# Evidence of subduction-related components in sapphirine-bearing gabbroic dykes (Finero phlogopite-peridotite): Insights into the source of the Triassic-Jurassic magmatism at the Europe-Africa boundary

Tommaso Giovanardi, Alberto Zanetti, Luigi Dallai, Tomoaki Morishita, Christophe Hémond, Maurizio Mazzucchelli

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Tommaso Giovanardi, Alberto Zanetti, Luigi Dallai, Tomoaki Morishita, Christophe Hémond, et al. Evidence of subduction-related components in sapphirine-bearing gabbroic dykes (Finero phlogopite-peridotite): Insights into the source of the Triassic-Jurassic magmatism at the Europe-Africa boundary. Lithos, 2020, 356, 10.1016/j.lithos.2020.105366. insu-03684661

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This is the peer reviewd version of the followng article:

Evidence of subduction-related components in sapphirine-bearing gabbroic dykes (Finero
phlogopite-peridotite): Insights into the source of the Triassic-Jurassic magmatism at the Europe-Africa
boundary / Giovanardi, Tommaso; Zanetti, Alberto; Dallai, Luigi; Morishita, Tomoaki; Hémond, Christophe;
Mazzucchelli, Maurizio. - In: LITHOS. - ISSN 0024-4937. - 354-355:(2020), pp. 1-16.
[10.1016/j.lithos.2020.105366]

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19/09/2025 16:25
1 Evidence of subduction-related components in sapphirine-bearing gabbroic dykes (Finero

2 Phlogopite–Peridotite): Insights into the source of the Triassic–Jurassic magmatism at the

3 Europe–Africa boundary

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5 Tommaso Giovanardi¹,²#, Alberto Zanetti³#, Luigi Dallai⁴, Tomoaki Morishita⁵, Christophe

6 Hémond⁶, Maurizio Mazzucchelli²*

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8 ¹DIPARTIMENTO DI SCIENZE DELLA TERRA E DELL'AMBIENTE, UNIVERSITA' DEGLI

9 STUDI DI PAVIA, VIA FERRATA 1, 27100 PAVIA, ITALY

10 ²DIPARTIMENTO DI SCIENZE CHIMICHE E GEOLOGICHE, UNIVERSITA' DEGLI STUDI

11 DI MODENA E REGGIO EMILIA, VIA CAMPI 103, 41121 MODENA, ITALY

12 ³ISTITUTO DI GEOSCIENZE E GEORISORSE-C.N.R. PAVIA, VIA FERRATA 1, 27100

13 PAVIA, ITALY

14 ⁴ISTITUTO DI GEOSCIENZE E GEORISORSE-C.N.R. PISA, VIA MORUZZI 1, 56124 PISA,

15 ITALY

16 ⁵FRONTIER SCIENCE ORGANIZATION, KANAZAWA UNIVERSITY, 920-1192

17 KANAZAWA, JAPAN

18 ⁶LABORATOIRE GÉOSCIENCES OCÉAN, UMR6538, INSTITUT UNIVERSITAIRE

19 EUROPÉEN DE LA MER, UNIVERSITÉ DE BREST & CNRS, PLACE NICOLAS COPERNIC,

20 F-29280, PLOUZANÉ, FRANCE.

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22 # these authors are equal first authors of this work

23 * Corresponding author: maurizio.mazzucchelli@unimore.it

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25 Keywords: sapphirine; Finero; mantle peridotite; dykes; Ivrea Verbano

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**Abstract**

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A gabbroic dyke swarm containing magmatic sapphirine occurs in the Finero phlogopite-peridotite (FPP), one of the major mantle massifs in the Ivrea-Verbano Zone (IVZ; western Southern Alps).

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Sapphirine is part of a particular mineral assemblage, including plagioclase, titanian pargasite, titanian phlogopite, and Cl-rich apatite; the latter mineral hosts calcite inclusions. The dykes cut the mantle foliation at a high angle, are bounded by orthopyroxenite layers, and show symmetric internal banding, represented by two outer hornblendite selvages and an inner leucogabbro band.

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The sapphirine occurs in up to 3 cm-thick irregular patches in both hornblendite selvages, along with Al-rich amphibole and green spinel.

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We present major and trace elements of minerals and bulk rock, as well as mineral O, Sr, and Nd isotopic compositions of dykes and the host peridotite from two different outcrops in the FPP area.

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Our data show that early melt migration developed through porous flow within cm-thick channels and was characterised by orthopyroxene dissolution. Following progressive percolation and reaction, the melt became silica saturated with segregation of orthopyroxenite in the centres of the channels. The banded internal structure of the dykes was caused by three different evolutionary stages, involving opening and enlargement of the conduits. The sapphirine and green spinel segregation took place at T > 1,000 °C, in the presence of melt with transient composition, which interstitially migrated and reacted with the cumulus minerals to form the hornblendite layers. The mineral chemistry of the newly-formed amphiboles indicates that the sapphirine parental melt was Al-rich, depleted to strongly depleted in Hf, Zr, Nb, Ta, Ti, Sc, V, and middle and heavy rare earth elements, and characterised by a positive Eu anomaly and (Zr/Hf)N < 1. These data suggest a parental melt with a significant amount of normative plagioclase. However, the studied veins do not show evidence of plagioclase assimilation, and we argue that this process could have occurred in magmatic bodies that are not outcropping today to the surface or in the melt source.

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The δ¹⁸O values of vein amphiboles and plagioclases vary from 6.9 to 8.6‰ SMOW, which is well above the mantle range, even when considering fractionation upon cooling. Given that

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orthopyroxene from the wall has “normal” mantle δ¹⁸O values (5.8‰), reaction with the host metasomatised peridotite cannot be responsible for the heavy δ¹⁸O signature, and the latter must have been imparted by crustal components deeper in the mantle.

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Our petrographic and geochemical evidence demonstrates that the northern IVZ records an extremely prolonged release, from the Variscan orogenic cycle to the Mesozoic exhumation, of K-H₂O-rich mantle-derived melts, mixed with subduction-related components. This finding provides valuable insights into the Triassic–Jurassic magmatism and the geodynamic environment at the Europe–Africa boundary.

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# 1. Introduction

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It is well known that the subcontinental lithospheric mantle (SCLM) may record significant heterogeneities in terms of both lithology and geochemical composition, resulting from the development of different tectono-magmatic events over a long time span (Mukasa and Shervais, 1999; Rivalenti et al., 2007a, b; Mazzucchelli et al., 2009; 2010; 2016; Borghini et al., 2017; Princivalle et al., 2014; Ponce et al., 2015; Rocco et al., 2017; Consuma et al., 2019). The petrochemical records of the lithospheric mantle and uprising melts can be very different based on the geochemical affinity of the melt, the composition of the peridotite mantle column, the modalities of melt migration, and the P, T, fO₂, and fH₂ conditions of the system. An exceptional locality where this can be studied in great detail is the phlogopite-peridotite mantle unit of Finero (Cawthorn, 1975; Siena & Coltorti, 1989). The Finero phlogopite-peridotite (FPP) is one of the most studied mantle massifs on Earth. It crops out in the northernmost part of the Ivrea-Verbano Zone (IVZ, western Southern Alps, Italy; Fig. 1) and consists of an association of phlogopite-bearing amphibole harzburgites and dunites, both locally associated with phlogopite-bearing amphibole pyroxenites. This is the result of an important episode of pervasive to channelled porous flow migration of melts containing a large volume of crustal components that induced virtually complete metasomatic recrystallisation of an older sequence of harzburgite and pyroxenite
associations (Lensch, 1968; Cawthorn, 1975; Griffin & O'Reilly, 1986; Siena & Coltorti, 1989; Shervais and Musaka, 1991; Hartmann & Wedepohl, 1993; Zanetti et al., 1999, 2016; Grieco et al., 2001, 2004; Morishita et al., 2003, 2008; Raffone et al., 2006; Selverstone & Sharp, 2011; Mazzucchelli et al., 2014; Tommasi et al., 2017). The crustal affinity of these metasomatic melts is implied by high δ¹⁸O values and by the Nd, Sr, Pb, H, S, Cl, and noble gas isotopic compositions of the FPP rocks (Hunziker and Zingg, 1982; Voshage et al., 1987, 1988; Cumming et al., 1987; Hartmann and Wedepohl, 1993; Obermiller, 1994; Seitz and Woodland, 2000; Downes, 2001; Matsumoto et al., 2005; Selverstone and Sharp, 2011). The melt migration processes have been mainly attributed to supra-subduction environments (e.g. Zanetti et al., 1999; Grieco et al., 2001, 2004; Morishita et al., 2003, 2008; Matsumoto et al., 2005), but also to extensional settings (Garuti et al., 2001; Zaccarini et al., 2004).

The available geochronological data indicate that this main event of melt migration took place in the Palaeozoic (Zanetti et al., 2016; Malitch et al., 2017), but many field, petrochemical, and geochronological data point to the development of several other tectono-magmatic events in the Mesozoic as well (Stähle et al., 1990, 2001; Grieco et al., 2001; Matsumoto et al., 2005; Morishita et al., 2003, 2008; Zanetti et al., 1999; 2016; Malitch et al., 2017).

Giovanardi et al. (2013) studied the occurrence of a late sapphirine-bearing gabbroic dykelet swarm previously reported by Siena & Coltorti (1989) and discovered by M. Mazzucchelli (Giovanardi et al., 2013). Based on a detailed petrographic and geochemical survey, they provided evidence for a multi-stage dyke formation, which involved fractional crystallisation associated with different patterns of melt-rock interactions.

With the aim of better constraining the nature of the primitive melt and petrological processes and inducing the segregation of sapphirine in the dykes, in this paper, new and detailed petrochemical data are reported for dykes and host peridotites close to the area studied by Giovanardi et al. (2013) and on a second finding in an adjacent area. Particular care has been dedicated to characterising the compositional variability in terms of major and trace element mineral chemistry, as well as the
mineral O, Nd, and Sr isotopic composition, to document the geochemical fractionation of the flowing melt following reaction with the host peridotite and early cumulates. The genetic and temporal relationships between the dyke emplacement and other events of melt migration recorded by the FPP unit, as well as the related geodynamic scenarios, are also addressed.

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## 2. Geological setting

In the Finero area (Fig. 1), the IVZ crops out in a pseudo-antiform structure with the FPP mantle unit at the core, flanked by what has been called the Finero Mafic Complex (Cawthorn, 1975; Siena & Coltorti, 1989; Lu et al., 1997a, b; Zanetti et al., 2013, 2014; Giovanardi et al., 2014; Mazzucchelli et al., 2014; Langone et al., 2017). The Mafic Complex is divided into three different units: the stratigraphically lower Layered Internal Zone unit (LIZ), in contact with the FPP; the Amphibole-Peridotite (Amphibole-Pd); and the External Gabbro (EG). The EG is placed in tectonic contact with the Amphibole-Peridotite by a Mesozoic high-T shear-zone (Langone et al., 2018). A tectonic contact also characterises the transition toward the amphibolite-facies metasediments and metabasites of the Kinzigite Formation (KF), which represent the metamorphic basement of the Adria Plate. Septa of KF rocks are embedded in the EG.

The Finero mafic-ultramafic sequence presents several differences with respect to the southern and central sectors of the IVZ (i.e. the Baldissero and Balmuccia peridotites and the Val Sesia Complex; Quick et al., 1995; Correia et al., 2012; Mazzucchelli et al., 2014 and references therein), providing evidence of a different geological evolution (Zanetti et al., 2013, 2014, 2016; Langone et al., 2017, 2018). In particular, the FPP has been completely recrystallised by several events of migrating melts with crustal affinity (Zanetti et al., 1999; 2016; Mazzucchelli et al., 2014 and references therein), and the relationships with the parental melts of the surrounding Mafic Complex are far from being understood (Giovanardi et al., 2014). Conversely, the peridotites cropping out in the central and southern part of IVZ did not undergo similar melt-induced recrystallisation and were emplaced in

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the Kinzigite Formation as tectonic slivers before the intrusion of the central and southern Mafic Complex (Quick et al., 1995; Mazzucchelli et al., 2014 and references therein).

The FPP is mainly represented by phlogopite-bearing amphibole harzburgites and associated phlogopite-bearing pyroxenites (Cawthorn, 1975; Siena & Coltorti, 1989; Hartmann & Wedephol, 1993; Zanetti et al., 1999; Raffone et al., 2006; Selverstone & Sharp, 2011; Mazzucchelli et al., 2014; Giovanardi et al., 2018). These lithologies derived from pervasive metasomatism of a depleted peridotite, which formed secondary orthopyroxene, amphibole, and phlogopite (Zanetti et al., 1999; Tommasi et al., 2017). Channelled migration events formed dunite bodies containing stratiform to podiform chromitites and, rarely, pyroxenite and hornblendite layers (Grieco et al., 2001, 2004; Zanetti et al., 2016).

Late stages of porous-flow melt migrations crystallised apatite- and dolomite-bearing wehrlites and apatite-rich, orthopyroxene-bearing peridotites, which sometimes contain carbonate-bearing domains showing marked modal and geochemical gradients with the host rocks (Zanetti et al., 1999; Morishita et al., 2003, 2008; Matsumoto et al., 2005; Raffone et al., 2006). U-Pb analyses on apatite and isotopic noble gas data have yielded Triassic ages (Morishita et al., 2008).

Apatite and calcite also occur in dykes of nepheline-bearing syenites of the Triassic age, associated with hornblendites (Stähle et al., 1990, 2001). An Early Jurassic U-Pb zircon age was determined for an alkali pegmatite (Grieco et al., 2001).

The rocks of the FPP related to the main event and the dunite bodies show similar geochemical features and absence of geochemical gradients. The harzburgite–pyroxenite association and the chromitite and pyroxenite layers in dunite bodies are depleted in Nb and high field strength elements and significantly enriched in K, Rb, Ba, Sr, and light rare earth elements (LREEs) (Hartmann & Wedephol, 1993; Zanetti et al., 1999, 2016; Mazzucchelli et al., 2014). The mineralogical and compositional features have been considered by several authors as evidence of the role of a slab-derived crustal component in the percolating melts (Hartmann & Wedephol, 1993; Mazzucchelli et al., 1995, 2016; Rivalenti et al., 1995, 2007a; Zanetti et al., 1999, 2016; Grieco et
al., 2001, 2004; Morishita et al., 2003, 2008; Ponce et al., 2015; and others). Thus far, a variety of different geochemical components has been identified based on isotopic data. In particular, the Hf (in zircon) and O (in zircon and pyroxene) isotopic compositions of chromitites suggest the presence of large volumes of continental crust in the migrating melts (Zanetti et al., 2016; Malitch et al., 2017), whereas the hydrogen, oxygen, and chlorine isotopic compositions of amphibole and phlogopite from the harzburgite-pyroxenite association show variability (δD from -29 to -86‰, δ¹⁸O from 4.9 to 6.1‰, and δ³⁷Cl from -2.0 to +2.1‰) consistent with mixtures of magmatic fluids contaminated by seawater (Hartmann & Wedepohl, 1993; Selverstone & Sharp, 2011).

Giovanardi et al. (2013) reported the occurrence in the FPP of late sapphirine-bearing gabbroic dykes. These dykes crosscut at high angles the pervasive mantle foliation and other lithologies, showing different mineralogical and major element mineral chemistry features with respect to other FPP rocks.

## 3. Samples and petrography

Two different dykes and their host peridotites from the FPP unit were investigated. The first dike (sample FI09C06, Fig. 2) was collected along the road that connects National Road 631 to a peridotite quarry located on the right flank of the Rio Creves valley (less than 100 m from the outcrops studied by Zanetti et al., 1999). The second dike (sample FI9664) is from a boulder along the Rio Creves, about 30 m upstream of its intersection with the Rio Cannobino. The host peridotite was collected 8 cm from the FI09C06 dyke, and close to the contact with the dyke represented by sample FI9664.

The gabbroic dykes are centimetric in thickness (mostly 2–5 cm; Fig. 2). They show variable strikes, usually crosscutting the harzburgite-pyroxenite association at high angles (Fig. 2).

A Mesozoic age for these sapphirine-bearing dykes is constrained by the observation that they also crosscut the foliation of protomylonites in the external domains of Mesozoic shear zones (Matysiak & Trepmann, 2015 and references therein), but are themselves deformed in a few cm-wide
182 mylonitic to ultramylonitic bands parallel to the protomylonite foliation (Tommasi et al., 2017).

183 Such shear zones were active at different crustal levels over a very long, Triassic–Jurassic time interval (235–180 Ma; Langone et al., 2018 and references therein). This finding is also consistent with the observation that all the late intrusive bodies, which are discordantly intruding the pervasive mantle foliation by hydraulic fracturing, do not show ages older than Triassic (Stähle et al., 1990, 2001; Grieco et al., 2001; Matsumoto et al., 2005).

188 The gabbroic dykes contain sapphirine (Giovanardi et al., 2013) and show a banded symmetric structure consisting of melanocratic zones at the peridotite contacts and a leucocratic zone representing the dyke core (Fig. 2). The melanocratic zones, in turn, can be divided in three different bands.

192 From the host peridotite to the dyke core, the following layers can be recognised:

193 1) An orthopyroxenite zone (hereafter, the Orthopyroxene Zone), established within the ambient peridotite. Orthopyroxene locally shows recrystallised rims with the growth of fine neoblasts of orthopyroxene and rarely olivine. The orthopyroxene sometimes presents exsolution lamellae.
195 Locally, black spinel and phlogopite occur as accessory phases. Phlogopite is concentrated in interstitial positions but rarely fills fractures within orthopyroxene crystals (Fig. 2). Giovanardi et al. (2013) reported the occurrence of a sapphirine- and amphibole-bearing recrystallisation front in the Orthopyroxene Zone.
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200 2) A first melanocratic zone (hereafter, the Early Amphibole Zone) inside veins formed by dark-brown amphibole (up to 1 cm long, named 'early amphibole') and associated small plagioclase grains, spinel, phlogopite, and apatite; in these zones, the magmatic texture is preserved, as evidenced by amphibole twinning. Apatite and spinel mainly occur as rounded inclusions in amphibole, whereas phlogopite is in interstitial position.
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205 3) A second melanocratic zone (hereafter the Late Amphibole Zone), consisting of light-brown to green amphibole (named 'late amphibole'), green spinel, sapphirine, and phlogopite. Sapphirine occurs in three textural positions: inclusions within late amphibole, coronas rimming spinel, and
isolated/aggregated crystals in interstitial positions (Fig. 2). Phlogopite is an accessory phase in interstitial positions. Late amphibole is smaller than early amphibole and is euhedral to anhedral in shape. Intermediate green-brown amphibole, often associated with spinel and sapphirine, is recognised near and through the Early Amphibole Zone-Late Amphibole Zone contact. Recrystallisation zones with a fine-grained texture occur. The Late Amphibole Zone is not continuous through the dykes. It forms patches, which can also occur only on one side of the leucocratic core of the dyke, or that can extend up to the Orthopyroxene Zone. The Late Amphibole Zone is more developed in sample FI9664 (up to about 1.5 cm in thickness) than in sample FI09C06 (up to 1 cm in thickness). In addition, sapphirine crystals in sample FI9664 can reach up to 1.5 mm in size, whereas in sample FI09C06, they are commonly < 0.2 mm.

4) A leucocratic zone (Leucocratic Zone, hereafter) formed by plagioclase and subordinate amphibole (both magmatic and relict from melanocratic zones). Apatite occurs as an accessory phase, sometimes included in plagioclase or in relicts of brown early amphibole. The apatite rarely contains calcite inclusions. Plagioclase shows twinning (mainly pericline and subordinately crossed albite and pericline), which is often partially or totally erased by recrystallisation induced by late deformation. Magmatic amphibole has greenish pleochroism, whereas relict amphibole, ripped from the melanocratic zones, is brown. Amphibole often forms single-crystal alignments parallel to the dyke strike, like in a flow texture. Phlogopite is rare and is associated with amphibole. Recrystallisation zones show fine-grained equigranular texture. Rarely, Fe-oxides, Fe-Ni sulphides, and pyrite occur.

The host peridotite away from the contact is a hornblende-harzburgite in modal composition with a porphyroclastic texture (Giovanardi et al., 2018). It is characterised by the presence of olivine and orthopyroxene porphyroclasts, with a secondary, undeformed mineral assemblage dominated by amphibole (15% by volume), in association with orthopyroxene, spinel, phlogopite, and clinopyroxene, strictly similar to the dominant peridotite type described by Zanetti et al. (1999) and Tommasi et al. (2017). A detailed petrographic inspection highlights the occurrence of a reacted
peridotite zone approaching the orthopyroxene zone. It is characterised by the presence of a
secondary mineral assemblage, modally dominated by long (up to 5 mm) phlogopite lamellae, with
associated subordinate amounts of undeformed orthopyroxene, spinel, amphibole, and
clinopyroxene. In the reacted zone, olivine was stable; conversely, the modal orthopyroxene content
is slightly lower than that in the peridotite far from the vein. Primary (e.g. olivine) and secondary
minerals in the reacted zones display elongation sub-parallel to the present-day vein strike.

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241 **4. Analytical methods**

Sample FI9664, representing the gabbroic dyke and the contact host harzburgite, was analysed for
whole-rock major and trace elements. Whole-rock major elements and Sc, V, Cr, Co, Ni, Cu, and
Zn were analysed by X-ray fluorescence spectrometry (XRF), whereas Li, Rb, Sr, Y, Zr, Nb, Ba,
rare earth element (REE), Hf, Ta, and Pb contents were analysed by inductively coupled plasma-
mass spectrometry (ICP-MS). Analyses were performed following the methods described by
Mazzucchelli et al. (2010) (data reported in Supplementary Material A).

Samples FI9664 and FI09C06 were analysed for mineral major elements with the electron
microprobe JEOL 8200 Super Probe housed at the University of Milano. Analyses were performed
following methods described by Ponce et al. (2015). Data are reported in Supplementary Material
A.

Mineral trace elemental contents were determined (as in Rivalenti et al., 2007b) with a LA-ICP-MS
housed at Istituto di Geoscienze e Georisorse, Consiglio Nazionale delle Ricerche, Sede Secondaria
di Pavia (data are reported in Supplementary Material A) using a Perkin Elmer SCIEX DCR-e
coupled with a solid-state laser source (Q-switched Nd:YAG, Quantel Brilliant). Data reduction was
performed using the GLITTER software (Griffin et al., 2008). NIST SRM 610 was used as an
external standard. Ca was used as an internal standard for clinopyroxene, amphibole, and
plagioclase, Si for olivine, orthopyroxene, and phlogopite, and Mg for spinel. Precision and
accuracy were assessed by repeated analysis of the standard BCR-2g and were found to be better

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than 10% at the ppm concentration level. Further information is reported by Giovanardi et al. (2017).

Sample FI09C06 was selected for O, Nd, and Sr isotopic analyses in mineral separates, after the combination of preliminary electron microprobe and LA-ICP-MS analyses suggested a more primitive nature for its parental melt.

O isotopes of pure mineral separates from sample FI09C06 were analysed at the I.G.G.-C.N.R., Pisa, by conventional laser fluorination with a Finnigan Delta Plus mass spectrometer. For each phase, 1–1.5 mg aliquots were necessary to measure the oxygen isotope composition. Analyses were performed following the methods described by Perinelli et al. (2011). Data are reported in Table 1.

Sr and Nd isotopic ratios of amphibole and plagioclase separates from sample FI09C06 and the host harzburgite were analysed at the Marine Environmental Sciences Laboratory (LEMAR) of the Institut Universitaire Européen de la Mer (IUEM), Université de Bretagne Occidentale. Analyses were carried out after dissolution and chromatographic separation using a TRITON Thermo- Ionisation Mass Spectrometer (TIMS), following the procedure described by Janin et al. (2012). Analyses were corrected for the reference material NBS987 ($^{87}$Sr/$^{86}$Sr = 0.710241 ± 0.000019, n = 6) for Sr and the La Jolla standard ($^{143}$Nd/$^{144}$Nd = 0.511847 ± 0.000009, n = 3) for Nd. Data are reported in Table 2.

## 5. Bulk rock chemistry

The gabbroic dyke FI9664 presents higher TiO₂, Al₂O₃, CaO, Na₂O, and K₂O contents and lower FeO and MgO contents than the host harzburgite collected at the contact.

The gabbroic dyke has trace element abundances that are one order of magnitude higher than those of the host harzburgite; only Rb and Li contents are similar in both rocks (Fig. 3; Supplementary Material A). The two rock types show some similarities in element fractionation: marked enrichments in LREEs with respect to middle-heavy rare earth elements (M-HREE; (La/Yb)N
286 values are 5.5 and 16–30 for the dyke and host harzburgite, respectively; primitive mantle data from
287 McDonough and Sun, 1995); (Th/U)N and (Zr/Hf)N always < 1; and positive Ba and Pb anomalies.
288 Conversely, the (Nb/Ta)N values are 1.6 for the gabbroic dyke and 0.7 for the host harzburgite.

289 Linearly fractionated LREE-enriched patterns were also shown by the nepheline-bearing alkaline
dykes described by Stähle et al. (2001). However, the latter had significantly higher trace element
290 contents than those of gabbroic dyke FI9664. The dykes reported by Stähle et al. (2001) also
291 displayed significant peculiarities in terms of fractionation of highly incompatible trace elements,
292 such as U, Th, Nb, Ta, Ba, and Pb, with respect to REEs.
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294 The spidergram of the peridotite close to the contact with dike FI9664 shows some relevant
differences with respect to those reported by Hartmann & Wedephol (1993): large positive Pb and
295 Hf anomalies, greater Rb, U, Ba, Ta, and Nb contents, and slightly lower LREE contents (Fig. 3).
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298 **6. Major element mineral chemistry**

299 **6.1. Host harzburgites**

300 The mineral composition of the host harzburgite far from the veins is similar to that of the
harzburgite-pyroxenite association reported by Zanetti et al. (1999). The amphibole is pargasitic in
301 composition, with only one analysis of Mg-hornblende composition. The mica is phlogopite, and
302 the clinopyroxene is diopside.
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304 Minor differences include a slightly Fe-enriched composition shown by olivine (Fo = 100 ×
Mg/(Mg + Fe²⁺tot), molar ratio of 90.4–91.1), orthopyroxene, spinel, and amphibole (Fig. 4).
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306 Orthopyroxene and spinel (hercynite to spinel in composition) are also richer in Al. Because of their
307 low Cr#, the FI09C06 oxides straddle the hercynite–spinel boundary, whereas oxides in the FPP are
308 chromite. Lower Cr content is shown by amphibole, which is also characterised by lower Na and
309 higher K contents.
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310 At the contact with the vein, the Mg# does not change significantly in the minerals. Aluminium
decreases in orthopyroxene, spinel, amphibole, and phlogopite, whereas Ti content is higher in both
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312 amphibole and phlogopite. In amphibole, Na decreases, balanced by higher K content (Fig. 5).

313 Conversely, the Na/K ratio is very variable in phlogopite.

314 Orthopyroxene from the Orthopyroxene Zone has lower Mg# and CaO content and higher Al₂O₃

315 content (Fig. 4). Similarly, its composition is different from orthopyroxene from the sapphirine-
316 bearing rock found in the LIZ northern unit of the Finero Mafic Complex by Sills et al. (1983). The
317 latter has lower Mg# and higher Al₂O₃ content (Fig. 4).

318

319 6.2. Gabbroic dykes

320 Notwithstanding the similar internal banding, the two studied dykes show marked differences in
321 terms of major elemental mineral chemistry. Significant compositional changes are also shown for
322 the different types of amphibole (i.e. the Early and Late Amphibole Zones and the Leucocratic
323 Zone; Fig. 6). The unique common feature of the amphiboles of both gabbroic dykes is the higher
324 Al content exhibited by the late amphibole (Giovanardi et al., 2013). Amphibole from gabbroic
325 dykes is mostly pargasitic in composition (unit formula calculated according to Ridolfi et al., 2018),
326 but sometimes Al substitution for Si in late amphibole is higher than 2 a.p.f.u., thereby entering the
327 sadaganaite compositional field.

328 In particular, amphibole from sample FI09C06 has distinctly lower Mg# (0.73-0.82) than sample
329 FI9664 (0.85-0.87). In sample FI09C06, Mg# is lower in early amphibole than in late and
330 leucocratic amphibole, whereas it is exactly the reverse in sample FI9664 (Fig. 6).

331 FI09C06 amphibole also displays the lowest CaO and Na₂O and the highest TiO₂ and K₂O contents
332 (TiO₂ 0.27-2.19 wt.% and K₂O 0.31-1.28 wt.% respectively) (Fig. 6).

333 In sample FI09C06, TiO₂ linearly increases with the decrease in Mg#. Instead, TiO₂ is higher in the
334 FI9664 early amphibole than in the leucocratic amphibole, but distinctly higher TiO₂ contents are
335 shown by late amphibole. Moreover, in sample FI9664 early amphibole, the increase in Mg# is
336 associated with increasing Cr and Ca, and with decrease in Na and Al contents.
As a whole, amphibole from the gabbroic dykes has higher $Al_2O_3$ content and lower Mg# and $Cr_2O_3$ content with respect to others FPP lithologies, resulting in similarity to the amphibole from the sapphirine-bearing rock in the LIZ (Sills et al., 1983) (Fig. 6). Amphibole from the two gabbroic-dykes samples shows lower CaO and $K_2O$ and higher $Na_2O$ contents than the respective host amphibole (Fig. 6).

Phlogopite from sample FI9664 presents a narrow range of major elemental contents with respect to sample FI09C06. In sample FI9664, phlogopite has higher Mg# values than phlogopite from sample FI09C06 (0.92–0.94 and 0.70–0.90, respectively) and commonly lower $TiO_2$ contents (Fig. 5). Phlogopite from the gabbroic dykes shows higher $Al_2O_3$ content and lower Mg# than the phlogopite from the host peridotite and the harzburgite-pyroxenite association (Fig. 5).

Spinel is mainly found in the Late Amphibole Zone. Unlike the spinel from the host harzburgite, the harzburgite-pyroxenite association reported by Zanetti et al. (1999), and the chromitite layers in dunite bodies (Grieco et al., 2001; 2004; Zanetti et al., 2016), the spinel from the gabbroic dykes does not contain $Cr_2O_3$ and can be classified as spinel (Mg# 0.46–0.75).

Plagioclase from sample FI9664 is commonly more anorthitic than the plagioclase from sample FI09C06 (An contents of 82–93 and 36–87, respectively). In both samples, some reversely zoned plagioclases are recognised.

The sapphirine composition falls near the 7:9:3 composition on the $SiO_2$–(FeO + MgO)–($Al_2O_3$ + $Cr_2O_3$ + $Fe_2O_3$) diagram (Fig. 7). Sapphirine from sample FI9664 is higher in $Al_2O_3$ content with respect to sample FI09C06, whereas it is lower in $SiO_2$ content.

Scanning electron microscope (SEM) investigations at the Centro Interdipartimentale Grandi Strumenti laboratories of the Università di Modena e Reggio Emilia, performed with an Environmental SEM Quanta-200 (Fei Company-Oxford Instruments), suggest that apatites from different zones of the dykes are Cl-apatites.

## 7. Trace elemental compositions
363 **7.1. Host harzburgite**

364 Amphibole and clinopyroxene from harzburgite FI09C06, 8 cm away from the contact with the
365 gabbroic dyke, are characterised by LREE-enriched linearly fractionated patterns (Figs. 8, 9 and
366 10). Their compositions are similar to those of the FPP harzburgite-pyroxenite association (Zanetti
367 et al., 1999). Proceeding toward the contact (i.e. sample FI9664), amphibole and clinopyroxene
368 show upward-convex REE patterns characterised by high variability in absolute contents (Figs. 8, 9
369 and 10). These patterns are similar to those from apatite-rich domains reported by Zanetti et al.
370 (1999) (Figs. 9 and 10).

371

372 **7.2. Gabbroic dykes**

373 The two gabbroic dykes show different trace element compositions. In sample FI09C06,
374 amphiboles from the Early Amphibole Zone and Leucocratic Zone have L-MREE-enriched
375 upward-convex patterns, similar to those of harzburgite near the contact (Fig. 8). In the Late
376 Amphibole Zone, some crystals show more fractionated patterns characterised by lower M-HREE
377 contents and positive Eu anomalies (Fig. 8). These variations are associated with marked depletion
378 in Ta, Zr, Hf, Y, Sc and V (Fig. 10, V is not shown). As a whole, amphiboles from different zones
379 of sample FI9664 show more fractionated REE patterns characterised by enrichment in LREEs and
380 depletion in HREEs (Fig. 8) with a nearly flat pattern in the LREE region. Amphibole from the
381 Early Amphibole Zone of sample FI9664 presents a small positive Eu anomaly ((Eu/Eu*)<sub>N</sub> within
382 1.04-1.65), which becomes more evident in the amphibole from the Late Amphibole Zone
383 ((Eu/Eu*)<sub>N</sub> within 1.29-1.62). Amphibole from sample FI9664 is also more enriched in Th, U, and
384 Pb (and to lesser extent in Sr) compared with amphibole from sample FI09C06 (Fig. 10).

385 Similar to the amphibole, REE patterns of plagioclase from sample FI9664 are more fractionated
386 than those of plagioclase from sample FI09C06, (Fig. 8). In both samples, the trace elemental
387 abundances of the plagioclase cores are greater than those of the plagioclase rims.
Apatite displays the typical LREE enrichment and high Th, U and Pb contents (12.31–17.35 ppm, 388 4.35–5.92 ppm, and 2.18–4.19 ppm, respectively). Nb, Ta, Zr, Hf, Ti, and Sc show negative 389 anomalies with values that are often below the detection limit. 390

No systematic trace elemental variations were found in phlogopite. 391

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## 8. O isotopes

The $\delta^{18}$O in silicates of sample FI09C06 shows a steady increase from the contact (Orthopyroxene 394 Zone) to the leucocratic gabbro in the vein core, through the Early and Late Amphibole Zones (Fig. 395 8, 10, and 11 and Table 1). In particular, it varies from 5.81‰ (standard deviation (std. dev.) 0.11) 396 in orthopyroxene from the Orthopyroxene Zone to 6.9‰ (std. dev. 0.05) in amphibole from the 397 Early and Late Amphibole Zones, and to 8.60‰ (std. dev. 0.01) in the plagioclase of the gabbroic 398 core. 399

The $\delta^{18}$O values of orthopyroxene from the Orthopyroxene Zone lie within the mantle range. They 400 are higher than the $\delta^{18}$O value reported by Hartmann & Wedephol (1993) for clinopyroxene from 401 the phlogopite-bearing amphibole harzburgite (Fig. 11), and for orthopyroxene from olivine- 402 chromitites of the FPP, but significantly lower than the orthopyroxene from orthopyroxene- 403 chromitites (Zanetti et al., 2016). 404

Vein amphibole and plagioclase have $\delta^{18}$O values significantly higher than the mantle range (Fig. 405 11). They are close to the highest values found in amphibole, phlogopite, orthopyroxene, 406 clinopyroxene, and zircon from the phlogopite-bearing amphibole harzburgite-pyroxenites 407 (Hartmann & Wedephol, 1993; Selverstone and Sharp, 2011) and chromitite layers of the FPP 408 (Zanetti et al., 2016). 409

The $\delta^{18}$O values obtained for the green spinel associated with sapphirine are markedly lower than 410 those of the associated late amphibole ($\delta^{18}$O = 4.38‰, std. dev. 0.10), as expected based on crystal- 411 chemical constraints (see Bindeman, 2008 and references therein). 412

413
414 **9. Sr and Nd isotopes**

415 The trace elemental concentrations of the mineral separates of early and late amphibole match the
416 differences highlighted by LA-ICP-MS analysis of thin sections. The Sr and Nd isotopic
417 compositions of early amphibole and late amphibole from the FI09C06 dyke are coincident within
418 analytical uncertainty (Table 2). Plagioclase from the FI09C06 Leucocratic Zone shows the same Sr
419 isotopic composition of the amphiboles, and only a slightly more radiogenic ¹⁴³Nd/¹⁴⁴Nd ratio.
420 Similar Sr isotopic compositions were documented in amphibole and apatite of discordant veins
421 from the FPP, which were characterised by higher ¹⁴³Nd/¹⁴⁴Nd ratios (Morishita et al., 2008). The
422 isotopic composition of the dyke minerals is enriched with respect to the depleted mantle reference
423 value and the mid-oceanic ridge basalt (MORB) field; these minerals fall within the ocean island
424 basalt (OIB) field (Fig. 12). In particular, they lie between the isotopic compositions of FPP
425 hornblende-syenite dykes (bulk rock from Stähle et al., 1990, 2001) and peridotites (amphibole and
426 clinopyroxene from Obermiller, 1994). Amphiboles separated from the host harzburgite (collected
427 far from the contact) show the most radiogenic Sr and unradiogenic Nd values, but have never
428 before been documented in the literature for FPP rocks (Fig. 12).

429

430 **10. Discussion**

431

432 **10.1. Petrogenesis of the sapphirine-bearing gabbroic dykes**

433 The two veins studied here share exactly the same internal banding as the sample described by
434 Giovanardi et al. (2013). This suggests that all three veins were part of an interconnected swarm,
435 which allowed them to record the same sequence of petrologic events. The symmetric structure of
436 the dykes suggests that the intrusion occurred through several events of melt percolation. In the
437 following section, we discuss, in detail, the different stages and processes that produced the
438 gabbroic dyke swarm.

439
440 **10.1.1. Melt percolation through the peridotite**

441 According to the process governing the emplacement of other dyke swarms in mantle sectors (e.g.
442 Mazzucchelli et al., 2010 and references therein), the melts flowing in the metasomatic haloes and
443 in the conduit were originally similar, undergoing severe fractionation owing to assimilation and
444 fractional crystallisation.

445 The characterisation of the wall peridotite has revealed some peculiarities. It is a common
446 observation that metasomatic haloes of cm-to-m-scale surround late dykes and veins formed by
447 melt segregation under mantle conditions. In most cases, the metasomatic haloes show marked
448 geochemical and mineralogical gradients as a function of the distance from the dike contact (Zanetti
449 et al., 1999; Ionov et al., 2002, Mazzucchelli et al., 2010; Borghini et al., 2016, 2017), which are
450 interpreted as the result of porous flow percolation of melt escaping from the open conduits.
451 Conversely, no progressive mineralogical or compositional variation is documented in the wall
452 peridotite of sample FI9664 at a variable distance from the vein (see Supplementary Material B).
453 Newly-formed clinopyroxene actually shows trace element heterogeneities, but these are randomly
454 distributed. Conversely, the trace element composition of amphibole is very homogeneous,
455 indicating late crystallisation from a unique melt.

456 The particular L/MREE-enriched convex-upward patterns shown by clinopyroxene and amphibole
457 from the wall peridotite of sample FI9664 also demonstrate that their parental melt could not have
458 been the same as that from which the adjacent early amphibole of dyke FI9664 crystallised (which
459 show linearly-fractionated REE patterns). However, based on the amphibole trace element
460 composition, the melt recorded by the wall peridotite of sample FI9664 must have been quite
461 similar to the melt that produced the Early Amphibole Zone of dyke FI09C06.

462 The textural evidence that primary and secondary minerals from the metasomatic haloes display
463 elongation sub-parallel to the present-day vein strike confirms that the development of melt
464 migration channels was associated with some local deformation.
These observations can be reconciled by assuming that the metasomatic haloes were not produced by melt escaping normally from the conduit. They were likely established during an early stage of melt migration via focused porous flow along channels developed in correspondence with structural weakness, with a direction roughly parallel to the strike of the present-day veins.

The absence of resorption textures in olivine of the reacted wall peridotite indicates that this mineral was stable. Conversely, the modal content of orthopyroxene is slightly lower than in the peridotite far from the vein. This evidence leads us to consider that the orthopyroxene saturation of the melts, indicated by the segregation of the reactive Orthopyroxene Zone and the presence of few orthopyroxenes in the veins (Giovanardi et al., 2013), was not a clear, primary characteristic, but possibly a consequence of an early stage of reactive melt migration characterised by olivine precipitation and orthopyroxene dissolution (see Piccardo et al., 2017).

The mineralogical mismatch between the nearly monomineralic, hydrous-mineral-free Orthopyroxene Zone and the Early Amphibole Zone suggests that they were segregated by compositionally different melts. This hypothesis is confirmed by large differences in terms of O isotopic compositions recorded by the orthopyroxene (5.8‰ standard mean ocean water (SMOW)) from the Orthopyroxene Zone and the amphibole from the Early Amphibole Zone (6.9‰ SMOW) within the vein. The lighter O isotopic composition of the Orthopyroxene Zone would support a hybrid composition of its parental melt caused by reactive porous flow through the FPP, with buffering of O isotopic composition at mantle values.

This conclusion led us to consider the possibility that the Orthopyroxene Zone formed before the segregation of the Amphibole Zones, possibly at the centre of the migration channels, as observed elsewhere in the FPP. It was successively split into two parts by the opening of the fracture.

### 10.1.2. Melt segregation in the open conduit

The Early Amphibole Zone was segregated when melts started flowing in open fractures. It is likely that amphibole crystallisation was triggered by the presence of an ultramafic wall, suggesting that
plagioclase was unstable in contact with the orthopyroxenite layers. The development of amphibole-rich selvages where hydrous gabbroic rocks come in contact with ultramafic layers (e.g. in the LIZ of the Finero Complex; Mazzucchelli et al., 2014) has previously been documented in the literature.

The two gabbroic dykes of this study show significant differences in terms of major and trace element mineral compositions. In particular, amphibole from the FI9664 vein shows an unusual relative enrichment in highly incompatible trace elements (U, Th, LREEs, and Na) and compatible elements (Mg, Ca, and Cr) with respect to those from the FI09C06 vein. Consistently, phlogopite is more enriched in Mg and Na, and plagioclase is more anorthitic. Assuming that the dyke swarm was produced by the injection of a unique melt, this relationship cannot be reconciled based on a melt evolution governed only by fractional crystallisation. Mazzucchelli et al. (2010) documented the same correlation between compatible and incompatible elements in minerals from cm-thick diorite dykes, demonstrating that such geochemical features can be modelled considering an assimilation of host minerals concomitant with fractional crystallisation (i.e. assimilation and fractional crystallisation (AFC) processes). Accordingly, we argue that the parental melt of vein minerals in sample FI9664 can be considered more evolved through AFC processes than that of dyke FI09C06. The concomitant enrichment of Na and Ca in the vein minerals of sample FI9664 evidences the role of amphibole in the assimilated component. This is also confirmed by the trace element patterns of FI9664 vein amphiboles, in which highly incompatible elements (such as LREEs, Th, and U) increase and moderately incompatible element (such as HREEs, Ti, and Y) decrease, moving toward the fractionation shown by those of host harzburgite away from the contact. The decrease of moderately incompatible elements with the progression of the melt evolution may be indicative of $^{Amphibole/LD}$ higher than 1 (Tiepolo et al., 2007). It is noteworthy that the Al content in sample FI9664 early amphibole is comparable with to slightly lower than in those from the FI09C06 vein, suggesting that the assimilation of host minerals was not effective in boosting the Al concentration of the evolving melt.
517

10.1.3. *Formation of sapphirine-bearing patches*

The mm-to-cm-thick sapphirine-bearing patches and stripes are randomly distributed from the internal end of the hornblendite seam to the Orthopyroxene Zone, although they are generally concentrated towards the centres of the veins. They are apparently the results of interstitial migration of a melt in chemical disequilibrium with the early cumulus minerals, as indicated by the recrystallisation fronts inside large early amphiboles. Both interstitial and recrystallised patches show that the injected melt was saturated in late amphibole, sapphirine, and spinel, whereas the saturation in phlogopite is uncertain, because it occurs only interstitially and the textural relationships are equivocal.

The following series of petrographic and geochemical features suggest that the sapphirine-bearing mineral assemblages were not simply related to interaction with the parental melt of the Leucocratic Zone, but were segregated because of the injection in the dyke swarm of an additional melt component coming from outside the system:

1) Early amphibole at the contact with or embedded by the plagioclase of the Leucocratic Zone does not show any evidence of comparable reaction.

2) Notwithstanding the significantly different major elemental composition of early amphibole documented in the two veins of this study, the major elemental chemistry of sapphirine-associated late amphibole converges toward specific Al and Mg# values. Phlogopite shows consistent variation, even though its attribution to early or late amphibole assemblages is doubtful because of its interstitial position. The data reported by Giovanardi et al. (2013) confirm this trend. This finding indicates that a unique component with a specific composition was responsible for the sapphirine segregation in the three sectors of the dyke swarm.

3) Normalised patterns of late amphibole from sapphirine-bearing areas show particular features: stronger LREE/HREE fractionation, depletion in M-HREEs, Zr, Hf, Nb, Ta, Ti, Sc,
543

and V, and the appearance of a positive Eu anomaly. The results also show inversion of the $(Zr/Hf)_N$ values, which are > 1 in the early amphibole, but < 1 in the late amphibole. The overall fractionation mimics that normally exhibited by plagioclase.

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545

546

The latter evidence, consistent with the apparent increase of the Al content in the system, suggests that the injected melt was characterised by a high normative plagioclase content.

547

548

The O, Nd, and Sr isotopic compositions of early and late Amphibole are very similar, suggesting that the plagioclase component belonged to the same magmatic cycle that produced the dyke swarm.

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550

551

However, petrographic surveying has not provided evidence of plagioclase assimilation and/or replacement by amphibole or another mineral, excluding plagioclase components derived by local assimilation. Thus, we suggest that the plagioclase signature was acquired by the injected melt through assimilation of cumulus plagioclase in hidden magmatic bodies and/or addition in the melt source, where it was likely present as a high-P metamorphic equivalent of pristine plagioclase.

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### 10.1.4. Segregation of the Leucocratic Zone

558

The petrographic evidence confirms that the formation of the Leucocratic Zone was a high energy event, which determined the partial disaggregation of the Early Amphibole Zone, with evident breaking of large early amphibole crystals. It can be locally recognised that two parts of a formerly unique crystal apparently lie on different sides of a vein. Giovanardi et al. (2013) suggested that this stage was accompanied by enlargement of the conduits.

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561

562

563

The Nd and Sr isotopic compositions of the plagioclase indicate that the parental melt had a cognate origin with those of the hornblendite selvages. This inference is also confirmed by the trace element composition of the leucocratic amphibole, which is strictly similar to that of the associated early amphibole.

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Nevertheless, the observation that the major elemental chemistry of the leucocratic amphibole is intermediate between that of early and late amphibole leads us to conclude that the parental melt of the sapphirine-bearing assemblages could be still present in diluted proportions.

## 10.2. Nature of the parental melts

Evaluating the nature of the parental melt of magmatic segregates belonging to narrow, cm-thick veins/dykes intruding mantle peridotites at high-T-P mantle conditions must be done with particular caution, with an effort first to characterise the modifications imparted by fractional crystallisation and the reaction/assimilation processes with the host rock (e.g. Mazzucchelli et al., 2010 and references therein). As discussed in the previous section, sample FI09C06 shows that its parental melt underwent the lowest degree of interaction with the host peridotite among the known dykes, and thus is the best candidate to obtain constraints on the nature of the parental melts.

An important contribution in defining the geochemical components in the melts is provided by the marked zoning in terms of O isotopic composition documented in sample FI09C06 between the Orthopyroxene Zone, the hornblendite selvage, and the Leucocratic Zone.

The $\delta^{18}$O values of early and late amphibole (6.9‰ SMOW) well exceed the mantle range (Bindeman, 2008; Polat et al., 2018), suggesting a large contribution by crustal components in the parental melt of sample FI09C06 hornblendite. The proportion of the crustal component is even larger in the leucocratic band, in which a $\delta^{18}$O value of 7.8‰ SMOW can be calculated for the amphibole in hypothetical equilibrium with the plagioclase composition (8.6‰ SMOW). This is done by conservatively considering a low closure T for the FPP of ~850 °C (provided by pyroxene-solvus geothermometry), an average value of An70 for plagioclase, and the $\delta^{18}$O mineral fractionation values reported by Bindeman (2008).

Published data on FPP minerals show a pronounced heterogeneity of the O isotopic composition, with $\delta^{18}$O values from below to well above the mantle range (Hartmann and Wedephol, 1993;
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Selverstone and Sharp, 2011; Zanetti et al., 2016). This heterogeneity may be due to multiple melt-migration events.

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Values of δ¹⁸O approaching those of early and late amphibole have sometimes been documented in harzburgites and chromitites (Hartmann & Wedephol, 1993; Zanetti et al., 2016), but the values of the leucocratic layers are markedly out of range.

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597

In addition, the δ¹⁸O value shown by orthopyroxene (5.81‰ SMOW) from the Orthopyroxene Zone provides a very important constraint, indicating how migrating melts, after prolonged interaction with the FPP, may have oxygen isotopic compositions buffered to mantle values. The decrease of δ¹⁸O values in the minerals from the hornblendite selvages with respect to those from the dyke core (i.e. the Leucocratic Zone) confirms such a buffering effect.

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We conclude that the high δ¹⁸O shown by dyke minerals cannot be the result of interaction between melts and the FPP but is a primary feature of the melts, which indicates the occurrence of large amounts of crustal components, either added to the melt in the source regions or at greater mantle depths.

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A subduction-related component was identified based on noble gas isotopic composition in Late Triassic apatite-rich layers (Matsumoto et al., 2005, Morishita et al., 2003, 2008), overprinting the crustal affinity of the host FPP. Amphiboles from such apatite layers show the same radiogenic Sr isotopic composition as the minerals of sample FI09C06, coupled to a slighter more radiogenic Nd isotopic composition.

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The isotopic compositions of early amphibole, late amphibole, and plagioclase from the FI09C06 dyke are significantly richer in radiogenic Sr and unradiogenic Nd isotopes compared with those reported from nepheline-bearing Triassic intrusions cutting the FPP (Stähle et al., 1990, 2001), which were interpreted as derived from alkaline melts of OIB affinity.

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Combining data from nepheline-bearing alkaline dykes, sapphirine-bearing dykes, and the FPP, we can define a trend at low radiogenic Nd and high radiogenic Sr isotopic ratios, suggesting mixing between asthenospheric components (OIB, according to Stähle et al., 1990; 2001 and Schaltegger et

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al., 2015) with materials derived from the continental crust (see compositional fields of Casetta et al., 2018a). We speculate that the crustal components were located at lithospheric deep levels after the Variscan orogenic cycle (e.g. Bonadiman et al., 1994) and mobilised by asthenospheric magmatism. This scenario is supported by the evidence of continental crustal metasomatism exhibited by mantle bodies involved in the Variscan collisional orogeny, such as the FPP, the Ulten Zone (Sapienza et al., 2009), and some of the Bohemian Massif (Becker et al., 1999; Schulmann et al., 2014).

In this scenario, it can be noted that the Nd and Sr isotopic compositions of sample FI09C06 lie on the mixing line between the composition of the Triassic alkaline dykes and the host harzburgite away from the contact. According to AFC simulation, assimilation of significant degrees of the metasomatised peridotite material (15%) into OIB must be considered to account for the composition of the parental melt of the FI09C06 veins.

Comparison of the trace elemental patterns of the amphibole from the vein apparently less contaminated by the host peridotite (FI09C06) with those of amphiboles segregated by primary hydrous alkaline basalts (e.g. Demeny et al., 2005) shows similarity in REE fractionation, but the early amphibole has an absolute content that is nearly one order of magnitude higher. In addition, a correlation with primary alkaline mantle melts is not directly supported by the negative Nb, Ta, and Ti anomalies shown by the early amphibole.

Equilibrium liquids calculated from the trace elemental composition of early amphibole from the FI09C06 dyke and amphibole-melt partition coefficients experimentally determined at 1,015 °C in the presence of moderately polymerised melts (dataset T2 1015; Tiepolo et al., 2007) match the REE contents and fractionation exhibited by shoshonitic rocks of the Triassic magmatism of the Dolomites area (Casetta et al., 2018a; 2018b; see Supplementary Material C). The calculated melts also show consistent negative Ti anomalies, $(Nb/La)_N$ values close to 1, and slight positive U, Th, and Pb anomalies; the latter are more pronounced in the shoshonitic melts. However, a best match in terms of U, Th, and Pb concentrations is shown by equilibrium liquids calculated in equilibrium
with amphiboles from 9,664 dykes. These observations suggest that the sapphirine-bearing gabbroic dykes may record the deep mantle input of Triassic K-rich calc-alkaline to shoshonitic melts erupted into the eastern Southern Alps, also documenting the fractionation trends responsible for the enrichment in Th, U, and Pb. This hypothesis is also supported by the Sr and Nd isotopic compositions of Triassic K-rich calc-alkaline to shoshonitic intrusive rocks and lavas of the Eastern Alps (Casetta et al., 2018a; Figure 12), which approach the isotopic trend defined by our model.

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## 10.3. Constraints on the geodynamic evolution of the Europe–Africa boundary

652

### 10.3.1. P–T constraints on the FPP

Geothermobarometric estimates constrain the intrusion of sapphirine-bearing dykes at very high-P and T conditions. Ab-initio calculations indicate that the primary crystallisation field of sapphirine in the MAS diagram (SiO₂–MgO–Al₂O₃) becomes considerably larger over 1.0 GPa and shrinks at 2.0 GPa (Belmonte et al., 2014). This evidence confirms the experimental results of Liu & Presnall (1990, 2000) and Milholland & Presnall (1998), indicating that magmatic sapphirine in FPP veins likely crystallised at P ≥ 1 GPa. Equilibrium T estimates based on the sapphirine–spinel Mg–Fe²⁺ exchange thermometer of Sato et al. (2006) and the amphibole–plagioclase thermometer of Holland & Blundy (1994) are mostly higher than 1,000 °C (up to 1,085 °C), confirming the T estimates of Giovanardi et al. (2013). The melt T was thus significantly higher than that in the host harzburgite, which shows solidus T typically corresponding to the water-oversaturated peridotite (965 °C at P = 1.1 GPa; Giovanardi et al., 2013). The absence of partial melting evidence in the host FPP confirms that the source of the uprising melts was at greater mantle depths. The high T, in combination with the high water and volatile contents, may have allowed the melt to migrate via porous flow along the direction of structural weakness (see Tommasi et al., 2017) before the opening of the conduits. The high-P emplacement conditions are consistent with the scenario where the FPP, after Palaeozoic metasomatism, remained at depths greater than the large lherzolitic mantle bodies south

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of the Anzola-Val Grande high-T shear zone (namely, from north to south, Premosello, Balmuccia, and Baldissero; Quick et al., 1995), until its exhumation at shallower levels at ~180 Ma (Zanetti et al., 2013, 2016; Langone et al., 2017, 2018; Decaris et al., 2017, Malitch et al., 2017; Petri et al., 2019).

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### 10.3.2. *Constraints on the Mesozoic mantle sources at the Africa–Europe Boundary*

The outcomes of this study indicate that in the Mesozoic, melts extremely rich in volatiles (H₂O, P, CO₂, and Cl), K, Na, and highly incompatible elements rose from mantle depths toward the surface.

The large crustal component present in the melts, as indicated by the O isotopic composition, leads us to consider that the high Al content, enrichment in large-ion lithophile elements and LREEs, and enriched Nd and Sr isotopic compositions are a primary feature inherited from the source.

This finding confirms the extreme complexity of the tectono-magmatic scenario recorded by the FPP. In particular, the northern IVZ records an extremely prolonged release of K-H₂O-rich mantle-derived melts mixed with subduction-related components, which lasted from the Variscan orogenic cycle to the Mesozoic exhumation of lithospheric mantle at shallower levels. This explains why the FPP records many generations of phlogopite-bearing mineral assemblages, showing variable field relationships, geochemical signatures, and ages (Hartmann & Wedephol, 1993; Zanetti et al., 1999; 2013; 2016; Stähle et al., 1990; 2001; Greco et al., 2001; 2004; Morishita et al., 2003, 2008; Malitch et al., 2017). It also provides a new interpretative frame for previous data indicating the emplacement of melts with subduction-related components in the Triassic (Mastumoto et al., 2005; Morishita et al., 2008, Malitch et al., 2017).

This magmatism-bearing subduction component appears to have roughly overlapped with the ascent of silica-undersaturated alkaline melts of OIB affinity (Stähle et al., 1990, 2001, Schaltegger et al., 2015), which likely also have some counterparts in the central IVZ (Fiorentini et al., 2018; Galli et al., 2019).

694
The FPP developed in a concomitant extensional-transtensional tectonic regime, the origin and geodynamic scenario of which are still strongly debated (Cassinis et al., 2008; Zanetti et al., 2013; Casetta et al., 2018a, b). Further investigations are needed to address whether the crustal components are remnants of old subduction events (Bonadiman et al., 1994), possibly located at lithospheric levels, reactivated by asthenospheric magmatism with OIB or depleted mantle affinity; represent crustal material recycled into asthenospheric mantle sources (Locmelis et al., 2016); or were added by active Mesozoic subduction (Cassini et al., 2008; Schmid et al., 2008; Morishita et al., 2008; Selverstone and Sharp, 2011; Zanetti et al., 2013).

It is a matter of fact that the major elements (in particular the enrichment in K and Al), trace elements, and Sr and Nd isotopic compositions of the sapphirine-bearing gabbroic rocks approach the geochemical affinity shown by Triassic K-rich calc-alkaline to shoshonitic magmatism of the Dolomites area (Eastern Alps) (Casetta et al., 2018a, b). Thus, the study of magmatic events at the roots of the continental crust of the Adria Plate can provide a unique opportunity to constrain the tectono-magmatic evolution of the Europe-Africa boundary.

709

## 11. Concluding remarks

A detailed survey of Mesozoic sapphirine-bearing gabbroic dykes within the FPP unit led us to recognise different stages of melt migration (from porous-flow migration in peridotite channels to flow in open conduits) and constrain the presence in the parental melts of a large continental crustal component acquired in the source region or at deeper lithospheric mantle levels.

The reaction between dyke melts and the strongly metasomatised FPP enhanced the crustal signature.

The high Al content of these melts allowed the segregation of magmatic sapphirine, which is a marker of high-P-T intrusive conditions.
This dyke swarm possibly represents a record of the input to the mantle of the K-rich calc-alkaline
719
to shoshonitic melts that characterised the magmatism widespread during the Triassic in the
720
Southern Alps area.
721

722

723 Acknowledgements

This work was supported by Ministero dell'Istruzione dell'Università e della Ricerca (MIUR),
724
Progetti di Rilevante Interesse Nazionale (PRIN) 2015 Protocol 20158A9CBM. We thank Jean
725
Yves Cottin and Dejan Prelević for their constructive comments and Elsevier Language Editing
726
Services for improving the quality of the text.
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998 Zanetti, A., Mazzucchelli, M., Sinigoi, S., Giovanardi, T., Peressini, G., Fanning, M., 2014.

999 Erratum Insights into the Melt-Lower Crust Interplay in Subduction-related Setting and the Mesozoic Geodynamic Evolution of the Southern Alps: evidence from the Finero Mafic Complex (Ivrea-Verbano Zone). Journal of Petrology 55, 1239-1240.

1002 Zanetti, A., Giovanardi, T., Langone, A., Tiepolo, M., Wu, F.-Y., Dallai, L., Mazzucchelli, M., 2016. Origin and age of zircon-bearing chromitite layers from the Finero phlogopite peridotite (Ivrea-Verbano Zone, Western Alps) and geodynamic consequences. Lithos 262, 58-74.

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1020 **Figure captions**

1021
1022 Figure 1: Geological map of the Finero area, modified after Mazzucchelli et al. (2014).

1023
1024 Figure 2: A) Sample FI09C06 crosscutting the host harzburgite foliation. The centre of the dyke is formed by the Leucocratic Zone, whereas the melanocratic zones (i.e. the Orthopyroxene Zone, Early Amphibole Zone, and Late Amphibole Zone) are indistinguishable; B) phlogopite vein cutting the Orthopyroxene Zone; C–F) occurrences of sapphirine in sample FI09C06 (C and D) and sample FI9664 (E and F). The figures show the increase of sapphirine size, from µm (C) to mm (D and E) and cm scale (F).
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1030
1031 Figure 3: Primitive mantle (PM)-normalised bulk rock trace elemental patterns of a gabbroic dyke and host harzburgite. PM values are from McDonough and Sun (1995). Literature values from the harzburgite-pyroxenite association of (1) Hartmann and Wedephol (1993) and nepheline-bearing hornblende syenitic dykes of (2) Stähle et al. (2001) are reported for comparison.
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1035
1036 Figure 4: Orthopyroxene major elemental contents from host harzburgites, dykes, and literature data. Plotted literature data are the harzburgite-pyroxenite association orthopyroxene from (1) Zanetti et al. (1999) and orthopyroxene from sapphirine-bearing rock from the Mafic Complex from (2) Sills et al. (1983).
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1040
1041 Figure 5: Phlogopite major elemental contents from host harzburgites, dykes, and literature data. Plotted literature data are: harzburgite-pyroxenite association phlogopite from (1) Zanetti et al. (1999).
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1045 Figure 6: Amphibole major elemental contents from host harzburgites, dykes, and literature data. Plotted literature data are: harzburgite-pyroxenite association amphibole compositions from (1) Zanetti et al. (1999) and (2) Morishita et al. (2008), amphibole from apatite-rich veins in the FPP
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1048 from (1) Zanetti et al. (1999) and (2) Morishita et al. (2008), and amphibole from sapphirine-
bearing rock from the Mafic Complex from (3) Sills et al. (1983).
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1050
1051 Figure 7: Sapphirine compositions plotted in the (MgO + FeO)–(Cr₂O₃ + Fe₂O₃ + Al₂O₃)–SiO₂ diagram (mol. %). Literature data are from sapphirine of the Finero Mafic Complex (Sills et al., 1983) and sapphirine 1 of Higgins et al. (1979).
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1054
1055 Figure 8: REE patterns of clinopyroxene and amphibole from the gabbroic dykes and the respective
1056 host rocks divided for samples and position. Values are normalised to Chrondrite-I (CI, values from
1057 Lyubetskaya and Korenaga (2007); δ¹⁸O values (normalised to SMOW) from phases of sample
1058 FI09C06 are reported near the REE patterns according to their positions in the sample.
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1060 Figure 9: PM-normalised trace elemental patterns of clinopyroxene from host harzburgite. PM
1061 values are from McDonough and Sun (1995). Plotted literature data are: clinopyroxene and
1062 amphibole average compositions from the harzburgite-pyroxenite association from (1) Zanetti et al.
1063 (1999) and (2) Morishita et al. (2008), clinopyroxene and amphibole average from apatite-rich veins
1064 in the FPP from (1) Zanetti et al. (1999) and (2) Morishita et al. (2008), reported as apatite veins.
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1066 Figure 10: PM-normalised trace elemental patterns of amphiboles from gabbroic dyke and host harzburgite. PM values are from McDonough and Sun (1995).
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1068
1069 Figure 11: δ¹⁸O normalised to SMOW. Values from literature are the field of mantle and variation
1070 of MORB from Rollinson (1993), and mantle ultramafics and mantle-derived melts (*) from
1071 Bindeman (2008). Finero literature data are reported from the harzburgite-pyroxenite association of
1072 Hartmann and Wedephol (1993) and Selverstone and Sharp (2011) and from chromitite in dunite
1073 bodies from Zanetti et al. (2016).
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Figure 12: ¹⁴³Nd/¹⁴⁴Nd vs ⁸⁷Sr/⁸⁶Sr recalculated at 225 Ma for amphibole and plagioclase from the various zones of sample FI09C06 (host and dyke): (1) amphibole data of the FPP from Obermiller (1994); (2) bulk rock data of the Finero Mafic Complex from Lu et al. (1997b); (3) data from an alkaline dyke in the FPP from Stähle et al. (1990); (4) data from an alkaline dyke in the FPP from Stähle et al. (2001); (5) amphibole and apatite data from discordant apatite-bearing veins in the FPP from Morishita et al. (2008); (6) Mesozoic shoshonitic magmatism (SS: silica saturated; US: undersaturated in silica) in the Predazzo area from Casetta et al. (2018a, b); Depleted MORB Mantel (DMM) from Workman and Hart (2005). The mixing model was calculated between the two end-members: the alkaline dyke S9 of Stähle et al. (1990) and the hosting peridotite FI09C06. Data for the melt in equilibrium with the alkaline dyke S9 are: ¹⁴³Nd/¹⁴⁴Nd = 0.512607, ⁸⁷Sr/⁸⁶Sr = 0.703720 (from Stähle et al., 1990; recalculated at 225 Ma), Nd = 8.333 ppm, and Sr = 830 ppm (calculated from LA-ICP-MS plagioclase analysis of albitite dykes similar to the dyke of Stähle et al., 1990, using the K_d of Dohmen and Blundy, 2014; average Nd = 0.25 ppm, Sr = 5810 ppm). Data for the melt in equilibrium with the host FPP peridotite are ¹⁴³Nd/¹⁴⁴Nd = 0.512130, ⁸⁷Sr/⁸⁶Sr = 0.708501 (recalculated at 225 Ma), Nd = 67.5 ppm, and Sr = 909 ppm (calculated using the K_d of Ionov et al., 2002).
<figure><img src="image_2.png" /><figcaption>Figure 1
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**Figure 2**
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<figure><img src="image_3.png" /></figure>
<figure><img src="image_4.png" /><figcaption>Figure 3
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<figure><img src="image_5.png" /><figcaption>Figure 4
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<figure><img src="image_6.png" /><figcaption>Figure 5
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<figure><img src="image_7.png" /><figcaption>Figure 6
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<figure><img src="image_8.png" /><figcaption>**Figure 7**
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<figure><img src="image_9.png" /><figcaption>Figure 8
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<figure><img src="image_10.png" /><figcaption>Figure 9
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<figure><img src="image_11.png" /><figcaption>Figure 10
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<figure><img src="image_12.png" /><figcaption>Figure 11
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<figure><img src="image_13.png" /><figcaption>Figure 12
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<table><caption>**Table 1**
Click here to download Table: Table 1_rev.docx</caption><thead><tr><th>Sample</th><th>Phase</th><th>δ<sup>18</sup>O</th><th>std. dev.</th></tr></thead><tbody><tr><td>FI09C06</td><td>Early Amp</td><td>6.94</td><td>0.00</td></tr><tr><td>FI09C06</td><td>Late Amp</td><td>6.86</td><td>0.05</td></tr><tr><td>FI09C06</td><td>P1</td><td>8.60</td><td>0.01</td></tr><tr><td>FI09C06</td><td>Opx</td><td>5.81</td><td>0.11</td></tr></tbody></table>

<caption>Table 1: O isotopic composition of minerals from the Spr-bearing gabbroic dykes.</caption>
<table><caption>Table 2: Sr and Nd isotopic composition of minerals from sample FI09C06 from the Spr-bearing gabbroic dykes and the host peridotite.</caption><thead><tr><th>Rock</th><th>Phase</th><th>Rb</th><th>Sr</th><th><sup>87</sup>Sr/<sup>86</sup>Sr</th><th>2SE</th><th><sup>87</sup>Rb/<sup>86</sup>Sr</th><th>Nd</th><th>Sm</th><th><sup>143</sup>Nd/<sup>144</sup>Nd</th><th>2SE</th><th><sup>147</sup>Sm/<sup>144</sup>Nd</th></tr></thead><tbody><tr><td>Host</td><td>Amp</td><td>6.5</td><td>281.2</td><td>0.708713</td><td>0.000008</td><td>0.066372</td><td>19.7</td><td>4.2</td><td>0.512317</td><td>0.000008</td><td>0.126796</td></tr><tr><td>Dyke</td><td>Early Amp</td><td>6.8</td><td>324.4</td><td>0.704722</td><td>0.000008</td><td>0.060834</td><td>24.8</td><td>8.1</td><td>0.512646</td><td>0.000008</td><td>0.197559</td></tr><tr><td>Dyke</td><td>Late Amp</td><td>8.6</td><td>189.5</td><td>0.704971</td><td>0.000008</td><td>0.131167</td><td>16.8</td><td>4.0</td><td>0.512559</td><td>0.000006</td><td>0.142593</td></tr><tr><td>Dyke</td><td>Pl</td><td>0.6</td><td>3109.7</td><td>0.704519</td><td>0.000008</td><td>0.000568</td><td>2.2</td><td>0.2</td><td>0.512482</td><td>0.000006</td><td>0.061238</td></tr></tbody></table>