# Eruptive activity on the western flank of Piton de la Fournaise (La Réunion Island, Indian Ocean): insights on magma transfer, storage and evolution at an oceanic volcanic island

G Boudoire, Yann-Aurélien Brugier, A. Di Muro, G. Wörner, I. Arienzo, N.

Metrich, V. Zanon, N Braukmüller, A. Kronz, Y. Le Moigne, et al.

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## ► To cite this version:

G Boudoire, Yann-Aurélien Brugier, A. Di Muro, G. Wörner, I. Arienzo, et al.. Eruptive activity on the western flank of Piton de la Fournaise (La Réunion Island, Indian Ocean): insights on magma transfer, storage and evolution at an oceanic volcanic island. Journal of Petrology, 2019, 60 (9), pp.1717-1752. 10.1093/petrology/egz045. insu-02407731

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# Eruptive Activity on the Western Flank of Piton de la Fournaise (La Réunion Island, Indian Ocean): Insights on Magma Transfer, Storage and Evolution at an Oceanic Volcanic Island

G. Boudoire<fnref n="1" />,²*, Y.-A. Brugier³,⁴,⁵, A. Di Muro²,⁶, G. Wo¨ rner⁷, I. Arienzo⁸, N. Metrich⁶, V. Zanon⁶,⁹, N. Braukmu¨ ller D¹⁰, A. Kronz⁷, Y. Le Moigne³,⁴,⁵ and L. Michon<fnref n="1" />

Received October 17, 2017; Accepted October 14, 2019

## ABSTRACT

Petrological and geochemical (major element, trace element, Sr–Nd isotope) data for recent (<5 kyr old) basalts that sporadically erupt on the western flank of Piton de la Fournaise (PdF), one of the most active volcanoes on Earth, allow the tracking of magma transfer and evolution from mantle to crustal depths. In the western peripheral area of PdF we document the broadly synchronous erup-tions of (1) primitive olivine and olivine–clinopyroxene transitional basalts with tholeiitic affinity that are closely associated in space with (2) transitional olivine basalts with alkaline affinity, and (3) hybrid lavas, intermediate between the ‘alkaline’ and the ‘tholeiitic’ end-members. The compos-ition of the latter overlaps with that of the lavas frequently erupted from the conduit system feeding the main summit cone. AlphaMELTS modelling, and fluid inclusion and clinopyroxene barometry, constrain the conditions of magma storage at 10–30 km, and the ascent of magma from the upper mantle to the shallow crustal plumbing system. Variable degrees of mantle melting, together with minor source heterogeneity and contamination with cumulate-derived partial melts, contribute to the diversity of PdF magmas. However, all these processes do not represent the dominant factors that produce the large variability we found in major element composition. Indeed, the composition of basalts erupted from PdF peripheral centers is strongly controlled by polybaric olivine–clinopyr-oxene fractionation at pressures higher than 3 kbar. Crystal textures and geochemical modelling suggest that fast magma ascent is critical to prevent clinopyroxene dissolution. Conversely, long-lasting magma stagnation promotes pyroxene resorption and magma differentiation. ‘Central’ eruptions occurring close to the PdF summit cone emit variably more evolved melts, which result from olivine–clinopyroxene–plagioclase differentiation at intermediate–shallow pressure (<3 kbar

<footnote n="1">¹Laboratoire Géosciences Réunion, Université de La Réunion, Institut de Physique du Globe de Paris (IPGP), Sorbonne Paris-Cite, UMR 7154 CNRS, Saint-Denis F-97715, France; ²Observatoire Volcanologique du Piton de la Fournaise (OVPF), Institut de Physique du Globe de Paris (IPGP), Sorbonne Paris-Cite, UMR 7154 CNRS, Université Paris Diderot, Bourg Murat, France; ³Institut des Sciences de la Terre d'Orleans (ISTO): UMR 7327, Université d'Orléans, 45071 Orléans, France; ⁴ISTO, UMR 7327, CNRS, 45071 Orléans, France; ⁵ISTO, UMR 7327, BRGM, BP 36009 Orléans, France; ⁶Institut de Physique du Globe de Paris (IPGP), Sorbonne Paris-Cite, UMR 7154 CNRS, Université Paris Diderot, Paris F-75005, France; ⁷GZG, Abt. Geochemie, Université t Go¨ ttingen, Goldschmidtstr. 1, Go¨ ttingen 37077, Germany; ⁸Istituto Nazionale di Geofisica e Vulcanologia (INGV), Osservatorio Vesuviano, Via Diocleziano 328, Napoli 80124, Italy; ⁹Instituto de Investigación,–ao em Vulcanologia e Avaliação,–ao de Riscos, Universidade dos Açores, Rua M-ae de Deus, Ponta Delgada 9500-321, Portugal; ¹⁰Instituto für Geologie und Mineralogie, Universitä t zu Ko¨ ln, Zu¨ lpicher Str. 49b, Ko¨ ln 50674, Germany</footnote>

<footnote>*Corresponding author. Present address: Observatoire Volcanologique du Piton de la Fournaise (OVPF/IPGP), Impasse de l'Observatoire, Bourg Murat 97418, France. Telephone: þ262 (0)2 62275292. Fax: þ262 (0)2 62591204. E-mail: guillaume.boudoire@gmail.com</footnote>
and in most cases <1 kbar). Deep and extensive magma mixing before injection into the crustal magma conduit system, located below the summit region, results in the apparent homogeneity of basalts erupted from the central area. As regards 'peripheral' eruptions, deep-seated stagnation of basaltic melts and differentiation at the mantle–crust transition zone (c. 4 kbar) produces a range of magma compositions. We demonstrate that rapid magma ascent from deep-seated reservoirs can bypass the central plumbing system. The eruptions of these magmas both in the central area and on the densely populated flanks have major consequences in terms of volcanic hazard at PdF.

Key words: transitional basalts; clinopyroxene crystallization and resorption; source heterogeneities; barometry; mantle–crust underplating

INTRODUCTION

A better knowledge of the conditions of deep magma storage systems below active volcanoes, as well as of the dynamics of magma migration towards both shal-low reservoirs and the surface, is critical for under-standing and forecasting volcanic activity. The physico-chemical evolution of magmas in the plumbing system of large, dominantly basaltic volcanoes occurs in a highly dynamic environment, potentially involving mul-tiple storage zones of different sizes and depths and a plexus of connecting magmatic pathways (Eaton & Murata, 1960; Ryan, 1988; Tilling & Dvorak, 1993; Pietruszka et al., 2015; Schipper et al., 2015).

On large volcanic edifices, central and peripheral vents can be active contemporaneously and may be fed from the following: (1) a central magma reservoir that drains at variable depths from radial dykes [see Rivalta et al. (2015) for a review]; (2) long-distance sheet dykes, which may extend for kilometers during final magma ascent (Tilling & Dvorak, 1993; Ishizuka et al., 2008); (3) a deep feeding region in the mantle offset with respect to the central area (Gonzales et al., 2013; Aizawa et al., 2014; Michon et al., 2015); (4) different magma batches stored in independent reservoirs (Rittmann & Sato, 1973; Hildner et al., 2011; Zanon & Frezzotti, 2013). In any of these scenarios, contrasting magma rheological properties and the stress distribution in both basement and edifice (e.g. owing to topographic effects as well as to edifice load, regional tectonics, and the occurrence of adjacent volcanic centers) will exert a strong control on the architecture and dynamics of plumbing systems (Roman & Jaupart, 2014). An additional important par-ameter is the melting and magma supply rate from the mantle source to shallower magma storage regions (Vlasteľic et al., 2018).

Rift zones across large volcanic edifices provide a window into their magmatic plumbing system and allow the link between central and peripheral volcanic activity to be addressed (Garcia et al., 1989). In some cases, such as at Kilauea (Hawaii), magma transfer in-side rift zones occurs only at shallow levels. Small vol-umes of variably evolved melts can be stored close to the surface and can be then mobilized during an erup-tion following lateral dike injection (Ryan, 1988; Poland

et al., 2014). In other cases, rift zones provide direct access to the deep storage system (e.g. Madeira, Desertas Islands and the Azores Archipelago) and testify to magma ascent from multiple storage levels located both in the crust and in the uppermost mantle (Schwarz et al., 2004; Zanon & Pimentel, 2015).

Piton de la Fournaise (PdF) comprises a central sum-mit cone built inside the Enclos Fouque´ (EF) caldera (Fig. 1) and rift zones that propagate outside the central area. The area located inside the EF caldera is the focus of the ‘central’ eruptive activity, which is the most fre-quent type at PdF (Villeneuve & Bache` lery, 2006; Michon et al., 2013). Outside the EF caldera, ‘peripheral’ eruptions occur along three radially oriented rift zones (Chevallier & Bache` lery, 1981; Villeneuve & Bache` lery, 2006; Michon et al., 2015): the NE rift zone (NERZ), the SE rift zone (SERZ) and the NW rift zone (NWRZ). Peripheral eruptions of PdF are infrequent (less than 3% of the historical eruptions) and poorly documented. Relatively few data exist on their magma compositions, storage conditions and ascent pathways, and the pos-sible link with the central conduit system (Bache` lery, 1981; Bureau et al., 1998, 1999; Sigmarsson et al., 2005). A better understanding of this peripheral activity is therefore crucial to interpret seismicity and gas emis-sions, and also for volcanic risk assessment, because of the presence of 250 000 inhabitants in areas threatened by a potential peripheral eruption. Both the NERZ and the SERZ, which are at present almost aseismic, have been structurally related to the seaward displacement and spreading of the eastern flank of the PdF edifice. These two rift zones represent shallow pathways for lat-eral magma transfer from the central area, analogous to Kilauea and Etna (Acocella et al., 2003; Boivin & Bache` lery, 2009; Poland et al., 2014; Michon et al., 2015). Conversely, deep seismicity [>11 km below sea level (b.s.l.); Michon et al., 2015] is frequent below the NWRZ and potentially tracks deep magmatic pathways. The NWRZ rift possibly lies at the top of a region of deep magma storage that could promote progressive lateral magma transfer towards the central area (Michon et al., 2015). This model implies a 10–20 km off-set (NW to SE) between the deep magma ponding zone and the shallow crustal plumbing system that feeds the
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Fig. 1. Map showing the location of the samples collected from Piton de La Fournaise, the main geomorphological features of the volcano and the general trends of the three rift systems (dashed lines) converging towards the central area. The NW rift zone (NWRZ) is characterized by deep seismicity located between 11 and 20 km b.s.l., whereas the NERZ and SERZ are aseismic (Michon et al., 2015). Names of samples with ages less than 10 ka (most <5 ka): PdC, Piton de Caille; PC, Petit Crate` re; CC, Crate` re Commerson; PRE, Piton de la Rivie` re de l'Est; PR14, Piton Rampe 14; PsG, Piton sous le Gr` te; HA, Hau` y; DP, Demi-Piton; CHI, Chisny; FL, Formica Le` o; CS, Crate` re Se` ry; 1998 (K), March 1998 eruption at Kapor; 1998 (H), March 1998 eruption at Hudson; 1998 (A), August 1998 eruption; 2007 (F), February 2007 eruption; 2007 (M), March 2007 eruption; 2007 (A), April 2007 eruption. The black square (NdB) is the 'Nez de Boeuf' outcrop (Albare` de et al., 1997). Black triangles (TB, Trous Blancs; PBI, Piton Bleu; PBe, Piton Bellevue) are complementary samples of lavas (>10 kyr old) emitted in the peripheral area analysed only for Sr-Nd isotopes. Gabbroic clasts from the major Bellecombe explosive event (Bc) were sampled in the Plaine des Sables. Enclaves ranging from dunite to wehrlite are widespread in CHI, DP and HA lavas, but are very rare in lavas emitted inside the central area. Rare and small fine-grained dunite enclaves have been reported in April 2007 oceanite lavas (Welsch et al., 2013). Complete nomenclature and description of studied eruptions are provided in Supplementary Data Electronic Appendix 1. Whereas the central area is inhabited, the population is thinly spread on the volcano flanks in the peripheral area.

frequently active PdF central summit area (Liuzzo et al., 2015; Boudoire et al., 2017).

Here, we integrate bulk-rock chemistry (major and trace element data, Sr-Nd isotopes), mineral chemistry, barometry and modeling for recent (<5 kyr old) lavas from both the peripheral (NWRZ) and central (EF cal-dera) areas of PdF to (1) compare the geochemical signatures of magmas linked to the central and peripheral activity, (2) reconstruct the conditions of magma stor-age and ascent, and (3) better constrain magma transfer from mantle to crustal reservoirs prior to eruptions. Because of the debated role of clinopyroxene fraction-ation to produce PdF magma variability (Kornprobst et al., 1979; Albare` de et al., 1997; Boivin & Bache` lery, 2009; Famin et al., 2009; Brugier, 2016), we focus on a set of recent clinopyroxene-bearing basalts from both the peripheral and central areas and compare these with the volumetrically dominant olivine-bearing lavas. We argue that clinopyroxene-bearing basalts have particular significance in providing a better understanding of processes of deep magma evolution and transfer at PdF.

## GENERAL SETTINGS

### Geological context

La Re´ union is the youngest and most active volcanic island in the Indian Ocean (Fig. 1) related to the hotspot that has been proposed to have originated the Deccan flood basalt province in western India 65 Myr ago (Courtillot et al., 1988; Duncan et al., 1989; Mahoney et al., 2002). The island is built on a Late Cretaceous to Paleocene oceanic crustal block bounded by transform faults, which shape the Mascarene Basin (Bonneville et al., 1988). The island rises 7 km above the ocean floor with the seismic Moho discontinuity being located between 10 and 13 km b.s.l. (Gallart et al., 1999; Li et al., 2003; Fontaine et al., 2015). Below this depth a jump in seismic wave velocity is possibly caused by the presence of a large, <3 km thick layer of dense material, probably related to underplating and crystallization of ascending melts (Charvis et al., 1999; Fontaine et al., 2015).

La Re´ union Island is composed of two large volcanic edifices (Piton des Neiges, PdN: 3070 m a.s.l.; Piton de la Fournaise, PdF: 2632 m a.s.l.) that are made of
geochemically similar magmas belonging to the transitional series (Upton & Wadsworth, 1966). PdF began its activity c. 530 kyr ago on the southwestern flank of the largest and now quiescent volcanic edifice, PdN (Gillot & Nativel, 1989). Eruptive activity of PdF mostly occurred inside the EF caldera (indicated as intracaldera or 'central activity'). These eruptions have occurred on average once every 9 months since 1985, with an estimated bulk lava output rate of 0 15–0 38 m³ s⁻¹ (Roult et al., 2012). Seismic hypocentres beneath the central area are mostly confined to crustal depths between 0 5 and 11 km b.s.l. (Michon et al., 2015). This crustal seismicity is related to the central magma plumbing system, composed of multiple reservoirs at variable depths (Nercessian et al., 1996; Albare` de et al., 1997; Aki & Ferrazzini, 2000; Battaglia et al., 2005; Famin et al., 2009; Peltier et al., 2009; Le´nat et al., 2012; Di Muro et al., 2014, 2015; Vlaste´ lic & Pietruszka, 2016).

The EF caldera is open seaward to the east and connected by two aseismic rift zones (NERZ and SERZ), which act as pathways for shallow lateral migration of magma outside the caldera, to feed the peripheral activity (Villeneuve & Bache` lery, 2006; Michon et al., 2015). Lava output rates of the SERZ and NERZ are low (0 07–0 09 m³ s⁻¹; Michon et al., 2015) and the chemical and mineralogical compositions of the lavas match those of magmas erupted inside the central area (Boivin & Bache` lery, 2009). A third and volcanically less active zone of peripheral activity is the NWRZ (0 03–0 06 m³ s⁻¹ of lava emission over c. 50 kyr; Michon et al., 2015), which connects the western flank of PdF to the older PdN edifice (Fig. 1). The N120 orientation of the NWRZ has been related to regional NNE–SSW extension, which reactivates deep lithospheric faults (Michon et al., 2007, 2015; Boudoire et al., 2017). The NWRZ differs from the NERZ and SERZ by its frequent and deep seismicity (>11 km b.s.l., which is within the lithosphere-ic mantle), and the large volume of many pyroclastic cones and associated lavas (Michon et al., 2015). As we will document below, the lavas erupted from the NWRZ zone differ in their chemistry from those emitted from both the other rift zones and the central area (Bache` lery, 1981; Le´nat et al., 2012).

## Magma compositions

Magmas erupted at La Re´ union Island originate from partial melting (10%) of a slightly heterogeneous source in the upper mantle (Fisk et al., 1988; Graham et al., 1990; Staudacher et al., 1990; Hanyu et al., 2001; Fretzdorff & Haase, 2002; Putirka, 2008a; Valer et al., 2017b). The radiogenic isotope systematics indicate the occurrence of limited compositional heterogeneities in the mantle source (Bosch et al., 2008; Pietruszka et al., 2009; Di Muro et al., 2014; Valer et al., 2017b), associated with small degrees of crustal contamination (Luais, 2004; Pietruszka et al., 2009). With the exception of some old plagioclase-bearing alkali basalts, lavas at PdF are mainly transitional basalts (Upton &

Wadsworth, 1966; Valer et al., 2017a), subdivided into 'steady-state basalts' (SSB; Albare` de et al., 1997) and the 'abnormal group basalts' (AbG; Famin et al., 2009; Salau¨ n et al., 2010; Le´nat et al., 2012; Valer et al., 2017b).

'Steady-state basalts' (SSB) range from aphyric to weakly porphyritic with olivine (6 clinopyroxene 6 plagioclase) as the main phenocrysts; that is, minerals that formed directly from the magma in which they are hosted (Davidson et al., 2007). SSB are particularly rich in olivine (>20%) and are also known as 'oceanites' (Lacroix, 1923). These lavas are emitted in both the central and peripheral areas during phases characterized by high magma output rate (Albare` de et al., 1997; Famin et al., 2009; Di Muro et al., 2014). Olivine crystals in these oceanites have been interpreted to be ante-crysts; that is, crystals related to the magmatic system but not directly crystallized from the host magma (Davidson et al., 2007). Antecrysts can be derived from either (1) the disruption of a 'crystal mush' within some intracrustal reservoirs (Clocchiatti et al., 1979; Famin et al., 2009; Welsch et al., 2013; Di Muro et al., 2014) or (2) crystallized cumulate bodies located at variable depths (Albare` de et al., 1997; Bureau et al., 1998).

'Abnormal Group Basalts' (AbG) were first documented by Albare` de et al. (1997) in the roughly hun-dreds of thousand years old peripheral Nez de Boeuf Series (Fig. 1). These basalts have a mildly alkaline af-finity, low CaO and SiO₂ contents, and are enriched in both compatible and incompatible trace elements with respect to SSB. AbG are very common among the lavas emitted in the peripheral area (Le´nat et al., 2012) and rarer in the central area (Boivin & Bache` lery, 2009). Occasionally, SSB- and AbG-type lavas may erupt sim-ultaneously from distinct craters, as observed during the 1998 central eruption (SSB Kapor lavas; AbG Hudson lavas; Fig. 1). Magmas erupted in 1998 were thought to be end-member compositions whose mixing originated the magmas erupted during the 1999–2007 period (Di Muro et al., 2014). Pressures recorded by melt and fluid inclusions in olivine from SSB lavas are usually low and on average <1 kbar (Famin et al., 2009; Di Muro et al., 2016). This strongly contrasts with the AbG Hudson basalts, which carry olivine-hosted melt and fluid inclusions that record pressures up to 5 kbar (Bureau et al., 1999).

## SAMPLES

For this study, we mainly sampled products erupted from the poorly studied peripheral scoria cones (i.e. outside the EF caldera and near its rim, inside the central area). We studied the following: (1) small monogen-etic strombolian cones in the peripheral area (e.g. Formica Le´o, Piton sous le Gi`te, Piton Rampe 14, Piton de la Rivie` re de l'Est, Petit Crate` re, Crate` re Commerson), located as far as 7–8 km from the summit cone of PdF and emitting predominantly clinopyroxene-bearing basalts (Fig. 1; Supplementary Data Electronic
Appendix 1; *supplementary data* are available for downloading at http://www.petrology.oxfordjournals.org; (2) large polygenetic cones in the peripheral area (e.g. Chisny, Demi-Piton, Hau" y) that emitted olivine-bearing, aphyric to olivine-phyric basalts (Fig. 1). Dunite, wehrlite and gabbroic enclaves are common in lavas and lapilli beds emplaced during large eruptions in the peripheral area (e.g. Upton & Wadsworth, 1972; Bache' lery, 1981), but they are very rare in the central area. These enclaves represent slowly cooled cumu-lates in magmatic reservoirs. The age of the gabbroic cumulates is unknown, but some have interstitial melt and thus may be related to the recent magmatic history (<5 kyr old; Upton et al., 2000). Examples of dunite and wehrlite enclaves erupted from the polygenetic Chisny cone (Bureau et al., 1998; Brugier, 2016) and gabbroic clasts in the Bellecombe Ash Member (Upton et al., 2000) have been analyzed for comparison with phenoc-rysts and antecrysts in lavas.

Our new data are compared with the literature on the 1972–2010 lavas from the central area (Albare` de et al., 1997; Vlaste´ lic et al., 2007; Peltier, 2007; Boivin & Bache` lery, 2009; Pietruszka et al., 2009; Salau" n et al., 2010; Di Muro et al., 2014, 2015). In particular, we focus on the lavas from the two largest recent central erup-tions of 1998 and 2007 that show significant compos-itional variability and record deep magma inputs (Bureau et al., 1999; Di Muro et al., 2014).

## ANALYTICAL METHODS

### Major and trace elements and isotope analyses of bulk samples

Thirty-eight lava samples were crushed and finely ground in agate. Major and trace element analyses were performed at the SARM (Service d'Analyse des Roches et des Mine' raux de Nancy, France; Supplementary Data Electronic Appendix 2), by induct-ively coupled plasma optical emission spectrometry (ICP-OES) and inductively coupled plasma mass spec-trometry (ICP-MS), respectively [see Carignan et al. (2001) for a review of the protocols and of reference materials]. Relative errors are <1% (SiO₂, Al₂O₃), <2% (MgO, FeO, CaO) and <5% (Na₂O, K₂O, TiO₂) for major elements. Relative errors are <15% (Sc), <10% (Th, Cr) and <5% for other trace elements used in this study.

Ten analyses of Sr and Nd isotopic ratios were performed at INGV-OV (Istituto Nazionale di Geofisica e Vulcanologia-Osservatorio di Napoli, Italy). Isotopic compositions were determined on c. 0 1 g of whole-rock powders after dissolution with ultrapure acids (HF–HNO₃–HCl mixtures). Sr and Nd were separated from the matrix through conventional ion-exchange proce-dures [standard separation scheme as described by Arienzo et al. (2013) and in references therein] and measured statically by thermal ionization mass spec-trometry (TIMS; Thermo Finnigan Triton TI). During col-lection of the isotopic data, replicate analyses of NIST-

SRM 987 (SrCO₃) and La Jolla international reference standards were measured. Mean measured values for ⁸⁷Sr/⁸⁶Sr for the NIST-SRM 987 and ¹⁴³Nd/¹⁴⁴Nd for the La Jolla standards were 0 710219 6 0 000018 (2r, n ¼ 101) and 0 511845 6 0 000010 (2r, n ¼ 52), respectively. Measured ⁸⁷Sr/⁸⁶Sr and ¹⁴³Nd/¹⁴⁴Nd ratios were normalized for within-run isotopic fractionation to ⁸⁶Sr/⁸⁸Sr ¼ 0 1194 and ¹⁴⁶Nd/¹⁴⁴Nd ¼ 0 7219 and were normalized to the recommended values of ⁸⁷Sr/⁸⁶Sr ¼ 0 71025 and ¹⁴³Nd/¹⁴⁴Nd ¼ 0 511850 for the NIST-SRM 987 and La Jolla standards, respectively (Supplementary Data Electronic Appendix 2).

### Electron microprobe analyses of mineral phases

Electron microprobe analyses of mineral phases were performed using two instruments. The first was a CAMECA SX100 electron microprobe at Camparis (Paris, France) operated with a 15 kV accelerating volt-age and a 20 nA current focused beam, with counting times varying from 10 to 50 s depending on the elem-ent. Relative errors were 0 5–1 5% (SiO₂, MgO), 2–4% (Al₂O₃, FeO, CaO), 5–15% (TiO₂, K₂O, Na₂O) and 20–45% (NiO, MnO, Cr₂O₃). For olivine, the relative error on CaO was 6–10%. The second instrument was a JEOL JXA-8900RL electron microprobe (University of Go¨ ttingen, Germany) operated with a 15 kV accelerating voltage and a 15 nA current defocused beam, with 5–15 s count-ing times per element in minerals. Relative errors were

0 5–1 5% (SiO₂, MgO, FeO, CaO, Al₂O₃), 1–4% (TiO₂), 5–15% (Na₂O, Cr₂O₃, NiO) and 5–35% (K₂O, MnO). For olivine, the relative error on CaO was 5–10%. No systematic difference was found between the two electron microprobe datasets.

### Fluid and melt inclusion barometry

Fluid inclusions were analyzed at room temperature at IPGP (Institut de Physique du Globe de Paris) using a Horiba T64000 Raman spectrometer manufactured by Jobin-Yvon with a 1800 grooves mm⁻¹ grating, offering a high resolution (<1 cm⁻¹) for peak identification. Fluid inclusions were excited using an Arb laser at 488 nm wavelength. Pressures were estimated using the CO₂ density deduced from the Fermi diad split (van den Kerkhof & Olsen, 1990; Frezzotti et al., 2012) between the 1388 and 1285 cm⁻¹ bands (D). We used the new densimeter of Lamadrid et al. (2016) calibrated for a Jobin-Yvon Horiba spectrometer using a 1800 grooves mm⁻¹ grating with a 514 nm wavelength (see Supplementary Data Electronic Appendix 3 for a complete description of the procedure). The total error related to the calibration and the analytical conditions

results in a maximum overestimation of CO₂ densities of 0 05 g cm⁻³ (i.e. less than 0 7 kbar).

Additionally, we calculated the saturation pressures of volatiles dissolved in melt inclusions using the Di Muro et al. (2016) database and the H₂O–CO₂ solubility model of Papale et al. (2006). This model is best suited for high pressures (>5 kbar).
# RESULTS

## Petrography

Two main petrographic groups characterize the lavas emitted in the peripheral area. One is defined by the presence of clinopyroxene phenocrysts next to olivine. The second lacks clinopyroxene phenocrysts, but has abundant olivine phenocrysts, as typically observed in PdF lavas (Welsch et al., 2009, 2013).

The small monogenetic strombolian cones built on the NWRZ of PdF (Crater near Crate' re Commerson pits, Petit Crate' re, Piton sous le Gite, Piton de la Rivie' re de l'Est, and Formica Le´ o; Fig. 1) have erupted lavas with unusually abundant clinopyroxene (up to 1 mm long) that have previously been partially described in the lit-erature (Bache' lery, 1981). These lavas, which we term here 'clinopyroxene-bearing' basalts, have more abun-dant clinopyroxene phenocrysts than olivine, and only a small amount of plagioclase (Fig. 2a). They contain skeletal olivine crystals up to 1 mm (Fig. 2b), with long dendritic overgrowths (up to 500 lm long) and sector-zoned clinopyroxene phenocrysts and microphenoc-rysts (Fig. 2c). Clinopyroxenes are euhedral and exhibit complex oscillatory zoning at their rims. Euhedral plagioclase microphenocrysts have only weak zoning. Rare inclusions of chromite are present in olivine and clinopyroxene. In all of these lavas, the groundmass is holocrystalline and composed of microlites of olivine, clinopyroxene, plagioclase and magnetite. Lavas erupted from different small peripheral cones show a minimal difference in their average crystal sizes and in the abundance of anhedral to subhedral olivine (Table 1). Anhedral olivine may be included in sector-zoned clinopyroxenes (Fig. 2d), suggesting that an early olivine growth was followed by cotectic olivine-clino-pyroxene crystallization.

Large polygenetic cones from the peripheral NWRZ (Chisny, Hau' y, Demi-Piton, Piton de Caille) erupted lavas in which olivine is the dominant mineral phase, as commonly observed in PdF lavas. They are defined here as 'olivine-bearing' lavas. The olivine ranges from up to 2 mm in size to microphenocrysts and microlites. Many olivine crystals are rounded and densely frac-tured (Fig. 2e). Clusters of polyhedral olivine crystals, like those described by Welsch et al. (2013), are rare and rounded (e.g. Fig. 2f). Olivine microphenocrysts are scattered in the groundmass. Chromite inclusions are common in olivine crystals, whereas magnetite micro-lites occur sparsely in the groundmass. Minor amounts of anhedral clinopyroxene microlites are confined to the groundmass, together with olivine, plagioclase and magnetite (Table 1).

Dunitic and wehrlitic enclaves are abundant only in 'olivine-bearing' lavas from large peripheral cones (e.g. Chisny). These ultramafic enclaves are mainly made of anhedral to subhedral olivine and clinopyroxene. Olivine crystals sometimes exhibit equilibrated 120 grain boundaries (Fig. 3a and b) and typically show pat-terned extinction indicating deformation bands.

Chromite inclusions are common in olivine. Fragments of gabbroic cumulates with variable grain size and amounts of evolved interstitial glass are widespread in the deposits of the Bellecombe Ash Member (Upton et al., 2000). In these gabbroic clasts, olivine and clino-pyroxene are generally euhedral to subhedral. Clinopyroxene exhibits complex patchy, concentric and sector-zoning patterns (Fig. 3c). The latter resemble those observed in 'clinopyroxene-bearing' lavas, but differ by the absence of oscillatory zoning and the oc-currence of corrosion textures. Plagioclase in gabbros is smaller than olivine and clinopyroxene, and has tabu-lar shapes and well-developed concentric and oscilla-tory zoning. Small gabbroic and dunitic fragments were also found in 'clinopyroxene-bearing' basalts (Fig. 3d).

Recent (1972–2010) lavas from the central area vary from aphyric basalts to cotectic (clinopyroxene-plagioclase 6 olivine) basalts to olivine porphyritic basalts (e.g. Welsch et al., 2013). In olivine basalts, rounded clinopyroxenes with dissolution textures have been reported (Welsch et al., 2009). The 1998 lavas vary from aphyric basalts to olivine–clinopyroxene basalts (Hudson vent; small volume; Fig. 1) to rare basalts enriched in plagioclase crystals (Kapor vent; large vol-ume; Boivin & Bache' lery, 2009). In particular, the '1998 clinopyroxene-bearing' lavas of the Hudson vent, previ-ously considered as the archetype of AbG lavas (Famin et al., 2009), display unusual textures with the largest variability in crystal sizes and shapes found in our sam-ple set. In the Hudson samples, euhedral, anhedral and skeletal olivine crystals coexist with clinopyroxene phe-nocrysts (rarely sector-zoned) and plagioclase micro-phenocrysts (Table 1). In contrast, the 2007 lavas vary from aphyric basalts (February) to cotectic basalts (March) to olivine-rich basalts that lack clinopyroxene phenocrysts (April–May; Di Muro et al., 2014).

## Mineral compositions

### Olivine

'Clinopyroxene-bearing' basalts from peripheral erup-tions share a similar composition of equilibrium olivine (Fig. 4a). Olivine core compositions are Fo83 5–81 5, whereas rims and dendrites are iron-rich (<Fo81 5; Fig. 4a). In 'olivine-bearing' lavas from peripheral cones, the olivine composition is polymodal with higher modes of MgO content and a wider range compared with 'clinopyroxene-bearing' basalts (Fig. 4b; Brugier, 2016). In cognate cumulate enclaves (Brugier, 2016), the average MgO content of olivine decreases from dunites to wehrlites to gabbros (Fig. 4c). The olivine compos-ition is bimodal in dunites (Fo87 5 and Fo84). Wehrlites lack MgO-rich olivine crystals, but share a common composition of olivine with dunites (Fo84). Gabbroic clasts from the Bellecombe Ash Member contain MgO-poor olivine crystals (Fo<82) with a main mode at Fo78, consistent with previous studies (Upton et al., 2000). The olivine composition in 1977–2010 lavas from the central area is bimodal (Fig. 4d) and is best represented
## Clinopyroxene-bearing lavas

<figure><img src="image_4.png" /></figure>

Fig. 2. Transmitted light photomicrographs of Piton de la Fournaise lavas from both the central and peripheral areas, representative of the petrographic features described in this study. (a) Skeletal olivine microphenocrysts, euhedral sector-zoned clinopyroxene microphenocrysts and euhedral plagioclase microphenocrysts in Piton sous le Gite lava. (b) Skeletal olivine phenocrysts with den-dritic overgrowths and (c) euhedral sector-zoned clinopyroxene phenocrysts in Petit Crate' re lava. Dark sectors have high Mg-Si contents, whereas bright sectors are enriched in Fe-Al-Ti. (d) Subhedral olivine antecryst with a high MgO content (Fo84) and abundant fluid inclusions. (e) Rounded olivine phenocrysts in Chisny lava. (f) Rounded polyhedral aggregate of olivine crystals in Chisny lava. Ol, olivine; Cpx, clinopyroxene; Pl, plagioclase; Sp, spinel.

by the products of the 2007 eruptions. The main and MgO-richest mode centered at Fo84 5–84 0 is best represented by phenocrysts in olivine-rich lavas from the large-volume April 2007 eruption (Fig. 4e). The second mode is centered at Fo81 5 as observed in the cotectic basalts from the small-volume March 2007 eruption (Fig. 4e; Di Muro et al., 2014). Olivine crystals from ‘1998 clinopyroxene-bearing’ lavas from the Hudson vent show a wide compositional range (centered at Fo86), with unusually high MgO contents (up to Fo88 2; Fig. 4f), not in equilibrium with the host-rock. MgO-rich olivine crystals (Fo>84 5) found in lavas from central eruptions are neither in equilibrium with the host groundmass and bulk-rocks, nor with the most

magnesian glasses found in pyroclasts from central eruptions (MgO: 8.6 wt % in April 2007 and 8.7 wt % in December 2008; Villemant et al., 2009; Di Muro et al., 2015).

Our data show that olivine populations can also be discriminated on the basis of their concentrations of Ca, which diffuses more slowly than Fe and Mg (Coogan et al., 2005). Two trends are recognized here (Fig. 5a and b) confirming previous findings by Di Muro et al. (2015). In trend 1, found in ‘2007 olivine-bearing’ lavas from the central area, CaO increases with decreasing Fo. In trend 2, CaO exhibits little variation over a range of Fo, as shown by the ‘1998 clinopyroxene-bearing’ lavas at the Hudson vent. The composition of olivine in
<table><caption>Table 1: Relative abundances and textures of olivine, clinopyroxene and plagioclase occurring as phenocrysts, microphenocrysts, microlites and in the groundmass in lavas emitted in the central and peripheral areas</caption><thead><tr><th rowspan="2"></th><th colspan="4">Euhedral or anhedral olivine</th><th colspan="4">Skeletal olivine</th><th colspan="4">Clinopyroxene</th><th colspan="4">Plagioclase</th></tr><tr><th>Pheno.</th><th>MPheno.</th><th>Micro.</th><th>G.</th><th>Pheno.</th><th>MPheno.</th><th>Micro.</th><th>G.</th><th>Pheno.</th><th>MPheno.</th><th>Micro.</th><th>G.</th><th>Pheno.</th><th>MPheno.</th><th>Micro.</th><th>G.</th></tr></thead><tbody><tr><td colspan="17"><strong>Peripheral (NWRZ) eruptions</strong></td></tr><tr><td colspan="17"><strong>Clinopyroxene-bearing lavas</strong></td></tr><tr><td>PC</td><td>X</td><td></td><td>X</td><td>X</td><td>XX</td><td>XX</td><td></td><td></td><td>XX</td><td>XX</td><td>X</td><td>X</td><td></td><td>XX</td><td>X</td><td>X</td></tr><tr><td>PsG</td><td>X</td><td></td><td>X</td><td>X</td><td>XX</td><td>XX</td><td></td><td></td><td>XX</td><td>XX</td><td>X</td><td>X</td><td></td><td>X</td><td>X</td><td>X</td></tr><tr><td>PRE</td><td>X</td><td></td><td>XX</td><td>XX</td><td>X</td><td>X</td><td>XX</td><td>XX</td><td></td><td>X</td><td>XX</td><td>X</td><td></td><td>X</td><td>XX</td><td>X</td></tr><tr><td>FL</td><td>X</td><td></td><td>X</td><td>X</td><td></td><td>XX</td><td>XX</td><td></td><td>X</td><td>XX</td><td>XX</td><td>X</td><td></td><td></td><td>XX</td><td>X</td></tr><tr><td colspan="17"><strong>Olivine-bearing lavas</strong></td></tr><tr><td>CHI</td><td>XX</td><td></td><td>X</td><td>X</td><td></td><td></td><td>X</td><td></td><td></td><td>X</td><td>X</td><td></td><td>X</td><td>X</td><td>X</td><td>X</td></tr><tr><td>HA</td><td>XX</td><td></td><td>XX</td><td>X</td><td></td><td></td><td>X</td><td></td><td></td><td>X</td><td>X</td><td></td><td>X</td><td>X</td><td>X</td><td>X</td></tr><tr><td colspan="17"><strong>Central eruptions</strong></td></tr><tr><td colspan="17"><strong>Clinopyroxene-bearing lavas</strong></td></tr><tr><td>March 1998 (H)</td><td>X</td><td></td><td>XX</td><td>XX</td><td>X</td><td>XX</td><td>XX</td><td>X</td><td>X</td><td>X</td><td>X</td><td>X</td><td>X</td><td>XX</td><td>X</td><td>X</td></tr><tr><td>CS</td><td>X</td><td></td><td>XX</td><td>XX</td><td>X</td><td>XX</td><td>XX</td><td>X</td><td></td><td>XX</td><td>XX</td><td>X</td><td>X</td><td>X</td><td>XX</td><td>X</td></tr><tr><td colspan="17"><strong>Olivine-bearing lavas</strong></td></tr><tr><td>3–5 April 2007</td><td>XX</td><td></td><td>X</td><td>X</td><td>X</td><td></td><td></td><td></td><td></td><td>X</td><td></td><td></td><td></td><td></td><td>X</td><td>X</td></tr></tbody></table>

Pheno., phenocrysts (>500 lm); MPheno., microphenocrysts (500–100 lm); Micro., microlites (<100 lm); G., groundmass (<50 lm). Olivine-bearing lavas from the peripheral area: CHI, Chisny; HA, Hau¨ y. Clinopyroxene-bearing lavas from the peripheral area: PC, Petit Crate` re; PsG, Piton sous le Gi` te; PRE, Piton Rivie` re de l’Est; FL, Formica Le´ o. Olivine-bearing lavas from the central area: 3–5 April 2007. Clinopyroxene-bearing lavas from the central area: CS, Crate` re Se´ ry; March 1998 (H), Hudson. XX marks the main size class for each min-eral phase.
Cumulates

<figure><img src="image_5.png" /></figure>

Fig. 3. Transmitted light photomicrographs (a, b, d) and BSE image (c) of enclaves embedded in lavas emitted along the NWRZ. (a) Dunite and (b) wehlite enclaves from the Chisny cone show anhedral to subhedral olivine, spinel and clinopyroxene crystals (Brugier, 2016). (c) Sector-zoned clinopyroxene phenocrysts with resorbed core from a gabbroic cumulate from the Bellecombe Ash Member. (d) Small gabbroic fragment in Petit Crate' re lava. Ol, olivine; Cpx, clinopyroxene; Pl, plagioclase; Sp, spinel.

'clinopyroxene-bearing' lavas from peripheral eruptions lies on (or between) these two trends (Fig. 5a). In con-trast, olivine from 'olivine-bearing' lavas in the periph-eral area shows a wide range in CaO contents at a given Fo (Fig. 5b). The CaO content progressively increases from core to rim, whereas Fo decreases. High CaO con-tents (up to 0 45 wt %) are measured mostly in olivine

Fo84. Conversely, low CaO contents (<0 1 wt %) are found in the cores of many olivine crystals with a large compositional range (Fo82–86). MgO-rich olivine crystals in dunites and wehrlites present a compositional spec-trum similar to that observed in ‘olivine-bearing’ lavas from peripheral eruptions (Fig. 5c). These anomalous low CaO contents, however, could result from a process of sub-solidus diffusive re-equilibration (Brugier, 2016). Conversely, in Bellecombe gabbros, olivine compos-ition falls on Trend 2; that is, like those found in ‘clino-pyroxene-bearing’ lavas from peripheral eruptions.

Pyroxene

'Clinopyroxene-bearing' basalts that erupted in the per-
ipheral area (Table 2) contain mainly augite phenoc-rysts
(XW0 ¼ 0 40 6 0 03; International Mineralogical Association
nomenclature, Morimoto, 1988). Diopside (XW0 ¼ 0 45 6 0
02) with augite rims is rarely observed (e.g. Formica Le´ o).
Only iron-rich augite microlites occur in the groundmass of
'olivine-bearing' lavas emit-ted in the peripheral area.
Diopside is more common in gabbroic cumulates (XW0 ¼ 0
45 6 0 01) and the most calcic clinopyroxenes occur in
dunitic and wehrlitic

enclaves (XW₀ ¼ 0 47 60 01). In basalts erupted in the
central area, augites and Mg-augites have been reported
(Albare` de et al., 1997; Boivin & Bache` lery, 2009; Di Muro
et al., 2015; Peltier, 2007) but diopside is scarce (Clocchiatti
et al., 1979; Peltier, 2007).

Sector-zoning in clinopyroxenes can potentially af-fect
their classification and their use for barometric pur-poses.
Two clinopyroxene compositions are found in
'clinopyroxene-bearing' basalts from peripheral erup-tions
and in some Bellecombe gabbros, corresponding to distinct
sectors of the crystals (Fig. 6a). Bright zones in back-
scattered electron (BSE) images are enriched in Ti and Al,
whereas dark zones are enriched in Mg and Si (Fig. 6a;
Table 2). Meanwhile, ratios such as Mg/Fe and Ti/Al do not
significantly differ in growth zones (Table 2; Supplementary
Data Electronic Appendix 2) and may thus be reliable to
compare the clinopyroxene composition in lavas from
various eruptive vents (Fig. 6b). Using a Kd ¼ 0 27 for Fe-
Mg partitioning be-tween melt and clinopyroxene (Putirka,
2008b; Brugier, 2016), we calculated the average Mg-
numbers of clino-pyroxene crystals in equilibrium with their
host lavas, for both peripheral and central areas. The
calculated (88 1–82 1) and measured (85 7–80 5) Mg-
numbers (100) are in good agreement (Supplementary
Data Electronic Appendix 4). The only exceptions are the
microlites in the groundmass of 'olivine-bearing' lavas from
peripheral eruptions and the microphenocrysts in 1998
'clinopyroxene-bearing' lavas, which are too iron-rich and in
equilibrium only with more evolved (MgO
<figure><img src="image_6.png" /></figure>

Fig. 4. Distribution of forsterite (Fo) content (%) in olivine phenocrysts and microphenocrysts from lavas and enclaves of Piton de la Fournaise. (a) ‘Clinopyroxene-bearing’ basalts (Petit Crate’ re, Formica Le’ o) from peripheral eruptions contain olivine crystals with cores of a unimodal forsterite composition (main mode at Fo83 and a tail towards lower Fo). (b) ‘Olivine-bearing’ basalts (Chisny, Hau’ y; Brugier, 2016; this study) and olivine-rich basalts (Piton de Caille; Bureau et al., 1998) emitted in the peripheral area contain olivine crystals with a broad polymodal forsterite content distribution (up to Fo88 8) and main compositional modes in the Fo83 5–86 5 range. (c) Bimodal olivine composition (Fo84 and Fo87 5) is observed in dunites (Brugier, 2016). The Fo content of the main mode decreases from dunites to wehrlites (Fo84; Brugier, 2016), to gabbros (Fo78; this study). (d) Bimodal distribution of olivine composition (modes at Fo84–84 5 and Fo81 5) in typical lavas emitted inside the central area during the 1977–2010 central activity. (e) Distinct distribution and modes of olivine composition in March 2007 aphyric lavas and April 2007 aphyric to olivine-rich lavas. (f) Olivine composition in March 1998 lavas of the Kapor Crater (SSB) and of the Hudson Crater (AbG). Arrows mark the expected compositions of olivine in equilibrium with the bulk-rock (or with the most magnesian glasses for April 2007), calculated with the partition (Mg/Fe) coefficient Kdmelt/ol ¼ 0 306 defined by Fisk et al. (1988) for La Re’ union Island lavas. Dataset of olivine compositions in basalts emitted in the central area is compiled from Boivin & Bache’ lery (2009), Di Muro et al. (2014, 2015) and Peltier (2007). Grey transparent boxes highlight the two main modes of forsterite content in olivine from central lavas (1977–2010). Abbreviations in parentheses refer to sample locations (see Fig. 1).

<5 wt %) melts (Fig. 6b; Supplementary Data Electronic Appendix 4). The highest Mg-numbers (94–96) and the lowest Ti/Al ratios (down to 0 11) are in the clinopyroxene of dunites (Mg-numbers 87–94; Ti/Al 0 11–0 22). The absence of Al enrichment in low Ti/Al

clinopyroxenes from dunites and wehrlites is consistent with previous findings on magnesian clinopyroxene at PdF (Fisk et al., 1988; Fretzdorff & Haase, 2002). Clinopyroxene in gabbroic cumulates from the Bellecombe deposit has a low Mg-number (78–90) and
its Ti/Al (0 18–0 32) overlaps the upper range found in clinopyroxene in lavas from both the central and peripheral areas (for the same Mg-number; Table 2).

## Feldspars

Plagioclase in 'clinopyroxene-bearing' lavas from peripheral centres is mostly not zoned, with a composition ranging from An72 1Ab27 4Or0.5 to An53 5Ab44 2Or2.3. Plagioclase (An68 3Ab30 3Or1.4 and An45 7Ab51 7Or2.6) is also present in the matrices of 'olivine-bearing' lavas erupted from the peripheral area. However, An contents are lower than those in 'clinopyroxene-bearing' lavas and consistent with crystallization from more alkaline melts. Plagioclase in Bellecombe gabbros records the highest anorthite content measured in this study (An84 8Ab14 8Or0.4). There, rare alkali feldspars up to An3 1Ab62 5Or34.4 are also observed. Plagioclase in lavas from the central area ranges from An79 0Ab20 2Or0.8 to An60 5Ab38 5Or1.0 and complex zoning is observed in both phenocrysts and microlites (Boivin & Bache' lery, 2009; Di Muro et al., 2015). In '1998 clinopyroxene-bearing' lavas, plagioclase is sodic, ranging from An67 4Ab31 7Or0.9 to An57 7Ab40 8Or1.5 and is generally not zoned.

## Barometry

### Pyroxene barometry

The clinopyroxene-based barometry is sensitive to the jadeite component (NaAlSi₂O₆), which increases with pressure, but is also influenced by temperature, water content and sector-zoning (Fig. 6a). As expected, data resulting from sector-zoned clinopyroxene crystals gave large pressure ranges for different sectors of the same crystal. We found that the clinopyroxene-melt equilibrium barometer [equation (30), Putirka et al., 2003] returns systematically higher and more variable pressures (DP 2 kbar) for different growth zones compared with the clinopyroxene-only barometer [DP <1 kbar; equation (32b), Putirka, 2008b]. For this reason, we use the clinopyroxene-only barometer of Putirka (2008b), which allows integration of a larger dataset (Fig. 7), and reduces the uncertainty related to the bar-ometer calibration (Putirka, 2008b). Crystallization pressures for several crystals in each sample, from both the peripheral and central areas (Fig. 7), are calculated. An uncertainty (1σ) of 2.6 kbar is reported from experimental calibrations, but for a large and consistent analytical dataset for many equilibrium clinopyroxene-melt pairs, this uncertainty is considerably reduced by averaging calculated pressures (Putirka, 2017). We thus report averaged pressure values with a 1σ uncertainty. Following Hammer et al. (2016), we assume that sectors low in Na and Al (dark sectors; Fig. 2b) provide minimum pressure estimates for sector-zoned pyroxenes (Fig. 7).

In clinopyroxene phenocrysts from 'clinopyroxene-bearing' lavas from peripheral eruptions, dark sectors provided a minimum average pressure of 3.1–6.0–6 kbar.

<figure><img src="image_7.png" /></figure>

Fig. 5. CaO content (wt %) relative to Fo content (%) in olivine crystals from (a) 'clinopyroxene-bearing' lavas and (b) 'olivine-bearing' lavas from both peripheral and central areas, and (c) cumulates. Two trends are observed in olivine of lavas erupted from the central area. Trend 1 (Fo decrease and CaO enrich-ment) is best represented by olivine in the March 1998 (Kapor), March and April 2007 lavas (Boivin & Bache' lery, 2009; Villemant et al., 2009; Di Muro et al., 2014) labeled 1998 (K), 2007 (M) and 2007 (A), respectively. Trend 2 (Fo decrease and almost constant CaO content) is represented by olivine in March 1998 (Hudson) lavas labeled 1998 (H). Olivine crystals in 'clinopyroxene-bearing' lavas from the peripheral area (PC, FL) fall on these trends. Scattered CaO contents occur in olivine from 'olivine-bearing' lavas emitted in the peripheral area (CHI, HA; Brugier, 2016; this study). Red shaded box highlights the expected range of forsterite contents of olivine in equilibrium with 'olivine-bearing' bulk-rock compositions (CHI, HA). Olivine in dunites and wehrlites spans a broad CaO range (Brugier, 2016). Most olivine crystals from the Bellecombe olivine–clino-pyroxene gabbros fall on the low-Ca Trend 2.

The highest pressures are recorded in clinopyroxene from dunite–wehrlite enclaves erupted as lithic fragments from the polygenetic Chisny cone, with an average of 5.6–6.0–8 kbar. A lower average pressure
<table><caption>Table 2: Representative composition (wt %) of clinopyroxene phenocrysts in lavas from central and peripheral areas, and clinopyroxene microlites in lavas emitted in the peripheral area</caption><thead><tr><th rowspan="3">Sample:</th><th colspan="12">Central eruptions*</th><th colspan="12">Peripheral eruptions (NWRZ)</th></tr><tr><th colspan="2">June 2000</th><th colspan="4">March 1998 (Kapor)</th><th colspan="4">March 1998 (Hudson)</th><th colspan="12">Formica Le'o (FL)</th></tr><tr><th>Pheno Core</th><th>Pheno Core</th><th>983-103 Pheno Core</th><th>Pheno Core</th><th>Pheno Rim</th><th>Pheno Rim</th><th>983-113 Pheno Core</th><th>Pheno Core</th><th>Pheno Rim</th><th>Pheno Rim</th><th>Pheno Core</th><th>Pheno Bright</th><th>Pheno Core</th><th>Pheno Bright</th><th>Pheno Core</th><th>Pheno Bright</th><th>Pheno Core</th><th>Pheno Bright</th><th>Pheno Core</th><th>Pheno Bright</th><th>Pheno Core</th><th>Pheno Bright</th><th>Pheno Core</th><th>Pheno Bright</th><th>Pheno Core</th><th>Pheno Bright</th></tr></thead><tbody><tr><td>SiO2</td><td>51.6</td><td>50.7</td><td>48.4</td><td>50.5</td><td>51.9</td><td>47.4</td><td>46.3</td><td>47.8</td><td>49.0</td><td>48.8</td><td>52.1</td><td>49.1</td><td>52.0</td><td>48.4</td><td>51.0</td><td>49.3</td><td>52.5</td><td>48.8</td><td>52.3</td><td>49.5</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Al2O3</td><td>2.4</td><td>2.9</td><td>5.2</td><td>3.4</td><td>2.3</td><td>6.6</td><td>6.0</td><td>5.9</td><td>5.1</td><td>5.6</td><td>2.7</td><td>5.3</td><td>2.7</td><td>5.8</td><td>3.1</td><td>4.8</td><td>2.4</td><td>5.1</td><td>2.3</td><td>4.9</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>TiO2</td><td>1.2</td><td>1.2</td><td>2.3</td><td>1.4</td><td>1.1</td><td>3.2</td><td>2.6</td><td>2.5</td><td>2.1</td><td>2.4</td><td>0.9</td><td>1.7</td><td>1.1</td><td>2.0</td><td>1.2</td><td>1.6</td><td>0.9</td><td>1.8</td><td>0.9</td><td>1.5</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>CaO</td><td>19.1</td><td>19.5</td><td>17.1</td><td>14.0</td><td>14.5</td><td>18.8</td><td>20.0</td><td>20.2</td><td>20.7</td><td>19.2</td><td>21.5</td><td>21.7</td><td>20.5</td><td>21.6</td><td>20.4</td><td>21.5</td><td>19.5</td><td>22.0</td><td>20.8</td><td>21.6</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Na2O</td><td>0.2</td><td>0.2</td><td>0.3</td><td>0.2</td><td>0.1</td><td>0.4</td><td>0.3</td><td>0.3</td><td>0.3</td><td>0.4</td><td>0.2</td><td>0.3</td><td>0.3</td><td>0.3</td><td>0.3</td><td>0.3</td><td>0.2</td><td>0.3</td><td>0.2</td><td>0.3</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>MgO</td><td>16.5</td><td>16.5</td><td>16.3</td><td>18.5</td><td>18.7</td><td>13.5</td><td>14.2</td><td>14.3</td><td>15.2</td><td>14.6</td><td>16.2</td><td>14.5</td><td>16.6</td><td>14.0</td><td>16.0</td><td>14.9</td><td>17.6</td><td>14.6</td><td>16.9</td><td>14.7</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>FeO</td><td>7.7</td><td>7.6</td><td>10.4</td><td>10.7</td><td>10.2</td><td>8.8</td><td>9.3</td><td>9.3</td><td>8.6</td><td>9.7</td><td>5.9</td><td>6.7</td><td>6.7</td><td>7.2</td><td>7.2</td><td>6.1</td><td>6.6</td><td>6.4</td><td>6.1</td><td>6.4</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Cr2O3</td><td>0.3</td><td>0.2</td><td>0.1</td><td>0.1</td><td>0.1</td><td>0.0</td><td>0.2</td><td>0.1</td><td>0.1</td><td>0.1</td><td>0.5</td><td>0.6</td><td>0.3</td><td>0.3</td><td>0.2</td><td>1.1</td><td>0.5</td><td>0.7</td><td>0.5</td><td>0.8</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Total</td><td>99.2</td><td>99.0</td><td>100.4</td><td>99.0</td><td>99.3</td><td>99.0</td><td>99.2</td><td>100.6</td><td>101.3</td><td>101.0</td><td>100.1</td><td>100.1</td><td>100.4</td><td>99.7</td><td>99.6</td><td>99.7</td><td>100.5</td><td>99.8</td><td>100.3</td><td>100.1</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Mg-no.</td><td>79.4</td><td>82.8</td><td>79.7</td><td>77.3</td><td>76.8</td><td>73.1</td><td>83.2</td><td>80.6</td><td>84.1</td><td>77.2</td><td>83.9</td><td>84.3</td><td>83.6</td><td>81.4</td><td>82.6</td><td>84.7</td><td>84.1</td><td>85.4</td><td>85.0</td><td>84.0</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Al(z)</td><td>2.6</td><td>3.7</td><td>6.8</td><td>4.0</td><td>2.3</td><td>6.4</td><td>8.5</td><td>7.5</td><td>6.8</td><td>6.4</td><td>3.0</td><td>6.6</td><td>3.2</td><td>6.8</td><td>3.7</td><td>6.2</td><td>2.9</td><td>6.6</td><td>2.9</td><td>6.1</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Ti/Al</td><td>0.30</td><td>0.26</td><td>0.29</td><td>0.26</td><td>0.32</td><td>0.31</td><td>0.28</td><td>0.27</td><td>0.26</td><td>0.28</td><td>0.22</td><td>0.20</td><td>0.25</td><td>0.22</td><td>0.25</td><td>0.21</td><td>0.23</td><td>0.23</td><td>0.24</td><td>0.20</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>X(En)</td><td>0.48</td><td>0.47</td><td>0.47</td><td>0.53</td><td>0.54</td><td>0.42</td><td>0.42</td><td>0.42</td><td>0.44</td><td>0.43</td><td>0.46</td><td>0.43</td><td>0.47</td><td>0.42</td><td>0.46</td><td>0.44</td><td>0.50</td><td>0.43</td><td>0.48</td><td>0.43</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>X(Wo)</td><td>0.40</td><td>0.40</td><td>0.36</td><td>0.29</td><td>0.30</td><td>0.42</td><td>0.43</td><td>0.43</td><td>0.43</td><td>0.41</td><td>0.44</td><td>0.46</td><td>0.42</td><td>0.46</td><td>0.42</td><td>0.46</td><td>0.40</td><td>0.46</td><td>0.43</td><td>0.46</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>X(Fs)</td><td>0.13</td><td>0.12</td><td>0.17</td><td>0.17</td><td>0.16</td><td>0.15</td><td>0.15</td><td>0.15</td><td>0.14</td><td>0.16</td><td>0.09</td><td>0.11</td><td>0.11</td><td>0.12</td><td>0.12</td><td>0.10</td><td>0.10</td><td>0.11</td><td>0.10</td><td>0.11</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr></tbody></table><table><thead><tr><th rowspan="3">Sample:</th><th colspan="12">Peripheral eruptions (NWRZ)</th></tr><tr><th colspan="4">Petit Crate` re (PC)</th><th colspan="4">Petit Crate` re (PC)</th><th colspan="4">Chisny (CHI)</th><th colspan="2">Hau` y (HA)</th></tr><tr><th>Pheno Core Dark</th><th>Pheno Core Bright</th><th>Pheno Core Dark</th><th>Pheno Core Bright</th><th>PC4814 Pheno Core Bright</th><th>PC4814 Pheno Core Dark</th><th>Pheno Core Bright</th><th>Pheno Rim</th><th>Pheno Rim</th><th>Pheno Rim</th><th>16C1 Micro</th><th>Micro</th><th>REU49 Micro</th><th>Micro</th></tr></thead><tbody><tr><td>SiO2</td><td>52.8</td><td>49.4</td><td>52.7</td><td>48.9</td><td>52.8</td><td>48.7</td><td>52.9</td><td>50.3</td><td>49.4</td><td>48.6</td><td>51.0</td><td>45.4</td><td>44.7</td><td>44.5</td><td>44.6</td></tr><tr><td>Al2O3</td><td>2.2</td><td>5.5</td><td>2.4</td><td>6.1</td><td>2.4</td><td>5.8</td><td>2.3</td><td>5.0</td><td>4.6</td><td>4.7</td><td>2.5</td><td>4.8</td><td>5.8</td><td>11.3</td><td>9.6</td></tr><tr><td>TiO2</td><td>0.6</td><td>1.4</td><td>0.5</td><td>1.4</td><td>0.7</td><td>1.7</td><td>0.5</td><td>1.1</td><td>1.4</td><td>1.9</td><td>1.2</td><td>3.8</td><td>4.7</td><td>4.0</td><td>3.9</td></tr><tr><td>CaO</td><td>17.2</td><td>19.6</td><td>17.7</td><td>19.2</td><td>17.7</td><td>20.5</td><td>17.5</td><td>19.8</td><td>20.5</td><td>20.5</td><td>19.6</td><td>19.8</td><td>20.7</td><td>15.7</td><td>14.4</td></tr><tr><td>Na2O</td><td>0.2</td><td>0.3</td><td>0.3</td><td>0.4</td><td>0.2</td><td>0.3</td><td>0.2</td><td>0.3</td><td>0.3</td><td>0.3</td><td>0.2</td><td>0.5</td><td>0.6</td><td>1.0</td><td>0.8</td></tr><tr><td>MgO</td><td>18.9</td><td>15.6</td><td>18.7</td><td>15.3</td><td>18.4</td><td>14.9</td><td>19.0</td><td>15.8</td><td>14.9</td><td>13.5</td><td>15.5</td><td>9.3</td><td>10.2</td><td>11.8</td><td>13.9</td></tr><tr><td>FeO</td><td>7.1</td><td>7.0</td><td>6.7</td><td>7.3</td><td>7.2</td><td>6.8</td><td>6.7</td><td>6.3</td><td>7.6</td><td>9.7</td><td>9.0</td><td>15.1</td><td>12.1</td><td>11.3</td><td>11.3</td></tr><tr><td>Cr2O3</td><td>0.4</td><td>0.6</td><td>0.6</td><td>0.4</td><td>0.3</td><td>0.7</td><td>0.5</td><td>0.9</td><td>0.3</td><td>0.1</td><td>0.0</td><td></td><td></td><td></td><td></td></tr><tr><td>Total</td><td>99.7</td><td>99.6</td><td>99.8</td><td>99.1</td><td>99.9</td><td>99.4</td><td>99.8</td><td>99.6</td><td>99.0</td><td>99.4</td><td>99.2</td><td>99.3</td><td>99.1</td><td>100.3</td><td>99.1</td></tr><tr><td>Mg-no.</td><td>83.5</td><td>82.7</td><td>84.8</td><td>81.5</td><td>82.4</td><td>83.3</td><td>84.5</td><td>82.3</td><td>81.0</td><td>75.1</td><td>77.2</td><td>56.5</td><td>64.1</td><td>67.4</td><td>72.0</td></tr><tr><td>Al(z)</td><td>2.5</td><td>6.5</td><td>2.9</td><td>6.9</td><td>2.5</td><td>7.1</td><td>2.6</td><td>5.6</td><td>5.5</td><td>5.6</td><td>2.8</td><td>5.8</td><td>6.6</td><td>10.9</td><td>10.0</td></tr><tr><td>Ti/Al</td><td>0.17</td><td>0.16</td><td>0.14</td><td>0.15</td><td>0.17</td><td>0.18</td><td>0.15</td><td>0.15</td><td>0.19</td><td>0.26</td><td>0.31</td><td>0.51</td><td>0.52</td><td>0.23</td><td>0.26</td></tr><tr><td>X(En)</td><td>0.54</td><td>0.46</td><td>0.53</td><td>0.46</td><td>0.52</td><td>0.45</td><td>0.54</td><td>0.47</td><td>0.44</td><td>0.40</td><td>0.45</td><td>0.29</td><td>0.32</td><td>0.40</td><td>0.45</td></tr><tr><td>X(Wo)</td><td>0.35</td><td>0.42</td><td>0.36</td><td>0.42</td><td>0.36</td><td>0.44</td><td>0.36</td><td>0.42</td><td>0.44</td><td>0.44</td><td>0.41</td><td>0.44</td><td>0.47</td><td>0.38</td><td>0.34</td></tr><tr><td>X(Fs)</td><td>0.11</td><td>0.12</td><td>0.11</td><td>0.12</td><td>0.12</td><td>0.11</td><td>0.11</td><td>0.11</td><td>0.13</td><td>0.16</td><td>0.15</td><td>0.26</td><td>0.21</td><td>0.21</td><td>0.21</td></tr></tbody></table>

(continued)
<table><caption>Table 2: Continued</caption><thead><tr><th rowspan="3">Sample:</th><th colspan="12">Cumulates</th></tr><tr><th colspan="2">Dunites</th><th colspan="4">Wehrlites</th><th colspan="8">Gabbros</th></tr><tr><th>Pheno Core</th><th>Pheno Core</th><th>Pheno Core</th><th>Pheno Core</th><th>Pheno Core</th><th>Pheno Rim</th><th>Pheno Core Dark</th><th>Pheno Core Bright</th><th>Pheno Core Dark</th><th>Pheno Core Bright</th><th>Pheno Core Dark</th><th>Pheno Core Bright</th><th>Pheno Rim</th><th>Pheno Rim</th><th>Pheno Rim</th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>52.0</td><td>53.0</td><td>49.8</td><td>51.5</td><td>51.2</td><td>53.3</td><td>51.3</td><td>47.1</td><td>51.5</td><td>48.2</td><td>51.1</td><td>48.5</td><td>47.7</td><td>51.0</td><td>50.1</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>2.4</td><td>2.0</td><td>3.9</td><td>2.8</td><td>2.7</td><td>1.8</td><td>2.8</td><td>6.7</td><td>2.9</td><td>5.8</td><td>2.8</td><td>5.6</td><td>5.8</td><td>3.3</td><td>3.8</td></tr><tr><td>TiO<sub>2</sub></td><td>0.6</td><td>0.3</td><td>1.2</td><td>0.7</td><td>0.7</td><td>0.5</td><td>1.2</td><td>2.4</td><td>1.1</td><td>2.2</td><td>1.2</td><td>1.8</td><td>2.3</td><td>1.1</td><td>1.4</td></tr><tr><td>CaO</td><td>22.4</td><td>23.7</td><td>21.7</td><td>23.1</td><td>22.3</td><td>22.2</td><td>21.0</td><td>21.1</td><td>20.7</td><td>21.7</td><td>21.1</td><td>21.6</td><td>21.8</td><td>21.6</td><td>21.8</td></tr><tr><td>Na<sub>2</sub>O</td><td>0.3</td><td>0.5</td><td>0.4</td><td>0.3</td><td>0.3</td><td>0.2</td><td>0.3</td><td>0.4</td><td>0.3</td><td>0.4</td><td>0.2</td><td>0.4</td><td>0.4</td><td>0.3</td><td>0.3</td></tr><tr><td>MgO</td><td>17.2</td><td>16.7</td><td>15.7</td><td>16.7</td><td>16.3</td><td>17.4</td><td>16.2</td><td>13.9</td><td>16.3</td><td>13.9</td><td>15.9</td><td>14.1</td><td>14.0</td><td>15.5</td><td>14.9</td></tr><tr><td>FeO</td><td>4.1</td><td>3.5</td><td>5.2</td><td>4.5</td><td>4.4</td><td>5.1</td><td>6.6</td><td>7.5</td><td>6.7</td><td>7.0</td><td>6.3</td><td>6.7</td><td>6.9</td><td>6.5</td><td>6.6</td></tr><tr><td>Cr<sub>2</sub>O<sub>3</sub></td><td>0.9</td><td>1.0</td><td>0.9</td><td>1.0</td><td>1.0</td><td>0.8</td><td>0.2</td><td>0.5</td><td>0.2</td><td>0.4</td><td>0.2</td><td>0.4</td><td>0.5</td><td>0.2</td><td>0.2</td></tr><tr><td>Total</td><td>100.0</td><td>100.8</td><td>98.9</td><td>100.7</td><td>99.4</td><td>101.4</td><td>99.8</td><td>99.6</td><td>99.9</td><td>99.8</td><td>98.9</td><td>99.3</td><td>99.6</td><td>99.7</td><td>99.3</td></tr><tr><td>Mg-no.</td><td>94.3</td><td>96.4</td><td>90.8</td><td>95.2</td><td>92.3</td><td>89.4</td><td>85.3</td><td>83.7</td><td>84.4</td><td>83.2</td><td>83.5</td><td>83.2</td><td>85.1</td><td>84.6</td><td>83.7</td></tr><tr><td>Al(z)</td><td>3.9</td><td>3.2</td><td>5.4</td><td>4.8</td><td>4.2</td><td>2.8</td><td>3.4</td><td>8.4</td><td>3.5</td><td>7.0</td><td>3.2</td><td>6.7</td><td>7.4</td><td>3.9</td><td>4.6</td></tr><tr><td>Ti/Al</td><td>0.15</td><td>0.11</td><td>0.20</td><td>0.15</td><td>0.17</td><td>0.17</td><td>0.27</td><td>0.23</td><td>0.24</td><td>0.24</td><td>0.26</td><td>0.20</td><td>0.25</td><td>0.22</td><td>0.23</td></tr><tr><td>X(En)</td><td>0.48</td><td>0.47</td><td>0.46</td><td>0.47</td><td>0.47</td><td>0.48</td><td>0.46</td><td>0.42</td><td>0.47</td><td>0.42</td><td>0.46</td><td>0.42</td><td>0.42</td><td>0.45</td><td>0.43</td></tr><tr><td>X(Wo)</td><td>0.45</td><td>0.48</td><td>0.46</td><td>0.46</td><td>0.46</td><td>0.44</td><td>0.43</td><td>0.46</td><td>0.43</td><td>0.47</td><td>0.44</td><td>0.47</td><td>0.47</td><td>0.45</td><td>0.46</td></tr><tr><td>X(Fs)</td><td>0.06</td><td>0.06</td><td>0.08</td><td>0.07</td><td>0.07</td><td>0.08</td><td>0.11</td><td>0.13</td><td>0.11</td><td>0.12</td><td>0.10</td><td>0.11</td><td>0.12</td><td>0.10</td><td>0.11</td></tr></tbody></table>

<footnote>*Published data for olivines in lavas emitted in the central area (Peltier, 2007; Boivin & Bache' lery, 2009). Pheno, phenocryst; Micro, microlite.</footnote>
<figure><img src="image_8.png" /></figure>

Fig. 6. Composition of clinopyroxenes from Piton de la Fournaise lavas and enclaves. (a) Proportion of aluminum cati-ons occupying tetrahedral sites [Al(z)] linearly correlates with TiO₂ content (wt %) in the cores of clinopyroxene phenocrysts. The composition of dark and bright sectors in sector-zoned clinopyroxene cores in 'clinopyroxene-bearing' lavas from per-ipheral vents, and gabbroic xenoliths (peripheral) show a large scatter, which overlaps the fields of magnesian augites and augites found in lavas emitted in the central area (1977–2010). The Na distribution spans a large range in sector-zoned clino-pyroxene, whereas a normal distribution is observed in clino-pyroxene from lavas emitted in the central area and, to a lesser extent, in the Bellecombe gabbros. (b) Ti/Al ratio (thermobar-ometer proxy; Bouhmedi, 1988; Cundari & Salviulo, 1989) versus Mg-number (temperature proxy) in pyroxenes. Mg-number [Mg²⁺/(Mg²⁺ + Fe²⁺)] is calculated from Fe²⁺ deter-mination in clinopyroxene based on Droop (1987). Mg-number decrease correlates with Ti/Al increase in rims (normal zoning) and microlites. (1) Clinopyroxenes in dunite and wehrlite xeno-liths (Brugier, 2016) have the highest Mg-numbers and some of the lowest Ti/Al ratios. (2) Clinopyroxene in 'clinopyroxene-bearing' lavas from peripheral vents and in Bellecombe gab-bros have a composition that overlaps with that of the most magnesian clinopyroxenes found in lavas from the central area (1977–2010; Peltier, 2007; Boivin & Bache' lery, 2009; Di Muro et al., 2015). (3) Clinopyroxenes in 1998 'clinopyroxene-bear-ing' lavas have high Ti/Al ratios and lower Mg-numbers, con-sistent with their crystallization in relatively evolved melts (Boivin & Bache' lery, 2009). (4) Microlites from 'olivine-bearing' basalts emitted in the peripheral area (CHI, HA) have the high-est Ti/Al ratios and lowest Mg-numbers, consistent with late crystallization from evolved melts.

(2 8 60 6 kbar) is documented from crystals in Bellecombe gabbros. In 1977–2010 lavas from the cen-tral area, Na contents are close to those found in clino-pyroxene from the Bellecombe gabbros (Fig. 6a; Table 2). There, the clinopyroxene-only hydrous barom-eter provides a main mode at 2 6 6 1 1 kbar.

Clinopyroxene compositions in 1998 'clinopyroxene-
bearing' lavas returned an average pressure of 3 6 6 0 9
kbar, in agreement with H₂O–CO₂ saturation pressures
estimated from melt and fluid inclusions in MgO-rich olivine
(Bureau et al., 1999). Similar relative pressure indicators
were found when considering the correlation between
clinopyroxene Mg-number and Ti/ Al ratio (Fig. 6b). For a
given Mg-number, and at con-stant oxygen fugacity, Ti
content and thus Ti/Al de-crease with increasing pressure
(Adam & Green, 1994; Pichavant et al., 2016), whereas
isobaric melt evolution (traced by Mg-number) can produce
variable Ti/Al ratios depending on the other crystallizing
phases (Boivin & Bache` lery, 2009; Di Muro et al., 2015).
Average Ti/Al ratios at constant Mg-number increase
(mostly related to Ti variations in our case) from
clinopyroxene crystals in 'clinopyroxene-bearing' lavas of
peripheral eruptions to those in Bellecombe gabbros and in
lavas emitted from the central area (Fig. 6b).

In summary, our results suggest that clinopyroxene
crystallization occurred over a large pressure range, but
always <9 kbar and mostly <6 kbar. Clinopyroxene
crystallization pressures partially overlap in lavas from the
central and peripheral areas, but the main modes of 3–6
kbar in peripheral basalts and ultramafic enclaves (dunites
and wehrlites) are systematically higher than those for the
lavas from the central area and in the Bellecombe gabbros.
It is known that clinopyroxene barometry needs additional
experiments to be refined; however, we underline that both
fluid inclusion data and AlphaMELTS modelling corroborate
our clinopyroxene-derived pressures (see below).

Fluid inclusions

Primary fluid inclusions in crystals are rarely observed in 'clinopyroxene-bearing' lavas from peripheral erup-tions (Fig. 8a and b). In this group of lavas, trails of sec-ondary fluid inclusions (Bodnar, 2003) are found only in rare subhedral olivine antecrysts (Fig. 8c). In olivine crystals from 'olivine-bearing' lavas from peripheral eruptions, rare primary fluid inclusions are present, but trails of secondary fluid inclusions are frequent. Conversely, crystals in wehrlites contain both abundant primary and secondary fluid inclusions (Fig. 8d). In gen-eral, the CO₂ density in primary fluid inclusions in oliv-ine (0 28–0 53 g cm⁻³) is higher than that in clinopyroxene (0 10–0 41 g cm⁻³). We suspect that this difference reflects, at least in part, advanced re-equilibration of fluid inclusions in clinopyroxene com-pared with those in olivine. This process might be favored by the tendency of clinopyroxene to deform and to crack causing poorer sealing capability com-pared with olivine.

In crystals from 'clinopyroxene-bearing' lavas from
peripheral eruptions, CO₂ densities in fluid inclusions range
from 0 18 to 0 58 g cm⁻³ (equivalent to 0 7–2 9 kbar
pressure). Higher-pressure fluid inclusions (up to 4 0 kbar)
were observed in rare Fo₈₄ subhedral olivine
<figure><img src="image_9.png" /></figure>

Fig. 7. Clinopyroxene crystallization pressures in Piton de la Fournaise magmas. (a, d) Clinopyroxene-melt geothermobarometry [equation (30), Putirka et al., 2003] based on iterative determination of pressure and temperature. Only clinopyroxene cores in equi-librium with the bulk-rocks (Kd 1/4 0 27 60 03) are reported here (Putirka, 2008b). Clinopyroxene microlites from 'olivine-bearing basalts' emitted in the peripheral area (Chisny, Hau' y) and in 1998 'clinopyroxene-bearing' lavas emitted at the Hudson vent in the central area are in disequilibrium with the bulk-rock (equilibrium with <5 wt % MgO melts). (b, c, e) Pressure distribution from clino-pyroxene-only geothermobarometry [equation (32b), Putirka, 2008b] was calculated for hydrous (H₂O 1/4 1 wt %; average in Piton de la Fournaise melts; Di Muro et al., 2016) melts with equation (32c). Consistent and similar pressure ranges are obtained with equa-tion (32a) calibrated for dry melts. Temperature is estimated from bulk-rock compositions based on the MgO thermometer devel-oped by Helz & Thornber (1987). For gabbroic cumulates, temperature was fixed at 1200 C in accordance with the average temperature of Piton de la Fournaise magmas (Pichavant et al., 2016). A higher temperature (1250 C) was computed for dunites and wehrlites corresponding to the average temperature of clinopyroxene appearance (see Modeling section). Average pressures (and 1r variations on the average) are reported for each distribution.

antecrysts (Fig. 4a; CO₂ density 0 68 g cm⁻³). In contrast, micro-Raman data on fluid inclusions in olivine phenocrysts in 'olivine-bearing' lavas from peripheral erup-tions display the highest pressures (mostly 1 7–4 2 kbar; Table 3). Our new results are consistent with the pres-sure range (3 2–5 7 kbar) previously recorded in olivine crystals from oceanites of the peripheral Piton de Caille eruption (Bureau et al., 1998). Fluid inclusions from crystals in dunitic and wehrlitic enclaves record a large range of densities from 0 31 to 0 53 g cm⁻³ (1 3–2 5 kbar) for olivine and from 0 11 to 0 45 g cm⁻³ for clinopyrox-ene (0 4–2 0 kbar). These values are consistent with

thermo-barometric results obtained by Bureau et al. (1998) on Chisny dunites (0 9–4 8 kbar). A review of pub-lished data (recalculated from Famin et al., 2009) shows that fluid inclusions in olivine crystal found in lavas from the central area have very low densities, consist-ent with trapping and/or re-equilibration at a pressure range lower than 0 9 kbar, and in most cases <0 7 kbar. This range is consistent with the low CO₂–H₂O satur-ation pressures recorded by melt inclusions hosted in olivine from lavas emitted from the central area (Table 4; Di Muro et al., 2016). The 1998 'clinopyroxene-bearing' lavas at the Hudson vent represent a rare
<figure><img src="image_10.png" /><figcaption>Fig. 8. Transmitted light photomicrographs of fluid inclusions (FI) in crystals from Piton de la Fournaise lavas. (a) Primary fluid in-clusion in an olivine dendrite from a ‘clinopyroxene-bearing’ lava from the peripheral area (e.g. Petit Crate` re). (b) Primary fluid inclusions in a clinopyroxene rim from a ‘clinopyroxene-bearing’ lava from the peripheral area (Formica Le´ o). (c) Trail of secondary fluid inclusions in an Fo84 olivine antecryst from a ‘clinopyroxene-bearing’ lava from the peripheral area (Formica Le´ o). (d) Abundant primary and secondary fluid inclusions in anhedral olivine within wehrlitic enclaves from Chisny in the peripheral area.</figcaption></figure>

(e) Raman spectra of the Fermi dyad of CO2 in fluid inclusions exhibit well measurable shifts from low density (e.g. secondary fluid inclusion in olivine from wehrlite) to high density (e.g. secondary fluid inclusion in subhedral olivine antecryst).

exception, featuring fluid inclusions with high CO2 density in olivine crystals in lavas from the central area. There, fluid inclusions in MgO-rich olivine crystals sug-gest that magma crystallized up to a maximum pres-sure of 4 4 kbar (Table 4; Bureau et al., 1999).

Bulk-rocks

Lavas from both peripheral and central areas are transi-
tional basalts (Fig. 9a), with slight tholeiitic or alkaline
affinities, whereas only rare peripheral eruptions pro-duced
basalts with an alkaline affinity. Our data partly contradict
the conclusions drawn by Salau¨ n et al. (2010)

and Le´ nat et al. (2012), who proposed that all lavas from
the peripheral area display an alkaline signature.

Peripheral area

'Clinopyroxene-bearing' basalts (n ¼ 6) from peripheral eruptions are compositionally variable (Figs 9-11) and span the full range from transitional basalts with a tho-leiitic affinity [sample Petit Crate` re (PC); MgO ¼ 7 5 wt %, K2O ¼ 0 58 wt %] to transitional basalts with an alka-line affinity [sample Crate` re Commerson (CC); MgO ¼ 9 3 wt %, K2O ¼ 0 81 wt %]. However, the majority of these basalts have a tholeiitic affinity (Fig. 9a) with a CaO/Al2O3 ratio greater than 0 75 (Fig. 9), no Eu
<table><caption>Table 3: Density of fluid inclusions in olivines and clinopyroxenes of Piton de la Fournaise lavas (Lamadrid et al., 2016)</caption><thead><tr><th rowspan="3">Sample</th><th rowspan="3">Host mineral</th><th colspan="8">Number Fluid inclusion barometry</th><th colspan="2">Clinopyroxene barometry [equation (32b)]</th></tr><tr><th colspan="2">Dmin (cm<sup>-1</sup>)</th><th colspan="2">Dmax (cm<sup>-1</sup>)</th><th colspan="2">dmin (g cm<sup>-3</sup>)</th><th colspan="2">dmax (g cm<sup>-3</sup>)</th><th rowspan="2">P (kbar)</th><th rowspan="2">Z (km b.s.l.)</th><th rowspan="2">P (kbar)</th><th rowspan="2">Z (km b.s.l.)</th></tr><tr><th></th><th></th><th></th><th></th><th></th><th>pure CO<sub>2</sub></th><th>pure CO<sub>2</sub></th><th>p 10% H<sub>2</sub>O</th><th>p 10% H<sub>2</sub>O</th></tr></thead><tbody><tr><td colspan="15"><strong>Peripheral eruptions (NWRZ)</strong></td></tr><tr><td colspan="15"><strong>Clinopyroxene-bearing lavas</strong></td></tr><tr><td>Petit Crate<sup>*</sup> re</td><td>ol</td><td>3</td><td>104.0</td><td>104.2</td><td>0.49</td><td>0.55</td><td>0.51</td><td>0.58</td><td>24-2.960.7</td><td>6.7-11.0</td><td>33605</td><td>8.5-11.6</td></tr><tr><td>Formica Le´o</td><td>cxpx</td><td>6</td><td>103.1</td><td>104.1</td><td>0.17</td><td>0.53</td><td>0.18</td><td>0.55</td><td>0.7-2.760.7</td><td>&lt;10.4</td><td>22603</td><td>5.7-7.6</td></tr><tr><td>Formica Le´o</td><td>ol (antecryst)</td><td>12</td><td>103.8</td><td>104.5</td><td>0.41</td><td>0.65</td><td>0.43</td><td>0.68</td><td>1.8-4.060.7</td><td>2.1-14.4</td><td></td><td></td></tr><tr><td colspan="15"><strong>Olivine-bearing lavas</strong></td></tr><tr><td>Chisny</td><td>ol</td><td>16</td><td>103.7</td><td>104.1</td><td>0.39</td><td>0.62</td><td>0.41</td><td>0.65</td><td>1.7-3.960.7</td><td>1.8-14.1</td><td></td><td></td></tr><tr><td>Hau´y</td><td>ol</td><td>5</td><td>104.4</td><td>104.5</td><td>0.63</td><td>0.65</td><td>0.65</td><td>0.69</td><td>3.9-4.260.7</td><td>9.8-15.0</td><td></td><td></td></tr><tr><td>Piton de Caille*</td><td>ol</td><td></td><td></td><td></td><td>0.57</td><td>0.77</td><td></td><td>0.81</td><td>3.1-5.7</td><td>9.4-17.5</td><td></td><td></td></tr><tr><td colspan="15"><strong>Cumulates</strong></td></tr><tr><td>Gabbros</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>28606</td><td>6.7-10.4</td></tr><tr><td>Wehrlite</td><td>ol</td><td>14</td><td>103.3</td><td>104.1</td><td>0.26</td><td>0.51</td><td>0.28</td><td>0.53</td><td>1.1-2.560.7</td><td>0.3-9.7</td><td></td><td></td></tr><tr><td>Wehrlite</td><td>cxpx</td><td>15</td><td>102.9</td><td>103.7</td><td>0.10</td><td>0.39</td><td>0.10</td><td>0.41</td><td>0.4-2.060.7</td><td>&lt;8.2</td><td>56608</td><td>14.7-19.7</td></tr><tr><td>Dunite*</td><td>ol</td><td></td><td></td><td></td><td>0.27</td><td>0.71</td><td></td><td></td><td>1.3-4.860.7</td><td>0.3-16.4</td><td></td><td></td></tr><tr><td colspan="15"><strong>Central eruptions</strong></td></tr><tr><td>March 1998 (H)*</td><td>ol</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>3.5-4.4</td><td>8.5-15.6</td><td>36609</td><td>9.1-13.8</td></tr><tr><td>2005</td><td>ol</td><td></td><td>102.0</td><td>103.2</td><td>0.00</td><td>0.21</td><td>0.00</td><td>0.23</td><td>0.0-0.960.7</td><td>&lt;4.8</td><td>26611</td><td>3.7-11.2</td></tr></tbody></table>

Corresponding depth (Z) below sea level (b.s.l.) was calculated assuming an average density of 2.70 g cm⁻³ for the Piton de la Fournaise edifice (Rousset et al., 1989) extending to 4 km b.s.l. (Fontaine et al., 2015), 2.89 g cm⁻³ for the 6 km thick oceanic crust (Carlson & Raskin, 1984; Gallart et al., 1999; Fontaine et al., 2015) and 3.30 g cm⁻³ for the upper lithospheric man-tle (Wang, 1970). Densities of fluid inclusions in olivines from lavas emitted during the 2005 central eruption were measured by Raman spectroscopy (Famin et al., 2009) and were here recalculated (see Analytical Methods). The maximum error (1r) on pressure estimates from fluid inclusions is reported (0.7 kbar; see Supplementary Data Electronic Appendix 3). Clinopyroxene barometric estimates (average pressure) with the variability (1r) on their average are shown for comparison.

<footnote>*Density of fluid inclusions in crystals from dunites, Piton de Caille and 1998 Hudson lavas were measured by microthermometry (Bureau et al., 1998, 1999).</footnote>
<table><caption>Table 4: Compilation of barometric data available for Piton de la Fournaise lavas</caption><thead><tr><th rowspan="3">Sample</th><th colspan="2">Barometry</th><th colspan="2">Average pressure and<br/>1r (kbar) from cpx barometry</th><th colspan="2">Maximum pressure (kbar)<br/>from fluid and melt inclusions</th><th rowspan="3">Clinopyroxene composition<br/>CaO/Al<sub>2</sub>O<sub>3</sub> ratio<br/>of clinopyroxenes</th></tr><tr><th rowspan="2">Ti/Al ratio of<br/>clinopyroxenes</th><th colspan="2">Cpx-bulk-rock</th><th colspan="2">Cpx-only</th><th rowspan="2">FI</th><th rowspan="2">MI</th></tr><tr><th></th><th></th><th></th><th></th></tr></thead><tbody><tr><td colspan="8"><strong>Peripheral eruptions (NWRZ)</strong></td></tr><tr><td colspan="8"><strong>Clinopyroxene-bearing lavas</strong></td></tr><tr><td>Petit Crate<sup>*</sup></td><td>0.17</td><td>6005</td><td>3.8</td><td>6.06</td><td>3.3</td><td>6.05</td><td>ol</td><td>29607</td><td>—</td><td>—</td><td>51617</td></tr><tr><td>Formica Le´o</td><td>0.22</td><td>6002</td><td>4.0</td><td>6.05</td><td>2.2</td><td>6.03</td><td>px</td><td>27607</td><td>—</td><td>—</td><td>55619</td></tr><tr><td>Formica Le´o (olivine antecryst)</td><td>—</td><td></td><td>—</td><td></td><td>—</td><td></td><td>ol</td><td>40607</td><td>—</td><td>—</td><td></td></tr><tr><td colspan="8"><strong>Olivine-bearing lavas</strong></td></tr><tr><td>Chisny<sup>*</sup></td><td>—</td><td></td><td>—</td><td></td><td>—</td><td></td><td>ol</td><td>39607</td><td>—</td><td>—</td><td>—</td></tr><tr><td>Hau´y<sup>*</sup></td><td>—</td><td></td><td>—</td><td></td><td>—</td><td></td><td>ol</td><td>2607</td><td>—</td><td>—</td><td>—</td></tr><tr><td>Piton de Caille</td><td>—</td><td></td><td>—</td><td></td><td>—</td><td></td><td>ol</td><td>7</td><td>—</td><td>2.9</td><td>—</td></tr><tr><td colspan="8"><strong>Cumulates</strong></td></tr><tr><td>Gabbros</td><td>0.25</td><td>6003</td><td>—</td><td></td><td>2.8</td><td>6.06</td><td>—</td><td>—</td><td>—</td><td>—</td><td>60616</td></tr><tr><td>Wehrlites</td><td>0.16</td><td>6003</td><td>—</td><td></td><td>5.6</td><td>6.08</td><td>ol</td><td>25607</td><td>—</td><td>—</td><td>88620</td></tr><tr><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>cpx</td><td>20607</td><td>—</td><td>—</td><td></td></tr><tr><td>Dunites</td><td>0.13</td><td>6003</td><td>—</td><td></td><td>—</td><td></td><td>ol</td><td>48607</td><td>—</td><td>—</td><td>107605</td></tr><tr><td colspan="8"><strong>Central eruptions</strong></td></tr><tr><td>March 1998 (H)<sup>*</sup></td><td>0.28</td><td>6002</td><td>—</td><td></td><td>3.6</td><td>6.09</td><td>ol</td><td>4.4</td><td>ol</td><td>5.1</td><td>40604</td></tr><tr><td>1977–2010</td><td>0.24</td><td>6005</td><td>3.8</td><td>6.11</td><td>2.6</td><td>6.11</td><td>ol</td><td>09607</td><td>ol</td><td>0.5</td><td>55621</td></tr></tbody></table>

Pressure from fluid inclusions was calculated from CO₂ density (see Supplementary Data Electronic Appendix 3). Volatile saturation pressure in melt inclusion was recalculated using the H₂O–CO₂ solubility model of Papale et al. (2006). Pyroxene barometry used clinopyroxene-only equation (Putirka, 2008b) and clinopyroxene–bulk-rock equilibrium equation (Putirka et al., 2003). CaO/Al₂O₃ ratio in clinopyroxenes is lower in lavas (3.6–5.5) than in cumulates (6.0–10.7), whose pyroxenes have the lowest Ti/Al ratio. No significant variation in the alkalinity index (AI) is observed between clinopyroxenes. Lavas have a tholeiitic affinity, except those marked with *, which have an alkaline affinity. The maximum error (1r) on pressure estimates from fluid inclusions is reported (0.7 kbar; see Supplementary Data Electronic Appendix 3).
<figure><img src="image_11.png" /></figure>

<figure><img src="image_12.png" /></figure>

<figure><img src="image_13.png" /></figure>

Fig. 9. Bulk-rock composition (major elements) of Piton de la Fournaise lavas emitted in both peripheral (this study) and central areas (1972–2010; Vlasté lic et al., 2007; Peltier, 2007; Boivin & Bache` lery, 2009; Salau¨ n et al., 2010; Di Muro et al., 2014, 2015). Black star is the primary melt composition estimated by Fretzdorff & Haase (2002). (a) SiO₂ vs K₂O þ Na₂O showing a slight alkaline affin-ity preferentially in ‘olivine-bearing’ lavas emitted in the peripheral area with respect to ‘clinopyroxene-bearing lavas’ and 1972– 2010 lavas emitted in the central area. (b) MgO vs Al₂O₃ showing a slight Al enrichment in ‘olivine-bearing’ lavas emitted in the per-ipheral area. (c) MgO vs CaO/Al₂O₃ showing SSB- and AbG-type lavas identified at Piton de la Fournaise (Boivin & Bache` lery, 2009; Famin et al., 2009; Le¨ nat et al., 2012) and discussed in the text. Compositions of the Bellecombe gabbros and Chisny wehrlites (arrows pointing outside the plots) are from Upton et al. (2000) and show a large scatter.
<figure><img src="image_14.png" /></figure>

Fig. 10. Bulk-rock composition (trace elements) of Piton de la Fournaise lavas emitted in both peripheral (this study) and central areas (1972–2010; Vlast´e lic et al., 2007; Peltier, 2007; Boivin & Bache´ lery, 2009; Salau¨ n et al., 2010; Di Muro et al., 2014, 2015). Symbols as in Fig. 9. (a) ‘Olivine-bearing’ basalts from peripheral vents are preferentially depleted in trace elements that are com-patible in clinopyroxene (such as Sc). (b) Eu/Eu* ratio [Eu/N/(Sm0.15N Gd0.55N)] where EuN, SmN and GdN are normalized to chondrite; Sun & McDonough, 1989). (c) ‘Olivine-bearing’ basalts from peripheral vents are preferentially enriched in incompatible trace ele-ments such as Th. Trace element composition in Piton de la Fournaise olivine and clinopyroxene crystals from Salau¨ n et al. (2010). Dashed lines represent the potential contamination trend with either residual evolved melts (e.g. April 2007 trachytic pumice; Salau¨ n et al., 2010) invoked to explain the peculiar AbG composition of the 1998 ‘clinopyroxene-bearing’ lavas emitted at the Hudson vent in the central area, or melts from low melting degree of wehrlite cumulates (this study; Pietruszka et al., 2009).
<figure><img src="image_15.png" /></figure>

Fig. 11. Bulk-rock composition (incompatible trace element ratios) of Piton de la Fournaise lavas emitted in both peripheral (this study) and central areas (1972–2010; Vlaste' lic et al., 2007; Peltier, 2007; Boivin & Bache' lery, 2009; Salau' n et al., 2010; Di Muro et al., 2014, 2015). Symbols as in Fig. 9. (a) REE normalized to typical OIB composition (Sun & McDonough, 1989). 'Olivine-bearing' basalts from peripheral vents are preferentially depleted in MREE and HREE with respect to 'clinopyroxene-bearing lavas' and 1972–2010 lavas emitted in the central area. (b) Sr–Nd isotopic variability in Piton de la Fournaise lavas emitted in the peripheral area spans the narrow range identified in previous studies of La Re' union lavas (Fisk et al., 1988; Bosch et al., 2008; Pietruszka et al., 2009). Black triangles represent complementary analyses from other peripheral vents (unpublished data from Trous Blancs, Piton Bleu and Piton Bellevue; Fig. 1). Lowest ⁸⁷Sr/⁸⁶Sr values are found in the 'clinopyroxene-bearing' lavas emitted in the peripheral area having a composition close to that of the lava emitted on 5 April 2007 in the central area. The most radiogenic Sr isotopic ratio is observed in 'olivine-bearing' lavas emitted in the peripheral area. In the peripheral area, intermediate compositions are found overlapping with that documented in recent (1972–2010) lavas emitted in the central area. (c) Positive correlation between Sr iso-tope ratios and ratios of variably incompatible trace elements is consistent with mixing between two end-members with slightly distinct isotopic and trace element signatures, as previously discussed for lavas emitted in the central area by Pietruszka et al. (2009) and Di Muro et al. (2014). Owing to a better compatibility of Yb in clinopyroxene than in olivine, variable Th/Yb ratios at con-stant ⁸⁷Sr/⁸⁶Sr can also record the effect of clinopyroxene fractionation and resorption (e.g. early and late stages of Chisny, CHI and CHI*; see discussion in the text).
depletion, low Th contents and high Sc, middle rare earth element (MREE) and heavy REE (HREE) contents (Figs 10 and 11a). The highest ¹⁴³Nd/¹⁴⁴Nd and the lowest ⁸⁷Sr/⁸⁶Sr compositions are observed in these basalts (Fig. 11b and c).

'Olivine-bearing' basalts (n ¼ 30) from the peripheral area are also compositionally variable (Figs 9–11). For instance, at Chisny, early stage lavas (first phases of activity) are less alkaline and more calcic (sample CHI; K₂O <0.8 wt % and CaO/Al₂O₃ ¼ 0.76–0.04) than late-stage lavas (last phases of activity: sample CHI*; K₂O >0.9 wt % and CaO/Al₂O₃ ¼ 0.68–0.01). 'Olivine-bearing' basalts from the peripheral area have on average higher alkali contents (K₂O from 0.70 to >1 wt %), lower SiO₂ contents and CaO/Al₂O₃ ratios down to 0.66 for a given MgO, compared with the 'clinopyroxene-bearing' basalts from the peripheral area (Fig. 9). They are also characterized by high Th contents and low Sc, MREE and HREE with respect to 'clinopyroxene-bearing' basalts and do not show Eu depletion (Figs 10 and 11a). They are marked by the lowest ¹⁴³Nd/¹⁴⁴Nd and the highest ⁸⁷Sr/⁸⁶Sr found in PdF subaerial lavas (Fig. 11b and c).

We note that some of the 'clinopyroxene-bearing' and many of the 'olivine-bearing' lavas from peripheral eruptions, marked by low CaO/Al₂O₃ ratios [e.g. samples from Formica Le'o (FL) and Demi-Piton(DP*) in Figs 10 and 11] do not systematically display low Th and high MREE and HREE contents. Similarly, the Sr and Nd isotope ratios do not correlate with CaO/Al₂O₃. For example, the early and late-stage lavas of Chisny (CHI and CHI* samples, respectively) are isotopically similar, despite differences in major and trace element abundances and ratios (Figs 8–11b). Sr and Nd isotopes are roughly anti-correlated, but do not correlate significantly with elemental ratios of LREE/HREE, as previously observed by Albare de et al. (1997). We only observe a weak positive correlation between ⁸⁷Sr/⁸⁶Sr and ratios of variably incompatible elements (i.e. Th/Yb; Fig. 11c) as previously observed (Pietruszka et al., 2009; Di Muro et al., 2014).

## Central area

The major and trace element compositions, as well as Sr–Nd isotopes, of the 1972–2010 lavas from central eruptions overlap partly with values for ‘clinopyroxene-bearing’ and ‘olivine-bearing’ basalts from peripheral eruptions (Figs 9–11). However, the 1972–2010 lavas display a slight Eu depletion (Eu/Eu* down to 0.96; Fig. 10b) when compared with basalts from the peripheral area (Eu/Eu* ¼ 1.04–6.01). In particular, the 1998 ‘clinopyroxene-bearing’ lavas from the Hudson vent have CaO/Al₂O₃ ratios, Th and REE contents, and isotope compositions similar to those of some ‘olivine-bearing’ lavas from peripheral eruptions (e.g. CHI*;

Boivin & Bache lery, 2009; Famin et al., 2009; Villemant et al., 2009; Salau n et al., 2010) but are marked by slight-ly stronger Eu depletion (Eu/Eu* ¼ 0.98; Fig. 10b). In contrast, the 2007 ‘olivine-bearing’ lavas erupted on 5 April have a composition (MgO ¼ 8.6 wt % and K₂O ¼ 0.63 wt % in glass) similar to those of some ‘clinopyroxene-bearing’ lavas from the peripheral area with respect to major and trace elements as well as Sr–Nd isotopes.

## Modeling of magma evolution

To constrain the polybaric evolution of PdF magmas we used the AlphaMELTS code (e.g. Ghiorso & Sack, 1995; Ghiorso et al., 2001; Asimow et al., 2003). This thermodynamic model takes into account the role of temperature, pressure, water content and redox state in the evolution of basaltic melts. For the calculations, we used the composition of the primary melt (Supplementary Data Electronic Appendix 5) in equilibrium with the MgO-richest olivine crystals found at PdF (Foglio 4) reported by Fretzdorff & Haase (2002). Among several possible compositions of primary melts estimated for the RE union magmas (Ludden, 1978; Sobolev & Nikogosian, 1994; Villemant et al., 2009), this is the most primitive melt (MgO ¼ 17.2 wt %) and thus the most likely to represent the parental melt composition for deep magmatic crystallization (9 kbar; Fretzdorff & Haase, 2002). We fix the initial temperature at 1350 C using the MgO thermometer of Helz & Thornber (1987), which has been shown to be appropriate for PdF magmas (Di Muro et al., 2016). This value is close to the 1330 C estimate of Fretzdorff & Haase (2002).

In a first set of MELT models (Simulation 1; Supplementary Data Electronic Appendix 5), we simulated isobaric equilibrium crystallization at four pressures (9, 6, 4 and 3 kbar) spanning the range constrained by clinopyroxene and fluid inclusion barometry (Fig. 12a). Model 1 (9 kbar) corresponds to the deepest potential storage zone suggested by Albare de et al. (1997) and Fretzdorff & Haase (2002) and to the highest pressure estimated by pyroxene barometry (Supplementary Data Electronic Appendix 5). Model 2 (6 kbar) corresponds to the highest pressure recorded by fluid inclusions in crystals from lavas from the peripheral area (Table 4). Model 3 (4 kbar) is close to the pressure recorded by the majority of fluid inclusions and is consistent with the depth range of the underplating layer identified by geophysical measurements (Charvis et al., 1999; Fontaine et al., 2015). Finally, Model 4 (3 kbar) accounts for pressures recorded in ‘clinopyroxene-bearing’ lavas from the peripheral area. An initial oxygen fugacity of NNO – 0.5 (where NNO is nickel–nickel oxide buffer) was selected in accordance with the recent review of Pichavant et al. (2016). Simulations were also run changing the initial dissolved water content from (1) dry conditions to (2) intermediate water contents (0.8 wt %), according to the average concentrations measured in most melt inclusions at PdF (Di
<figure><img src="image_16.png" /></figure>

<figure><img src="image_17.png" /></figure>

<figure><img src="image_18.png" /></figure>

Fig. 12. AlphaMELTS modeling of Piton de la Fournaise magma evolution as a function of (a) oxygen fugacity, water content and pressure. Primary melt (black star) estimated by Fretzdorff & Haase (2002) for La Re´ union Island is based on equilibrium with the most MgO-rich olivine found in submarine lavas (Fo90 4). Temperature of clinopyroxene first appearence and modeled cotectic oliv-ine composition are indicated in italics for each simulation. (b) CaO/Al₂O₃ and (c) Al₂O₃ evolution paths fitting the composition of ‘clinopyroxene-bearing’ lavas (3–4 kbar) and ‘olivine-bearing’ lavas (6 kbar) emitted in the peripheral area are reproduced starting from a primitive melt (white star). Primitive melt composition results from 19 wt % olivine fractionation from the primary melt of Fretzdorff & Haase (2002). The composition of this primitive melt (10 3 wt % MgO; 1 0 wt % H₂O; NNO – 0 5) is consistent with those of the most magnesian melt inclusions found in olivines (Bureau et al., 1998; Di Muro et al., 2016). Olivine–clinopyroxene crystal-lization (>3 kbar) fits with the composition of lavas emitted in the peripheral area. Differentiation at 3 kbar can reproduce the evolved glass composition found in the Bellecombe gabbros (Upton et al., 2000). The addition of olivine antecrysts controls part of the variability of bulk-rock compositions. MgO-rich olivine antecrysts (Fo84 as modeled at 4 kbar, consistent with the pressure recorded by fluid inclusions) are indeed observed in Formica Le´ o lavas and part of the geochemical variability of lavas from periph-eral vents can be explained by addition of <10% of olivine antecrysts (this study).
Muro et al., 2016) and (3) high water contents (13 wt %) as estimated by Di Muro et al. (2014) for the 5 April 2007 deep magma input into the central area. The range of water concentration is slightly higher than that estimated by Putirka (2008a) for the primary magmas of La Re´ union (H₂O ¼ 0.52 wt %). However, Putirka's calculations were made for hot (1573 C) asthenospheric melts produced by 13.5% mantle partial melting, and did not take into account potential water enrichment of melts by crystallization in the pressure range between 40 and 10 kbar. Five runs were performed at 9 kbar to assess the influence of water content and oxidation state on phase equilibria (Fig. 12a; Supplementary Data Electronic Appendix 5, Simulation 1). In all runs, the first liquidus phase to form is olivine, followed by clinopyroxene. At 9 kbar, in dry basalts, Fo₈₉ olivine appears at 1350 C and clinopyroxene (augite) at 1330 C. Decreasing pressure to 3 kbar delays pyroxene appearance and expands significantly the field of olivine stability. At 3 kbar, coticetic Fo₈₄ olivine and clinopyroxene (augite) form at 1140 C. Our modeling predicts that the molar Ti/Al ratio in clinopyroxenes increases from 0.14 to 0.25 with decreasing pressure. An increasing amount of dissolved water (up to 13 wt %) further delays clinopyroxene crystallization by about 90 C and thus may play an important role. Changes in oxidation state have only a minor effect and change only slightly the olivine composition (Supplementary Data Electronic Appendix 5). Regardless of the assumed pressure, early olivine is very magnesian (Fo₉₁–₈₉). Such MgO-rich olivine crystals, however, are only rarely found in PdF lavas (Di Muro et al., 2016). Small MgO-rich olivine grains in AbG Hudson lavas from the central area (Bureau et al., 1999) and in lavas from the Hau´y peripheral vent (Fig. 5b) show a composition up to Fo₈₈ but are not in equilibrium with their host (bulk-rocks and matrix glasses).

To explore the potential effect of deep magma stor-age and early olivine fractionation, we performed a se-cond set of runs starting with a lower MgO content for the parental magma (Simulation 2; Supplementary Data Electronic Appendix 5). In these runs, we fixed the start-ing melt composition (MgO ¼ 10.3 wt %; H₂O ¼ 1.0 wt %; Fig. 12b) to that of a primitive melt that had already fractionated 19 wt % of MgO-rich olivine at 9 kbar. This starting composition is close to that of the most primi-tive melt inclusions found in olivine (Bureau et al., 1998; Di Muro et al., 2016). At 6 kbar (Fig. 12b and c), cooling and crystallization of primary magma produces olivine Fo₈₇–₈₅ similar to that of the MgO-richest olivine found in ‘olivine-bearing’ lavas from peripheral eruptions and dunites. In this scenario, we find that the ‘early’ CaO/Al₂O₃ decrease at high MgO recorded by some ‘olivine-bearing’ lavas from peripheral eruptions (e.g. samples from Chisny, CHI* in Fig. 12b and c) can be reproduced only by predominant crystallization of pyroxene at 6 kbar and temperatures lower than 1270 C. This process drives the initial melt composition after fraction-ation of 4 wt % olivine and 12 wt % clinopyroxene close to that of the associated mafic glasses (c. 8 wt % MgO

for Chisny; Fig. 12b and c). At 4 kbar, the simulation reproduces coticetic pyroxene together with an olivine composition (Fo₈₄) like that observed in the natural ‘clinopyroxene-bearing’ lavas from both the peripheral and central areas of PdF. Computed magma composi-tions are close to the composition of ‘clinopyroxene-bearing’ lavas from the peripheral area. Our modeling also shows that low-MgO olivine (<Fo₈₄) (i.e. the dom-inant type of olivine erupted in the central area of PdF) typically forms at pressures lower than 3 kbar from evolved melts (MgO < 6.3 wt %). Plagioclase is never stable in mafic melts (MgO > 5 wt %) in the pressure range 3–9 kbar at temperatures between 1140 and 1350 C and water contents between 0.8 and 1.3 wt %.

DISCUSSION

Polybaric magma evolution at Piton de la Fournaise

Barometric data allow us to formulate a refined model of the magma plumbing system at PdF (see Table 4 and Fig. 13a for a summary). Results of fluid inclusion stud-ies and pyroxene barometry closely agree for lavas from peripheral vents whose crystal shapes indicate fast crystal growth (Fig. 1). Relatively to these products, the majority of primary fluid inclusions in olivine phe-nocrysts record a maximum pressure of 2.96.0.7 kbar, consistent with the average pressure estimated by clinopyroxene barometry (3.16.0.6 kbar; Fig. 7b). Higher pressure values (up to 4.0.60.7 kbar) are recorded only by trails of secondary fluid inclusions in MgO-rich olivine antecrysts (Table 4). In lavas from the central area, previous studies have found that the majority of melt and fluid inclusions trapped in olivine Fo₈₀–₈₅ indicate pressures lower than 0.9 kbar (Bureau et al., 1998, 1999; Famin et al., 2009; Di Muro et al., 2014, 2016), and in most cases below 0.7 kbar (Famin et al., 2009). The pressures estimated using fluid and melt inclusion barometry are always much lower than the values derived from clinopyroxene geothermobar-ometry (2.7.6.1.1 kbar; Fig. 7e). The observed discrep-ancy may result from a combination of several processes such as the following: (1) an overestimation of pressures from the clinopyroxene-based barometry of Putirka (2008b) owing to kinetic effects during crystal growth on Mg-Fe partitioning (Mollo et al., 2010); (2) the incorporation of olivine antecrysts (Salau´n et al., 2010) from shallow cumulate bodies, similar to those encountered by the Grand Brule´ drillcore below the eastern flank of the volcano (<2 kbar; Rancon et al., 1989); (3) late crystallization of low-MgO olivine (Fo<84) during low-pressure magma evolution (Di Muro et al., 2015, 2016) after the deep crystallization of MgO-rich olivine (Fo>84) and clinopyroxene; (4) late-stage re-equilibration of inclusions and diffusive loss of volatiles during magma storage in the shallow part of the central plumbing system (Wanamaker et al., 1990; Famin et al., 2009). Option (1) is not realistic in our case as we found
<figure><img src="image_19.png" /></figure>

<figure><img src="image_20.png" /></figure>

<figure><img src="image_21.png" /></figure>

Fig. 13. Compilation of the most notable petrological and geochemical markers in Piton de la Fournaise melts that are discussed in this study. (a) Comparison between clinopyroxene and fluid inclusion barometry. Pressure ranges are provided for each sample taking into account for fluid inclusions the natural variability and the maximum analytical error (r 5 0 7 kbar) and for clinopyroxene the 1r variation on the average. Variations in (b) major elements (MgO vs CaO/Al₂O₃) and (c) trace elements (Th/Yb vs CaO/Al₂O₃) in Piton de la Fournaise lavas from both central and peripheral areas. The geochemical variability of Piton de la Fournaise lavas has been attributed to the following: (1) the melting of small-scale, slight heterogeneities in the mantle plume source to variable degrees (Albare` de & Tamagnan, 1988; Bosch et al., 2008; Pietruszka et al., 2009; Di Muro et al., 2014; Vlaste´ lic et al., 2018); (2) evolu-tion of magnesian parental melt controlled by clinopyroxene fractionation (Albare` de et al., 1997; Boivin & Bache´ lery, 2009; Famin et al., 2009; Salau¨ n et al., 2010); (3) olivine cannibalism by evolved magmas (Salau¨ n et al., 2010); or (4) contamination or mixing of magnesian melts with differentiated melts (Salau¨ n et al., 2010) or with low-degree melts of wehrlite cumulates (this study; Pietruszka et al., 2009). These hypotheses are numbered in the figure and are discussed further in the text.
a good agreement between clinopyroxene and fluid in-
clusion barometry for the peripheral samples. Hence, this
feature validates the pressure range computed for samples
from the central area. A significant effect of antecryst
incorporation (Option 2) can be also discarded as the
majority of olivine crystals in lavas from the cen-tral area
are phenocrysts in equilibrium with their host (Fig. 4). Late
olivine crystallization (Option 3) at shallow levels produces
low-MgO olivine (Fo81 5) that is indeed found in many
evolved lavas from the central area. There, MgO-richest
olivines (Fo84–88), resulting from earlier and deeper magma
differentiation, are poorly represented, possibly because of
diffusive re-equilibration during longer storage in the central
con-duit (Option 4). Consequently, we suggest that the
clinopyroxene-only barometry tracks the early and deep
stages of high-pressure evolution of these mag-mas from
the central area, whereas this information is not preserved
in fluid inclusions owing to re-equilibration and later long-
term magma ponding at a shallower level. Similar
differences between fluid inclu-sion and clinopyroxene
barometries have been also reported for lavas from the
Canary Islands and Madeira, and have been attributed to
the resetting of partially decrepitated inclusions during
storage at shallow depth (Hansteen et al., 1998; Schwarz
et al., 2004).

High pressures estimated based on fluid and melt in-
clusion barometry are ubiquitous in lavas emitted in the
peripheral area. Bureau et al. (1998) reported the high-est
pressures in melt inclusions (up to 2 9 kbar) and fluid
inclusions (up to 5 7 kbar) in MgO-rich (Fo87–85) olivines
from Piton de Caille oceanites. These researchers also
documented high pressures (up to 4 8 kbar) in fluid
inclusions in MgO-rich olivine from dunitic enclaves carried
by Chisny lavas. In our study, ‘olivine-bearing’ lavas emitted
in the peripheral area and dunite and wehrlite enclaves
record similar high pressures up to 4 2 kbar (Table 4) in fluid
inclusions in olivine crystals whose composition falls in the
range between the MgO-richest mode (Fo87–85) measured
in Piton de Caille oce-anites (Bureau et al., 1998) and the
MgO-richest mode found in lavas from the central area
(Fo84). In contrast, ‘clinopyroxene-bearing’ lavas emitted in
the peripheral area record only intermediate pressures up
to 2 9 kbar (Table 4) and their olivine compositions (Fo83 5–
81 5) fall between the two main modes found in lavas from
the central area. Clinopyroxene barometry of Bellecombe
gabbros and of lavas emitted in the central area con-strain
olivine–clinopyroxene–plagioclase crystallization to an
intermediate to shallow depth range (<2 7
6 1 3 kbar pressure). The 1998 ‘clinopyroxene-bearing’
lavas emitted from the Hudson vent in the central area
differ from these observations with unusually MgO-rich
olivine (up to Fo88 2) hosting inclusions recording a
maximum pressure of c. 4 4 kbar in fluid inclusions (Bureau
et al., 1999) and c. 5 1 kbar in melt inclusions (Bureau et
al., 1999; Di Muro et al., 2015). In this specific case, this
range agrees with the average pressure we estimate from
clinopyroxene barometry (3 6 6 0 9 kbar;

Table 4) and highlights the peculiarity of the 1998 ‘clino-
pyroxene-bearing’ lavas erupted from the Hudson vent with
respect to other lavas in the central area (high pres-sures,
MgO-richest olivine, alkaline enrichment).

Magma variability at Piton de la Fournaise

We document variable crystallization pressures of mag-
mas emitted in the peripheral area of PdF (2 7–5 7 kbar),
associated with a relatively large geochemical variabil-ity
(Figs 9–11). Magmas erupted from the northeastern and
southeastern rift zones (NERZ, SERZ) are compos-itionally
only slightly variable and overlap with those of lavas emitted
in the central EF caldera, in contrast to the original
suggestion of Le´ nat et al. (2012). In contrast, lava
compositions from the northwestern rift zone (NWRZ) are
compositionally diverse and only partly overlap with lava
compositions found in the central area. Previous studies
proposed several processes to explain the compositional
variability of PdF lavas, which we synthesize in Fig. 13 (see
corresponding numbers in circles on the figure): (1) variable
melting degree of minor mantle source heterogeneities
(Albare` de & Tamagnan, 1988; Pietruszka et al., 2009; Di
Muro et al., 2014; Vlaste´ lic & Pietruszka, 2016); and/or
(2a) a plagio-clase effect (Boivin & Bache` lery, 2009;
Famin et al., 2009; Di Muro et al., 2015; Valer et al., 2017a);
and/or (2b) a clinopyroxene effect (Kornprobst et al., 1979;
Albare` de et al., 1997); and/or (3) crystal recycling and re-
sorption (Albare` de et al., 1997; Salau¨ n et al., 2010); and/
or (4) mixing of mantle-derived magmas together with
differentiated trachyte melts (Villemant et al., 2009; Salau¨
n et al., 2010) or with low melting-degree partial melts
derived from hydrothermally altered cumulates (Pietruszka
et al., 2009). In the light of our new results, we now discuss
in detail these processes that may ex-plain the geochemical
variability in PdF basalts.

Variable melting degree of minor mantle source
heterogeneities

We document a small range of ⁸⁷Sr/⁸⁶Sr and ¹⁴³Nd/¹⁴⁴Nd
isotope ratios in the recent lavas in the per-
ipheral area that is, however, significantly larger than the
isotopic variability previously reported for central lavas at
PdF (Fisk et al., 1988; Albare` de et al., 1997; Bosch et al.,
2008; Pietruszka et al., 2009; Di Muro et al., 2014). The
slight variations in Sr–Nd isotope composi-tions correlate
only with ratios of variably incompatible elements (Fig. 11)
and are consistent with minor mantle source
heterogeneities. These can be related either to
compositionally distinct components intrinsic to the La Re´
union plume itself or to the incorporation of the am-bient
Indian mantle asthenosphere into the upwelling plume, in
agreement with some previous interpreta-tions (Pietruszka
et al., 2009; Di Muro et al., 2014; Vlaste´ lic & Pietruszka,
2016). ⁸⁷Sr/⁸⁶Sr and 1/Sr are lin-early correlated,
suggesting a mixing process between two isotopically
distinct components [e.g. Petit Crate´ re (PC) and Chisny
(CHI) end-members]. For instance,
mixing between melts resulting from variable melting degrees of an isotopically heterogeneous mantle source can produce the observed ranges in Sr–Nd iso-topes and part of the variations in incompatible trace element ratios [e.g. Formica Le´o (FL); Figs 11 and 13]. This model was previously proposed for the central area to explain the variability observed during the major 2007 eruption episode and during the entire 1998–2006 eruptive period (Di Muro et al., 2014). The oc-currence of extensive mixing in the central crustal magma storage system is also supported by the limited range in the Pb isotopic composition of central lavas (Vlaste´ lic et al., 2009).

We argue, however, that slight mantle source effects explain only part of the geochemical variability of melts. In fact, we demonstrate that at PdF transitional magmas with variable compositions of both major (CaO/Al₂O₃, alkalinity) and trace elements (MREE, HREE) can be emitted within the same vent [e.g. early (CHI) and late stage of Chisny (CHI*); Fig. 9] without variation of Sr–Nd isotopic signature (Figs 11 and 13) and with similar variations in Y (by a factor of 0 3), HREE (by a factor of 0 3–0 4) and highly incompatible element abundances (K₂O varies by a factor of 0 3).

## Plagioclase fractionation

The average Eu/Eu\* ratio is near unity for PdF lavas in both central and peripheral areas (Eu/Eu\* ¼ 1 04 60 01), despite significant variations in Al₂O₃ and CaO/Al₂O₃ (Figs 9 and 10). This rules out a major role for plagioclase fractionation in the petrogenesis of tran-sitional magmas with alkaline affinity (Famin et al., 2009). Our conclusion is consistent with the rare occur-rence of plagioclase (Boivin & Bache´ lery, 2009; Di Muro et al., 2015) and the results of the AlphaMELTS model-ing, which confines gabbroic evolution to intermediate– shallow depths (Fig. 12). Lavas emitted in the central area and the interstitial glasses from the Bellecombe gabbros best represent this shallow (<3 kbar) gabbroic fractionation and confirm that plagioclase does not play a major role at higher pressures and mainly forms in evolved cool (and/or degassed) melts. Recent reviews have shown that plagioclase microlites or microphenoc-rysts in lavas emitted in the central area are less anor-thitic, and their abundance increases in evolved lavas with MgO <6–7 wt % (Boivin & Bache´ lery, 2009; Di Muro et al., 2015). More calcic plagioclase (An84) is observed in the Bellecombe gabbros and rare and older plagioclase-rich lavas found at the base of the PdF edi-fice (Valer et al., 2017a). Modeling suggests that these calcic plagioclase compositions should be in equilib-rium with an intermediate melt with 6 wt % MgO over a large pressure range from 1 to 9 kbar for water contents from 0 8 to 1 5 wt % (Lange et al., 2009). However, plagioclase-rich lavas are observed only in specific eruptive phases of PdF that correspond to decreasing magma supply, promoting plagioclase crystallization (Valer et al., 2017a). Experiments on relatively evolved

melt compositions from the Kerguelen Islands (5 wt % MgO and 1 2 wt % H₂O; Scoates et al. 2006) also demon-strate that plagioclase tends to disappear at a pressure >5 kbar. By combining our AlphaMELTS simulations and these experimental data on similar lavas we may conclude that significant plagioclase fractionation at PdF starts at pressures <3 kbar; that is, at crustal levels. Therefore, the general absence of Eu depletion in per-ipheral lavas excludes a major role for plagioclase crys-tallization during PdF magma differentiation. Olivine and clinopyroxene are the dominant fractionating phases over a large range of depths, in accordance with the pyroxene barometry performed in this study.

## The role of clinopyroxene in controlling compositional variations

Prior to this study, clinopyroxene phenocrysts were thought to occur only rarely at PdF, except in a prehis-toric ankaramite sample (Albare` de et al., 1997), in Bellecombe gabbros and in cumulates from the Grand Brule´ drillcore (Rancon et al., 1989; Upton et al., 2000). These products were interpreted as the result of shallow magma fractionation. Instead, our new petrological and geochemical data, as well as our modeling results, em-phasize the polybaric nature of clinopyroxene fraction-ation during melt evolution. This explains the CaO/Al₂O₃ and Sc depletion, and also partially accounts for the MREE–HREE depletion and alkali enrichment observed in most of lavas emitted in the peripheral area [e.g. Chisny (CHI) and Crate` re Commerson (CC); Figs 9–12].

A major role for clinopyroxene fractionation has been reported in other ocean island basalt (OIB) suites (e.g. Madeira, El Hierro island; Schwarz et al., 2004; Stroncik et al., 2009) and is widely invoked to explain the decrease in CaO, SiO₂, Sc, MREE and HREE, and the increase in FeO and K₂O in OIB in general (Naumann & Geist, 1999; Mattsson & Oskarsson, 2005). Interestingly, the distribution of transitional basalt compositions at PdF, with either tholeiitic or alkaline affinity, reflects that of the shield- and post-shield stages at Hawaii (Clague, 1987). There, the Hualalai Volcano shows a temporal evolution from an early tholeiitic shield stage to a post-shield alkaline stage, with lavas that also carry abun-dant dunite, wehrlite and gabbro xenoliths (Peterson & Moore, 1987). This temporal geochemical evolution has been explained by a decrease in the magma production rate, consequent advanced solidification of shallow magma bodies with limited recharge and migration of the magma storage zone to deeper levels, where clino-pyroxene fractionation occurred (Clague, 1987; Chen et al., 1991). However, at PdF, transitional melts with al-kaline and tholeiitic affinities are produced synchron-ously during its whole eruptive history (Albare` de et al., 1997; Michon et al., 2015; this study), ruling out a tem-poral evolution of source conditions, as at Hawaii. The heterogeneities we document in peripheral lavas at PdF could, at least partly, stem from variable depths of magma differentiation following dunitic and wehrlitic
trends. For instance, using the composition of natural clinopyroxenes (Salau¨ n et al., 2010; this study) and the melt composition of the 5 April 2007 eruption (similar to the parental magma composition modeled above), we calculate that 10–12 wt % of clinopyroxene crystalliza-tion could lead to alkali increase (p0 4 wt % on K₂O þ Na₂O) together with CaO/Al₂O₃ decrease (–0 2) and ex-plain part of the trace element variability (Figs 9 and 10) observed in lavas emitted in the peripheral area, at con-stant Sr–Nd isotopic signature (e.g. CHI, CC).

Meanwhile, even if olivine–clinopyroxene differenti-ation
in the uppermost mantle lithosphere is a common process
below volcanic islands (Hildner et al., 2011) and plays a
major role in the geochemical variability of PdF magmas,
we highlight that our AlphaMELTS modeling shows some
differences from experimental results. First, recent
crystallization experiments on SSB basalt compositions
have shown that clinopyroxene may crystallize before
olivine at pressures higher than 3–4 kbar (Brugier, 2016),
whereas our study suggests that at all pressures (up to 9
kbar in this study), olivine is the first phase that crystallizes
in PdF magmas, followed by clinopyroxene. The evolved
composition of the starting material used for the
experimental investigations (MgO

¼ 7 8 wt %; Brugier, 2016) contrasts with the compos-ition
of the starting melt used in our modeling (MgO

>10 3 wt %). The widespread mild alkali enrichment to-
gether with low CaO/Al₂O₃ ratios that we document in bulk-
rocks and in matrix glasses of peripheral eruptive lavas is
not observed in the most magnesian melt inclu-sions found
in olivine (MgO >10 wt %; Fig. 9a). This sug-gests that these
compositional features are attained later in the
differentiation process and this justifies the use of a more
primitive (MgO >10 wt %) starting com-position in our model
compared with the experiments. Second, our modeling
does not reproduce the crystal-lization of diopside that we
observe in ultramafic enclaves and gabbros. This
observation could either re-flect the limitations of the
AlphaMELTS code or be evi-dence for distinct melt
evolution processes involved in forming ultramafic enclaves
and gabbros.

Crystal recycling and resorption

In ‘olivine-bearing’ lavas from peripheral eruptions there is
an extensive recirculation of olivine crystals, as shown by
the common presence of clusters of polyhe-dral crystals
(Fig. 2f), abundance of wehrlite and dunite enclaves (Fig.
3a and b), polymodal forsterite distribu-tion (Fig. 4b), and
low Ca contents in olivine overlap-ping that found in the
olivine of ultramafic enclaves (Fig. 5b). The incorporation of
5–10 wt % olivine can ex-plain the observed discrepancy
between the compos-ition of matrix glasses and bulk-rocks
in ‘olivine-bearing’ lavas (e.g. CHI in Fig. 13) as previously
reported (Albare` de et al., 1997; Salau¨ n et al., 2010). An
interesting result of our study is that ‘olivine-bearing’ lavas
from peripheral vents do not contain pyroxene phenocrysts
despite (1) large amounts of wehrlite

enclaves and olivine antecrysts and (2) geochemical
evidence for clinopyroxene fractionation (alkaline affin-ity
and depletion in CaO, SiO₂, Sc, MREE, HREE). This
represents the ‘pyroxene paradox’ (Francis, 1986) and
suggests the role of ‘cryptic’ clinopyroxene fractionation.

With respect to the results discussed above, we pro-
pose that clinopyroxene crystals can be fractionated and
potentially resorbed and redissolved by mafic melts during
their ascent, as also proposed for Hawaii (Frey et al., 1990;
Wanless et al., 2006) and Iceland (Furman et al., 1991;
Mattsson & Oskarsson, 2005). At PdF, petrological
arguments in favor of deep clinopyr-oxene resorption were
also previously presented by Fretzdorff & Haase (2002) for
old submarine lavas carry-ing MgO-rich olivine (Fo90). The
2005 olivine-rich lavas in the central area also contain
resorbed sector-zoned clinopyroxenes (Welsch et al.,
2009), and we observed partially resorbed clinopyroxenes
in gabbroic cumu-lates from the Bellecombe Ash Member.

To evaluate clinopyroxene stability conditions during
magma ascent and storage we simulated the reaction
between (1) a solid assemblage of olivine–clinopyrox-ene
formed at 6 kbar and 1190 C (predicted by the mod-eling of
CHI–CHI* paths from an initial melt with 10 wt
% MgO and 1 wt % H₂O; Fig. 13) and (2) the residual melt
(8 wt % MgO; 1 18 wt % H₂O) at 1190 C ponding at 4 kbar;
that is, the depth of the underplating layer (Supplementary
Data Electronic Appendix 5, Simulation 3). AlphaMELTS
modeling suggests that clinopyroxene should indeed
dissolve rapidly when the assemblage is stored at shallower
pressures. Similarly, experimental studies on basaltic
compositions at a pressure and tem-perature range
relevant for PdF (e.g. Shaw & Dingwell, 2008; Chen &
Zhang, 2009) have shown that the dissol-ution rate of
clinopyroxene is more sensitive to tem-perature (i.e. degree
of superheating) and melt composition (volatile content
increase, silica–alkali en-richment, and calcium depletion)
in contrast to olivine. Interestingly, an experiment at 1250 C
and 5 kbar (Chen & Zhang, 2009; close to the conditions
discussed in this study) on a basaltic melt showed that
clinopyroxene dissolution is at least two times faster than
olivine dis-solution and may occur in a time span as short
as a few days. These studies are consistent with our results
and demonstrate that (1) melts generating ‘olivine-bearing’
lavas are generally richer in MgO and thus hotter than those
generating ‘clinopyroxene-bearing’ lavas (Helz & Thornber,
1987) and (2) early deep olivine–clinopyrox-ene
fractionation tends to increase the water contents of the
residual melts, together with increasing the alkali contents.
The clinopyroxene resorption process would thus be
favored by magma storage at intermediate depths, where
more evolved and possibly slightly water-enriched magmas
may stagnate and/or mix with a more primitive and hotter
melt depressing the clino-pyroxene liquidus. Conversely,
fast magma ascent with-out intermediate storage may favor
the survival of clinopyroxene–olivine equilibrium, as
observed in
‘clinopyroxene-bearing’ lavas that display fast-growth crystal textures and no evidence of re-equilibration processes (e.g. comparison between fluid inclusions and pyroxene barometry). Our results confirm that crys-tal recycling and resorption in basaltic melts also con-tribute to the geochemical variability of the magmas erupted (Costa & Dungan, 2005).

Mixing of mantle-derived magmas together with differentiated trachytic melts and/or with low melting-degree partial melts derived from hydro-thermally altered cumulates

With respect to the effect of the magmatic processes described above, we note that the composition of some eruptive products may not be related to either mantle source heterogeneities or crystal fractionation. This is the case, for instance, for some lavas from the late stage of the Chisny eruptive activity (e.g. CHI*) that are much more enriched in incompatible trace elements (Fig. 10c), but have similar Sr-Nd isotopic compositions and CaO/Al₂O₃ ratios (Figs 10c and 13c) compared with those emitted during the first phases at Chisny (e.g. CHI).

Such ‘anomalous’ enrichment in incompatible ele-ments in PdF melts was previously reported in the 1998 ‘clinopyroxene-bearing’ lavas erupted from the Hudson vent and in some other rare lavas in the central area (Bureau et al., 1998; Pietruszka et al., 2009; Salau¨ n et al., 2010). In the case of the 1998 ‘clinopyroxene-bearing’ Hudson lavas, which are representative of the AbG sig-nature (Famin et al., 2009), mixing between compos-itionally distinct, mafic and evolved melts was proposed (Figs 9–11; Salau¨ n et al., 2010). For instance, mixing of 6 wt % of a trachytic melt (Salau¨ n et al., 2010) into an olivine-bearing melt, such as produced during the 5 April 2007 eruption, may explain most of the chemical variability in major and trace elements of these lavas (Fig. 13). This is also consistent with the observed Eu depletion and decrease in CaO/Al₂O₃ (Fig. 10b).

Although some lavas erupted from peripheral vents (CHI*, DP) share some similarities with the 1998 ‘clino-pyroxene-bearing’ lavas (AbG) such as CaO/Al₂O₃ de-pletion and alkaline and incompatible trace element enrichment (Figs 9 and 10), they strongly differ in their petrological characteristics (Table 1), distinct olivine compositions (Figs 5 and 6), the absence of negative Eu anomalies (Fig. 10b) and different MREE/LREE (light REE) ratios (Fig. 11a). Mixing with a strongly differenti-ated (and low-temperature-high-viscosity) trachytic melt thus appears unlikely. On the other hand, Pietruszka et al. (2009) suggested that contamination with small amounts of low melting-degree partial melts derived from hydrothermally altered cumulates may ex-plain the chemistry of some unusual historical lavas in the central area. To test this hypothesis, we have simu-lated (Figs 10c and 13c) contamination of a melt with a low CaO/Al₂O₃ ratio (after clinopyroxene fractionation;

e.g. CHI) by partial melts derived from wehrlite cumu-lates ('94.G1' with about 60% of clinopyroxene, 35% of olivine and 5% of interstitial melt; Upton et al., 2000). We have preferred wehrlite cumulates instead of gab-broic cumulates because contaminated melts do not show Eu depletion (e.g. CHI*; Fig. 10b). For this simula-tion we use the melt-olivine-clinopyroxene partition coefficients of Salters & Longhi (1999) and Adam & Green (2006) and assumed simple batch melting (Shaw, 1970). We calculate that less than 1% of partial melting of such wehrlitic cumulates is sufficient to produce a melt with CaO/Al₂O₃ ¼ 0.56, Th ¼ 12 ppm and Th/Yb ¼ 7.1 and a major element composition close to the inter-stitial glasses from the Bellecombe gabbro cumulates (Upton et al., 2000). Successive mixing between the composition of low CaO/Al₂O₃ (0.71) CHI basaltic melts and 5% of such a partial melt fits the composition of the CHI* melts (Figs 10c and 13c). This amount of mixed partial melt is similar to those calculated by Pietruszka et al. (2009).

We suggest that the generation of low-degree partial melts from cumulates may be favored by the intrusion of ‘olivine-bearing’ magmas into the cumulate pile at the depth of the mantle-crust underplating layer, con-sistent with evidence for abundant inclusions of cumu-lates and recycled crystals. With respect to the above considerations on clinopyroxene stability in PdF melts, this supports the role of clinopyroxene resorption in producing some contaminant liquids and constitutes a reasonable explanation for the variability observed in some incompatible trace elements, at constant Sr-Nd isotopic signature, at similar major element compositions.

Linking petrological and geophysical
observations

Our study, combined with recent geophysical (Michon et al., 2015) and geochemical (Liuzzo et al., 2015; Boudoire et al., 2017) data, provides further insights into the plumbing system of PdF (Fig. 14).

At La Re´ union, the occurrence of a deep zone of magma ponding within the mantle lithosphere at 30 km b.s.l. (c. 8–10 kbar pressure) has been proposed on the basis of seismic tomography (Fig. 14a; Charvis et al., 1999; Fontaine et al., 2015). This depth range overlaps with that of the deepest seismic swarms (c. 30 km b.s.l.) that started to be recorded just after the major 2007 eruption (Massin et al., 2011). This deep level may also correspond to a level of neutral buoyancy, allowing deep magma stagnation and the first stage of high-pressure olivine–clinopyroxene crystallization to pro-ceed as proposed by Albare` de et al. (1997). Rare MgO-rich olivine (Fo>88) and high-pressure clinopyroxene documented in submarine lavas by Fretzdorff & Haase (2002) may derive from this early, deep stage of differ-entiation. The occurrence of a deep level of melt differ-entiation in the lithospheric mantle has also been reported for other OIB suites and could facilitate the
<figure><img src="image_22.png" /><figcaption>Fig. 14. Conceptual model of the deep plumbing system of Piton de La Fournaise showing magma transfer and storage below the NW rift zone (see text for explanation). Large filled rectangles represent the main depth ranges constrained by barometric data, obtained from fluid inclusion analysis (red for mantle pressures; blue for crustal pressures). Partially transparent rectangles represent the pressure range recorded by fluid inclusions. Lateral magma migration in the upper mantle is based on a review of recent seismic data (Michon et al., 2015). Fast ascent of deep magmas below the NWRZ permits the preservation of a deeper chemical signature and represents a major source of potential volcanic hazard on the densely populated volcano flanks. Pressure-depth relations are from Table 3.</figcaption></figure>

early evolution of melt from tholeiitic to transitional compositions (Naumann & Geist, 1999; Stroncik et al., 2009). Between these high pressures and the lower (maximum) pressures (c. 4 8–5 7 kbar) preserved in some olivine-hosted fluid and melt inclusions (Table 4; Fig. 13) there is a gap for which there is no pressure re-cord, which could testify to fast magma ascent, with limited cooling and crystallization, from the deep reser-voir at 30 km towards the base of the oceanic crust. The pressure record of 3 9–4 2 kbar by olivine-hosted fluid inclusions from ‘olivine-bearing’ lavas erupted in the peripheral area (3 9–4 2 6 0 7 kbar) is consistent with magma ponding and fractionation between 15 and 10 km b.s.l. A similar pressure (4 kbar) recorded by fluid inclusions in MgO-rich olivine antecrysts hosted in ‘clinopyroxene-bearing’ lavas confirms stagnation and crystallization at the crust–mantle boundary (Fig. 14b). In 1996 and 2014, clusters of seismic events were identified at this intermediate depth (15–16 km b.s.l.; 5 kbar) close to the NW border of EF caldera. These events marked the onset of unrest leading to the central volcanic activity in 1998 and 2014 (Battaglia et al., 2005; Boudoire et al., 2017).

A <3 km thick layer of dense crustal underplated rocks was identified at the depth range from 10 to 13 km b.s.l., based on seismic studies, corresponding to a pressure of 3 3–4 2 kbar (Charvis et al., 1999; Gallart et al., 1999; Fontaine et al., 2015; Fig. 14c). At least some of the dunitic and wehrlitic cumulate lithic fragments from Chisny may derive from this layer. The absence of

clinopyroxene phenocrysts in ‘olivine-bearing’ lavas emitted in the peripheral area and carrying and recy-cling abundant wehrlites and dunites suggests that the underplating layer might also be the zone in which re-sorption of clinopyroxene occurs. Fast ascent from this stagnation level is supported by voluminous peripheral cones that contain large (>10 cm), dense xenoliths. Fast ascent rates were also proposed by Sigmarsson et al. (2005) based on short-lived radionuclide disequilibria in NWRZ lavas. The frequency of recent volcanic cones along the narrow NWRZ (<50 kyr; one eruption every 200 yr; Michon et al., 2015) suggests that magma ascent is facilitated along zones of weakness that, according to Bache` lery (1981) and Michon et al. (2007), may be related to the structural inheritance of the underlying oceanic lithosphere. These observations imply that pathways of fast magma ascent from the upper mantle can open below the western volcano flank.

‘Clinopyroxene-bearing’ basalts (2 4–2 7 6 0 7 kbar on average) and the Bellecombe gabbros represent magma storage at intermediate pressures at a deep crustal level close to the Moho. Rapid magma ascent from such relatively deep levels can occur not only in the peripheral but also in the central summit area (Fig. 14d). This pressure range is consistent with a high concentration of abundant micro-earthquakes observed during the seismic swarm of March 1998 preceding the eruption of the 1998 ‘clinopyroxene-bearing’ basalts at the Hudson vent (e.g. Battaglia et al., 2005; Michon et al., 2015). These seismic events trace a subvertical
seismic zone from 7.5 km depth towards the summit and represent magma ascent within the main central conduit system. Many geochemical and geophysical studies have focused on the shallow (crustal) part of the central plumbing system of PdF (<1 kbar; Aki & Ferrazzini, 2000; Battaglia et al., 2005; Fukushima et al., 2005; Famin et al., 2009; Di Muro et al., 2014, 2015). These studies emphasize the presence of small-volume magma reservoirs at or above sea level (Nercessian et al., 1996; Battaglia et al., 2005; Sigmarsson et al., 2005; Famin et al., 2009; Peltier et al., 2009; Welsch et al., 2009; Vlast´ lic & Pietruszka, 2016). Low pressures recorded by most olivine-hosted melt and fluid inclu-sions in lavas from the central area and the rare occur-rence of low Ti/Al pyroxenes in these lavas support these interpretations. Moreover, the compositions of most magmas that were processed in and erupted from the central plumbing system follow a gabbroic differen-tiation trend that is confined to shallow depths (e.g. Bellecombe gabbros; <2 7 kbar). This shallow plumbing system is conceived of as a network of magma pockets and dykes (Albare` de et al., 1997; Di Muro et al., 2014) hosting differentiated and differentiating magmas that feed the eruptions of the summit cone but can also erupt laterally on the NE and SE rift zones (Fig. 14e).

## CONCLUSIONS

Extrusive and intrusive products recently (<5 kyr) emit-ted on the western flank of Piton de la Fournaise (PdF) reveal petrological and geochemical variations that pro-vide new insights into the deep magma plumbing sys-tem, and link the central and peripheral magmatic systems of this highly active volcano.

We document chemical and Sr–Nd isotopic varia-tions in erupted lavas, closely associated in time and space, from both the peripheral and the central area of the volcano. The lavas emitted in the peripheral area ex-hibit Sr–Nd isotopic variations that overlap (and exceed) those of lavas emitted in the central area. Sr–Nd isotop-ic variations and their correlation with high to moder-ately incompatible element ratios suggest melting of small-scale heterogeneities in the mantle source region. To better identify possible end-members and their ori-gin in the mantle further work is required. As an ex-ample, it is unknown whether the Pb isotopic composition of peripheral lavas correlates with Sr–Nd variations. Apart from variations in the mantle source, mixing between basaltic melt end-members with dis-tinct Sr-isotope signatures, and variable ratios of incom-patible trace elements, this study highlights compositional variations caused by polybaric clinopyr-oxene crystallization–resorption and olivine crystalli-zation–recycling to produce the major element compositional variability of PdF basalts.

Barometric data highlight the early magma crystal-lization and differentiation below La Re´ union Island at mantle depths (8–10 kbar), whereas magma evolution is enhanced during ponding over a pressure range of

3–6 kbar and at shallower levels within the volcanic edi-fice itself (Fig. 14). The scarcity of MgO-rich olivine (Fo>88) in erupted magmas and the absence of true primitive compositions in subaerial lavas is thus well explained by early, deep (>8 kbar) fractionation of pri-mary melts. Deep (3–6 kbar) cotectic crystallization of olivine (Fo<88) and low Ti/Al clinopyroxene controls the composition of most ‘olivine-bearing’ lavas emitted in the peripheral area. These basalts record the geochem-ical signature of high-pressure olivine–clinopyroxene fractionation at lithospheric mantle depths, but lack pyr-oxene phenocrysts despite the presence of abundant wehrlite enclaves and olivine antecrysts. This discrep-ancy is thought to result from clinopyroxene resorption during magma storage at intermediate depths, where more evolved and possibly slightly water-enriched magmas may pond, and/or by mixing with more primi-tive and hotter melts that depress the clinopyroxene liquidus. In addition, some of these magmas show evi-dence of contamination with small amounts (up to 5%) of low-degree partial melts derived from cumulates. Conversely, ‘clinopyroxene-bearing’ lavas from the per-ipheral area of Piton de la Fournaise display widespread textures of fast crystal growth and record olivine–clino-pyroxene crystallization at intermediate crustal depths (3 kbar), followed by rapid magma ascent that is essen-tial to prevent complete clinopyroxene dissolution. The conduit system centered below the EF caldera feeds the NE and SE rift zones with relatively evolved magmas (<7–8 wt % MgO). These basalts, produced by shallow ‘gabbroic differentiation’ (<3 kbar and in most cases <1 kbar), contain MgO-poor olivine (Fo<84) and high Ti/ Al clinopyroxenes.

In this study, we argue for multiple magma ponding zones within the uppermost mantle beneath PdF and within the crust, with potential mixing between slightly isotopically distinct melts. We demonstrate that fast as-cent of deep-seated magmas occurs over both the cen-tral area and the densely populated volcano flanks, with major implications in terms of volcanic hazard. These new findings should be carefully taken into account in volcano monitoring and for civil protection.

## ACKNOWLEDGEMENTS

F. Albare` de is gratefully acknowledged for stimulating discussions on PdF magmatism. His previous work has inspired this study, contributing to improve the under-standing of PdF activity and raising important questions relevant for volcano monitoring. K. Putirka is gratefully acknowledged for detailed and constructive hints on the use of clinopyroxene barometry and related sources of errors and uncertainties. We acknowledge the assist-ance of L. Grousset, M. Fialin, G. Labeau and D. Neuville during microanalysis. Constructive comments by Wendy A. Bohrson, Aaron J. Pietruszka and one an-onymous reviewer are acknowledged for improving the quality of the paper.
## FUNDING

The Université de La Reunion and the ANR 'STRAP' (ANR-14-CE03-0004) have funded this work. The MARES project (PTDC/GEO-FIQ/1088/2014), funded by the Fundacao para a Ciecia e Tecnologia (FCT), Portugal, provided financial support for the scientific meetings of V.Z. in France. This is IPGP contribution number 4083.

## SUPPLEMENTARY DATA

Supplementary data are available at Journal of Petrology online.

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