# The role of mantle inheritance and refertilization during breakup at magma-poor rifted margins: What can we learn from the SW-Australia margin?

Mélanie Ballay, Marc Ulrich, Gianreto Manatschal

► To cite this version:

Mélanie Ballay, Marc Ulrich, Gianreto Manatschal. The role of mantle inheritance and refertilization during breakup at magma-poor rifted margins: What can we learn from the SW-Australia margin?. Earth and Planetary Science Letters, 2025, 653, 10.1016/j.epsl.2025.119212. insu-05082253

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# The role of mantle inheritance and refertilization during breakup at magma-poor rifted margins: What can we learn from the SW-Australia margin?

Mélanie Ballay* ID, Marc Ulrich ID, Gianreto Manatschal

Institut Terre et Environnement de Strasbourg-EOST, CNRS-UMR 7063, Université de Strasbourg, 5 Rue René Descartes, 67084 Strasbourg, France

<figure><img src="image_3.png" /></figure>

## ARTICLE INFO

**Keywords:**

Magma-poor rifted margin

Refertilization

Mantle-melt interaction

Magmatic processes modeling

## ABSTRACT

The study of mantle-melt interactions at magma-poor rifted margins, accounting for about half of the world margins, is so far limited by the access to samples. Present concepts are therefore mainly based on studies of the Western Tethys and southern North Atlantic domain, and it remains unclear if they are applicable to other magma-poor rifted margins. Here, we present a study of well-preserved peridotites dredged along the south-western Australian margin (Diamantina Zone). Our petrological data show that two types of peridotites occur: spinel lherzolite, which represents the inherited subcontinental mantle (SCLM) before the separation of Australia and Antarctica, and plagioclase peridotite formed by syn-rift melt entrapment at low pressure in the plagioclase stability field during mantle exhumation. The observed mantle types as well as the mantle-melt processes identified in the Diamantina peridotites are very similar to those documented in the Alpine-Apennine and Iberia margins. Hence, the Alpine/Iberia model of magma-poor rifted margins may be globally applicable and does not depend on the SCLM inheritance. Remnants of mantle inheritance may be preserved in the refertilized mantle, suggesting a progressive transition from an inherited to a refertilized mantle during breakup. Finally, this study highlights the role of inheritance and refertilization as the main features controlling final rifting and breakup at magma-poor rifted margins.

## 1. Introduction

Understanding plate boundaries and the processes linked to their formation are among the major remaining challenges in plate tectonics. While processes at divergent plate boundaries at Mid Ocean Ridges (MOR) are relatively well constrained, those at Ocean Continent Transitions (OCTs), where lithospheric breakup occurs, are much less understood. Since the thermodynamic conditions and processes responsible for the formation of a new plate boundary likely differ from those acting at steady-state MORs, the latter cannot be considered as analogues to the former. Thus, characterizing the tectonic, magmatic, and hydrothermal processes and their feedback, and constraining the P-T-t conditions at which new divergent plate boundaries form are of primary importance in the understanding of lithospheric breakup processes. In the last two decades, several studies have demonstrated that mantle and magmatic rocks derived from OCTs record a complex history (Chazot et al., 2005; Ferrari et al., 2022; Manatschal and Müntener, 2009; Montanini et al., 2006; Müntener et al., 2010; Müntener and

Manatschal, 2006; Picazo et al., 2016; Rampone et al., 2020). However, there are only a handful of sites worldwide where OCTs are accessible to sampling. Most of these sites belong to the Alpine/Iberian system that resulted from the breakup of Pangea, with a subcontinental lithospheric mantle (SCLM) carrying a strong Variscan inheritance. The study of the Alpine/Iberian mantle rocks led to the development of an influential generic model that was applied to all magma-poor rifted margins (Fig. 1), which account for half of the world present-day rifted margins (Fig. 2A). A main highlight of the Alpine/Iberian model is that mantle inheritance plays a key role (Ferrari et al., 2022; Picazo et al., 2016). Therefore, a legitimate question is whether the Alpine/Iberia model can also be extrapolated to magma-poor margins that experienced different pre-rift histories? However, answering this question and testing the role of mantle inheritance require access to mantle rocks from OCTs with different pre-rift histories is required.

In this paper, we focus on mantle inheritance and mantle-melt reactions recorded in poorly investigated samples dredged along the Diamantina zone, located at the magma-poor southwest Australian

<footnote>* Corresponding author.</footnote>

<footnote>E-mail address: mballay@unistra.fr (M. Ballay).</footnote>
margin. We present new results from well-preserved mantle rocks that result from rifting and breakup along the Australia-Antarctica margins, overprinting a Proterozoic lithosphere. The aim of our study is twofold: we first discuss the inherited mantle characteristics and breakup-related mantle-melt processes recorded in mantle rocks exhumed along the Diamantina OCT; then, we evaluate the applicability of the Alpine/ Iberian model by comparing mantle-melt processes associated with breakup of a Proterozoic SCLM with those from the Iberia-Western Tethys Paleozoic SCLM. Our results are finally used to test whether the depletion/refertilization hypothesis controlling the magma-budget and magma-retardation required to form magma-poor rifted margins

can be globally extrapolated.

**2. Mantle-melt interactions recorded in OCTs: state of the art**

**2.1. The OCT: a recording area for breakup processes**

OCTs at magma-poor rifted margins form, by definition, the spatial and temporal transition between crustal and lithospheric breakup. They also form the limit between continental and oceanic lithospheres on either side and document the transition from tectonic extension to magmatic dominated accretion. While crustal breakup (i.e., the

<figure><img src="image_4.png" /><figcaption>**Fig. 1.** Spatial and temporal evolution of the SCLM from rifting to seafloor spreading showing a direct link between the different temporal stages and the evolution of the SCLM along cross-sections of the evolving rifted margin.</figcaption></figure>
separation of continental crust followed by exhumation of SCLM) is relatively easy to locate in space and time using seismic data (Chenin et al., 2024), lithospheric breakup, which corresponds to the onset of seafloor spreading, is more difficult to locate and requires a direct access to mantle and magmatic rocks. So far only few OCTs have been studied from a petrological point of view, among which the Iberia and its conjugate Newfoundland margins in the southern North Atlantic and the fossil remnants of the Western Tethys margins in the Alps are the best documented ones (Basch et al., 2019; Boillot et al., 1987; Borghini and Rampone, 2007; Chazot et al., 2005; Chenin et al., 2017; Ferrari et al., 2022; Manatschal and Müntener, 2009; McCarthy and Müntener, 2015;

Müntener et al., 2010; Müntener and Manatschal, 2006; Piccardo et al., 2014, 2007; Rampone et al., 2020; Rampone and Borghini, 2008; Tribuzio et al., 2004, 2016; Tucholke et al., 2004). Using a statistical approach on mantle rock data from the Western Tethys system, Picazo et al. (2016) defined three major mantle types. A mantle type 1 located close to the continent and corresponding to the SCLM sensu stricto, which may (1a) or may not (1b) have undergone depletion events pre-dating rift onset. A mantle type 2 that underwent melt percolation and trapping, forming a large domain of refertilized mantle. And finally, a mantle type 3, corresponding to a typical, depleted abyssal peridotite, formed at a steady-state MOR. Although this subdivision is of first order,

<figure><img src="image_5.png" /><figcaption>**Fig. 2.** A) Overview of the world's rifted margins (modified from Reston, 2009) and continental terranes with the age of the last major lithospheric overprint (data from Furnes et al., 2014). The yellow rectangle indicates the location of the Diamantina Zone. 90°E: Ninety-East Ridge; AAD: Australia-Antarctica Discordance; BR: Broken Ridge; CIR: Central Indian Ridge; Di: Diamantina Zone; GVL: Basins off George V Land; I: Iberia margin; KP: Kerguelen Plateau; LB: Labuan Basin; NF: Newfoundland margin; SEIR: South East Indian Ridge; SWIR: South West Indian Ridge; T: Tasmania; WT: Western Tethys. B) Bathymetric map of the Diamantina zone, South East Indian Ocean. Colored dots indicate the location of MD80 and MD110 dredges and the lithologies of dredged samples (Chatin et al., 1998; Beslier et al., 2004). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>
it nevertheless illustrates that two aspects need to be considered when describing mantle rocks from OCTs: (i) inheritance, i.e., the mantle state and its composition prior to onset of rifting, and (ii) mantle-melt reactions occurring during the breakup.

## 2.2. Inheritance vs. mantle-melt interactions during breakup: a petrological approach

Our knowledge about the processes and conditions operating during magma-poor breakup are mainly based on petrological studies that focused on three aspects: i) defining the nature and state (P/T) of mantle rocks recording the pre-rift state, i.e., the SCLM inheritance (Müntener and Manatschal, 2006); ii) characterizing the reactions between the mantle and melts percolating through the SCLM (Montanini et al., 2006; Müntener et al., 2010; Piccardo et al., 2007; Rampone et al., 1997, 2020; Rampone and Borghini, 2008), and iii) quantifying the P-T-t conditions related to exhumation during rifting (Dygert and Liang, 2015; Smye et al., 2017). These studies highlighted how partial melting, exhumation and/or refertilization processes during breakup are recorded by mantle rocks. Since the chemical composition of major and trace elements in mantle minerals is sensitive to P-T conditions, mineral chemistry is a major tool that can be used to characterize the evolution of mantle-melt reactions. For instance, Cr# in spinel is correlated to heavy rare earth elements (HREE) in co-existing clinopyroxene. Hence, it can be used directly as a proxy to quantify partial melting degree in absence of late metasomatic events (Hellebrand et al., 2001). Spinel with elevated titanium contents (TiO₂ > 0.2 wt%) is indicative of melt percolation or mafic intrusions within the peridotite, and therefore can be used to track mantle-melt reactions (Müntener and Manatschal, 2006; Picazo et al., 2016). These studies showed that mineral textures and chemical compositions of clinopyroxene and plagioclase also show contrasted variations whether peridotites are subjected to partial melting, sub-solidus re-equilibration and/or refertilization.

Although inherited SCLM is mainly formed of porphyroclastic, fertile sp-lherzolite (5–15 % clinopyroxene), harzburgites or significant proportions of spinel and garnet pyroxenites can be observed as well (up to 10%; Bodinier et al., 2008; Müntener et al., 2010). Clinopyroxene has high sodium (Na₂O ~ 1.5 wt%) and aluminum (Al₂O₃ = 5–9 wt%) contents. Plagioclase may occasionally be present in inherited SCLM at the rim of spinel grains, indicating a sub-solidus reaction of cpx + opx + sp = 2ol + pl (e.g., Borghini et al., 2010; Herzberg, 1978; Rampone et al., 1993). However, the presence of Pl-peridotite can also result from mantle refertilization by percolating melts (Kaczmarek and Müntener, 2008; Rampone et al., 1997; Seyler and Bonatti, 1997). In the latter scenario, plagioclase crystallizes together with pyroxenes and/or olivine due to porous flow migration of melts, leading to the partial replacement of the initial porphyroclastic assemblage (Kaczmarek and Müntener, 2008). A direct consequence of the plagioclase-clinopyroxene co-crystallization is the depletion of Na₂O and Al₂O₃ in pyroxenes. The trace element content of the latter is also deeply modified compared to clinopyroxene of inherited SCLM, showing light rare-earth elements (LREE) and Zr enrichments and Eu and Sr depletion (e.g., Piccardo et al., 2007; Rampone et al., 1993). Previous studies from the Western Tethys showed that up to 12% of melt can be trapped in the refertilized mantle (Müntener et al., 2010). In addition, inherited SCLM can show contrasting P-T paths. It can either equilibrate at relatively low temperature (850–950 °C; Borghini et al., 2011; Montanini et al., 2006; Müntener et al., 2010) in the spinel stability field and reach the seafloor following a cold exhumation path. Alternatively, in the infiltrated mantle domain it follows a hot exhumation path, recording a temperature increase (~1100 °C) due to melt percolation at shallow depths in the plagioclase stability field (Basch et al., 2019; Piccardo et al., 2009; Müntener et al., 2010; Rampone and Borghini, 2008).

# 3. Geological setting

## 3.1. The Australia–Antarctica conjugate margins

Fragmentation of southeastern Gondwana resulted in the separation of Australia and Antarctica and the formation of the SE Indian Ocean. This process was also responsible for the formation of a wide OCT of exhumed mantle along the SW Australian margin, referred to as the Diamantina zone (Beslier et al., 2004). This zone is assumed to represent the conjugate of the Labuan basin, located at the Antarctica margin (Fig. 2B; Munschy, 1998). Both margins show an architecture typical of a magma-poor rifted margin (Sutra et al., 2013), including segmented necking and hyperextended domains, bounded oceanward by an exhumed mantle domain. The abundant mantle rocks and sparse MORB-type magmatic rocks dredged from the Diamantina zone (Beslier et al., 2004; Munschy et al., 1994; Royer et al., 1998), as well as the basement architecture (Fig. 1) share many similarities with the Iberia-Newfoundland and fossil Alpine Tethyan margins (Alps). A major difference, however, is that Australia-Antarctica margins resulted from the rifting of a much older SCLM, including the Mesoproterozoic Fraser-Albany orogeny (Maier et al., 2016; Spaggiari et al., 2015) and the Archean 3.10 to 2.60 Ga old Yilgarn craton (Champion and Cassidy, 2007). Another important difference is the complex formation, rifting and breakup of the Broken Ridge and Kerguelen magmatic plateau, whose implication in the evolution of the study area remains debated (Magri et al., 2024).

## 3.2. The Diamantina Zone

The Diamantina Zone has been the focus site of the two oceanic cruises MD80 and MD110 that dredged the samples investigated in this study. The zone is located at the SW Australia margin, limited by the Naturaliste Plateau to the north, the Broken Ridge Plateau to the west and the Tasmania-George V Land to the east (Fig. 2A). The Diamantina zone shows E-W trending ridges ranging between 3000 and 6000 m water depth. The topography (Fig. 2B) is highly heterogeneous, undulating in a N090° to N130° direction (Munschy, 1998). To the north, seismic reflection lines show north dipping tilted blocks separated by south-dipping normal faults (Direen et al., 2007, 2013; Munschy, 1998). The Diamantina Zone is interpreted as a wide, segmented, hyperextended domain, bounded southward by a wide domain of exhumed subcontinental mantle and proto-oceanic crust (Beslier et al., 2004, Gillard et al., 2015; Sayers et al., 2001). The zone is divided into three segments separated by NW-SE trending fracture zones. The first segment, located in the western part, is limited along the Naturaliste fracture zone, interpreted as a Mesozoic transform fault (Chatin et al., 1998), from a central segment located between 111°E and 115°E. The latter is separated along the Leeuwin fracture zone from an eastern segment located between 115°E and 120°E, directly south of the Australian continent (Fig. 2B).

## 3.3. Debates and remaining questions

The rift evolution along the Diamantina zone is still debated. Recent studies suggest that the zone formed during Late Jurassic/Early Cretaceous rifting followed by mantle exhumation and early seafloor spreading that eventually failed (Gillard et al., 2015). A later event resulted in the breakup and separation of the Broken Ridge-Kerguelen magmatic plateau at about 45 Ma (Magri et al., 2024), and the formation of the present spreading system between Australia and Antarctica. While this post-45 Ma history is well described, little is known about the much older mantle exhumation stage. This is mainly due to the absence of clear magnetic anomalies, due to the Cretaceous quiet period. Out of all samples dredged during the two campaigns, only one age has been measured on a basalt from the dredge 7 of MD80 cruise (Chatin et al., 1998). It yielded an age of 93±4 Ma using the ⁴⁰Ar/³⁹Ar method on
plagioclase and bulk rock. Although the age of mantle exhumation in the Diamantine zone remains unclear, the fact that it took place during the magnetic quiet zone seems undeniable. Also the conditions at which mantle exhumation occurred are still largely unknown. Chatin et al. (1998) discussed the nature of mantle inheritance and suggested refer-tilization. However, no clear petrological and geochemical constraints have been presented. Nicholls et al. (1981) published limited petrolog-ical results from only three lherzolite samples that already showed marked variations in their mineral chemistry with notably high Mg\#. Here, we focus on a set of mantle samples that were exhumed to the seafloor during the Early Cretaceous predating breakup. Based on new petrological and mineral chemistry data, we document for the first time the conditions and processes of magma-poor mantle exhumation of a Proterozoic mantle lithosphere.

# 4. Samples and analytical methods

## 4.1. Origin of samples

The samples presented in this study were collected during the two successive MD80 and MD110 oceanic cruises in the 1990s, both oper-ated by the Marion Dufresne research vessel. These oceanic cruises were designed to constrain the mechanisms of crustal thinning and litho-spheric breakup and to explore when and how this breakup occurred. Numerous publications used the geophysical data collected during the two cruises (e.g., Bernard et al., 2005; Direen et al., 2007; Halpin et al., 2008; Magri et al., 2024; Rotstein et al., 2001), but only a few studies dealt with the dredged samples (Beslier et al., 2004; Chatin et al., 1998).

A total of 199 samples were recovered from 5 dredges during MD80 cruise, and 342 samples from 11 dredges during MD110 cruise, covering a wide area along the Diamantina zone (Fig. 2B). The dredged samples include ultramafic (peridotites, pyroxenites), magmatic (gabbros, ba-salts, diorites, granites) and metamorphic rocks (metasediments, gneiss, amphibolite), in addition to sediments, breccias and polymetallic nod-ules. Here, we focus on dredge sites where ultramafic rocks were collected, i.e., DR02, DR06 and DR09 sites from MD110 and DR03 from MD80. MD80-DR3 was dredged on a northward tilted block, west of the Naturaliste fracture zone. MD110-DR06 was located along the Leeuwin fracture zone, MD110-DR02 and DR09 were dredged further south on both sides of the Leeuwin fracture zone (Fig. 2B). All samples underwent hydrothermal alteration, with a degree of serpentinization varying from 25 to almost 100 %. However, relics of primary minerals (mainly py-roxene and spinel, less commonly olivine) were found in most of the samples presented in this study. A total of 47 samples were analyzed, including 10 samples from the MD110-DR02, 20 samples from the MD110-DR06, 10 samples from the MD110-DR09 and 7 samples from the MD80-DR3.

## 4.2. Analytical methods

Major element concentrations of minerals were measured by micro-X-ray fluorescence (μ-XRF) at the Institut Terre et Environnement de Strasbourg (ITES, France) and by electron microprobe at Service Com-mun de Microscopie Électronique et de Microanalyses (Nancy, France). Chemical maps were also obtained using μ-XRF at ITES. Trace element concentrations were analyzed using a laser ablation inductively coupled plasma mass spectrometer (LA-ICP-MS) at Georessources Nancy (France). Additional details about the analytical methods, including analytical conditions, standardization and the whole dataset are given in supplementary materials.

# 5. Results

## 5.1. Sample petrography

The nature of the dredged peridotites is largely dominated by

lherzolites of which two different types can be distinguished. A first type, widely observed (MD80-DR03 and MD110-DR02 and DR09), corresponds to Sp-lherzolites. Interestingly, numerous pyroxenites were dredged together with the Sp-lherzolites. Despite some samples showing proto-mylonitic textures (particularly for samples recovered from DR09 site), most Sp-lherzolites are unstrained, showing coarse-grained to porphyroclastic textures. Serpentinization is generally extensive (up to 95 %), but most samples keep their primary mineral fabric and contain relics of pyroxene and spinel (Fig. 3A and B). Olivine, however, was completely serpentinized. Orthopyroxene occurs as porphyroclasts, while clinopyroxene is generally smaller, both exhibiting cleavages fil-led with exsolution lamellae of the complementary pyroxene. Spinels are relatively large (up to 2 mm), show sub-euhedral shapes and are of a light brown color (Fig. 3B). Some spinels in MD80-DR3 show reaction coronas of chloritized plagioclase. It should be noted that only small grains of clinopyroxene and spinel were found in serpentinite clasts embedded in carbonates in MD80-DR3.

A second type, corresponding to Pl-lherzolites, was only found in MD110-DR06 (Fig. 3C to H). Compared to Sp-lherzolite, they are better preserved from serpentinization (30–50 % of serpentine), and thus contain more relics of primary minerals, including olivine grains pre-served in the core of serpentine mesh textures (Fig. 3D). Orthopyroxene shows porphyroclastic textures, with irregular, concave grain bound-aries filled with olivine (Fig. 3D). Clinopyroxene is small, with a rounded and lobate texture and occur interstitially in the primary porphyroclastic mineral assemblage (Fig. 3E and F). Interestingly, clinopyroxene is systemically associated with plagioclase (Fig. 4). The latter is largely replaced by chlorite (Fig. 3C-D-E-H) but preserved grains can be found (Fig. 3F), rarely surrounding spinel (Fig. 3H) but rather interstitial. Spinel displays heterogeneous shapes: most grains have irregular, holly-leaf shapes with a dark-brown color (Fig. 3H), whereas others, occurring as inclusions within orthopyroxene porphyroclast (Fig. 3G), share sim-ilarities in color and shape with spinel from Sp-lherzolites (Fig. 3B).

## 5.2. Mineral chemistry

### 5.2.1. Clinopyroxene

All clinopyroxenes have a composition corresponding to diopside, with chemical compositions that vary depending on the host rock. In Sp-lherzolite, clinopyroxene shows high Na₂O concentrations (0.8–1.8 wt %; Fig. 5A, C) with high Mg# (Mg/(Mg+Fe); 0.91 to 0.95; Fig. 5A). It also shows high Cr₂O₃ contents (0.7–2 wt%, Fig. 5B) with higher values in MD110-DR02 and lower values in MD80-DR3, and Al₂O₃ ranging from 4 to 8 wt% with homogenous values at the sample scale. No sig-nificant core-to-rim zonation is observed. Trace element patterns of clinopyroxene from Sp-lherzolite show distinct shapes depending on their respective dredge sites (Fig. 6). At MD110-DR02 and DR09, cli-nopyroxene is depleted in incompatible elements with Pb, Sr and Zr-Hf negative anomalies. It is characterized by LREE depletion ((La/Sm)N~0.1) and show flat HREE patterns. Clinopyroxenes from site MD80 are less depleted in LREE without any fractionation between MREE and HREE.

Clinopyroxene from Pl-lherzolites is characterized by lower Na₂O (<0.7 wt%; Fig. 5A, C) and Mg# (mean at 0.909±0.02) relative to Sp-lherzolites. Al₂O₃ compositions show a similar range (3 - 8 wt%; Fig. 5B, D) to that of clinopyroxene of Sp-lherzolites. A significant dif-ference exists, however, between porphyroclasts and neoblasts. The chemical maps show that porphyroclasts exhibit core-to-rim zonation (Fig. 4) with high Al₂O₃ concentrations at the core (up to 8 wt%, Fig. 5D) and lower at the rim (3 wt%). The opposite is observed for Cr₂O₃ con-centrations, leading to higher Cr# at the clinopyroxene rims than in the cores (0.2 vs. 0.1, respectively; Fig. 4H). Trace element patterns are also depleted in incompatible elements and show deeper Sr and Pb negative anomalies compared to clinopyroxene in Sp-lherzolites. REE patterns are slightly less depleted in LREE with humped M- to HREE patterns ((Gd/Yb)N>1) and a negative Eu anomaly (Eu/Eu* ~0.89).
<figure><img src="image_6.png" /></figure>

(caption on next page)
**Fig. 3.** Microphotographs of A) a preserved clinopyroxene (Cpx) porphyroclast (cross-polarized) (MD110-DR02-A27) and B) typical brown colored spinel (plan polarized) in highly serpentinized Sp-lherzolite (MD110-DR02-A66). C) Peritectic plagioclase (Pl) – orthopyroxene (Opx) assemblage replacing a primary clino-pyroxene popyroclast (Cpx I) in Pl-lherzolite (MD110-DR06-171). D) Orthopyroxene porphyroclast with corroded rims (holly-leaf texture, red arrows) with corrosion embayments filled by olivine (Ol) and plagioclase (here altered into chlorite, 'Pl'; MD110-DR06-161A). E) Secondary clinopyroxene neoblast (Cpx II) in assemblage with plagioclase ('Pl') near an amoeboid spinel (MD110-DR06-161A). F) Original, former clinopyroxene porphyroclast (Cpx I) consumed during mantle-melt reaction, forming typical orthopyroxene-plagioclase assemblage (MD110-DR06-162B). G) Spinel in inclusion within orthopyroxene in Pl-lherzolite, similar in shape to spinel found in Sp-lherzolite (MD110-DR06-161A). H) Typical dark-colored Cr-rich spinel in Pl-lherzolite, with plagioclase coronae (dashed line) suggesting sub-solidus spinel-to plagioclase transition (MD110-DR06-161A).(For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

### 5.2.2. Orthopyroxene

Orthopyroxene in Sp-lherzolite has high Mg# (ranging from 0.90 to 0.95 – mean at 0.92). Al₂O₃ and CaO contents vary from 3 to 6 wt% and 0.4 to 3 wt% respectively, and major element profiles do not show any core-to-rim zonation. Chondrite-normalized REE patterns show strong fractionation between L- and HREE with a strong depletion in LREE (usually below the limit of detection).

Orthopyroxene from Pl-lherzolites has lower Mg# (0.89 to 0.92) and strong core-to-rim zonations in Al₂O₃ and Cr₂O₃. Clinopyroxene porphyroclasts show similar trends (Fig. 5). Cores contain ~5 wt% Al₂O₃ while rims have Al₂O₃ concentrations <2.5 wt%. Cr₂O₃ concentrations show core-to-rim enrichment from ~0.6 to 0.8 wt%, leading to the same Cr# range than that reported in clinopyroxene porphyroclasts. SiO₂ and MgO also show increasing concentrations from core to rim. The REE concentrations are slightly higher in Pl-lherzolite orthopyroxene relative to those in Sp-lherzolite, with a similar strong L- to HREE fractionation. REE concentrations also show zonation, with a typical LREE depletion at the crystal rim relative to the core (see supplementary materials).

### 5.2.3. Spinel

Spinels show contrasting compositions between Sp- and Pl-lherzolites. Spinels from Sp-lherzolite have high Mg# (0.74–0.85) and low Cr# (0.09–0.26) (Fig. 7A and B). They also show low TiO₂ contents (<0.15 wt%). The degree of partial melting (F) calculated based on spinel Cr# (Hellebrand et al., 2001) shows that peridotites at MD110-DR02, DR09 and MD80-DR03 underwent relatively low degrees of melting (F<10%; Fig. 7C). In contrast, spinels from Pl-lherzolite have lower Mg# (0.52 to 0.79), higher Cr# (0.19 to 0.56) and high TiO₂ concentrations (>0.2 wt%, up to 0.79 wt%). Interestingly, some spinels found in inclusions within orthopyroxenes show compositions like those of spinels from Sp-lherzolite (Cr# 0.128 to 0.175; Mg# 0.77 to 0.85; TiO₂<0.15 wt%) and similar calculated F of 3–7% (Fig. 7C).

### 5.2.4. Olivine

Relics of olivine were only found in Pl-lherzolite (Figs. 3 and 4), while in Sp-lherzolite, olivine was completely altered into serpentine. Olivine shows high MgO (48.0–51.4 wt%) and low FeO (6.2–9.6 wt%), resulting in a relatively high Mg# in the range of 0.90–0.93. NiO abundances vary from 0.23 to 0.46 wt%. Trace element concentrations were almost systematically below the limit of detection.

### 5.2.5. Plagioclase

Fresh plagioclase only occurs in 2 samples (MD110-DR06-162B and MD110DR06-168). In other samples, plagioclase is replaced by chlorite (Fig. 3). Plagioclase has compositions corresponding to labradorite to bytownite (An71 ± 8). REE patterns show LREE enrichment (up to 2 times chondrite concentrations) and a strong depletion in HREE (< 0.1 times chondrite, heaviest REE being systematically < LOD; see supplementary figure). They also show a prominent positive Eu anomaly (Eu/Eu* ~ 7.8).

## 5.3. Geothermobarometry

To estimate the equilibrium temperature and pressure for the various peridotites from the Diamantina zone, mineral compositions were used to calculate several geothermobarometric estimates (e.g., Ca-in-

orthopyroxene (T<sub>Ca-in-opx</sub>) and 2 pyroxenes (T<sub>BKN</sub>; Brey and Köhler, 1990)). The only barometer available so far for spinel peridotite uses calcium concentration in olivine (D'Souza et al., 2019). However, no olivine was preserved in Sp-lherzolites, so a pressure of 15 kbars was assumed for temperature calculation. Temperatures (T<sub>BKN</sub>) were obtained for MD110-DR02 (944 ± 53 °C), MD110-DR09 (943 ± 85 °C) and MD80-DR3 (944±28 °C) respectively. Significantly higher temperatures were calculated using T<sub>Ca-in-opx</sub> (1230 ± 87 °C and 1139 ± 127 °C, 1238 ± 124 °C respectively). This may reflect the presence of clinopyroxene exsolution lamellae under the spot analysis of EPMA and µ-XRF. These higher temperatures may correspond to those reached prior to clinopyroxene exsolution (Cherniak and Liang, 2022; Jollands and Müntener, 2019;). For Pl-lherzolite, we combined thermometers with the FACE (Fumagalli et al., 2017) and the anorthite-based barometers (Borghini et al., 2010, 2011). We found a pressure of 7 kbar and a slightly higher T<sub>BKN</sub> of 1013 ± 46 °C compared to Sp-lherzolites. T<sub>Ca-in-opx</sub> for Pl-lherzolite are also more elevated (1178 ± 54 °C).

# 6. Discussion

## 6.1. Exhumed mantle peridotites along the Diamantina Zone: what do they record?

### 6.1.1. Two mantle types

Dredged mantle samples from the Diamantina Zone can be subdivided into two groups: Sp-lherzolite and Pl-lherzolite (Fig. 3). While Sp-lherzolites are widespread and occur at sites MD110-DR02, DR09 and MD80-DR3, Pl-lherzolites have only been found at MD110-DR06, located along the Leeuwin fracture zone (Fig. 2B). Considering the stability domains of spinel and plagioclase, the occurrence of two types of peridotites indicates that these rocks were formed at two distinct depths. Many pyroxenites were dredged with Sp-lherzolites at sites MD110-DR02 and DR09 but were very scarce at site MD110-DR06. At this site peridotites were almost exclusively Pl-lherzolite. This suggests that we are not only dealing with two different types of peridotite, but that we are also dealing with two different domains of the mantle.

### 6.1.2. Sp-lherzolite: a subcontinental lithospheric mantle inheritance

Sp-lherzolites found at the various dredging sites in the Diamantina Zone contain Na₂O-rich clinopyroxene (between 0.8 and 1.8 wt%) with high Mg# (> 0.90; Fig. 5) and spinel with low Cr# and high Mg# (Fig. 7). All these features suggest a sub-continental mantle origin. This result is consistent with geophysical modelling (Laat et al., 2023) and a geochemical composition assumed for SCLM of Mesoproterozoic age (Maier et al., 2016). The high Mg# (Mg# > 0.74) and low Cr# in spinel are indicative of a fertile composition of Sp-lherzolite. It also implies that the latter underwent relatively low degrees of partial melting (<10%) of unknown age (Fig. 7). Additional information about the melting conditions recorded by Sp-lherzolites can be obtained from the chemical composition of clinopyroxene. Fig. 8 shows the concentrations of Nd<sub>N</sub> vs. Na₂O in clinopyroxene. These two elements are of particular interest since both are sensitive to the degree of partial melting and have approximately the same incompatibility at low pressure ($D_{Na}^{Cpx/l} \approx D_{Nd}^{Cpx/l} \sim 0.2$; Blundy et al., 1995). However, at high pressure in the garnet stability field, Na becomes less incompatible than Nd
<figure><img src="image_7.png" /><figcaption>**Fig. 4.** A) Cross-polarized microphotographs of a Pl-lherzolite sample (MD110-DR06-161c). B) Phase map showing mineralogical distribution of A). C) Al₂O₃, D) Cr₂O₃ and E) Cr# compositions of orthopyroxene in Pl-lherzolite showing core-to-rim zonations. F), G), H) Same as C), D), E), respectively, for clinopyroxene. Red arrows in F) show unzoned Cpx neoblasts, while green arrows locate zoned, primary Cpx porphyroclasts. The phase and concentration maps were calculated based on µ-XRF chemical maps applying the approach of Ulrich et al. (2014). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>
<figure><img src="image_8.png" /><figcaption>**Fig. 5.** Major element compositions of clinopyroxene in Sp- and Pl-peridotites from the Diamantina zone. A) Na₂O (wt%) vs. Mg# B) Al₂O₃ (wt%) vs. Cr₂O₃ (wt%) C) TiO₂ (wt%) vs. Na₂O (wt%) D) Kernel density distribution of Al₂O₃ concentrations in clinopyroxene from Pl-lherzolites, showing that neoblasts are depleted in Al₂O₃ relative to primary porphyroclasts. Data from Australia-Antarctica (blue - Beyer, 2002; McCarthy et al., 2020; Delpech et al., 2021) and other domains (grey - passive margins compilation from Picazo et al., 2016; Müntener et al. 2004; 2010; Müntener and Manatschal, 2006; Rampone et al., 1996; and Lena Through peridotites from Hellebrand and Snow, 2003) are shown for comparison.</figcaption></figure>

($D_{Na}^{Cpx/l}$ ~0.5; Blundy et al., 1995). Applying the modeling approach of Hellebrand et al. (2002) and Müntener et al. (2010), we calculated the theoretical evolution of the Na₂O and Nd composition of clinopyroxene during the partial melting of a lherzolite. Na₂O concentrations of clinopyroxene in most Sp-lherzolites are higher than the depleted mantle (DM) clinopyroxene (0.89 wt%; Workman and Hart, 2005). Therefore, a primitive mantle-like composition was assumed for the starting material (McDonough and Sun, 1995). Two melting models were calculated assuming melting in the spinel or garnet stability field (curves A and B respectively, Fig. 8A). A third “polybaric” model was developed with 2, 4 and 6% of melting of a garnet peridotite followed by conversion of modal assemblages into spinel peridotite (Johnson et al., 1990) and additional melting in the spinel mantle field (curves C₂, C₄, C₆, Fig. 8A).

This latter model fits most of the data and highlights that Sp-peridotites along the Diamantina zone underwent partial melting degree that rarely exceeded 10 %. This result is in good agreement with estimates based on Cr# in spinel (Fig. 7). More specifically, our model shows that MD110-DR02 and some MD80DR3 Sp-lherzolites are residual after 2–4 % melting in the garnet stability field, followed by melting in the spinel stability field. These estimates are very similar to those previously made for the peridotites of the Central Alps (Müntener et al., 2010).

Interestingly, other samples from MD80-DR3 and MD110-DR09 are better explained by melting occurring almost exclusively in the garnet stability field (up to 8–9 %), with only low melting (<2 %) in the spinel mantle domain, making them more comparable to the Lena Trough peridotites described by Hellebrand and Snow (2003). However, this
<figure><img src="image_9.png" /><figcaption>**Fig. 6.** Rare earth and trace element patterns of pyroxenes in Sp- and Pl-lherzolites from the Diamantina zone. REE are normalized to C1 chondrites (Anders and Grevesse, 1989) and trace elements are normalized to the Primitive Mantle (McDonough and Sun, 1995).</figcaption></figure>

model failed at reproducing the low Na₂O/high NdN concentrations of MD110-DR06 Pl-lherzolites, which suggests that they derive from another magmatic process.

### 6.1.3. Pl-lherzolite: a tape recorder of mantle-melt reactions during exhumation

The Pl-lherzolites show very different textures from those observed in the Sp-lherzolites. These include corroded orthopyroxene with corrosion embayment filled by olivine (Fig. 3D), pre-existing clinopyroxene porphyroclasts partially replaced by orthopyroxene and plagioclase assemblage (Fig. 3C), and interstitial clinopyroxene neoblasts in close association with interstitial plagioclase (Fig. 3E). All these textures evidence the occurrence of melt impregnation rather than sub-solidus reactions (e.g., Kaczmarek and Müntener, 2008; Müntener et al., 2010; Rampone and Borghini, 2008). Olivine embayment on pyroxenes (Fig. 3D) are characteristic of reactive porous flow, occurring at spinel facies and known to crystallize olivine at the expense of pyroxene (Basch et al., 2020; Piccardo et al., 2007; Rampone et al., 2020). This process is distinct from the refertilization process, which consists of impregnation by percolating melts at the plagioclase facies (Fig. 3E, F).

Mantle-melt reactions are also indicated by mineral chemistry such as high TiO₂ concentrations in spinel and slight LREE enrichment and “humped shaped” middle to heavy REE patterns of clinopyroxene. The difference between Cr# of clinopyroxene and Cr# of spinel (Fig. 9) is also in favor of melt-rock interaction rather than subsolidus reequilibration (Müntener and Manatschal, 2006). Na₂O and Al₂O₃ depletion and Sr and Eu negative anomalies in clinopyroxene are indicative of co-crystallization with plagioclase, consistently to that was reported in Pl-lherzolites from the Alpine-Apennine ophiolites (Basch et al., 2019; Borghini and Rampone, 2007; Kaczmarek and Müntener, 2010; Müntener et al., 2004; Rampone et al., 1995). The refertilization assumption was tested in Fig. 8B by simple mass balance modelling the addition of N-MORB-type melts (1 % increment) in a depleted peridotite at shallow depth in the plagioclase stability field. The Na₂O and Nd compositions of residual peridotite were calculated after 4 % of melting in the garnet stability field, followed by 4% additional melting under spinel facies conditions. The elemental zonation reported in orthopyroxene from Pl-lherzolites, characterized by core-to-rim Al₂O₃ and CaO depletion and MgO enrichment (Fig. 4), indicates that this mineral experienced a transition from spinel-to-plagioclase facies (Borghini et al., 2010; Smye
<figure><img src="image_10.png" /><figcaption>**Fig. 7.** Major element compositions of spinels from the Diamantina zone. A) Cr# vs. TiO₂ (wt%); B) Cr# vs. Mg#; C) Comparison between partial melting degrees recorded by peridotites exhumed along the Diamantina zone with data from other margins worldwide. The partial melting degrees were calculated as a function of Cr# of spinels, applying the following equation: F (%) = 10 ln(Cr#) + 24, valid for Cr# values ranging from 0.1 to 0.6 (Hellebrand et al., 2001). Only spinels with TiO₂ < 0.15 wt% were considered, preventing effects of melt percolation on the chemical composition of spinel. Data from literature are the same as for Fig. 5.</figcaption></figure>

et al., 2017; Witt-Eickschen and Seck, 1991). This transition was considered in our model by the conversion of spinel into plagioclase (completely converted - light red star; 50% - dark red star) according to the following equation: Sp + Cpx + Opx = 2 Ol + Pl (Rampone et al., 1993). The crystallization modes during refertilization were assumed to be Zcpx/Zplg = 0.09/0.04 (Müntener et al., 2010). Our results show that small fractions of melt entrapment and crystallization lead to a steeper evolution of the Na₂O/NdN ratio in clinopyroxene compared to melting trends.

This evolution matches well with the composition of clinopyroxene in Pl-lherzolites. The amount of melt (up to 5%) stored in Pl-lherzolites of the Diamantina zone is also consistent with previous estimates calculated for impregnated peridotites from the Alpine-Apennine system or the Iberia margin (Abe, 2001; Müntener et al., 2010; Rampone et al., 2020). It should be noted that Müntener et al. (2010) performed additional tests about the nature of the refertilized agent (from depleted to

enriched melts). Their conclusions were that the nature of melts does not significantly affect the refertilization trends but only controls the amount of refertilization.

The refertilized domain could thus be considered as a hybrid domain where inherited SCLM is modified by a later melt percolation. Our calculations show that the chemistry of Pl-lherzolites is not controlled by the nature of the refertilizing agent but rather depends on the mantle chemical composition prior to refertilization. This result demonstrates that SCLM inheritance has a considerable incidence on the final chemistry of the refertilized mantle (especially the type of SCLM – that means T1a or T1b, which defines the fertility of the inherited mantle) but does not control the processes involved in the formation of magma-poor rifted margins.

Finally, the partial melting models developed in this study demonstrate that, when the starting point is aligned with the regional geological history (DMM source or primitive mantle), the same
<figure><img src="image_11.png" /><figcaption>**Fig. 8.** Na₂O vs. NdN concentrations in clinopyroxene of Diamantina zone peridotites. **A)** Melting modeling of clinopyroxene of a primitive-like mantle source starting at garnet [A] and spinel [B] stability fields. [C₂], [C₄] and [C₆] dashed blue curves simulate a "polybaric" melting with 2 %, 4 % and 6 % of melting in the garnet stability field followed by additional melting in the Sp-peridotite field. **B)** refertilization modeling of a depleted source after 4 % of melting in the garnet stability field, followed by 4 % melt extraction under spinel facies conditions and sub-solidus re-equilibration in the plagioclase stability field (the light red star indicates full Sp-to-Pl replacement while the dark red star considers only 50 % of Sp is converted into Pl, using the univariant reaction of Rampone et al., 1993). The refertilization agent is a N-MORB-type melt, black dots along the refertilization curves indicate 1 % melt increment. Melting and refertilization models are modified from Hellebrand and Snow (2003) and Müntener et al. (2010). Additional parameters are given in the text. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

<figure><img src="image_12.png" /><figcaption>**Fig. 9.** Cr#<sub>Cpx</sub> vs. Cr#<sub>Sp</sub> of the peridotites from the Diamantina Zone. Sp-peridotites from Lena Trough are from Hellebrand and Snow (2003); Sp- and Pl-lherzolites from Eastern Central Alps are from Müntener et al. (2004; 2010).</figcaption></figure>

processes are involved, i.e., an initial partial melting phase in a garnet-facies domain, then a transition to the spinel-facies domain and finally partial melting in the spinel-facies domain. In the same way, regardless of the chemistry of the refertilizing liquid or the crystallization modes, the chemistry of pyroxenes will be modified (in terms of REE, Na₂O, Al₂O₃ and Cr₂O₃ in particular) but not the process itself, which always takes place under similar PT conditions.

## 6.2. The Diamantina mantle rocks: how do they fit the alpine-iberia model?

The research in mantle rocks from OCTs is almost entirely based on studies from samples from the Iberia-Newfoundland and Western Tethys systems. These studies highlighted the heterogeneous composition of the SCLM (Picazo et al., 2016; Vauchez et al., 2000) and the importance of refertilization during lithospheric breakup (Müntener et al., 2004), which we refer to as the so-called Alpine-Iberia model. This model proposed that the inherited SCLM corresponding to mantle T1, was progressively refertilized and modified, giving rise to mantle T2. However, a direct genetic link between T1 and T2 is not so common in petrographic observations. It is also important to note, that the Alpine-Iberia model is based on a relatively regional data set and has not been constrained elsewhere in domains exhibiting a different and much older pre-rift inheritance. In this context, the mantle rocks derived from the Diamantina zone are of a particular importance, since they derive from the Mesoproterozoic Fraser orogenic domain, which separates the margin from the Archean Yilgarn craton. Thus, the subcontinental mantle exhumed along the SW Australia margin is necessarily older than the one from the Iberia-Newfoundland and Alpine domains and consequently their inheritance is likely different. Hence, a major question is whether the mantle rocks from the Diamantina zone provide new observations that either reject, support, or complete the Alpine-Iberia model.

### 6.2.1. Do mantle-melt interactions depend on the type of inherited mantle?

Picazo et al. (2016) compiled a large amount of petrological and geochemical data from the Iberia and Western Tethys systems, which allowed to map the distribution of mantle T1 and T2. Along the Diamantina Zone, the same mantle subdivision occurs. In both the Alps and the Diamantina examples, pyroxenites are commonly observed within the inherited SCLM (Müntener et al., 2004) but are rarely observed in the refertilized domains (Pl-lherzolite). Na₂O concentrations of clinopyroxene from the inherited mantle in the Diamantina Zone are in line
with those reported from the Alps. In contrast, CaO content is particu-
larly high compared to the Alpine equivalents. Partial melting estimates,
based on Cr# of spinel, show degrees < 11 %, which is much lower than
those reported from the Alps (8–16 %) or Newfoundland (up to 17 %),
but still higher than those from Iberia (< 3 %). In the refertilized
domain, clinopyroxene has lower Na₂O and Al₂O₃, and higher Cr₂O₃
contents compared to those from the Alps and Iberia. When comparing
the chemistry of these clinopyroxenes as a function of their domain in
the OCT notable differences can be seen in terms of Al₂O₃ and Cr₂O₃,
which are interpreted as variations of degrees of refertilization. This
may be the result of a continuous process that converts the inherited
mantle, with its intrinsic variability, into increasingly homogeneous
clinopyroxene in the refertilized mantle (Picazo et al., 2016). Thus,
variations in major elements in clinopyroxene from MD110-DR06
peridotite remains preserved only in some porphyroclasts. This is
particularly true for some elements (i.e., Al₂O₃). This may be attributed
to various degrees of melt infiltration and the related thermal evolution
(Borghini et al., 2023). However, the lack of variation in Na₂O con-
centration, which remains constant at low values, could be the sign of a
previous high partial melting rate. This observation may be character-
istic of the mantle type 1b, i.e., an inherited depleted subcontinental
mantle.

### 6.2.2. Can a direct genetic link between inherited and refertilized mantle type be proven?

In the Alpine-Iberia model, the link between the inherited sub-
continental mantle (T1) and the refertilized mantle (T2) has been
assumed (Muntener et al., 2010; Picazo et al., 2016), but it has never
been clearly demonstrated. Interestingly, some spinels are found
embedded in orthopyroxene porphyroclasts with contrasting color and
texture relative to darker spinels with holly-leaf textures that domi-
nantly occur in Pl-lherzolites (Fig. 3). Their high Mg# (up to 0.86), low
Cr# (<0.2) and TiO₂ (<0.2 wt%) differ from those affected by melt
percolations (Mg# <0.73; Cr# >0.28; TiO₂ >0.2 wt%). However, they
share similarities in texture and chemistry with spinel of Sp-lherzolites
(Fig. 7). Moreover, similar estimates of partial melting degree were
obtained from these spinels (5–7 %) like those from the Sp-lherzolites.
These embedded spinels could have been preserved from melting and
refertilization events and could thus represent the missing link between
Sp-lherzolite (mantle T1a) and Pl-lherzolite (mantle T2). Thus, the
Diamantina peridotites do not only follow the predictions of the
Alps-Iberia model, but they also complete the model by showing that
these spinels might be the missing link between the inherited SCLM and
its refertilization during exhumation.

### 6.3. The Iberia-Alpine mantle-melt interaction model: what did we learn from the SW-Australia margin?

About half of the world rifted margins show seismic characteristics of
magma-poor rifted margins (Fig. 2A; Reston, 2009). However, at present
only samples from a very small portion of these margins have been
sampled. This limits our understanding of the processes controlling the
formation of magma-poor systems and leaves questions such as how, and
to what extent inheritance of the SCLM can control the magmatic budget
during breakup difficult to answer. It remains also questionable if the
mantle-melt interaction processes described in the Alpine-Iberian model
can be considered as globally applicable to explain magma starvation or
delay of melt extrusion at all magma-poor rifted margins? Our results
show that mantle-melt processes at the Diamantina Zone, despite very
different age and inheritance of the SCLM, show similar processes during
breakup and the same mantle subdivision (T1 and T2) along the OCT.
These processes include the refertilization of the inherited SCLM, a
process that leads to the delay of melt extrusion at the seafloor. In this
context, the nature of the inherited mantle, particularly the lower SCLM,
may be of importance. Studies in the Iberia-Newfoundland and Western
Tethys systems show that mantle type 1 can be divided into two

subtypes. A fertile SCLM, referred to as mantle type 1a that is distinct
from a depleted mantle referred to as type 1b that has undergone partial
melting events occurring before the onset of rifting, probably during a
previous post-orogenic or subduction event (Picazo et al., 2016). The
presence of mantle T1b is fundamental since it may explain why these
margins evolved as so-called magma-poor margins. Indeed, the presence
of a depleted lower lithospheric mantle may explain why melts were not
directly extracted at the seafloor but were first stored within the deeper
and depleted SCLM. Thus, the occurrence of mantle T1b at other
magma-poor margins could provide further support for the idea that the
depleted nature of the lower SCLM could be a fundamental feature in the
formation of magma-poor margins. In this frame, our study shows that
the primary minerals found in Pl-lherzolites show evidence of an older
partial melting event, similar to the T1b-type mantle described by Picazo
et al. (2016). Our study further supports the idea that the presence of a
depleted lower SCLM may be at the origin of retardation of
magma-emplacement at the seafloor, favoring mantle exhumation
(Chenin et al., 2024). Thus, we propose that the occurrence of a depleted
lower SCLM may constitute a major control on the formation of
magma-poor rifted margins.

## 7. Conclusion

The petrology of the peridotites exhumed along the Australia-
Antarctica ocean continent transition zone demonstrates the occur-
rence of two mantle domains along the SW Australia OCT. Similarly to
what has been described earlier from the Iberia and Western Tethys
margins, Sp-lherzolites correspond to an inherited sub-continental
mantle while Pl-lherzolites are formed during melt infiltration at low
pressure in the plagioclase stability field during exhumation of the
SCLM. Based on an older heritage, the samples of the Diamantina zone
enable us to show that the Alpine-Iberia model of magma-poor rifted
margins is more broadly applicable. Traces of this inheritance are pre-
served in the refertilized mantle with relicts characterized by distinctive
textures and chemistry, demonstrating the link between mantle type 1b
(lower part of the SCLM that experienced an older depleted mantle) and
2. In addition, we show that the inherited nature of the SCLM could
control the formation of magma-poor margins. Indeed, melt storage in
the depleted lower SCLM during mantle exhumation may delay magma
extraction at the seafloor and may explain the formation of magma-poor
rifted margins.

### CRediT authorship contribution statement

**Mélanie Ballay:** Writing – review & editing, Writing – original draft, Methodology, Investigation, Formal analysis, Data curation, Conceptualization. **Marc Ulrich:** Writing – review & editing, Writing – original draft, Visualization, Validation, Supervision, Software, Resources, Project administration, Methodology, Investigation, Funding acquisition, Formal analysis, Data curation, Conceptualization. **Gianreto Manatschal:** Writing – review & editing, Writing – original draft, Visualization, Validation, Supervision, Software, Resources, Project administration, Methodology, Investigation, Funding acquisition, Formal analysis, Data curation, Conceptualization.

### 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

Marie Odile Beslier and Jacques Girardeau are gratefully acknowl-
edged for giving us full access to the samples from the MD110 oceano-
graphic campaign. The authors warmly thank John Moine for his help
during sample preparations and Meredith Morin for her preliminary work on MD110 samples during her Master internship. Thin section production and µ-XRF analyses were performed at the CoRTecS PeMCI plateform of ITES, University of Strasbourg. Chantal Pfeiffert is warmly thanks for her help during LA-ICP-MS measurements and Olivier Rouer for his help during EPMA measurements at Georesources Nancy. This study benefited of the financial support of the French Ministry of Higher Education and Research, and the M6 consortium (GM).

## Supplementary materials

Supplementary material associated with this article can be found, in the online version, at doi:10.1016/j.epsl.2025.119212.

## Data availability

Is in supplementary.

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