# Carbonate metasomatism at mantle wedge conditions, evidence from trace element and stable isotope (Fe, Zn) signatures of orogenic peridotites

Baptiste Debret, Paolo A Sossi, N. Malaspina, Adrien Gautier, N. Mattielli,

H. O'Neill, J. Villalobos-Orchard, Frédéric Moynier

► To cite this version:

Baptiste Debret, Paolo A Sossi, N. Malaspina, Adrien Gautier, N. Mattielli, et al. Carbonate metasomatism at mantle wedge conditions, evidence from trace element and stable isotope (Fe, Zn) signatures of orogenic peridotites. Chemical Geology, 2025, 680, pp.122649. 10.1016/j.chemgeo.2025.122649. hal-04974652

HAL Id: hal-04974652

https://hal.science/hal-04974652v1

Submitted on 3 Mar 2025

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altered oceanic crust (AOC; e.g., Hawkesworth et al., 1993; Plank and Langmuir, 1993; Miller et al., 1994; Elliott et al., 1997; Class et al., 2000; Elliott, 2004). However, whether or not the slab-derived components are directly transferred to the mantle wedge by fluids or indirectly, through an intermediate reservoir (e.g., serpentinite melange; Marschall and Schumacher, 2012), and their nature (supercritical/liquid/fluid) remain an active research frontier. Orogenic suprasubduction mantle peridotites are an alternative and complementary way to arc magmas for studying fluid/melt crust-to-mantle transfer in subduction zones. Petro-geochemical reconstructions show that these rocks, from a diversity of localities, contain high-pressure volatile-bearing minerals (e.g., phlogopite, amphibole, carbonate and sulfide) and anhydrous (e.g., pyroxene, garnet) phases enriched in LILE and light rare earth elements (LREE) with respect to heavy rare earth elements (HREE; e.g., Zanetti et al., 1999; Nimis and Morten, 2000; Zhang et al., 2000, 2007; Malaspina et al., 2006, 2009; Scambelluri et al., 2006; Selverstone and Sharp, 2011; Chen et al., 2017; Förster et al., 2024). These geochemical features are characteristic of slab-derived components. However, the exact nature of the metasomatic agent able to transport the elemental budget at high pressure (HP) remains unclear as the orogenic peridotites often record multistage metamorphic events, including protolith inheritance, HP-metasomatism and late exhumation processes, leading to variable trace element signatures from one geodynamic setting to another. Hence it remains an open question as to whether the geochemical records of metasomatized orogenic peridotites correspond to heterogeneous conditions in the suprasubduction mantle wedge worldwide or whether these relate to complex and specific P-T paths associated with regional and multiple metasomatic events (e.g., Malaspina et al., 2006, 2023; Scambelluri et al., 2006; Zhang et al., 2007; Pellegrino et al., 2020).

Metal stable isotope variations have the potential to complement trace element observations in mantle wedge environments. This is because the formation of melts, aqueous and supercritical fluids or brines can lead to contrasting isotope fractionation of elements that are sensitive to the speciation of the fluid/liquid phase, such as Zn and Fe (e. g., Williams et al., 2004; Debret et al., 2016, 2021; Liu et al., 2016; Pons et al., 2016; Moynier et al., 2017; Wang et al., 2017; Gerrits et al., 2019; Sossi and Debret, 2021; Li et al., 2024). In addition to chemical composition, this fractionation is also influenced by the redox state of the fluid or melt, since equilibrium isotope fractionation between different fluids and minerals is driven by contrasts in bonding conditions between chemical elements, including coordination degree and oxidation state (Ducher et al., 2016; Fujii et al., 2014; Polyakov and Mineev, 2000; Schauble, 2004; Sossi and O'Neill, 2017). Iron isotopes in arc magmas are systematically lighter than their MORB counterparts at a given Mg # (Mg/[Mg + Fe²⁺]) (Dauphas et al., 2009; Foden et al., 2018; Nebel et al., 2015; Yan et al., 2024), attesting to processes occurring in subduction zones that are distinct from those during the genesis of MORB. One proposed mechanism states that Fe is transported by slab-derived fluids, which is supported by the inverse correlation observed between the slab thermal parameter and the iron isotope composition of basaltic magmas from arcs worldwide (Foden et al., 2018). This is further supported by the sympathetic enrichment of B (relative to fluid-immobile Nb) in these lavas (Chen et al., 2023).

Another way of testing this hypothesis is through examination of the isotopic signature of metasomatized orogenic peridotites that provide complementary information on redox processes occurring in subduction zones. In detail, the direction and magnitude of the isotope fractionation of such metasomatized peridotites relative to those unaffected by subduction processes can provide insights into the nature of the fluid involved (i.e., melts, aqueous and supercritical fluids or brines) and the oxygen fugacity conditions in the mantle wedge. For example, Turner et al. (2018) reported light Fe isotope signature in peridotite xenoliths hosted in basaltic andesites on the Batan Island (Philippines) that are too light to have been formed by equilibrium partial melting alone. This requires the addition of slab fluid enriched in light Fe isotopes. Similarly, serpentinized forearc mantle wedge peridotites show large negative

deviations (down to $\delta^{56}\text{Fe} = -0.26 \text{ ‰}$; Debret et al., 2020) that are correlated with high bulk rock $\text{Fe}^{3+}/[\sum \text{Fe}]$ ratio, suggesting the transfer of oxidizing slab fluids with an isotopically light signature. Silicate melts are invariably heavier than the peridotites or eclogites from which they are derived owing to the greater incompatibility of $\text{Fe}^{3+}$ relative to $\text{Fe}^{2+}$ (Dauphas et al., 2009; Tiraboschi et al., 2023). Therefore, the isotopically light signatures observed in metasomatized peridotites are likely imparted by slab-derived fluids, in which ferrous iron, particularly in chloride form, is more soluble than ferric iron (e.g., Chou and Eugster, 1977).

To examine the nature of slab metasomatism and the oxygen fugacity ($\text{fO}_2$) record of mantle wedge peridotites, we combine trace element concentration with Fe and Zn isotope analyses in peridotites from Balmuccia and Finero (Ivrea Zone, Italian Alps). We show that metasomatized peridotites from Finero have high LREE/HREE relative to unmetasomatized samples from Balmuccia and are also 1 $\log_{10}$ unit more oxidised, which is consistent with the percolation of slab-derived fluids. A negative correlation between U/Th and LREE/HREE ratios is associated with isotopically light Zn signature and high Zr/Hf ratios in metasomatized samples, which cannot be achieved through simple metasomatism of peridotites by silicate melts (or fluids) at high pressures and temperatures. Rather, these constraints suggest the dissociation of an oxidised, carbonate rich metasomatic agent (either fluid or melt) probably during Finero massif exhumation.

## 2. Geological settings

Balmuccia and Finero massifs are the largest and most extensively studied ultramafic bodies outcropping in the northern part of the Ivrea Zone (Fig. 1a). These bodies are aligned at the northwestern margin of the lowest stratigraphic units of the Ivrea Zone, representing the lowermost unit of the alpine continental crust (Fountain, 1976; Mehert, 1975). Both Balmuccia and Finero massifs were tectonically emplaced into the Ivrea crust during the Alpine Orogeny (Lu et al., 1997;

<figure><img src="image_2.png" /><figcaption>**Fig. 1.** Generalized geologic maps of the Ivrea- zone (northern Italy). The red stars correspond to the studied ultramafic massifs. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>
Quick et al., 1995). Despite their similar tectonic locations, these massifs display strong petrological and geochemical differences suggesting that they represent different segments of the lithosphere, generated in different tectonic environments (Rivalenti and Mazzucchelli, 2000; Selverstone and Sharp, 2011).

The Balmuccia ultramafic massif is a 5 km-long lens bounded by the Lensch mafic complex. The massif is dominated by spinel lherzolite with minor harzburgite and dunite (associated with chromitites) often interlayered with pyroxenites, which occur in two series, the chrome diopside (Cr-Di) and the aluminium augite (Al-Aug) suites, and is crosscut by a wide range of later mafic dykes (Mazzucchelli et al., 2009; Shervais, 1979). The pyroxenites display Sr and Nd isotopic compositions similar to MORB and are indistinguishable from that of the host lherzolite, suggesting that the massif record a tectonic thinning of a subcontinental lithosphere prior the ocean formation (Shervais and Mukasa, 1991). The Balmuccia massif is characterized by extremely fresh peridotites and pyroxenites with negative loss-on-ignition values (Shervais, 1979; Shervais and Mukasa, 1991). The studied samples are taken from Sossi et al. (2018). They were collected along the Sesia river that bisects the massif at its southern tip. This sample suite is subdivided according to peridotite, including lherzolite, harzburgite and dunite, and pyroxenites (Cr-Di and Al-Aug) (Table 1).

The Finero ultramafic massif is an antiformal body that comprises

phlogopite- and amphibole-bearing peridotites at its core. The massif is surrounded by amphibolite- to granulite facies layered metagabbros (Zanetti et al., 1999). The abundance of hydrous phases (amphibole and phlogopite) is highly variable at both outcrop and massif scale (Giovanardi et al., 2014). Rare carbonate (dolomite) were also described by Zanetti et al. (1999) and Morishita et al. (2008). In contrast to Balmuccia, spinel harzburgite is the dominant rock type within the peridotite, with subordinate lherzolite, dunite and websterite. They are often interlayered with clinopyroxenite. Although the nature of the metasomatic agent (i.e., melt, aqueous or supercritical fluids) is debated in detail, there is a general consensus that the phlogopite- and amphibole-bearing peridotites formed in a subduction-like context (Hartmann and Hans Wedepohl, 1993; Zanetti et al., 1999; Selverstone and Sharp, 2011). Based on bulk rock composition of phlogopite- and amphibole-bearing peridotites and thermodynamic modelling, Selverstone and Sharp (2011) calculated P-T conditions of equilibration nearby 900 °C and 1.2–1.6 GPa, compatible with shallow mantle wedge conditions. The studied samples were collected along the private road to the Finero quarry and along the Cannobino river, from the Provola bridge to the Creves–Cannobino confluence (Tommasi et al., 2017). We subdivided the studied samples according to their mineralogy, distinguishing between peridotites and pyroxenites (Table 1). The sampled peridotites correspond to spinel harzburgites with variable amounts of

<table><caption>**Table 1**
Fe and Zn isotope analyses in ‰.</caption><thead><tr><th>Sample Name</th><th>Locality</th><th>Mineralogy</th><th>δ<sup>56</sup>Fe</th><th>2sd</th><th>δ<sup>57</sup>Fe</th><th>2sd</th><th>n</th><th>δ<sup>66</sup>Zn</th><th>2sd</th><th>δ<sup>68</sup>Zn</th><th>2sd</th><th>n</th></tr></thead><tbody><tr><td colspan="13" style="text-align: left; font-weight: bold;">Western Alps</td></tr><tr><td>BM-12</td><td>Balmuccia</td><td>Lherzolite</td><td>0.02</td><td>0.01</td><td>0.03</td><td>0.02</td><td>2</td><td>0.15</td><td>0.02</td><td>0.27</td><td>0.06</td><td>2</td></tr><tr><td>BM-14</td><td>Balmuccia</td><td>Lherzolite</td><td>0.06</td><td>0.01</td><td>0.10</td><td>0.01</td><td>2</td><td>0.16</td><td>0.06</td><td>0.31</td><td>0.12</td><td>1</td></tr><tr><td>BM19</td><td>Balmuccia</td><td>Dunite</td><td>0.04</td><td>0.01</td><td>0.06</td><td>0.02</td><td>2</td><td>0.12</td><td>0.02</td><td>0.23</td><td>0.04</td><td>2</td></tr><tr><td>BM22</td><td>Balmuccia</td><td>Lherzolite</td><td>0.03</td><td>0.02</td><td>0.04</td><td>0.03</td><td>2</td><td>0.18</td><td>0.06</td><td>0.37</td><td>0.08</td><td>2</td></tr><tr><td>BM-23</td><td>Balmuccia</td><td>Lherzolite</td><td>0.03</td><td>0.04</td><td>0.04</td><td>0.06</td><td>2</td><td>0.14</td><td>0.02</td><td>0.29</td><td>0.15</td><td>2</td></tr><tr><td>BM-26B</td><td>Balmuccia</td><td>Lherzolite</td><td>-0.04</td><td>0.01</td><td>-0.06</td><td>0.01</td><td>2</td><td>0.11</td><td>0.06</td><td>0.60</td><td>0.12</td><td>1</td></tr><tr><td>BM-27</td><td>Balmuccia</td><td>Lherzolite</td><td>0.01</td><td>0.04</td><td>0.01</td><td>0.06</td><td>6</td><td>0.14</td><td>0.04</td><td>0.28</td><td>0.09</td><td>2</td></tr><tr><td>BM28-Pd</td><td>Balmuccia</td><td>Dunite</td><td>0.01</td><td>0.01</td><td>0.02</td><td>0.02</td><td>2</td><td>n.d.</td><td></td><td>n.d.</td><td></td><td></td></tr><tr><td>BM29-Pd</td><td>Balmuccia</td><td>Al-Aug</td><td>n.d.</td><td></td><td>n.d.</td><td></td><td></td><td>0.10</td><td>0.06</td><td>0.20</td><td>0.12</td><td>1</td></tr><tr><td>BM30</td><td>Balmuccia</td><td>Lherzolite</td><td>0.00</td><td>0.03</td><td>0.01</td><td>0.04</td><td>2</td><td>0.15</td><td>0.06</td><td>0.30</td><td>0.12</td><td>1</td></tr><tr><td>BM31</td><td>Balmuccia</td><td>Dunite</td><td>-0.03</td><td>0.02</td><td>-0.04</td><td>0.04</td><td>4</td><td>0.16</td><td>0.04</td><td>0.33</td><td>0.08</td><td>1</td></tr><tr><td>BM35</td><td>Balmuccia</td><td>Lherzolite</td><td>0.03</td><td>0.02</td><td>0.05</td><td>0.03</td><td>4</td><td>0.12</td><td>0.06</td><td>0.25</td><td>0.12</td><td>1</td></tr><tr><td>BM-40</td><td>Balmuccia</td><td>Lherzolite</td><td>0.03</td><td>0.03</td><td>0.05</td><td>0.05</td><td>2</td><td>0.12</td><td>0.02</td><td>0.26</td><td>0.04</td><td>2</td></tr><tr><td>BM42</td><td>Balmuccia</td><td>Lherzolite</td><td>n.d.</td><td></td><td>n.d.</td><td></td><td></td><td>0.16</td><td>0.06</td><td>0.30</td><td>0.12</td><td>1</td></tr><tr><td>BMH-1</td><td>Balmuccia</td><td>Lherzolite</td><td>0.00</td><td>0.00</td><td>0.00</td><td>0.00</td><td>2</td><td>0.18</td><td>0.06</td><td>0.39</td><td>0.12</td><td>1</td></tr><tr><td>BMH-2</td><td>Balmuccia</td><td>Lherzolite</td><td>0.03</td><td>0.01</td><td>0.05</td><td>0.02</td><td>2</td><td>0.13</td><td>0.11</td><td>0.25</td><td>0.23</td><td>2</td></tr><tr><td>BMH-3Pd</td><td>Balmuccia</td><td>Lherzolite</td><td>-0.01</td><td>0.02</td><td>-0.02</td><td>0.03</td><td>2</td><td>0.11</td><td>0.06</td><td>0.22</td><td>0.03</td><td>2</td></tr><tr><td>BMH-5</td><td>Balmuccia</td><td>Lherzolite</td><td>0.02</td><td>0.01</td><td>0.03</td><td>0.02</td><td>2</td><td>0.12</td><td>0.06</td><td>0.25</td><td>0.12</td><td>1</td></tr><tr><td>BMH-9</td><td>Balmuccia</td><td>Lherzolite</td><td>0.03</td><td>0.02</td><td>0.04</td><td>0.03</td><td>2</td><td>0.16</td><td>0.06</td><td>0.33</td><td>0.12</td><td>1</td></tr><tr><td>BM28-Px</td><td>Balmuccia</td><td>Clinopyroxenite</td><td>0.13</td><td>0.01</td><td>0.19</td><td>0.02</td><td>2</td><td>0.21</td><td>0.06</td><td>0.42</td><td>0.12</td><td>1</td></tr><tr><td>BM29-Px</td><td>Balmuccia</td><td>Clinopyroxenite</td><td>n.d.</td><td></td><td>n.d.</td><td></td><td></td><td>0.18</td><td>0.06</td><td>0.37</td><td>0.12</td><td>1</td></tr><tr><td>BMH-3Px</td><td>Balmuccia</td><td>Clinopyroxenite</td><td>0.09</td><td>0.01</td><td>0.13</td><td>0.01</td><td>2</td><td>0.23</td><td>0.01</td><td>0.45</td><td>0.01</td><td>2</td></tr><tr><td>FIN12</td><td>Finero</td><td>Serpentinite</td><td>0.08</td><td>0.09</td><td>0.13</td><td>0.11</td><td>4</td><td>n.d.</td><td></td><td>n.d.</td><td></td><td></td></tr><tr><td>FIN2</td><td>Finero</td><td>Amp-Sp Harzburgite</td><td>-0.06</td><td>0.06</td><td>-0.09</td><td>0.14</td><td>3</td><td>-0.01</td><td>0.05</td><td>0.01</td><td>0.16</td><td>3</td></tr><tr><td>FIN3</td><td>Finero</td><td>Sp Harzburgite</td><td>0.00</td><td>0.03</td><td>-0.02</td><td>0.05</td><td>3</td><td>-0.21</td><td>0.03</td><td>-0.31</td><td>0.12</td><td>3</td></tr><tr><td>FIN4</td><td>Finero</td><td>Amp-Sp Harzburgite</td><td>0.00</td><td>0.03</td><td>0.01</td><td>0.07</td><td>2</td><td>0.14</td><td>0.03</td><td>0.30</td><td>0.14</td><td>5</td></tr><tr><td>FIN5</td><td>Finero</td><td>Phl-Sp Harzburgite</td><td>0.02</td><td>0.04</td><td>0.05</td><td>0.07</td><td>3</td><td>0.00</td><td>0.02</td><td>0.07</td><td>0.02</td><td>5</td></tr><tr><td>FIN6</td><td>Finero</td><td>Amp-Sp Harzburgite</td><td>0.01</td><td>0.04</td><td>-0.01</td><td>0.04</td><td>5</td><td>-0.07</td><td>0.04</td><td>-0.11</td><td>0.11</td><td>4</td></tr><tr><td>FIN8</td><td>Finero</td><td>Phl-Sp Harzburgite</td><td>-0.04</td><td>0.06</td><td>-0.07</td><td>0.07</td><td>2</td><td>-0.26</td><td>0.05</td><td>-0.33</td><td>0.10</td><td>3</td></tr><tr><td>FIN9</td><td>Finero</td><td>Amp-Sp Harzburgite</td><td>-0.04</td><td>0.00</td><td>-0.06</td><td>0.01</td><td>2</td><td>-0.01</td><td>0.02</td><td>0.10</td><td>0.07</td><td>5</td></tr><tr><td>FIN10</td><td>Finero</td><td>Amp-Sp Harzburgite</td><td>-0.01</td><td>0.01</td><td>0.00</td><td>0.03</td><td>4</td><td>-0.19</td><td>0.07</td><td>-0.28</td><td>0.15</td><td>2</td></tr><tr><td>FIN11</td><td>Finero</td><td>Clinopyroxenite</td><td>0.12</td><td>0.04</td><td>0.16</td><td>0.06</td><td>3</td><td>0.23</td><td>0.03</td><td>0.50</td><td>0.05</td><td>5</td></tr><tr><td>FIN7</td><td>Finero</td><td>Phl Clinopyroxenite</td><td>0.08</td><td>0.02</td><td>0.10</td><td>0.06</td><td>3</td><td>0.08</td><td>0.04</td><td>0.26</td><td>0.12</td><td>4</td></tr><tr><td colspan="13" style="text-align: left; font-weight: bold;">Standards</td></tr><tr><td>BCR-2</td><td></td><td></td><td>n.d.</td><td></td><td>n.d.</td><td></td><td></td><td>0.15</td><td>0.04</td><td>0.31</td><td>0.08</td><td>5</td></tr><tr><td>Preferred values<sup>1</sup></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>0.17</td><td>0.05</td><td></td><td></td><td></td></tr><tr><td>BHVO-2</td><td></td><td></td><td>0.10</td><td>0.01</td><td>0.20</td><td>0.06</td><td>2</td><td>0.33</td><td>0.04</td><td>0.60</td><td>0.07</td><td>2</td></tr><tr><td>duplicate</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>0.31</td><td>0.02</td><td>0.55</td><td>0.07</td><td></td></tr><tr><td>Preferred values<sup>2</sup></td><td></td><td></td><td>0.10</td><td>0.03</td><td>0.13</td><td>0.05</td><td></td><td>0.28</td><td>0.04</td><td></td><td></td><td></td></tr><tr><td>UB-N</td><td></td><td></td><td>0.03</td><td>0.04</td><td>0.07</td><td>0.07</td><td>3</td><td>0.29</td><td>0.07</td><td>0.59</td><td>0.13</td><td>3</td></tr><tr><td>Preferred values<sup>3</sup></td><td></td><td></td><td>0.06</td><td>0.03</td><td>0.08</td><td>0.05</td><td></td><td>0.38</td><td>0.09</td><td>0.73</td><td>0.3</td><td></td></tr></tbody></table>

1, Moynier et al. (2017); 2, Craddock et al. (2013); Telus et al. (2012)

Data for Balmuccia samples were reported in Sossi et al. (2016) (Fe) and Sossi et al. (2018) (Zn). When n = 1, the 2sd corresponds to long-term reproducibility.
clinopyroxene (< 10 %). The clinopyroxenes are often associated with large amount of fluid and mineral inclusions, such as olivine and orthopyroxene (Fig. 2a) suggesting a refertilization of the peridotite through metasomatic fluids. Amphibole and phlogopite are the dominant hydrous minerals (Fig. 2b).

# 3. Method

## 3.1. Mineral major element analyses

The major element composition (SiO₂, Al₂O₃, FeO, FeOTotal, MnO, MgO, CaO, Na₂O, K₂O, TiO₂) of spinels, olivines and orthopyroxenes from Finero massif was analyzed on a CAMECA SX100 electron microprobe at the Laboratoire Magmas et Volcans (Clermont-Ferrand, France) for oxygen fugacity estimates (see discussion). Quantitative analyses were performed using a monochromatic crystal filtering the polychromatic incident X-ray beam. For most analyses, the beam voltage was 15 kV and the intensity 15 nA. Results are displayed in Supplementary Table 1.

## 3.2. Bulk rock major element analyses

Major element analyses of Balmuccia ultramafic rocks are from Sossi et al. (2018). Major element analyses (SiO₂, Al₂O₃, Fe₂O₃, MnO, MgO, CaO, Na₂O, K₂O, TiO₂) of Finero ultramafic rocks were carried out at the Service of Analyses of Rocks and Minerals (SARM) in Nancy, France, using a Thermo Scientific iCap6500 ICP-OES. Sample digestions were performed on LiBO₂ fluxed fusions following the procedures described by Carignan et al. (2001). Results for reference material UB-N analyzed alongside the samples show a difference of less than 2 % with literature values (Supplementary Table 2).

## 3.3. Bulk rock trace element analyses

For trace element analyses of the Balmuccia samples, 100 mg of fine powder (crushed in a steel jaw crusher and ground in W—C mills) was dissolved on a hot plate in a concentrated HCl-HF-HNO₃ (1:0.5:0.2) acid mixture at 140 °C for 2 days. During subsequent evaporation, 15.7 M HNO₃ was added periodically to convert the residue to nitrates and avoid the formation of fluorides, which were then re-dissolved in 2 mL of 15.7 M HNO₃ with a few drops of 24 M HF. The Teflon beakers containing the solutions were placed in a PTFE Teflon-lined high-pressure bomb vessel and heated to 200–210 °C in order to dissolve chromite.

After drying down, they were re-dissolved in 2 mL of 15.7 M HNO₃. Typically, 50 µL of this solution was diluted 100×, to 5 mL, with 2 % (0.317 M) HNO₃ and spiked with 10 ppb each of Rh and Re, resulting in a total dilution factor of ≈ 1000× (Eggins et al., 1997). Calibration was performed by standardization against synthetic solutions diluted to the 500, 100, 10, 1, and 0.1 ppb levels. These solutions were used in conjunction with BHVO-2 to correct for differences in the matrix between samples and the standard solutions, as well as oxide corrections on the Rare Earth Elements (REE). Rhodium (< 138 amu) and Re (≥ 138 amu) internal standards corrected were used for mass bias and instrumental drift corrections. Samples were introduced by a glass nebulizer to a Peltier-cooled glass spray chamber coupled to an Agilent 7700 quadrupole ICP-MS at the Australian National University. Collision/reaction cell gases were not used. Detection limits were of the order of 0.1 ppt (0.1 ppb in the original powder) and the relative standard deviation was typically between 0.1 and 2 % and did not exceed 5 %.

Additional, trace element data on Finero ultramafic rocks were carried out at the Institut de Physique du Globe de Paris (IPGP) using an Agilent 7900 quadrupole ICP-MS. Elements with masses between that of sodium (23) and arsenic (75), as well as silver (107), cadmium (111) and gadolinium (157), were measured using a collision-reaction cell with helium gas (5 mL/min) to remove polyatomic interferences. All other elements were measured without collision gas. Reported uncertainties were calculated using error propagation equations and considering the combination of standard deviation on replicated consecutive signal acquisitions (n = 3), internal-standard ratio and blank subtraction. A JP1 peridotite and a UB-N serpentinite were used as external standard. The obtained values are in good agreement with literature data (Supplementary Table 2).

## 3.4. Bulk rock Fe and Zn stable isotope analyses

50 mg of whole rock powders from the Finero massif were dissolved using concentrated HF and HCl acids in 7 mL PTFE Teflon square digestion vessels with wrench top closures in an oven at 160 °C for 3 days. The prior study of Debret et al. (2018) compared this method to Parr bomb dissolution techniques and showed that it results in the digestion of refractory phases, such as spinels, present in peridotites. Following evaporation of the HF and HCl, the samples were subsequently treated with a 1:3 mix of concentrated HCl and HNO₃ and refluxed at 140 °C for 3 days in an oven. Finally, samples were brought into solution in 6 M HCl prior to column chemistry. Quantitative

<figure><img src="image_3.png" /><figcaption>**Fig. 2.** Thin section microphotographs of the Finero peridotites. (a) Microphotograph (cross polarized light) of a clinopyroxene containing a large olivine cluster. This micro-texture indicates a reaction between olivine and a fluid / melt to form clinopyroxene. (b) Microphotograph (plane polarized light) of a phlogopite observed in Finero peridotites.</figcaption></figure>
purification of Zn and Fe was achieved by chromatographic exchange, using 1 mL of AG1-x8 (200–400 mesh) and 1.2 mL Bio-Spin© Chromatography columns, following the procedure adapted from Maréchal et al. (1999) and developed by Sossi et al. (2015). All reagents used in the chemistry and mass spectrometry procedures were distilled and sub-boiled at the IPGP. The total procedural blank contribution was <18 ng for Zn and < 80 ng of Fe, which is negligible compared to the amount of Zn and Fe in the samples (blank contribution is <1 %).

Iron and zinc isotope analyses were performed by multiple-collector inductively-coupled-plasma mass-spectrometry (MC-ICP-MS; Thermo Scientific Neptune Plus) at the IPGP. Instrumental mass fractionation was corrected using the sample-standard bracketing technique. In the case of Fe analyses, solutions consisted of 3 ppm of natural Fe in 0.5 M HNO₃, which were introduced into the mass spectrometer using a quartz cyclonic spray chamber and PFA 100 µl/min nebulizer. The international Fe isotope standard IRMM-014 was used as both the bracketing and delta notation reporting standard. To resolve the polyatomic interferences ⁴⁰Ar¹⁶O, ⁴⁰Ar¹⁴N, and ⁴⁰Ar¹⁶OH⁺ that occur on masses ⁵⁶Fe, ⁵⁴Fe, and ⁵⁷Fe, respectively, the instrument was run in high-resolution mode, which gave a mass resolving power (mass/Amass) of ~12,000. The standard Fe beam intensities typically varied between 15 and 20 V for ⁵⁶Fe using a 10¹⁰ Ω resistor. Data were collected in 15 cycles of 8.388 s integrations. Isotope ratios are reported as δ⁵⁶Fe in permil notation relative to IRMM-014 external standard:

$$ \delta^{56}\text{Fe} = ((^{56}\text{Fe}/^{54}\text{Fe}_{\text{sample}})/(^{56}\text{Fe}/^{54}\text{Fe}_{\text{IRMM}} - 1) \times 1000 $$

$$ \delta^{57}\text{Fe} = ((^{57}\text{Fe}/^{54}\text{Fe}_{\text{sample}})/(^{57}\text{Fe}/^{54}\text{Fe}_{\text{IRMM}} - 1) \times 1000 $$

The δ⁵⁷Fe was given to demonstrate mass dependency of the measurements. All reported errors are 2SD. Mass dependence, long-term reproducibility, and accuracy were evaluated by analyzing an in-house FeCl₂ salt standard. The values obtained for repeated measurements of this standard solution yielded average δ⁵⁶Fe values of -0.75 ± 0.04 ‰ and δ⁵⁷Fe values of -1.12 ± 0.13 ‰ 2sd, n = 21. These average values are in excellent agreement with previous measurements (Williams and Bizimis, 2014). The international rock reference materials United States Geological Survey (USGS) BCR-2 (basalt), USGC PCC-1 (peridotite) and Centre de Recherches Pétrographiques et Géochimiques (CRPG, Nancy, France) UB-N (serpentinite) were analyzed in the same analytical sessions as the samples, yielding δ⁵⁶Fe and δ⁵⁷Fe values in good agreement with literature values (Table 1). Iron isotope analyses for the Balmuccia peridotites are taken from Sossi et al. (2016) using the procedures described in Sossi et al. (2015).

Zinc isotope compositions were measured on the purified sample solutions. These solutions were prepared at 200 ng/g of natural Zn. This gave a Zn intensity of ~4 V for ⁶⁴Zn. The solutions were introduced to the instrument using a PFA 100 µl/min nebulizer coupled to a quartz cyclonic spray chamber. Owing to a lack of polyatomic interferences across the mass range for Zn the machine was operated in low mass resolution mode. The intensity of ⁶²Ni was monitored to correct for the potential interference of ⁶⁴Ni on ⁶⁴Zn. The Zn isotope composition of the sample is presented as a delta value in permil notation relative to the JMC-Lyon isotopic standard.

$$ \delta^{66}\text{Zn} = ((^{66}\text{Zn}/^{64}\text{Zn}_{\text{sample}})/(^{66}\text{Zn}/^{64}\text{Zn}_{\text{JMC}} - 1) \times 1000 $$

$$ \delta^{68}\text{Zn} = ((^{68}\text{Zn}/^{64}\text{Zn}_{\text{sample}})/(^{68}\text{Zn}/^{64}\text{Zn}_{\text{JMC}} - 1) \times 1000 $$

The δ⁶⁸Zn was given to demonstrate mass dependency of the measurements. Due to a limited supply of the JMC-Lyon standard solution, samples were measured relative to an AA-ETH pure Zn solution. Mass dependence, long-term reproducibility, and accuracy were evaluated by analyzing the JMC-Lyon standard solution in the same analytical sessions as the samples. The difference of AA-ETH relative to JMC-Lyon gave an average δ⁶⁶Zn value of +0.32 ± 0.04 ‰, 2sd, n = 18. This average value is in excellent agreement with previous measurements

(Archer et al., 2017). The international rock reference materials USGS BCR-2 (basalt), BHVO-2 (basalt) and CRPG UB-N (serpentinite) were analyzed in the same analytical sessions as the samples, yielding δ⁶⁶Zn values in good agreement with literature (Table 1). Zinc isotope analyses for the Balmuccia peridotites are taken from Sossi et al. (2018) using the procedures described in Sossi et al. (2015).

# 4. Results

## 4.1. Major and trace elements

Results of whole rock analyses are reported in Supplementary Table 2 and illustrated in Figs. 3 and 4. Balmuccia peridotites are characterized by CaO, Al₂O₃ and MgO contents ranging from 0.20 to 7.21 wt%, 0.42 to 4.13 wt% and from 31.79 to 48.38 wt%, while pyroxenites are characterized by higher CaO (14.93–16.36 wt%), Al₂O₃ (4.50–12.24 wt%) and lower MgO (18.48–21.75 wt%) contents. All Balmuccia samples display mostly negative loss on ignition (L.O.I.) values (down to -0.39 wt% and up to 0.22 wt%), indicating weight gain during ignition due to oxidation of FeO to Fe₂O₃ and negligible loss of volatile compounds (H₂O, CO₂). Pyroxenites are characterized by high REE concentrations relative to peridotites but with a similar pattern. The REE patterns of Balmuccia ultramafic rocks display variable depletion in light REE (LREE; LaN/YbN ≤0.01–1; N: chondrite normalized) and a flat medium and heavy rare earth element (M-HREE) segment. Their extended trace element patterns are characterized by a positive Pb anomaly (PbN/Cen = 6–212, n: primitive mantle normalized) and variable alkali enrichments (CsN/Thn = 2–33) relative to neighbouring elements.

Finero peridotites display CaO, Al₂O₃ and MgO contents ranging from 0.20 to 4.13 wt%, from 0.19 to 1.51 wt% and from 38.16 to 45.93 wt%. The Finero pyroxenites are characterized by low CaO (3.74–16.24 wt%), Al₂O₃ (1.22–2.20 wt%) and high MgO (24.80–38.16 wt%) relative to Balmuccia pyroxenites. Finero samples are characterized by variable L.O.I. (b.d.l.-5.12 wt%), in agreement with the presence of hydrous phases in thin sections (amphibole or phlogopite, Fig. 2). One sample of serpentinite (FIN 12) display similar XMg (0.92), CaO (1.80 wt%) and Al₂O₃ (2.93 wt%) but high L.O.I. content (9.71 wt%) relative to fresh ultramafic rocks. The REE patterns are characterized by LREE enrichments (LaN/YbN = 1.4–11.5) and a variable Eu anomaly (Eu/Eu* = 0.8–2.2). As for Balmuccia, pyroxenites have similar patterns relative to peridotites but display higher REE concentrations. The extended trace element patterns are characterized by negative Zr, Hf (HfN/Smn < 1.1) and positive Sr (Sr/Sr* = 0.7–5.1) and Pb (PbN/Cen = 4–380) anomalies. They are also characterized by alkali (CsN/Thn = 79–854) enrichments relative to neighbouring elements.

## 4.2. Fe and Zn isotopes

The δ⁶⁶Zn and δ⁵⁶Fe of Balmuccia ultramafic rocks range from +0.10 ± 0.06 to +0.23 ± 0.01 ‰ and from -0.04 ± 0.01 to +0.06 ± 0.01 ‰, respectively, and overlap those of the estimated compositions of the primitive mantle (Fig. 5). Balmuccia pyroxenites display high δ⁵⁶Fe values (from +0.09 ± 0.01 to +0.13 ± 0.01 ‰) and δ⁶⁶Zn values (from +0.18 ± 0.06 to +0.23 ± 0.01 ‰) relative to peridotites (δ⁵⁶Fe from -0.03 ± 0.02 to +0.06 ± 0.01 ‰; δ⁶⁶Zn from +0.10 ± 0.06 to +0.18 ± 0.06 ‰; Sossi et al., 2018, Sossi et al., 2016). Finero peridotites display lower δ⁶⁶Zn and δ⁵⁶Fe values, from -0.26 ± 0.05 to +0.14 ± 0.03 ‰ and from -0.06 ± 0.05 to +0.06 ± 0.02 ‰, respectively (Fig. 5). Finero pyroxenites display high δ⁵⁶Fe (+0.08 ± 0.02 – +0.12 ± 0.04 ‰) and δ⁶⁶Zn (+0.08 ± 0.04 – +0.23 ± 0.03 ‰) values, similar to those of the Balmuccia pyroxenites.

# 5. Discussion

Mineralogical and geochemical variations affecting mantle peridotites are commonly interpreted to reflect variable degrees of melt
<figure><img src="image_4.png" /><figcaption>Fig. 3. Plots of (a) Al₂O₃ and (b) CaO versus MgO contents of Balmuccia and Finero samples (anhydrous contents in wt%). The boxes highlight a zoom in a small insert presenting peridotite data. Non-Modal batch melting model are shown in the small inserts, see main text and Supplementary Table 3 for more details about the model. Literature data are represented by small symbols (data from Hartmann and Wedepohl, 1993). Primitive mantle values are from McDonough and Sun (1995).</figcaption></figure>

extraction that can be followed by either fluid- and/or melt meta-
somatism (e.g., Herzberg (2004) and reference therein). The Balmuccia
ultramafic rocks correspond to fresh and unmetasomatised peridotites
composed of clinopyroxene, orthopyroxene, olivine and spinel (Shervais
and Mukasa, 1991; Hartmann and Hans Wedepohl, 1993). In contrast,
the Finero samples contain high-pressure hydrous (i.e., amphibole,
phlogopite) and carbon-bearing (i.e., dolomite) phases. Previous
geochemical studies have proposed that these phases result from a high-
pressure metasomatic event in a subduction related context (Cannaò
et al., 2022; Giovanardi et al., 2014; Matsumoto et al., 2005; Selverstone
and Sharp, 2011; Zanetti et al., 1999). This sample suite including both
unmetasomatized and metasomatized ultramafic rocks therefore pro-
vides an unmatched opportunity to identify and characterise the relative
effects of melting- and subduction related processes on the mantle
wedge.

5.1. *Partial melting of the mantle wedge*

Before addressing the role of subduction-derived components in the studied samples, we first investigate the effects of high temperature melting processes on peridotite composition prior their metasomatism at high pressure. It must be noted that late alteration, e.g., serpentinization, can affect the δ<sup>66</sup>Zn of peridotite, shifting them toward to heavy values relative to the primitive mantle (Debret et al., 2024). However, the Balmuccia samples are particularly fresh with Loss On Ignition values ranging from -0.4 to 0.2 wt%, while the Finero samples display low δ<sup>66</sup>Zn relative to the primitive mantle. Similarly, no correlation between Zn (or Fe) isotopes and indices of serpentinization (e.g., LOI) was observed (Supplementary Fig. 1). We hence considered that the isotopic values of the samples are not primarily controlled by late alteration processes.

The “fertility” of a peridotite, which refers to the degree to which it
<figure><img src="image_5.png" /><figcaption>Fig. 4. Bulk rock REE and trace element patterns of (a, c) Balmuccia and (b, d) Finero ultramafic rocks normalized to the chondrite and primitive mantle values (McDonough and Sun, 1995), respectively.</figcaption></figure>

<figure><img src="image_6.png" /><figcaption>Fig. 5. Fe and Zn isotopic ranges of Balmuccia and Finero ultramafic rocks compared with mantle peridotite and pyroxenite worldwide. Primitive mantle values are from Craddock et al. (2013) and Sossi et al. (2018) for Fe and Zn isotopes, respectively. Data are form: 1, Weyer and Ionov (2007); 2, Poitrasson et al. (2013); 3, Williams et al. (2005); 4, Huang et al. (2011); 5, Zhao et al. (2012); 6, Zhao et al. (2015); 7, Turner et al. (2018); 8, Zhang et al. (2019); 9, Su et al. (2015); 10, Debret et al. (2020); 11, Doucet et al. (2016); 12: Wang et al. (2017); 13: Beunon et al. (2020); 14: Debret et al., 2024; 15: Sossi et al. (2018).</figcaption></figure>
has experienced melt extraction and its capacity to produce partial melts, is often expressed using CaO, Al₂O₃ and MgO contents. Fertile lherzolites display high CaO or Al₂O₃ and low MgO relative to refractory harzburgites and dunites (e.g., Bodinier and Godard, 2013). The studied samples display a negative covariation between CaO or Al₂O₃ and MgO contents with pyroxenites displaying the highest CaO or Al₂O₃ and lowest MgO contents and peridotites showing lower CaO or Al₂O₃ and MgO contents relative to the primitive mantle (Fig. 3). This reflects a decrease in modal clinopyroxene passing from pyroxenites to peridotites. One peridotite sample (BM-14) displays intermediate composition with abnormally high CaO (7.21 wt%) or Al₂O₃ (4.13 wt%) and low MgO (31.79 wt%) contents, reflecting lithological heterogeneities at sample scale (i.e., mixing with pyroxenite) during sample crushing. It must be noted that Finero peridotites record a large range of fertility, from lherzolite to dunite -like compositions. This variability contrasts with the forearc peridotites that commonly experienced very high degrees (>20 %) of melt extraction and display pyroxene poor harzburgite to dunite like compositions (e.g., Parkinson and Pearce, 1998).

The change in major element contents of peridotites during partial melting can be quantified through a non-modal batch melting model developed by Shaw (1970). Although this approach is not strictly applicable, because the partition coefficient must change as a function of melt fraction, it suffices to demonstrate the concept. Using the partitioning and modal melting reaction coefficients determined at 3 GPa by Walter (1998, 2003), this model is presented in Fig. 3 along with peridotite CaO, Al₂O₃ and MgO contents. Mineral modes, partition coefficients and modal melting reaction coefficients can be found in Supplementary Table 3. It should be noted that pressure has minimal effect on partitioning and modal melting reaction coefficients below 3 GPa (Walter, 1998). This melt modelling successfully fits the CaO, Al₂O₃ and MgO compositions of the peridotite dataset, suggesting that the observed major element geochemical diversity can be explained by different degrees of partial melting. However, as this lithological heterogeneity occurs at the metre-scale, partial melting is not likely to have occurred in-situ by adiabatic decompression in the mantle, as the peridotite massifs are km-sized.

As shown in Fig. 5, the isotopically heavy composition of MORB relative to primitive mantle (Teng et al., 2013; Sossi et al., 2016), indicates that melting processes are expected to decrease both the δ⁵⁶Fe and δ⁶⁶Zn values in peridotites, resulting in isotopically light peridotites that experienced extensive melt extraction (Doucet et al., 2016; Williams et al., 2004). Interestingly, negative covariations between either δ⁵⁶Fe and δ⁶⁶Zn and indices of peridotites fertility are observed in our samples (Fig. 6). However, previous attempts to model the behaviour of δ⁵⁶Fe and δ⁶⁶Zn during partial melting (e.g., Williams et al., 2004; Weyer and Ionov, 2007; Williams and Bizimis, 2014; Dauphas et al., 2017; Sossi and O'Neill, 2017; Sossi et al., 2018; Beunon et al., 2020) have shown that melting has only a limited effect on the isotopic composition of residual peridotite. For example, 30 % of partial melting will shift the δ⁵⁶Fe and δ⁶⁶Zn of less than 0.1 ‰ in the residual peridotite at pressures <3 GPa. It must be noted that Doucet et al. (2016) suggested a higher Zn isotope fractionation during high degree of partial melting, which was reconsidered as unrealistic by most recent studies (e.g., Sossi et al., 2018; Beunon et al., 2020). Furthermore, Fang et al. (2022) recently duplicated four fertile lherzolites from Doucet et al. (2016) and concluded that previously highly fractionated δ⁶⁶Zn values for mantle peridotites were due to analytical problems.

Here we used the non-modal batch melting model developed by Sossi and O'Neill (2017) for isotope partitioning to model both δ⁵⁶Fe and δ⁶⁶Zn values for peridotites during partial melting at 1300 °C. These models, which are presented in Fig. 6, employ melt-mineral fractionation factors from Sossi and O'Neill (2017) and Sossi et al. (2018) for Fe and Zn, respectively (Supplementary Table 3). Consistent with previous studies (e.g., Nebel et al., 2015; Dauphas et al., 2017; Sossi et al., 2018), the isotopic composition of peridotite is minimally affected by melting processes (up to 0.03 ‰ lighter at f = 0.3 for Fe and Zn) because most of

<figure><img src="image_7.png" /><figcaption>Fig. 6. Plots of (a) δ⁵⁶Fe and (b) δ⁶⁶Zn versus MgO contents. Non-Modal batch melting model are shown in grey line, see main text and Supplementary Table 3 for more details about the model.</figcaption></figure>

the Fe and Zn remains in the source, limiting the amount of fractionation that can develop in the residue by the lever rule. The models provide an adequate description of the observed Zn isotope composition of Balmuccia peridotites, though some samples tend to display lower δ⁵⁶Fe values relative to that predicted by non-modal batch melting models. A similar observation was reported by Dauphas et al. (2017) who suggested that the used isotope fractionation coefficients between melt and minerals are not well enough documented to predict the full extent of Fe isotope fractionation during partial melting.

On the other hand, the δ⁵⁶Fe variations of Finero peridotites overlap with those of the Balmuccia samples. However, their Zn elemental and isotopic compositions show values that are too low (Figs. 5b & 6b) to be explained by non-modal batch melting. For a given MgO content, there is a consistent shift of up to 0.3 ‰ in δ⁶⁶Zn between peridotites from Finero and those from Balmuccia. While some variability is expected in mantle peridotites worldwide (Fig. 6), which might reflect some mantle heterogeneities at the global scale, Finero samples display isotopically light Zn compositions relative to mantle peridotites from other tectonic settings such as cratonic mantle and serpentinized forearc peridotites. This suggests that the isotope signatures of the Finero peridotites are incompatible with regional mantle heterogeneities with the most straightforward explanation being the influence of subduction-related component.

## 5.2. High pressure metasomatism

Metasomatism in a subduction setting can be related to C-O-S-H fluids and/or hydrous melts, which are generated in the wedge or in the slab. The Finero samples exhibit significant enrichment in the LREE relative to the HREE (La/N/Yb_N = 0.7–11.5), which is absent in unmetasomatized Balmuccia ultramafic rocks (Figs. 4 and 7). Such “spoon shape” fractionation is typically recorded by suprasubduction
<figure><img src="image_8.png" /><figcaption>**Fig. 7.** Plots of La/Yb versus (a) Th (ppm), (b) U (ppm) concentrations and (c) U/Th ratios in Balmuccia and Finero peridotites. Litterature date are in small symbol (data from Hartmann and Wedepohl, 1993).</figcaption></figure>

peridotites metasomatized by crust-derived fluid phases (e.g., Scam-belluri et al., 2006; Malaspina et al., 2006, 2023). It is likely due to the high modal amount of clinopyroxene that partitions LREE with respect to the other mineral phases. These enrichments in LREE are also corre-lated with an increase of LILE concentrations (e.g., Cs, Rb) in the bulk rock (Fig. 4). The later are commonly interpreted as a contribution of slab derived fluids. Indeed, through prograde metamorphism, the con-centration of LILE decreases in metasedimentary rock by progressive devolatilization (Bebout et al., 2007; Bebout et al., 1999). Similarly, slab-derived aqueous/supercritical liquids are expected to fractionate LREE elements relative to HREE. Hence the contribution of slab derived component in Finero peridotites can explain the observed enrichments in LREE and LILE. In agreement with this interpretation, the Finero geochemical patterns, characterized by a depletion in major elements,

but enriched in K, Rb, Ba, Sr and LREE, were first attributed by Zanetti et al. (1999) to a slab-derived “carbonatite” melt component where pyroxenites are regions of high concentrations of such melt. Similarly, Selverstone and Sharp (2011) proposed that metasomatism by fluids released from a downgoing slab produced the δ³⁷Cl composition of Finero phlogopite-bearing peridotite, while amphibole-bearing perido-tites require an additional melt component. To distinguish between metasomatism related to hydrous fluids or silicate melts at eclogite facies conditions it is instructive to examine the behaviour of a variety of trace elements, most notably U, Th and LREE among these phases (Kessel et al., 2005). In particular, at 900 °C, it is possible to discriminate fluids from melt or supercritical fluids based on excess of U over Th (U/ Th > > 1). Peridotite La/Yb ratios increase with U and Th concentra-tions, with Balmuccia samples having the lowest and Finero samples having the highest La/Yb ratios, U and Th concentrations (Fig. 7a-b). Lanthanum is relatively more mobile in fluids or melts than Yb, such that highly metasomatized samples should display the highest La/Yb ratios. Hence, these observations suggest a high degree of metasomatism in Finero relative to Balmuccia samples. However, Th is sparingly soluble in fluids (Kessel et al., 2005), making difficult to reconcile trace element enrichment in Finero peridotites with metasomatism by an aqueous fluid. Although U/Th ratios are relatively constant between the studied massifs (Fig. 6c), U/Th is negatively correlated with La/Yb ratios in Finero samples. This might suggest a slightly higher mobility of Th relative to U during metasomatic processes. In melts, the partitioning of Th is similar (U/Th ~ 1) at 900 °C to slightly higher (U/Th > 1) at 1000 °C to that of U (inferred peak temperature in Finero) while, at these conditions, an excess of Th over U is always expected in supercritical liquids (Kessel et al., 2005). Metasomatism by both melts and super-critical liquids can explain the composition of Finero peridotites. How-ever, the occurrence of supercritical liquids in subduction systems requires pressure > 2.3 GPa for sedimentary systems, > 3.4 GPa for basaltic systems and > 4 GPa for ultramafic systems (Ni et al., 2017). Hence, a direct metasomatism by supercritical fluid phases is unlikely due the low-pressure conditions record of Finero massif.

### 5.3. Deciphering the nature of Finero metasomatism using stable isotopes

During subduction, numerical models (e.g., Syracuse et al., 2010) suggests that top slab temperatures might be high enough to cross the wet solidus for hydrated sediments and altered oceanic crust in most arcs. However, both HP-slab derived sediments (δ⁶⁶Zn = +0.40 ± 0.13 ‰; Debret et al., 2021) or eclogite crust (δ⁶⁶Zn = +0.22 ± 0.13 ‰; Inglis et al., 2017) display high δ⁶⁶Zn values relative to mantle peridotites. Also, on the whole, silicate melts preferentially complex heavy zinc isotopes relative to the residual peridotite as shown, for example, by the heavy zinc signature of MORB relative to mantle peridotites (Fig. 5b). The metasomatism of the mantle wedge by slab-derived melts will therefore result in the transfer of a heavy zinc component and therefore shift mantle wedge peridotites toward heavy values, which is not observed in Finero, therefore calling for other mechanisms.

Zinc isotope fractionation during slab devolatilization is intimately linked to Zn complexation in metamorphic fluids. It has recently been proposed that the transfer of isotopically heavy zinc from the slab to the mantle wedge occurs by oxidizing fluids carrying, for example, car-bonate (e.g., Zn–CO₃; Debret et al., 2021) or sulfate (e.g., Zn–SO₄; Pons et al., 2016) ligands. More reduced fluids carrying, for example, Zn (HS)₂⁻ will have isotopically light Zn isotopic signatures (Debret et al., 2018). Theory also predicts that S-bearing fluids (either Fe-bearing sulfate- or sulfide-rich) will have isotopically light Fe isotope signa-tures (Fujii et al., 2014), compatible with that of the Finero peridotite (Fig. 5).

In order to explore any potential link between the Zn and Fe signa-tures of metasomatized samples and the transfer of a redox sensitive element (S or C), we investigated the fO₂ conditions recorded by mantle wedge peridotites. The fO₂ was calculated using the olivine-
orthopyroxene-spinel (OOS) equilibrium reaction:

$$6\text{Fe}_2\text{SiO}_4 + \text{O}_2 = 3\text{Fe}_2\text{Si}_2\text{O}_6 + 2\text{Fe}_3\text{O}_4 \quad (1)$$

Assessment of the oxygen fugacity requires treatment of the activity-composition relations of mineral solid solutions of the samples. For olivine and orthopyroxene, this is largely a function of Fe—Mg substitutions, which mix symmetrically and near-ideally (O'Neill et al., 2003; Wiser and Wood, 1991). The magnetite component in spinel is more complex; the experimental work by Wood (1990) shows that $Fe^{3+}/\Sigma Fe$ ($aFe_3O_4$) is a positive function of the Cr content as well as the $fO_2$. If not accounted for, this leads to a systematic overestimation of $fO_2$ as a function of Cr content of spinel (e.g. Mattioli and Wood, 1988), which increases from lherzolite to dunite. A comparison of the expressions of O'Neill and Wall (1987), Mattioli and Wood (1988) and Nell and Wood (1991), Wood (1990) showed that the Nell-Wood parameterization provided the most accurate $fO_2$, with a deviation of no more than 0.35 log units, and no systematic bias on Cr content. Therefore, the $fO_2$ estimates determined by the Nell and Wood (1991) calibration were used (Supplementary Table 4; see also Davis et al. (2017) for a recent assessment).

According to these calculations, Finero peridotites equilibrated at a $fO_2$ ranging between +0.2 and +1 log units above the Fayalite-Magnetite-Quartz buffer (ΔFMQ). These variations are close to that reported in metasomatized supra-subduction garnet peridotite samples from Sulu belt (China) and Ulten Zone (Italian Central Alps), along with UHP garnet peridotites from Bardane (Western Gneiss Region, Norway) by Malaspina et al. (2009) and (Malaspina et al., 2010) based on $Fe^{3+}$ contents of metamorphic garnet. Reported values range from -1.2 ± 1 (for the deepest samples of Bardane - 6 GPa) to +1.1 ± 0.4 (for the Sulu and Ulten peridotites - 5 GPa and 3 GPa, respectively) relative to FMQ buffer. They are also consistent with $fO_2$ calculated for spinel peridotite xenoliths from other subduction settings, typically ranging from -1 to +2 relative to FMQ buffer (Brandon and Draper, 1996; Cottrell et al., 2021; Parkinson and Arculus, 1999; Peslier et al., 2002). As already observed by Malaspina et al. (2009), the orogenic samples do not follow a trend of decreasing $fO_2$ with pressure, such as recorded by sub-cratonic xenoliths (Woodland and Koch, 2003). In contrast, OOS oxybarometry in Balmuccia peridotites record lower $fO_2$ (ΔFMQ = 0) than in the Finero peridotites, more in line with the oxygen fugacities recorded by mid-ocean ridge peridotites (ΔFMQ = 0 ± 0.4; Birner et al., 2018) or other massif peridotites (ΔFMQ ~0) (Davis et al., 2017; Woodland et al., 1992).

The more oxidised nature of the Finero peridotites is inconsistent with the circulation of reduced fluids with an isotopically light Zn isotope signature in the massif. However, it suggests a causal link between occurrence of metasomatic phases such as phlogopite, dolomite and amphibole therein and oxidation. These hydrous and carbonate rich phases require a metasomatism driven by fluids potentially equilibrated above the CCO (C + $O_2$ = $CO_2$) oxygen buffer, which, at 1.5 GPa defines an $fO_2$ within 0.2 log units of the FMQ buffer and becomes relatively more oxidizing as pressure increases (e.g., Frost and Wood, 1995). The stability of dolomite in these rocks is a further evidence for metasomatism at $fO_2$s above the FMQ buffer. Our $fO_2$ estimates show that this metasomatic event is such that the present equilibrium state of the Finero massif records ~ΔFMQ +1, at 900 °C and 1.2–1.6 GPa, hence being suggestive of chemical exchange at high $fO_2$ between mantle wedge peridotites and slab fluids. It must be noted that these high $fO_2$ conditions do not preclude the formation of Fe(II) species in fluids or rocks, as for example Fe(II)-rich meta-peridotites were reported at highly oxidizing conditions (i.e., FMQ + 3). This conclusion is also consistent with arc magma studies that predict, based on $Fe^{3+}/\Sigma Fe$ ratio measurements, a mantle wedge area equilibrated nearby ~ΔFMQ +1.5 ± 1 (Behrens and Gaillard, 2006; Carmichael, 1991; Gaborieau et al., 2020; Kelley and Cottrell, 2009). Although our $fO_2$ estimates in orogenic peridotites are suggestive of oxidation reactions associated with slab fluid transfer in subduction zone, Malaspina et al. (2009)

showed that the addition of components such as $K_2O$, $H_2O$ and $CO_2$, and therefore formation of phlogopite, amphibole and carbonates, can lead to variations in oxygen chemical potential (i.e., $fO_2$), unrelated to net whole-rock oxidation or reduction. The elevated $fO_2$ values recorded in the Finero samples are associated with either light $\delta^{56}Fe$ and $\delta^{66}Zn$ values relative to PM (Fig. 8), which suggests a concomitant oxidation and isotope fractionation of mantle wedge peridotites through metasomatic processes. In addition, the mobility of isotopically light Fe in slab fluids is inferred to result from the presence of Fe(II)-$CO_3$ or Fe(II)-$SO_4$ complexes equilibrated up to Hematite-Magnetite buffer (~ΔFMQ +3; Debret et al., 2016). Hence, association of light $\delta^{66}Zn$ and possibly $\delta^{56}Fe$ values with high $fO_2$ in refractory peridotites require an external oxidative agent. However, neither a direct transfer of aqueous slab fluids or silicate melts can simultaneously explain the trace element- and Fe-Zn isotopic signatures of the Finero peridotites.

Previous petrographic and geochemical studies have proposed that the Finero massif could record a multistage fluid evolution, with Finero phlogopite-bearing peridotite resulting from a carbon rich hydrous melt metasomatism that subsequently exsolves a carbon rich agent during its exhumation to the surface (Morishita et al., 2003). Evidence for a carbon rich metasomatism includes the occurrence of apatite and carbonate rich layers in Finero peridotites (Zanetti et al., 1999; Morishita et al., 2003). The formation of these layers was attributed to the crystallization of carbonatitic melts at upper mantle conditions. Similarly, carbon isotope ($\delta^{13}C$) investigations of Finero peridotites show enrichments in C relative to the deep mantle (Cannaò et al., 2022). Additional micro-Raman investigations by Cannaò et al. (2022) document the occurrence of graphite and carbonate bearing fluid inclusions, specifically in olivine and clinopyroxene, that are likely responsible for the high C concentrations reported in Finero peridotites.

Carbon rich fluid metasomatism in an open system is likely to produce high ratio of pyroxenes, as well as the crystallization of apatite, amphibole and/or phlogopite and the release of $CO_2$-rich fluids as the massif cross between 1.5 and 2 GPa the following reaction (e.g., Sapienza et al., 2009):

$$(1) \text{ enstatite} + \text{dolomite} = \text{forsterite} + \text{diopside} + \text{CO}_2$$

The high abundance of pyroxene, phlogopite and amphibole in Finero peridotites (Fig. 2) and the presence of carbon rich fluid inclusions are consistent with such a reaction. Similarly, carbon rich fluid metasomatism is expected to increase Nb/Ta and Zr/Hf ratios in mantle peridotite (e.g., Chakhmouradian, 2006; Rudnick et al., 1993). In agreement with this, Finero peridotites display high Zr/Hf ratios (33–66) relative to Balmuccia peridotites (Zr/Hf = 12–37).

The complexes $Zn-CO_3$ in fluid form are among the isotopically heaviest complexes that could form on Earth (Fujii et al., 2014). Hence, the formation of metasomatic minerals (e.g., phlogopite, pyroxenes or carbonates) through the dissociation of a carbonate rich agent (e.g., eq. (1)) will necessarily favour the formation of an isotopically light Zn peridotite and the release of an isotopically heavy Zn fluid (or melt). This is supported by a negative correlation between Zr/Hf ratios and $\delta^{66}Zn$ values in Finero peridotites (on a restricted amount of analyses, see Supplementary Fig. 2). To further test such a hypothesis, we attempted to model the evolution of $\delta^{66}Zn$ in peridotites through the dissociation of a carbonate rich fluid. The behaviour of Zn in fluid was approximated to that of $Zn-CO_3$ complexes in aqueous fluids using the ab initio calculations of Fujii et al. (2014) as, to our knowledge, the behaviour of $Zn-CO_3$ complexes in melts remains unknown. The elemental and isotopic behaviour of Zn in peridotite was modelled using the available ab initio calculations of Dacher et al. (2016) for smithsonite ($ZnCO_3$), hydrozincite ($Zn_5(CO_3)_2(OH)_6$) and hemimorphite ($Zn_4Si_2O_7H_2O$)) assuming to represent Zn behaviour in carbonates and silicate, respectively. The calculated fractionation factors between carbonate rich fluids and minerals ($\alpha_{CO_2-minerals}$) are displayed in Fig. 9a. At 900 °C, the 1000 x ln ($\alpha_{CO_2-minerals}$) ranges between 0.1 and 0.2, with
<figure><img src="image_9.png" /><figcaption>Fig. 8. (a) Zn and (b) Fe isotope ratios versus oxygen fugacity estimates relative to FMQ (Δ = log fO₂<sup>sample</sup> – log fO₂<sup>FMQ</sup>). The grey area correspond to primitive mantle values.</figcaption></figure>

<figure><img src="image_10.png" /><figcaption>Fig. 9. (a) Calculated Zn fractionation factors between ZnCO₃ in fluids and metasomatic minerals like phlogopite or amphibole, approximated to hemimorphite (Zn₄Si₂O₇(OH)₂H₂O), and solid carbonates such as hydrozincite (Zn₅(CO₃)₂(OH)₆) and smithsonite (ZnCO₃) based on ab initio calculations of Fujii et al. (2014) and Dacher et al. (2016). (b) Rayleigh distillation modelling the dissociation of a carbonate component during peridotite metasomatism.</figcaption></figure>

αCO₂-smithsonite > αCO₂-hydrozincite > αCO₂-hemimorphite. Hence, the formation of either carbonate or hydrous silicate during carbonate rich fluid metasomatism in an open system is expected to decrease the δ<sup>66</sup>Zn of peridotite. The δ<sup>66</sup>Zn evolution of peridotite was then modelled using a simple Rayleigh distillation model, according to the following equation:

$$ (2) \delta_f - \delta_i = 1000 (F^{(\alpha-1)} - 1) $$

where δf and δi are the initial and final isotopic value of the peridotite, F is the fraction of Zn that remains in the rock, ranging from 1 (unreacted) to 0 (all the Zn is lost in the fluid phase), and α is the fractionation factor

(αCO₂-minerals) at 900 °C, i.e., metamorphic peak P-T conditions of the Finero massif. The initial composition of the peridotite was set to a mantle like value (+0.1 ‰; Sossi et al., 2018). Results of this model are displayed in Fig. 9b. The δ<sup>66</sup>Zn values of Finero peridotites ranging from 0 to –0.26 ‰ can be modelled through the extraction of about 40 to 80 ‰ of the zinc. It must be noted that exchanges of Zn between peridotite and carbon rich fluids or melts in an open system has the potential to significantly affect the isotopic signature of the host peridotite as for example carbonatite produce by immiscibility processes display high Zn concentrations (up to 86 ppm, Martin et al., 2013) relative to mantle peridotites. However, such a process will have negligible impact on the δ<sup>56</sup>Fe of mantle peridotites as the solubility of Fe in either carbon bearing fluid or carbonatite melt is low (< 1 wt%; (Berkesi et al., 2020; Scambelluri et al., 2015) relative to the Fe content of mantle peridotites.

## 5.4. Conclusions and perspectives

Investigations of Balmuccia peridotites show that their geochemical signatures can be achieved through simple non-modal batch melting model. In contrast, both trace element concentrations and Zn stable isotopes of phlogopite- and amphibole-bearing peridotites in the Finero massif point toward the addition of a metasomatizing agent. Based on trace elements patterns of Finero peridotites, the parent metasomatizing agent was likely a melt carrying an excess of Th and LREE over U and HREE, respectively. The presence of supercritical fluids was ruled out due to the low-pressure record of the massif. Correlations between LREE/HREE ratios and LILE concentrations in Finero peridotites suggest an influence of a slab derived melt. This is consistent with previous trace element and radiogenic isotopic interpretations inferring the involvement of sedimentary component in the metasomatic overprint of the Finero Peridotite massif (Hartmann and Hans Wedepohl, 1993; Zanetti et al., 1999). Strangely, the light Zn isotope signature of Finero peridotites is incompatible with a direct metasomatism by slab derived melts, the latter being necessarily heavy relative to mantle peridotites. Such a contrasting signature between trace elements and Zn isotopes can be reconciled through a metasomatism by carbon rich fluids (or melts) in an open system. This process is likely to produce phlogopite or amphibole bearing peridotites with an abnormally light Zn isotope signature and the release of isotopically heavy CO₂-rich fluids.

## CRediT authorship contribution statement

**B. Debret:** Writing – original draft, Visualization, Validation, Supervision, Resources, Project administration, Methodology, Investigation, Funding acquisition, Formal analysis, Data curation, Conceptualization. **P.A. Sossi:** Writing – original draft, Methodology, Investigation, Data curation, Conceptualization. **N. Malaspina:** Writing – original draft, Conceptualization. **A. Gautier:** Writing – original draft,
Formal analysis, Data curation. **N. Mattielli:** Writing – original draft, Funding acquisition, Conceptualization. **H. O'Neill:** Writing – original draft, Data curation. **J. Villalobos-Orchard:** Writing – original draft, Formal analysis, Data curation. **F. Moynier:** Writing – original draft, Funding acquisition.

Declaration of competing interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

Acknowledgments

The authors thank H. Beunon, K. Koga and C. Nicollet for their
assistance during the Finero 2018 field trip, J. Devidal for its help during
microprobe analyses, P. Burkel, Tutu and D. Rigoussen for their precious
help during geochemical analyses. We thank Zuxing Chen and Tomo
Moroshita for critical comments on earlier version of this article and
careful editorial handling by Sonja Aulbach. This study was initiated by
an Institut de Physique du Globe de Paris Visiting position attributed to
N. Malaspina. This work was supported by the IPGP analytical platform
PARI, Region Ile-de-France SESAME Grants no. 12015908, and DIM
ACAV +, the ERC grant agreement No. 101001282 (METAL) (F.M.), the
Agence Nationale de la Recherche (ANR) CARBioNic "ANR-22-CE49-
0001-01" (B.D.) and the Fond National de la Recherche Scientifique de
Belgique (FNRS).

**Appendix A. Supplementary data**

Supplementary data to this article can be found online at https://doi.
org/10.1016/j.chemgeo.2025.122649.

Data availability

Data are available in supplementary material.

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