-migmatites intersects the biotite-dehydration melting reaction (Le Breton & Thompson, Reference Le Breton and Thompson1988; Spear, Reference Spear1994; Waters, Reference Waters2001; White et al. Reference White, Powell, Holland, Johnson and Green2014). Pressures obtained using the Grt–Crd barometer (Bhattacharya, Reference Bhattacharya1986) of 0.58 ± 0.01 GPa and fall within the P range defined by the peak assemblage of a similar cordierite-metapelite from the IVZ (Redler et al. Reference Redler, Johnson, White and Kunz2012) using thermodynamic modelling (orange field in Fig. 15). Temperatures obtained from orthopyroxene-bearing migmatites using the Ti-in-biotite thermometer, ranging between 700 and 830 °C, also fit with the peak assemblage and P–T conditions detected in similar samples in the IVZ by Redler et al. (Reference Redler, Johnson, White and Kunz2012) (blue field in Fig. 15) but at higher pressure (P = 0.8 GPa) respect to the cordierite-bearing ones. Temperatures yielded by black amphibolite using the Ti-in-Amp thermometer of Liao et al. (Reference Liao, Wei and Rehman2021) range between 800 and 930 °C. These values, ∼100 °C higher than metapelites agree with the destabilisation of hornblende at ∼T > 900 °C (dashed black line in Fig. 15), meaning that amphibolite recorded higher temperature conditions. Pressures obtained from black amphibolite for the D₂ stage applying the Amp–Pl geobarometer (Molina et al. Reference Molina, Moreno, Castro, Rodríguez and Fershtater2015) range between ∼0.7 and 1.2 GPa (Fig. 15).

The thermal conditions during the D₂ and D₃ stages in metapelites are similar since the Ti-in-Bt thermometry applied on biotite marking both S₂ and S₃ foliations yielded similar T °C values. In fact, these values should be considered only as minimum temperatures (up to 800 °C) related to the D₂ and D₃ phases, i.e., slightly before biotite destabilisation due to incongruent melting (D₂) or related to melt consumption and back reaction with peritectic minerals (D₃) producing biotite. The T °C range of black amphibolites coincides with one of the D₂ and D₃ stages of Manzotti & Zucali (Reference Manzotti and Zucali2013; MZ2 and MZ3 in Fig. 15) and the P–T peak conditions of similar cordierite- and orthopyroxene-bearing metapelites from the IVZ from Redler et al. (Reference Redler, Johnson, White and Kunz2012). Taking this into account is possible to constrain the P–T conditions for melt production and consumption. The T °C range is comprised between ∼800 and 900 °C and the P between ∼0.5 and 0.8 GPa. The field of orthopyroxene-bearing migmatites from the Val Strona (IVZ), determined by Redler et al. (Reference Redler, Johnson, White and Kunz2012, Reference Redler, White and Johnson2013), can be used as upper pressure limit.

The inferred metamorphic conditions and P–T paths for the VP lower continental crust are compatible with an evolution typical to many high-grade terranes related to extensional geodynamic settings: prograde to peak metamorphic conditions reaching and overstepping the suprasolidus conditions, followed by nearly isobaric heating (Jiao et al. Reference Jiao, Brown, Mitchell, Chowdhury, Clark, Chen, Chen, Korhonene, Huang and Guo2023 and references therein). This could explain the close P–T conditions and similar mineral assemblage stability during the D₂ and D₃ events. However, the detection of kyanite in garnets by Gardien et al. (Reference Gardien, Reusser and Marquer1994) shifting the early path to higher pressure conditions followed by decompression heating may highlight a more complex evolution probably linked to Variscan orogenic crustal thickening during the Carboniferous (Carosi et al. Reference Carosi, Petroccia, Iaccarino, Simonetti, Langone and Montomoli2020; Schulmann et al. Reference Schulmann, Edel, Catalán, Mazur, Guy, Lardeaux, Ayarza and Palomeras2022; Roda et al. Reference Roda, Spalla, Filippi, Lardeaux, Rebay, Regorda, Zanoni, Zucali and Gosso2023) and subsequent orogenic collapse and lithospheric extension during late Carboniferous-Permian (Lardeaux & Spalla, Reference Lardeaux and Spalla1991; Marotta & Spalla, Reference Marotta and Spalla2007; Schuster & Stuwe, Reference Schuster and Stüwe2008; Roda et al. Reference Roda, Regorda, Spalla and Marotta2018; Ewing et al. Reference Ewing, Rubatto, Lemke and Hermann2023).

9.e Geodynamic implications for the Adria lower continental crust: comparison with the Ivrea-Verbano Zone lower crustal section

The IVZ (Southern Alpine Domain) and the IIDK (Austroalpine Domain), both belonging to the Adria microplate, have several similarities with the VP in terms of rock types, P–T conditions, and ages of metamorphism.

Metapelites, felsic granulite (usually named ‘stronalites’; Redler et al. Reference Redler, Johnson, White and Kunz2012) and black amphibolite with localised partial melting (Kunz et al. Reference Kunz, Johnson, White and Redler2014) occur both in the VP and IVZ. The bulk rock compositions of both VP and IVZ metapelites and metabasite are also similar (Fig. 8). However, an important difference with the IVZ lies in the distribution of the three migmatite types (i.e., Grt-, Crd- and Opx-bearing) and in the map-scale continuity of the isograds, with higher-grade conditions approaching the Insubric Line (i.e., from the Ms- to Opx isograds). In fact, the VP rock types crop out without a clear crustal zonation unlike the IVZ, where the different parageneses reflect a clear metamorphic zonation. Moreover, prograde muscovite occurring in IVZ migmatites is lacking in the VP ones. The only remnant of prograde dehydration melting reactions is the relict biotite included within peritectic phases or deformed biotite grains scattered within the rock matrix.

The thermal state described in the IVZ with similar methods used in the present work (i.e., the Zr-in-Rt; Ewing et al. Reference Ewing, Hermann and Rubatto2013) is comparable with the one described in the VP. However, in the IVZ, higher temperatures reaching 1000 °C have been locally detected close to the contact with the Mafic Complex (Barbooza & Bergantz, Reference Barboza and Bergantz2000; Luvizotto & Zack, Reference Luvizotto and Zack2009). Conversely, the VP never reached the same UHT conditions, suggesting that no mafic bodies were in place close by and large enough to produce such temperatures in the VP rocks. For this reason, the VP represents a good case study to investigate the regional Permian metamorphism without the effects of contact metamorphism.

The few available ages for the VP metapelites obtained through U-Pb geochronology on zircon and monazite yielded late Carboniferous and Permian ages (Zucali et al. Reference Zucali, Manzotti, Diella, Pesenti, Risplendente, Darling and Engi2011; Pesenti et al. Reference Pesenti, Zucali, Manzotti, Diella and Risplendente2012; Kunz et al. Reference Kunz, Manzotti, von Niederhäusern, Engi, Darling, Giuntoli and Lanari2018), similar to those for the IVZ (Barbooza et al. Reference Barboza, Bergantz and Brown1999; Peressini et al. Reference Peressini, Quick, Sinigoi, Hofmann and Fanning2007; Ewing et al., Reference Ewing, Hermann and Rubatto2013; Williams et al., Reference Williams, Kelsey and Rubatto2022). In particular, Kunz et al. (Reference Kunz, Manzotti, von Niederhäusern, Engi, Darling, Giuntoli and Lanari2018) provided two groups of ages by dating zircons from VP migmatites of ∼290–280 Ma and ∼270 Ma. Permian ages are also found in the Arolla Series granitoids and orthogneiss (∼290 Ma; Bussy et al. Reference Bussy, Venturini, Hunziker and Martinotti1998; Manzotti et al. Reference Manzotti, Rubatto, Zucali, El Kohr, Cenki-Tok, Ballèvre and Engi2018). This geochronological link between the VP HT metamorphism and magmatism in the neighbouring Arolla Series could explain partial melting occurring in the VP lower crust and the emplacement of granitic intrusions in the Arolla Series upper crust (Kunz et al. Reference Kunz, Manzotti, von Niederhäusern, Engi, Darling, Giuntoli and Lanari2018).