# Protracted magma evolution and transcrustal magmatic plumbing system architecture at Erta Ale volcano (Afar, Ethiopia)

Juliette Pin, Gilles Chazot, Lydéric France, Bénédicte Abily, Andrey Gurenko, Hervé Bertrand, Alexandra Loppin

► To cite this version:

Juliette Pin, Gilles Chazot, Lydéric France, Bénédicte Abily, Andrey Gurenko, et al. Protracted magma evolution and transcrustal magmatic plumbing system architecture at Erta Ale volcano (Afar, Ethiopia). Journal of Petrology, 2024, 65 (12), pp.egae118. 10.1093/petrology/egae118. hal-04868889

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# Protracted magma evolution and transcrustal magmatic plumbing system architecture at Erta Ale volcano (Afar, Ethiopia)

Pin Juliette <fnref n="1" />, <fnref n="2" />, *, Chazot Gilles <fnref n="1" />, France Lydéric <fnref n="2" />, <fnref n="3" />, Abily Bénédicte <fnref n="1" />, Gurenko Andrey <fnref n="2" />, Bertrand Hervé <fnref n="4" />, Loppin Alexandra <fnref n="1" />

## Abstract :

The Afar region is one of the only places on Earth where magmatic continental rifting and associated ongoing break-up processes are exposed onshore. The several active magmatic segments there are characterized by contrasted morphologies, crustal thicknesses, magma production rates, and magma-tectonic styles. In the Erta Ale Range rift segment, extension is magmatically accommodated, making the range the ideal place to study the magmatic behavior of a mature rift segment. Erta Ale Range comprises sub-segments with magma compositions ranging from basalts to rhyolites, but only the Erta Ale Volcano (EAV) sub-segment is active, where only basaltic compositions have been reported so far. Here, we show for the first time protracted differentiation at EAV that is not expressed volcanically at the surface, but is rather accessible via unique cognate gabbroic and microgabbroic blocks, and recorded by mixing with erupted basaltic magmas. These cognate samples record previously unknown mushy and evolved parts of the EAV plumbing system. To constrain their origin and evolution, we measured the major and trace element compositions of the bulk rocks, interstitial glasses, and melt inclusions. We also measured the oxygen isotopic compositions of olivine crystals, interstitial glasses, and melt inclusions. By combining these results with textural relationships and oxy-thermo-barometry calculations, we discuss magma differentiation and storage conditions, as well as magmatic interactions during transport through the crust. Comparison of our results with rhyolite-MELTS thermodynamic models highlights that protracted fractional crystallization is the main process of magma evolution, and when associated with reactive porous flow is capable of forming the evolved compositions observed (up to 75 wt.% SiO2). We also use the model outputs to quantify distinct steps of igneous differentiation in both shallow and deep crustal reservoirs, and we highlight significant interactions with hydrothermally altered wall rocks. We discuss this model within the geological contexts of the Erta Ale Range rift segment and the larger Afar region, and highlight contrasts with mature oceanic systems to argue that the region is not in the final stages of continental break-up.

<footnote n="1">¹ Université de Bretagne Occidentale (UBO) Laboratoire Geo-Ocean (LGO), Institut Universitaire Européen de la Mer (IUEM), Place Copernic, 29280 Plouzané, France</footnote>

<footnote n="2">² Université de Lorraine, CNRS, CRPG, F-54000 Nancy, France</footnote>

<footnote n="3">³ Institut Universitaire de France (IUF), France</footnote>

<footnote n="4">⁴ Ecole nationale supérieure de Lyon (ENS Lyon), 15 parvis René Descartes, 69342 Lyon, France</footnote>

<footnote>* Corresponding author : Juliette Pin, email address : juliette.pin@univ-brest.fr</footnote>
**Keywords** : Erta Ale, magma reservoir, melt inclusions, oxygen isotopes, rifting
58 INTRODUCTION

59 The Afar region of the East African Rift is one of the only places on Earth where ongoing
60 magmatic continental rifting is exposed. There, several active magmatic segments (for
61 example, Asal, Manda Hararo, and the Erta Ale Range) are characterized by contrasting
62 morphologies, crustal thicknesses, magma production rates, and magma-tectonic styles.

63 Extensional rates along the Afar rift segments are similar to those of slow oceanic spreading
64 centers and are similarly characterized by intense magmatism associated with dyke
65 emplacement and normal faulting (Buck et al., 2005; Medynski et al., 2016). Due to these
66 similarities, and by contrast with other Afar rift segments, the Erta Ale Range has been
67 interpreted as representing the final stage of continental break-up based on three general
68 arguments. (1) Quaternary fissure eruptive volumes and crustal thicknesses in the region
69 indicate the late-stage thinning of the heavily intruded, weakened plate just before the onset of
70 seafloor spreading (Bastow & Keir, 2011; Keir et al., 2013). (2) Calculated magma fluxes and
71 parental basalt compositions indicate a greater role of plate stretching in mature rifts at the
72 onset of sea-floor spreading (Hutchison et al., 2018). (3) Shallow axial magma bodies in the
73 Erta Ale Range are typical of fast-spreading ridges (Moore et al., 2019). Accordingly, Erta
74 Ale Range lithosphere is considered to be extensively thinned and undergoing decompression
75 melting (Bastow & Keir, 2011), generating a high flux of parental tholeiitic basalts and
76 precluding magmatic evolution toward differentiated compositions (Hutchison et al., 2018).
77 Nevertheless, setting a precise definition of continental break-up remains elusive, as it
78 strongly depends on which geological characteristics is considered. From a geochemical point
of view, the break-up is complete when the magmas are coming from a depleted asthenospheric mantle source (DMM, Gale et al., 2013) similar to the source of the normal mid-ocean ridge basalts (MORB). Concerning the crust, the break-up is reached when no continental crust is left, and the whole axis structure is composed of juvenile material. In terms of geophysics and numerical models, the removal of the continental lithospheric mantle, the thinning of continental crust and its eventual replacement by oceanic one appears to represent the break-up. It is important to note here that extension is likely to be depth-dependent with a decoupling between upper and lower lithosphere, and continental crust break-up occurring before or after the continental lithospheric mantle one, depending on the strength of those various domains (Huismans & Beaumont, 2011; 2014). In any case, all those concepts defining the break-up are progressive; the whole transition period between stable continental crust and oceanic crust is thus referred herein as “ongoing breakup”. In magma-rich oceanic-continental transition, this very period corresponds to the formation of Seaward Dipping Reflectors (Geoffroy et al., 2005; Stica et al., 2014).

According to Hutchison et al. (2018), the presence of acidic lavas in older Erta Ale Range eruptive centers, and their lack in the more recent activity (relative age determined by the geological mapping of Barberi & Varet (1970), see figure 1a) indicates that magma fluxes have increased through time, and that break-up likely occurred only recently, with present-day magma compositions restricted to basalts at Erta Ale volcano (EAV). This nearly direct transfer of basaltic magma to the surface is presented as similar to that observed at oceanic spreading centers, although ocean ridge magmas are likely stored in mushy reservoirs and are not always transferred directly to the surface (Lissenberg et al., 2019; Boulanger & France, 2023); in addition, such direct magma transfer may only be relevant to EAV and not for the entire Erta Ale Range segment. Rather than considering the various Afar segments as representing different stages of rift evolution, their characteristics could also derive from
different styles of rifting, mainly controlled by the magma supply and the long-term interplay between the continental lithosphere and the underlying Afar plume (e.g., Stab et al., 2016; Brune et al., 2023). Indeed, Stab et al. (2016) attributed the difference between the wide, magma-rich crust in the Central Afar segments and the narrow, thin, relatively magma-poor crust in the Erta Ale Range segment to differences in magma supply. They argued that both segments are near break-up and that apparent crustal thicknesses cannot be used to predict the location and timing of break-up.

The Erta Ale Range comprises six eruptive centers with magma compositions ranging from basalt to rhyolite. Of those, only EAV is currently active, where solely basaltic magmas have been described so far (figure 1a). EAV has hosted permanent lava lakes for decades, and its eruptive activity consists of sporadic overflows of these basaltic lava lakes (figure 1b and 1c). Such persistent activity is interpreted as endogenous edifice growth (the growth of a volcanic edifice due to the intrusion of magma; Espín Bedón et al., 2024) by Francis et al. (1993), which Oppenheimer & Francis (1997) indirectly estimated to be by more than 50% of the total magmatic activity. Therefore, EAV is the best available location for assessing the plumbing system of an active near break-up rift.

Because magmatic processes control the morphologies and magma-tectonic styles of magmatic segments (e.g., Medynski et al., 2013; 2015; 2016), we here aim to address how the EAV magmatic plumbing system is organized, the range of magma compositions present at depth, and crustal magma storage conditions. We present the first direct evidence of endogenous growth and extreme melt evolution at EAV, with new cognate gabbroic and microgabbroic block samples giving us access to an as-yet undescribed part of the EAV plumbing system. Based on petrological and geochemical constraints, we highlight the presence of various levels of magma storage and differentiation within the EAV plumbing system; locally, these processes result in protracted magma evolution toward rhyolite
129 magmas. Based on our new results, we conclude that the EAV plumbing system architecture differs in many aspects from those at mid-ocean ridges and should therefore be considered as a typical rift-related igneous system.

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## 133 GEOLOGICAL SETTING

134 The Afar depression is both the triple junction between the African, Somalian, and Arabian plates and where the continental East African Rift meets the oceanic spreading centers of the Gulf of Aden and the Red Sea (McKenzie et al., 1970; Tazieff et al., 1972; Le Pichon & Francheteau, 1978; figure 1d). The depression started to form 25 Ma, synchronously with the formation of the Red Sea (Barberi et al., 1972b). Extension is accommodated tectonically by faults and magmatically by volcanism and dyking across several rift segments, including the EAR, Dabbahu-Manda Hararo, and Asal (figure 1d). Afar is also influenced by its eponymous mantle plume (Pik et al., 2006). Volcanism is principally expressed as dykes and eruptive fissures, although craters and cones form at large central volcanoes (Varet, 2018). Each of the three segments presented above has erupted in the last 50 years and is oriented NNW-SSE to NW-SE (figure 1d).

145 The Erta Ale Range is a NNW-oriented rift segment (Barberi & Varet, 1970) considered to be an *en echelon* southern propagation of the Red Sea spreading center (Tazieff et al., 1972; Manighetti et al., 1998). This segment is currently extending at ~20 mm/yr due to spreading between the Nubian Plate and the Danakil Microplate (McClusky et al., 2010). The Ethiopian plateau and the Danakil Alps lie to the east and west, respectively. Erta Ale Range lavas form a complete differentiation series from basalts to rhyolites (Barberi & Varet, 1970; Barrat et al., 1998) erupted from six different eruptive centers: Gada Ale, Alu-Dalafilla, Bora Ale, EAV, Hayli Gubbi, and Ale Bagu. The oldest known Erta Ale Range lavas were erupted
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1.2 ± 0.5 Ma (Barberi et al., 1972a). Beneath the range, the Moho is estimated to lie at about
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15 km depth (Hammond et al., 2011; Keir et al., 2013).

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EAV (figure 1b and 1c) is the most active eruptive center of the Erta Ale Range and
has hosted near-permanent lava lakes for at least 50 years (Varet, 1971). Activity at the
summit appears to have persisted since at least the beginning of the twentieth century
(Dainelli & Marinelli, 1907), with local oral traditions suggesting a much longer recent
eruptive history (Barnie et al., 2016). Lava lakes, fresh lava flows, and fumaroles were
reported from the 1940s (Mohr, 1962; Barberi & Varet, 1970). The first field campaigns to
EAV took place between 1968 and 1973 (Varet, 1971; Barberi et al., 1974; Tazieff, 1973; Le
Guern et al., 1979), during which the lava lake rose until it overflowed in 1973; since those
early campaigns, EAV has been monitored *via* remote sensing and several field campaigns
(Oppenheimer & Francis, 1997). The level of the lava lake remained fairly stable from around
1974 until it contracted between 2003 and 2004, at one point freezing over almost completely.
Lava lake activity restarted around 2005, eventually overflowing in 2010 (Field et al., 2012a).
EAV's most recent eruption was the 2017–2019 fissure eruption in the southern caldera,
associated with and followed by the contraction of the lava lake (Moore et al., 2019).

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SAMPLES
177 Most samples used herein were collected during a field campaign in March 2003. Lava flows, cognate blocs, Pele's hair, and direct samples of the lava lake (Table 1) were collected within the EAV caldera and from its close surroundings (figure 1b and 1c).
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179

180 Lava flow samples were collected both within (BC13, BC14, BC16, BC21) and beyond the EAV caldera (BC17, BC18) (figure 1c). BC17 and BC18 are thus older than the intracaldera flows, and the BC17 lava flow overlies the BC18 flow (figure 1b). Within the caldera, BC16 was sampled from a spatter cone of unknown age, and BC13 and BC14 from 181
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184 1974 lava flows that overlie the BC21 flow (figure 1c). All of these lava flows likely formed
185 during lava lake overflows.
186

186 Pele's hair and samples from the lava lake directly record the lava lake composition.
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192 Strands of Pele's hair are basaltic glass filaments that drift directly from the active lava lake. Pele's hair samples ERTA-01 to ERTA-06 were found lying on recent lava flows, and samples ERTA-07 to ERTA-10 were emitted during the 2003 field campaign; they are all therefore representative of the lava lake composition in March 2003. Direct samples of the active lava lake (LL1, LL2, and LL3) were collected during a 2001 field campaign by Jacques Durieux, and sample LL2011 was collected by Bernard Marty in 2011.
193

193 The 2003 EAV field campaign also allowed the sampling of several cognate blocks
194 (CBs; BC01 to BC12 and BC15). They are singular rock fragments that present a variety of
195 textures ranging from basaltic to gabbroic and microgabbroic (or doleritic). These samples
196 were found either atop lava flows or breaking the surface of it and cross-cutting the roof of
197 lava tubes (figure 1f), indicating that they were vigorously projected from the vent and
198 smashed the surfaces of lava flows, and are therefore younger than the flows on which they
199 were found. These decimetric (figure 1f) to metric (figure 1e) CBs may have originated from
200 an explosive phase at EAV; they are found as far as several kilometers beyond the caldera,
201 with the largest being closest to the caldera edge. Pinzuti (2006) used cosmogenic ³⁶Cl dating
on the topmost lava flow exposed in the caldera rim and an overlying CB. They obtained an age of 4.0 ± 0.5 ka for the lava flow, indicating the maximum possible age of caldera formation, and 2.4 ± 0.41 ka for the CB, and therefore for the explosive event during which it was emplaced. Because no pyroclastic deposit is present on EAV (figure 1a), this explosive event can tentatively be related to phase of caldera collapse or the formation of either of the intracaldera pit craters.

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## 209 ANALYTICAL TECHNIQUES

### 210 Whole-rock analyses

Bulk analyses were performed by X-ray fluorescence (XRF) and inductively coupled plasma mass spectrometry (ICP-MS). Major element bulk analyses were obtained by XRF at the Université de Lyon (France). Results published in the two Proficiency Geopt 1 (Thompson et al., 1996) and Geopt 2 (Thompson et al., 1998) showed that laboratory uncertainties are less than 1%, 1.5% and 2% for major elements above 60%, near 10% and below 1%, respectively. Trace element bulk analyses were conducted in 2007 by ICP-MS at ACME Analytical Labs (Whitehorse, Canada) with a precision of 5%.

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### 219 In-situ analyses

In-situ analyses were conducted by electron probe microanalysis (EPMA), laser ablation ICP-MS (LA-ICP-MS), micro X-ray absorption near edge structure spectroscopy (μXANES), and secondary ion mass spectrometry (SIMS).

EPMA analyses of Si, Al, Ti, Ca, Na, K, Mn, Mg, Fe, Cr, Ni, and P were performed during two different sessions. The first session was performed using a CAMECA SX100 equipped with five wavelength-dispersive spectrometers (WDS) at the Laboratoire Magmas et Volcans (LMV, Clermont-Ferrand, France), a 15 kV accelerating voltage, and a 15 nA beam
defocused to 20 µm for interstitial glasses or 5 or 10 µm for melt inclusions; a focused beam was used for other phases. Precision, asserted via repeat analyses of standards, was better than 5% for Cr and Ni and better than 3% for other elements. The second session was performed using a CAMECA SX100 equipped with five WDS detectors at Geo-Ocean (Brest, France), a 15 kV accelerating voltage, and 6 nA beam defocused to 10 µm for interstitial glasses and melt inclusions; for other phases, we used a focused 20 nA beam. Na has been analyzed first to prevent element migration. Information about the crystals used, the counting time and the standards used can be found in supplementary data 2. All melt inclusion compositions were corrected for post-entrapment crystallization using Petrolog3 (Danyushevsky & Plechov, 2011; see the detailed protocol in supplementary data 1 and 2). In the second session, we also acquired X-ray chemical maps of Al, Ca, Fe and Ti using a 20 kV, 100 nA beam defocused to 10 µm, a step size of 15 µm, and a dwell time of 0.06 seconds. In order to obtain quantitative chemical maps X-ray maps were cross-calibrated with quantitative in-situ mineral measurements following the approach presented in France et al. (2024), and by using XMapTools (Lanari et al., 2014). For this cross-calibration, we used 76 mineral phase measurements obtained within the mapped area, and distributed among the different phases composing the sample.

Trace element analyses of <sup>29</sup>Si, <sup>43</sup>Ca, <sup>44</sup>Ca, <sup>45</sup>Sc, <sup>49</sup>Ti, <sup>51</sup>V, <sup>53</sup>Cr, <sup>60</sup>Ni, <sup>85</sup>Rb, <sup>88</sup>Sr, <sup>89</sup>Y, <sup>90</sup>Zr, <sup>93</sup>Nb, <sup>137</sup>Ba, <sup>139</sup>La, <sup>140</sup>Ce, <sup>141</sup>Pr, <sup>146</sup>Nd, <sup>147</sup>Sm, <sup>153</sup>Eu, <sup>157</sup>Gd, <sup>159</sup>Tb, <sup>163</sup>Dy, <sup>165</sup>Ho, <sup>166</sup>Er, <sup>169</sup>Tm, <sup>172</sup>Yb, <sup>175</sup>Lu, <sup>178</sup>Hf, <sup>181</sup>Ta, <sup>208</sup>Pb, <sup>232</sup>Th, and <sup>238</sup>U in interstitial glasses, melt inclusions, and olivine crystals were performed by LA-ICP MS. Lava lake samples mountain on epoxy were analyzed at the LMV using a Q-switched Nd-YAG laser source operating at 266 nm coupled to a Fisons PQ2+ ICP-MS. Analyses were performed in a He atmosphere (He flux ranging from 450 to 700 mL/min). The laser was operated at a repetition rate of 6 Hz and the spot diameter was around 60 µm. Precision and accuracy were assessed on the BCR-2G reference
252 glass. Precision is between 4 and 6% for most of the elements and between 10 and 15% for
253 Sc, Rb and Pb. Accuracy compared to the average published values is between 5 and 10% for
254 Y, Nb, Th and most of the REE, between 10 and 15% for Zr, Ho and Hf and better than 5%
255 for the other elements (see supplementary data 2). Interstitial glasses and melt inclusions were
256 analyzed on thin sections at GeoRessources (Nancy, France) using a 193 nm Newave NWR
257 (ESI) LA system coupled to a quadrupole Agilent 7900 ICP-MS. Analyses were performed in
258 a He atmosphere (He flux of 650 mL/min). The laser was run at 10 Hz and a fluence of 6
259 J/cm². Glasses were analyzed with a 30 µm spot size and 30 s acquisition time for the
260 background signal followed by 40 s for the ablation signal. Olivine crystals were analyzed
261 with a 100 µm spot size and 120 s acquisition time for the background signal followed by 60 s
262 for the ablation signal. The SRM NIST 610 silicate glass was used as an external standard,
263 and the SRM NIST 612 silicate glass was measured as an unknown during the entire session
264 (values from Jochum et al., 2011; see supplementary data 2).

265 The oxygen isotopic compositions of olivine grains and interstitial glasses were
266 analyzed by SIMS, as well as melt inclusions in some olivine grains. Olivine crystals and
267 interstitial glass fragments were separated from the bulk rocks by hand crushing and picking;
268 melt inclusions were exposed by polishing. Olivine grains and glass fragments were mounted
269 in indium-based holders together with standards. The analyses were carried out using the
270 upgraded CAMECA IMS 1270 E7 instrument at the CRPG (Nancy, France) with a focused
271 ¹³³Cs⁺ primary beam that was rastered over an area of 10 × 10 µm to ensure a flat bottom in
272 the ablated pit. ¹⁸O/¹⁶O isotopic ratios were analyzed in multi-collection mode. Each olivine
273 grain and glass fragment were analyzed four times to ensure homogeneity of the samples. The
274 spot position was carefully examined under a microscope after analysis and the exact major
275 element composition of each SIMS spot analysis was determined by EMPA by analyzing
276 points as close as possible to SIMS spots. An internal standard of San Carlos olivine was used
277 to monitor for potential drift of Instrumental Mass Fractionation (IMF) values during the
278 analytical session; it was routinely reanalyzed up to four times at the beginning and end of
279 each session and every ~2 hours to bracket the unknown sample measurements. All ¹⁸O/¹⁶O
280 measurements produced during the analytical session were then normalized to those obtained
281 for San Carlos olivine (supplementary data 2). The matrix effect on IMF was determined by
282 analyzing a series of reference materials of known isotopic composition with various major
283 element compositions: (1) a series of experimental synthetic olivines with δ¹⁸O values from
284 8.35 to 9.04‰ and forsterite contents (Fo) ranging from Fo₀.₃ to Fo₉₀ (Synthetic Watson
285 olivine; supplementary data 2), and (2) a series of reference glasses ranging in chemical
286 composition from basalt through andesite to rhyolite (KL2G, ATHO-G, ML3B-G, StHs6-
287 80G, T1-G, GOR123, and BHVO-2G; Jochum et al., 2006). The matrix effect in olivine was
288 related to Fo content (supplementary data 2), whereas that in glass was mainly related to SiO₂
289 content (supplementary data 2). ¹⁸O/¹⁶O ratios are reported in delta-notation relative to the
290 accepted value of ¹⁸O/¹⁶O in standard mean ocean water ((¹⁸O/¹⁶O)SMOW = 0.00200520;
291 Baertschi, 1976) as δ¹⁸O = [((¹⁸O/¹⁶O)SIMS/((¹⁸O/¹⁶O)SMOW - 1)] × 1,000. The uncertainty on
292 each analysis is reported as the sum of the internal error from the SIMS (typically below
293 ±0.1‰) and the standard deviation calculated from the IMF values (typically below ±0.2‰)
294 obtained during replicate measurements of the San Carlos olivine internal standard.

295 Micro X-ray absorption near edge structure spectroscopy (µXANES) measurements
296 were conducted on glass samples from the Erta Ale lava lake by using ID21 beamline at the
297 European Synchrotron Radiation Facility (ESRF; France). A Si(111) single crystal was used
298 as a monochromator. Energy calibration was performed using the derivative of the XANES
299 spectrum of an iron foil. µXANES measurements were acquired at the iron K-edge (7112eV)
300 in the fluorescence mode. The pre-edge region of the XANES spectra is used as the centroid
301 position of this pre-edge is function of the Fe³+/Fe²+ ratio (or redox value) in the measured
sample. Standards with known and variable $Fe^{3+}/Fe_{total}$ ratio were used: B11 ($Fe^{3+}/Fe_{total}$ =0.19), B14 ($Fe^{3+}/Fe_{total}$ =0.08), B17 ($Fe^{3+}/Fe_{total}$ =0.60), and B20 ($Fe^{3+}/Fe_{total}$ =0.25)

(Botcharnikov et al., 2005). The fitting procedure used is the same as that of Wilke et al.

(2001). The 2σ uncertainty in the $Fe^{3+}/Fe_{total}$ ratio measured by XANES for this range of

oxidation states is ± 0.01 (Berry et al., 2003; Cottrell et al., 2009). µXANES detail method is

described in supplementary data 2.

## 309 RESULTS

### 310 Petrography & mineral major element compositions

Pele's hair strands are very thin glass filaments often attached to drops of glass (Pele's tears),

which can be up to 5 mm in diameter and sometimes contain small crystals of plagioclase,

augite, and olivine. Similarly, lava lake samples are composed of a glassy basaltic matrix with

bytownite phenocrysts up to one centimeter in size.

Lava flow samples (figures 2 and 3; supplementary data 3) are vesicular (BC13, ~15%

vesicles; BC14, BC17, and BC18, ~30%; BC16, ~5%; BC21, ~50%), fine-grained, and

contain euhedral to subhedral phenocrysts of plagioclase (An₄₆ in BC17 to An₈₅ in BC16; An

= Ca/(Ca + Na + K), clinopyroxene (Mg#56 in BC17 to Mg#85 in BC13; Mg# = Mg/(Mg +

Fe)), and olivine (Fo₄₅ in BC17 to Fo₈₅ in BC16). Phenocrysts are generally aggregated

(figure 2b) and represent 20–30% of the lava flow samples, except for BC17, which is

aphyric. Olivine and clinopyroxene crystals are as large as 500 µm and plagioclase as large as

3 mm. Plagioclase and olivine contain melt inclusions of basaltic composition (analyzed in

BC18; SiO₂ = 50–54 wt.%; MgO = 1.5–2.5 wt.%; supplementary data 3). Some plagioclase

phenocrysts have local spongy textures (figure 2c). Matrices are glassy (~80% glass in lava

flow samples, except for BC16 which is only ~5% glass) to microcrystalline (figure 2b). The
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microcrystalline phase assemblage is composed of 30–50% plagioclase, 5–20%

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clinopyroxene, 5–20% olivine, and 5–20% Fe-Ti oxides (ilmenite and titanomagnetite).

328

Cognate Blocks (CB) are massive samples (i.e., no vesicles present) that show a variety of textures ranging from microlitic (crystals smaller than 1 mm; BC6, BC7, BC15) to doleritic (crystals between 1 and 3 mm in size; BC2, BC3, BC5, BC9, BC10–BC12; figure 2g) to gabbroic (crystals larger than 3mm; BC1, BC4, BC8, BC10; figure 2a, 2h and 2i). In all CBs, plagioclase represents the main phase (50–70%), followed by clinopyroxene (5–40%), ferrous rounded olivine (0–40%), and Fe-Ti oxides (0–15%). BC8 and BC10 contain ~10% interstitial glass of trachytic and rhyolitic compositions, respectively (figure 2a; supplementary data 3).

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Crystals are mostly large (1–5 mm) euhedral to subhedral plagioclases (An₁₅ in BC4 to An₈₀ in BC8; figure 3a) with irregular rims, small (<1 mm) subrounded olivines (Fo₁₄ in BC4 to Fo₇₄ in BC6; figure 3b), clinopyroxenes (Mg#25 in BC4 to Mg#83 in BC6; figure 3c), and oxides. Plagioclases can present primitive cores that are sometimes spongy as in the lava flow samples (An₆₀ to An₈₀; figure 2i; supplementary data 3), and have normal, oscillatory, or complex optical zoning patterns with rims showing a large range of compositions from An₁₅ to An₇₀ depending on the sample; nonetheless, anorthite contents in plagioclase rims are systematically lower than those in cores of the same grains (figure 4a and 4b). Olivine in BC8 contains mineral inclusions of augite and plagioclase (figure 2d) as well as rhyolitic glass inclusions, in turn containing ilmenite and apatite (figure 2e, and 2k). Fe-Ti oxides (ilmenite and titanomagnetite) are interstitial, up to 1 mm in size, and represent 10–20% of the CB mineral assemblages (figure 4d and 4e); they can contain mineral inclusions (labradorite and augite) and sometimes show exsolution textures between ilmenite and titanomagnetite (BC4, BC10; figure 2j). Other CB accessory phases (~1–2% of the samples) are apatite (BC3, BC4, BC8; figure 2a and 4e), sulfide (chalcopyrite; BC8; figure 2f), mica (biotite; BC9, BC11),

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349

350
351 alkali feldspar (BC1, BC4), orthopyroxene (BC4, BC8, BC10; figure 4e), and amphibole (BC3, BC4, BC9, BC12; figure 2h). In BC8, orthopyroxene occurs as reaction rims between olivine and the interstitial trachytic melt (figure 2l and 4e). In BC4, amphibole is late-magmatic, replacing pyroxene (figure 2h). Most CB samples are fresh, but some show some alteration and oxidation (for example, ferrous alteration of olivine in BC1, BC2, BC5, BC7, and BC9).

357

358 **Geochemical data**

359 *Bulk-rock major element compositions*

360 Bulk-rock and glass compositions are presented in figure 5 and supplementary data 3; we use MgO content as a proxy of the degree of differentiation (driven initially by olivine fractionation). They have subalkaline affinity. Lava flow samples show limited major element variations: MgO contents vary from 5.93 wt.% in sample BC18 to 6.58 wt.% (BC16), SiO₂ from 49.68 wt.% (BC18) to 50.31 wt.% (BC16) (figure 5a), and CaO from 10.97 wt.% (BC17) to 12.36 wt.% (BC21) (figure 5e). BC17 is the only volcanic rock slightly off the main trend defined by the other samples; it has a lower Al₂O₃ content (13.37 wt.% vs. 16.30 wt.% on average for the other samples; figure 5b) and higher FeOₜ (12.60 wt.% vs. 9.94 wt.% on average; figure 5c), TiO₂ (2.95 wt.% vs. 2.09 wt.% on average; figure 5d), and P₂O₅ contents (0.42 wt.% vs. 0.30 wt.% on average; figure 5h).

370

371 The range of major element compositional variations in CBs is larger than that in the lava flows (figure 5). MgO contents vary from 3.04 wt.% in sample BC4 to 7.28 wt.% (BC8), SiO₂ from 49.83 wt.% (BC6) to 53.15 wt.% (BC4) (figure 5a), and CaO from 7.44 wt.% (BC4) to 11.06 wt.% (BC6) (figure 5e). Harker diagrams presenting the relationships of major elements vs. MgO (figure 5) highlight the positions of CB samples along the major element
trends defined by the lava flow samples. With decreasing MgO content (i.e., with further differentiation), this trend is characterized by increasing SiO₂, FeO, Na₂O, K₂O, and P₂O₅ contents, decreasing CaO and Al₂O₃ contents, and TiO₂ contents that first increase until MgO contents fall below ~4 wt.% and then decrease. The least differentiated samples in this trend are the lava flows, whereas the most differentiated is always BC4. Sample BC8 is slightly offset from the other samples, but its bulk composition is close to those of the lava flows.

The major element compositions of the lava flows overlap literature data for EAV (figure 5), which cover a limited range of basaltic compositions: for example, 5.6–6.3 wt.% MgO, 48.6–51.3 wt.% SiO₂ (figure 5a), and 10.8–12.4 wt.% CaO (figure 5e). CB samples cover a wider range of compositions, overlapping some of the wider range reported for the Erta Ale Range (e.g., 10.9–0.39 wt.% MgO, associated with high SiO₂ contents up to 71.6 wt.% and low CaO values down to 1.44 wt.%). These compositions are similar to the trachytic and rhyolitic compositions of melt inclusions and interstitial matrix glasses in the CB samples. Some Erta Ale Range literature data have compositions showing particularly low TiO₂, P₂O₅, and FeO contents compared to the rest of the dataset; those samples contain 5–3.5 wt.% MgO (symbols circled in black in figure 5).

Trace elements: lava flows, CBs, interstitial glasses, and melt inclusions

Trace element compositions are reported in supplementary data 3 and figure 6 and 7. As with their major elements, the bulk-rock trace element compositions of the CB and lava flow populations overlap. Lavas are slightly enriched in light REEs relative to heavy REE, with La/N/Yb_N values (the subscript 'N' indicating normalization to chondrite from Sun & McDonough (1989)) ranging from 6.4 to 7 and La concentrations from 19.7 to 29.8 ppm (figure 7). Sample BC17 is slightly enriched in trace elements compared to other lava flow
samples, but its pattern is parallel to those of other lava flows in figure 6. All lava flows have trace element patterns subparallel to those of CBs.

CB REE patterns are almost parallel, with La<sub>N</sub>/Yb<sub>N</sub> only ranging from 7 to 8 (figure 6a and 7). They show variable trace element compositions, with La contents ranging from 27.6 ppm in BC8 to 72.2 ppm in BC4. As with other incompatible elements such as Zr, Th, or Nb, La is well correlated with MgO in CBs (figure 7c).

Interstitial glasses and melt inclusions have similar heavy REE patterns but are enriched in light REEs compared to lava flows and CBs (figure 6c and 6d). They are strongly enriched in the most incompatible elements such as Ba, Rb, Th, U, and Nb, as well as in Zr and Hf, and show negative Eu and Sr anomalies (e.g., Eu/Eu* = Eu<sub>N</sub>/[Sm<sub>N</sub> × Gd<sub>N</sub>]<sup>0.5</sup> ranging from 0.2 to 0.8). Accordingly, interstitial glasses and melt inclusions have high La<sub>N</sub>/Yb<sub>N</sub> values, high La contents, and Zr and other incompatible element concentrations well correlated with La, despite the distinct Zr/La ratios in glasses (figure 7c).

413

414 *Oxygen isotopes*

Oxygen isotopic data are presented in figure 8a and 8b. The O isotopic compositions of lava lake glasses range from 4.4 ± 0.09‰ to 6.8 ± 0.09‰ (mean 5.8‰, *n* = 43), with glasses sampled in 2001, 2003, and 2011 spanning the same overall range (figure 8a). Interstitial glasses (11 measurements) and melt inclusions (2 measurements) from CB sample BC8 have δ<sup>18</sup>O values ranging from 4.7 ± 0.11‰ to 5.4 ± 0.09‰ (mean 5.1‰) (figure 8a). Fe-rich olivine (mean Fo<sub>48</sub>) from BC8 δ<sup>18</sup>O values range from 3.3 ± 0.09‰ to 5.5 ± 0.1‰ (mean 4.2‰, *n* = 44) (figure 8b). Olivines from lava flows (BC13 and BC18) have a mean δ<sup>18</sup>O value of 4.2‰ (23 measurements) (figure 8b). In BC18, olivine (Fo<sub>76–77</sub>) δ<sup>18</sup>O values range between 3.8 ± 0.08‰ and 4.4 ± 0.08‰ (figure 8b). In BC13, two olivine populations are
present: one with Fo₇₉ and mean δ¹⁸O = 5.3‰, and the other one with Fo₄₇ and mean δ¹⁸O = 3.7‰ (figure 8b).

426

427 µXANES

428 µXANES results are presented in supplementary data 2 and in figure 9b. Fe³⁺/Feₜotal was
429 converted to fO₂ using the parametrization of Kress & Carmichael (1991) and the ΔFMQ
430 buffered value were obtain from fO₂ using equations from Ballhaus et al. (1991) at 1150°C
431 and 400 MPa. We acquired 3 spectra on 2001 lava lake sample LL1, with an average
432 Fe³⁺/Feₜotal value of 0.195, which corresponds to log fO₂ of -7.30 (+1.44 ΔFMQ) (figure 9b).

433 We also acquired 3 spectra on 2001 lava lake sample LL3, with an average Fe³⁺/Feₜotal value
434 of 0.148, which correspond to log fO₂ of -7.87 (+0.86 ΔFMQ) (figure 9b). We measured 2
435 spectra on 2 distinct 2003 Pele's hairs (one spectrum on each), which gave a Fe³⁺/Feₜotal value
436 of 0.188 and 0.109, corresponding to a log fO₂ of -7.37 and -8.52, respectively (+1.36 and
437 +0.22 ΔFMQ) (figure 9b).

438

439 ESTIMATION OF INTENSIVE PARAMETERS

440 Oxy-thermobarometry

441 Our temperature estimation results (figure 9a; see detailed methods in supplementary
data 2) show the highest temperatures for lava lake glasses (1152–1184 °C; machine-learning
442 thermometer of Weber & Blundy (2024)). Bulk rock temperatures (1105–1190 °C; machine-
443 learning thermometer of Weber & Blundy (2024)) are very close to olivine-melt inclusion
444 equilibration temperatures (1077–1187 °C, mean 1099 °C for both olivine-liquid
445 thermometers; Petrolog3 thermometer from Danyushevsky & Plechov (2011) and
446 thermometer of Mallman & O'Neill, 2013) and the BC18 melt inclusion temperatures
447 determined using the machine-learning thermometer (1091–1107 °C; Weber & Blundy.,
448
2024). Those estimates likely represent the melt inclusion trapping temperature. We obtained lower temperatures of 935–1040 °C for plagioclase-melt equilibrium (thermometer of Putirka (2008)), titanomagnetite-ilmenite equilibrium (thermometer of Sauerzapf et al. (2008)), and BC8 melt inclusion temperatures, and BC8 and BC10 interstitial glass temperatures determined using machine-learning thermometer (Weber & Blundy., 2024). These lower temperatures are consistent with plagioclase rim formation at lower temperature, the rapid re-equilibration of Fe-Ti oxides upon cooling (Hou et al., 2021), and rhyolitic (BC10) and trachytic (BC8) melt compositions documenting the latest stages of evolution.

Redox estimates using the various oxybarometers are similar and largely overlap (figure 9b; see detailed methods in supplementary data 2). The overall mean fO₂ obtained is ΔFMQ +0.39 (reported in log units relative to the fayalite-magnetite-quartz oxygen buffer). Results obtained using the titanomagnetite-ilmenite oxybarometer (Sauerzapf et al., 2008) range from ΔFMQ +0.29 to +0.54, whereas µXANES, clinopyroxene-plagioclase (France et al., 2010), and olivine-liquid oxybarometry (Mallman & O'Neill, 2013) results span wider ranges of ΔFMQ +0.04 to +1.26, ΔFMQ -0.30 to +1.01, and ΔFMQ +0.3 to +1.3, respectively.

We estimated pressure (figure 9c) using the machine learning-based thermobarometer of Weber & Blundy (2024) with calculation parameters as described in supplementary data 2 for thermometry. Estimated pressures have a mean standard deviation of ±200.5 MPa. Bulk-rock lava flow and CB compositions pressures overlap within error (400.0–400.7 MPa and 400.3–600.2 MPa, respectively). Lava lake glasses return a wide range of pressures (100.7–500.4 MPa), but most are clustered near the median value of 400.4 MPa. This value is within the range recorded by BC8 olivine melt inclusions (400.3–400.7 MPa) and slightly above that recorded by BC18 olivine melt inclusion (300.6–300.7 MPa). Interstitial glasses in CB
samples (BC8 and BC10) record low pressures between 100.2 MPa and 300.0 MPa, with mean values of 200.6 MPa (BC8) and 100.3 MPa (BC10).

475

476 Thermodynamic calibration of the differentiation sequence

Here, we aim to test whether the fractional crystallization of a primitive Erta Ale Range basalt can explain the range of major element compositions observed at EAV. We used rhyolite-MELTS (Gualda et al., 2012) to model fractional crystallization during isobaric cooling at a constant fO₂ buffer. Among the 53 scenarios tested (see supplementary data 4 and 5), the model that best fit the data from both EAV and Erta Ale Range is a two-step fractional crystallization model (figure 10). In the first step (run 52; supplementary data 4), a theoretical basalt (composition B5) of composition similar to the most primitive basalts described in the Erta Ale Range by Barrat et al. (1998) and containing 0.1 wt.% water cooled from 1290 (liquidus) to 1200 °C at 400 MPa and ΔFMQ ~0.5. In the second step (run 24), the starting composition was the mean lava lake composition (composition B3; supplementary data 4) containing 0.2 wt.% water, which cooled from 1150 to 900 °C at 100 MPa ΔFMQ ~0.5.

Among the solid phases, clinopyroxene appears first at 1290 °C, then plagioclase at 1225 °C, olivine at 1148 °C), Fe-Ti oxide at 1108 °C, and finally apatite at 1038 °C (figure 10a). The model ultimately produces a solid phase assemblage comprising 51% clinopyroxene, 37% plagioclase, 3.5% olivine, 7% Fe-Ti oxide, and 0.6% apatite. The fractionation of these minerals modifies the liquid composition, decreasing its MgO content, and markedly increasing its SiO₂ content at the onset of Fe-Ti oxide crystallization at 4 wt.% MgO (figure 10b). The TiO₂ content in the liquid increases up to 3.8 wt.% until the onset of Fe-Ti oxide crystallization, and then decreases (figure 10c). We parameterized the relation between temperature and mineral composition obtained using rhyolite-MELTS and applied it to our
quantitative X-ray chemical maps (map presented in figure 4) to map temperature upon crystallization in figure 11 (see details in supplementary data 2).

Following Boulanger et al. (2020), we also used the relation between temperature and mineral composition to estimate composition-, temperature-, and pressure-dependent REE partition coefficients (D) at each step of fractional crystallization. We considered olivine, plagioclase, and clinopyroxene as the main mineral phases and derived D<sup>Cpx-Melt</sup>, D<sup>Ol-Melt</sup>, and D<sup>Pl-Melt</sup> from the internally consistent P-T-X-dependent models of Sun & Liang (2012, 2013, 2014) and Sun et al. (2017). These D values were used to calculate the evolving REE contents of the melt using a classic fractional crystallization model: C<sub>i</sub><sup>melt</sup> = C<sub>i,0</sub><sup>melt</sup> × F<sup>D<sub>i</sub>-1</sup>, with C<sub>i</sub><sup>melt</sup> the concentration of element i within the melt, C<sub>i,0</sub><sup>melt</sup> the initial concentration of element i within the melt, D<sub>i</sub> the global partition coefficient of element i, and F the remaining melt fraction. To investigate the extent of reactive porous flow during the crystallization sequence, we followed the approach developed for oceanic igneous systems (e.g., Boulanger et al., 2020; Sanfilippo et al., 2020; Ferrando et al., 2021) and modelled evolving mineral REE contents during assimilation-fractional-crystallization (AFC) using the equation of DePaolo (1981). The AFC model simulates melt evolution during assimilation of a pre-existing crystal matrix by a disequilibrium melt, and simultaneous crystallization of a new generation of minerals. We took the composition of the percolating melt to be the mean lava lake composition (supplementary data 6). Assimilated phases were 80% primitive plagioclase and 20% primitive clinopyroxene (supplementary data 6). The mass ratio of material assimilated to material crystallized (r) was set to 0.4. DePaolo (1981) proposed that 0.3 is the maximum value thermally consistent with the assimilation of a cold host rock (at ~150°C). Here we model the assimilation of a hot mush, and thus we explored r values between 0 and 1 (Boulanger & France, 2023) and finally selected 0.4 as the value allowing the best fit to our
521 data. Results of both the fractional crystallization and AFC REE models are presented in
522 figure 7a and supplementary data 6.

523

524 DISCUSSION

525 Magma differentiation and volcanic-plutonic connection

526 Evidence for protracted magma differentiation

527 The various datasets obtained herein allow us to reconstruct the liquid and crystal lines of
528 descent (LLD and CLD, respectively) of EAV lavas. Incompatible element contents and
529 fractionations (e.g., between strongly and weakly incompatible elements) are similar for all
530 bulk-rock data. The similarity between lava flows and CBs, as well as the lack of a
531 cumulative signature in the CBs, clearly highlight that their bulk compositions are
532 representative of their parental melt composition. Lava flow and CB compositions form a
533 single trend for most major elements (figure 5) and parallel normalized patterns for trace
534 elements (figures 6 and 7), indicating their cogenetic formation. Therefore, the observed
535 chemical trends can be considered as the LLD and can be compared with melt compositions
536 documented by interstitial glasses and melt inclusions.

537 The CLD is well documented by the mineral compositions (Fo in olivine, Mg# in
538 pyroxene, An in plagioclase), which span a large compositional range among the various lava
539 flow and CB samples, as well as within individual samples. In particular, plagioclase major
540 element compositions are not strongly affected by diffusive re-equilibration, and thus preserve
541 initial major element zonations (figures 3 and 4). The correlation between highly incompatible
542 elements like La, used as a fractionation index, and the contents of MgO and other highly
543 incompatible trace elements (figure 7) highlights that both the LLD and CLD can be
544 interpreted as differentiation trends. Based on the major element compositions, it appears that
545 differentiation was controlled by the crystallization of olivine (MgO decrease) and
546 clinopyroxene (CaO decrease), then plagioclase (Al₂O₃ decrease), Fe-Ti oxides (FeO and

547 TiO₂ decrease, SiO₂ increase), and finally apatite (P₂O₅ decrease) (figure 5). Mineral

548 zonations, specifically An in plagioclase (from high-An cores to low-An rims, as evidenced

549 by CaO zonations in figure 4b), strongly suggest that the differentiation process was not

550 equilibrium crystallization, but that fractional crystallization governed, at least partially, the

551 LLD and CLD. Although fractional crystallization has been widely proposed for Erta Ale

552 Range due to the presence of evolved magmas (Barberi & Varret, 1970; Bizouard et al., 1980;

553 Barrat et al., 1998; Hagos et al., 2016; Hutchison et al., 2018; Watts et al., 2020; 2023), its

554 impact has been considered to be only limited at EAV based on the lack of differentiated

555 lavas at this volcano (e.g., Barrat et al., 1998; Field et al., 2012a; Hagos et al., 2016).

556 However, the trachytic to rhyolitic compositions of the most evolved melt observed herein,

557 and the associated mineral phase compositions (plagioclase down to An₁₅, olivine down to

558 Fo₁₄, clinopyroxene down to Mg#25), contrast this view. Our results thus imply that

559 protracted magma differentiation proceeded at depth at EAV.

560

561 *Melt-mush reactions*

562 The CBs studied herein document the most evolved magmas ever reported at EAV. Another

563 peculiar feature is their doleritic to gabbroic textures, and the presence of interstitial glass in

564 some samples (figure 2 and 4). These coarse-grained textures and the presence of euhedral to

565 subhedral crystals likely document relatively slow cooling during the initial solidification of

566 the magma, and the presence of interstitial glass indicates a final quenching event related to

567 the transport of the CBs to the surface. The CBs therefore likely represent pieces of a

568 solidifying mush that was present at depth at the time of CB emplacement at the surface. Here

569 we consider a mush as a medium composed of crystals and melt (± gas), in which the crystals
570 form a continuous framework through which melt is distributed (e.g., Cashman et al., 2017; Boulanger & France, 2023).
571

572 In the CB samples, we observed mineral inclusions, irregular plagioclase boundaries, rounded olivine crystals, and mineral–liquid reactions (i.e., orthopyroxene reaction rims between olivine and interstitial melt in BC8; figure 4). Trace element data (figure 6 and 7) show that melt inclusions and interstitial glass have highly fractionated light/heavy REE ratios compared to the lavas (La/N/Yb/N value from 6.1 to 17.1). These petrographic and geochemical features have previously been interpreted as evidence for reactive porous flow (Coogan et al., 2000; Gao et al., 2007; Cooper et al., 2016; Boulanger et al., 2020, 2021; Sanfilippo et al., 2020; Zhang et al., 2021; Ferrando et al., 2021; Gleeson et al., 2021; Boulanger and France, 2023). Reactive porous flow causes magmatic evolution in mushy systems via the percolation of a liquid through a nearly immobile crystal framework, resulting in liquid–crystal disequilibrium, in turn causing dissolution–precipitation reactions (Bouilhol et al., 2015; Lissenberg et al., 2019; Sanfilippo et al., 2020; Boulanger & France, 2023). Our modeling results (figure 7a) and the observation of some resorbed mineral textures, support the role of such a reactive process within the EAV plumbing system. Therefore, the evolution of the EAV plumbing system is not solely governed by fractional crystallization, but also by reactive porous flow.

588

589 Evidence for interactions with hydrothermally altered reservoir margins: Insights from oxygen isotopes

591 The oxygen isotopic ratios of mantle rocks and mantle-derived magmas are rather homogeneous (δ¹⁸O = 5.2–5.8‰; Mattey et al., 1994; Eiler, 2001; Bindeman, 2008). Magmas with δ¹⁸O values that deviate from the mantle signature have been classically attributed to surface alteration, hydrothermal circulation, and/or crustal contamination (e.g., Gregory &
Taylor, 1981; Bindeman & Valley, 2001; Valley et al., 2005; Bindeman, 2008). Here, we track potential crustal contamination or interaction with hydrothermally altered reservoir margins by characterizing the oxygen isotopic compositions of olivine and glass in samples documenting the various stages of differentiation at EAV. For this, we follow the approach of Muehlenbachs & Byerly (1982), Bindeman (2008), and France et al. (2016). To directly compare our results with glass isotopic compositions, the oxygen isotopic compositions of the melt (δ¹⁸O) in equilibrium with olivine (δ¹⁸Oolivine) were calculated using fractionation factors (Δolivine-melt = δ¹⁸Oolivine - δ¹⁸Omelt). Because Δolivine-melt is dependent on olivine composition, we used the relation determined experimentally by Zhao & Zheng (2003) (figure 8c).

Among the processes that allow δ¹⁸O values to deviate from the mantle value during magma evolution, we modeled three various scenarios: fractional crystallization (FC; Eq. 9.46 in White, 1997), equilibrium crystallization (EC; Eq. 9.66 in White, 1997), and assimilation-fractional crystallization (AFC; Eq. 9.67 in White, 1997). Input values for these three equations are the melt fraction and the amount of each crystallizing mineral phase, which was taken step-by-step from the rhyolite-MELTS fractional crystallization model output.

Fractionation factors of crystallizing mineral phases were taken from Zhao & Zheng (2003), and we considered their variation with temperature (calculated at each step of the rhyolite-MELTS model). We used a MORB starting composition (Eiler, 2001). The AFC model also requires the mass ratio of material assimilated to material crystallized (r) and the δ¹⁸O value of the assimilated material. We tested isotopic compositions of the assimilated material ranging from 2‰ to 5‰, and mass ratios from 0 (corresponding to pure fractional crystallization) to 0.3, as those values are the best fit to our data. Modeling results are presented in figure 8c and supplementary data 7.

Our AFC modelling shows that olivine and glass oxygen isotopic compositions from EAV can be reproduced when r = 0–0.3 for an assimilated material with δ¹⁸O = 2‰. Pure
fractional crystallization produces only a minor $\delta^{18}$O$_{melt}$ increase, related to the fractionation of mineral phases (olivine, clinopyroxene, and plagioclase) that have relatively small fractionation factors. When Fe-Ti oxides start to crystallize (i.e., at MgO = 4.5 wt.%), the increase of $\delta^{18}$O$_{melt}$ becomes significant because oxides have larger isotopic fractionation factors than silicates (here fixed at $\Delta_{oxide-melt} = -4$; Zhao & Zheng, 2003). In contrast, for an assimilated material with the same $\delta^{18}$O value (2‰), AFC lowers $\delta^{18}$O$_{melt}$ values down to 4‰ for the most evolved compositions, reproducing almost all the EAV data.

Such a low $\delta^{18}$O component is likely representative of hydrothermally altered reservoir margins (e.g., Troch et al., 2020). Indeed, the high-temperature hydrothermal circulation of low-$\delta^{18}$O meteoric fluids at active volcanic systems causes alteration, generating low-$\delta^{18}$O volcanic crust (Norton and Taylor, 1979; Gregory & Taylor, 1981; Grimes et al., 2013). The assimilation of such a component was also proposed along with fractional crystallization for the Erta Ale Range (figure 8c; Barrat et al., 1998), and for several active segments in the Main Ethiopian and Afar Rifts (Barrat et al., 1993; 2003; Hutchison et al., 2018). Although the exact amount and composition of this assimilated material is not yet precisely constrained, it is clear that EAV (and possibly the larger Erta Ale Range) magmas are isolated from any high-$\delta^{18}$O continental crustal materials during their transfer through the plumbing system.

In figure 8c, the large range of $\delta^{18}$O values of the primitive (MgO ≈ 6.5 wt.%) lava lake glass fragments (4.4–6.8‰) is notable. Although the mean $\delta^{18}$O value is 5.8‰, the extreme values are quite far from the basaltic values expected for mantle derived magmas. This may be due to the peculiar dynamics of the lava lake system. Indeed, the EAV lava lake is known for its continuous and important gas flux (Oppenheimer et al., 2004; Sawyer et al., 2008), which could strongly alter the basalts in the rim of the lava lake pit crater. Indeed, altered basalts can have high $\delta^{18}$O values up to 20‰ (Eiler, 2001). Given the large lava-lake level fluctuations of several to tens of meters at EAV (Varet, 1971; Barberi et al., 1974;
Tazieff, 1973; Le Guern et al., 1979; Oppenheimer & Francis, 1997; Harris et al., 2005; Field et al., 2012a), the assimilation of altered pit-crater basalts into the lava lake could potentially lead to large δ¹⁸O heterogeneities in lava lake samples.

### Evidence for hybridization of magmas en route to the surface

Lava flow samples contain relatively poorly evolved olivine (Fo₇₀-₈₃; figure 3) consistent with their basaltic bulk compositions (figure 5). Nevertheless, some samples (BC13, BC17) contain a second olivine population with Fo₄₄-₅₃, similar to the range of olivine compositions observed in the more evolved CB samples. In BC13, the primitive olivine group has δ¹⁸O values between 5.9‰ and 6.3‰, whereas the evolved olivines have lower δ¹⁸O values between 4.7‰ and 5.2‰ (figure 8c) that are consistent with interactions with hydrothermally altered reservoir margins according to AFC models (figure 8c). The coexistence of two distinct olivine generations within a single magma clearly highlights that some mixing has occurred at depth within the plumbing system between basaltic magmas with mantle-like δ¹⁸O values and highly differentiated magmas with low, altered δ¹⁸O values. Such mixing may have occurred during the ascent of the basalt through the more evolved shallow reservoirs from which the CBs formed. A similar mixing process has likely occurred at the larger scale of the Erta Ale Range because some magmas clearly display bulk rock compositions that deviate from the general differentiation trends of Erta Ale Range and EAV lavas (black circles in figure 5). Those samples clearly plot on mixing lines in the Fe-Mg, Ti-Mg, P-Mg, and Si-Mg graphs, and thus document mixed compositions involving both basaltic and acidic magmas (figure 5). Those same samples are also enriched in Al, which may indicate that magma mixing was associated with plagioclase assimilation, as proposed by Hutchison et al. (2018) for other active rift segments in Afar and in the Main Ethiopian Rift.
670 *Inferences from thermodynamic modeling and oxy-thermobarometry*

671 To constrain the temperature, pressure, redox condition, water content, and starting
672 composition of the fractional crystallization and reactive porous flow processes that govern

673 both the LLD and CLD at EAV, we performed oxy-thermobarometry calculations and

674 thermodynamic modelling in rhyolite-MELTS (see subsection **Thermodynamic calibration of the differentiation sequence**' and supplementary data 2, 4 and 5).

676 Previously calculated $fO_2$ values for EAV lavas range from $\Delta FMQ$ -0.6 to +0.3 (figure
677 9b; Bizouard et al., 1980; de Moor et al., 2013). Our oxy-barometer calculations from

678 ilmenite-magnetite equilibrium and clinopyroxene-plagioclase equilibrium in BC8 and BC10

679 (CBs) are slightly above those values and indicate similar redox conditions as those obtained
680 from our XANES measurements. The overall mean value is $\Delta FMQ$ +0.35, with values

681 ranging from $\Delta FMQ$ -0.3 to +1.26, which are both consistent with previous data and in the
682 range of mid-ocean ridge basalts (Jugo et al., 2010), suggesting that oxygen fugacity is quite

683 steady at EAV. The lava lake samples, which are likely partially degassed, exhibit similar

684 redox conditions to the other samples, indicating that degassing does not significantly affect
685 the magma redox state at EAV.

686 Temperatures previously calculated for EAV lavas range from 900 °C to 1230 °C
687 (Giggenbach & Le Guern, 1976; Field et al., 2012a; de Moor et al., 2013), and various authors

688 have directly measured lava lake temperatures up to 1200 °C (Le Guern et al., 1979; Burgi et
689 al., 2002; Oppenheimer & Yirgu, 2002; Oppenheimer et al., 2004; Sawyer et al., 2008;

690 Spampinato et al., 2008; figure 9a). Our thermometry calculations highlight that ilmenite-

691 titanomagnetite pairs equilibrated at similar temperatures to plagioclase rims in equilibrium
692 with their adjacent interstitial melt (966–1015 °C, mean 988 °C; figure 9a). This is consistent

693 with temperature estimates obtained by applying the machine-learning thermometer to

694 interstitial glasses and melt inclusions in CB samples (935–1040 °C; figure 9a). Those results
correspond to the temperature at which the last phases stabilized in CB sample BC8, and thus document the last stages of the protracted differentiation recorded in EAV samples. Indeed, Fe-Ti oxides are among the last phases to crystallize in the studied crystallization series, and the silicic melt is also produced at the end of the protracted differentiation.

Temperature estimates obtained with the Petrolog3 (Danyushevsky & Plechov, 2011) document earlier (thus hotter) conditions corresponding to melt inclusion entrapment. This calculation was applied to two samples (CB BC8 and lava flow BC18) and gave similar temperature estimates of 1077–1100 °C (mean 1088 °C; figure 9a). Olivine crystallizes relatively early in the CLD and continues to crystallize as increasingly Fe-rich olivine through the last stages of evolution. Olivine–melt inclusion equilibrium temperatures thus likely document the differentiation stage at which those melt inclusions were trapped. However, BC8 and BC18 temperatures are similar, even though BC18 contains olivine and melt inclusion with less evolved compositions than BC8 (figures 3b and 5). This suggests that, although BC8 evolved toward more evolved compositions than BC18, its minerals started to crystallize at higher temperatures, closer to the beginning of differentiation. This is consistent with the temperatures recorded by the bulk compositions of the lava flow and CB samples (recording melt–mush equilibrium at 1105–1190 °C before their extraction) and is consistent with plagioclase compositions up to An₈₀ in both samples (figure 3a).

These results indicate that as the bulk magma was differentiating, BC8 olivine hosts and their melt inclusions re-equilibrated through diffusion with the surrounding magma to achieve evolved (iron-rich) compositions. Indeed, olivine achieves chemical equilibration relatively quickly (Chakraborty, 2010), and the lack of zoning in BC8 olivines (figure 4c) strongly suggests that they re-equilibrated with the surrounding evolved melt (which requires only a few years for olivine grains smaller than 100 µm; see later subsection ‘Preliminary constraints on the kinetics of igneous processes’). Furthermore, olivine-hosted melt inclusions
also re-equilibrate quickly by diffusion through their host (Gaetani & Watson, 2000; Cottrell et al., 2002). Indeed, the similarity of melt inclusion and interstitial glass compositions in BC8 (figure 5) suggests re-equilibration between the melt inclusions and the interstitial melt surrounding their host olivines. Finally, the highest temperatures are recorded by lava lake glasses (1152–1184 °C), in agreement with direct measurements of the lava lake. Thus, the various thermometers record different parts of the cooling path between ~1200 °C and ~900 °C (figure 11).

Radar interferometry data indicate the presence of shallow axial magma chambers (~1–2 km deep) that produce dykes that feed eruptions (Moore et al., 2019; Xu et al., 2017, 2020). These depths are consistent with our lowest pressure estimate from CB interstitial glasses (100.2–300.0 MPa; figure 9c). However, bulk-rock and melt inclusion compositions indicate pressures of 300.6–600.2 MPa, requiring transcrustal rather than shallow magma reservoirs.

733

## 734 Plumbing system architecture: An integrated model

By combining our and previous results, we can reconstruct the plumbing system organization and related processes that govern magma differentiation at EAV. We summarize these results in the schematic plumbing system model presented in figure 12. Seismic reflection/refraction data suggest a crustal thickness of ~15 km (Hammond et al., 2011; Keir et al., 2013). Our thermodynamic models suggest that magma differentiation at EAV initiates in relatively primitive basaltic magma reservoirs in the deep crust (~400 MPa kbars), and our machine-learning barometry calculation suggests that bulk rocks, lava lake glasses, and melt inclusions record initial pressures between 300.6 and 600.2 MPa, placing their origin in the lower portion of the 15-km-thick crust. Those mafic magmas likely fed both eruptions at the surface and shallower reservoirs where further differentiation produced the CBs. The bulk rock
745 compositions of the CBs, which represent the mobilized melt composition, document the
746 composition of such magmas intruded at shallower levels. Indeed, the most evolved CB bulk-
747 rock composition (BC4) equilibrated at conditions similar to those at which the most
differentiated plagioclase crystals in lava flows equilibrated (e.g., BC17, BC18; figure 13).
748

749 The compositional range from the most primitive lava flows and CBs to the most
750 evolved bulk CB composition (or to the most evolved mineral rim in a lava flow) likely
documents the varied extents of differentiation occurring in the deep crustal reservoirs. There,
751 magmas differentiate mostly by fractional crystallization in a crystal-poor environment and
752 are thus mobilizable. Once intruded into cooler, shallower reservoirs (likely below 300.0 MPa
753 based on CB interstitial glasses), those magmas either inject dikes toward the surface (e.g., Xu
754 et al., 2017), or undergo further differentiation, as recorded by the intense zonations in various
755 CB minerals (figures 3, 4, 11 and 13). This protracted differentiation stage ultimately forms
756 rhyolitic interstitial melts that are rich in water and incompatible elements. The partial
757 solidification of magmas within shallow reservoirs forms crystal-rich mushes (uneruptible
758 igneous media) in which interstitial melts can only migrate via reactive porous flow (Cooper
759 et al., 2016; Gleeson et al., 2021; Boulanger & France, 2023; Mourey et al., 2023).
760

761 At EAV, petrographic and geochemical evidence of reactive porous flow are restricted
762 to CB interstitial melts, and not observed in bulk CB and lava flow compositions, strongly
763 supporting that this process is restricted to mushy igneous environments. Despite local
764 migration through mushes in shallow reservoirs, Si-rich interstitial melts likely never
765 aggregate sufficiently to feed surface eruptions, because no such evolved compositions have
766 yet been effused at EAV. Nonetheless, similar magmas have been emitted elsewhere in the
767 EAR, although they require further study to clarify whether they represent the extraction of
768 similar interstitial Si-rich melts or rather the last products of magmatic cycles at various
eruptive centers. In any case, at EAV and at the larger scale of the EAR, the presence of such
769
evolved magmas is also attested by the hybridization of mafic magmas (lava flows) during their transfer to the surface (figures 3, 10 and 12). Indeed, lava flows record evidence of both magma mixing (figure 12) and mush cannibalism (figures 3, 5 and 8). The key features of these processes are the clear mixing trends in the bulk-rock compositions, Al enrichments in hybridized melts that indicate plagioclase assimilation, and the presence of evolved minerals similar to CB mushy assemblages within basaltic lava flow samples. Our results highlight that, although they do not erupt, the presence of such evolved melts at depth clearly influences the composition of some erupted magmas.

Beyond its transcrustal architecture, oxygen isotopes highlight that the EAV plumbing system is isolated from the continental crust, implying a well-developed juvenile igneous crust. Our results do not necessarily imply that continental crust has been fully replaced by mafic crust on a large scale, but rather that in the center of this continental rift segment, crystallized plutonic rocks are sufficiently abundant to shield the plumbing system from any continental crust influence (figure 12). This fact alone highlights the significant role of endogenous growth at EAV. Indeed, interactions with hydrothermally altered magmatic rocks from previous igneous stages have been reported in the Erta Ale Range (Barrat et al., 1998), elsewhere in Afar (Manda Hararo segment, Barrat et al., 2003; Tadjoura Gulf, Barrat et al., 1993), at mid-ocean ridges (e.g., France et al., 2010, 2021; Grimes et al., 2013), in Iceland (e.g., Bindeman et al., 2012; Schattel et al., 2014; Caracciolo et al., 2022), and at Yellowstone (e.g., Bindeman & Valley, 2001).

Application of our thermodynamic LLD to samples related to lower crustal differentiation documents temperature variations from >1200 °C to 1080 °C, consistent with thermometry estimates from bulk lava flow and CB compositions, as well as lava flow plagioclase compositions (figure 13). Magmas injected into the shallow upper crust thus equilibrated at ~1200–1080 °C, then underwent protracted differentiation in mushy reservoirs
down to ~900 °C before their ultimate solidification or partial remobilization during basaltic hybridization (figures 12 and 13). Because the entirety of the shallow differentiation stage (i.e., from ~1080 to ~900 °C) is not expressed by eruptions at the surface, it thus represents significant endogenous growth at EAV, consistent with earlier propositions by Francis et al. (1993) and Oppenheimer & Francis (1997).

800

## 801 Implications

### 802 Preliminary constraints on the kinetics of igneous processes

We have presented clear evidence that magma is stored within the crust to form mushes and very evolved compositions represented petrographically by the CB samples. However, the rarity of CBs erupting creates a sampling bias because only more primitive basaltic lava flows are described in the literature (e.g., Barrat et al., 1998; Field et al., 2012a; Hagos et al., 2016).

Field et al. (2012a) estimated the residence times of lava flow olivines based on the absence of Ca zoning in olivine, requiring a minimum homogenization time of ~10 years, and the preservation of P zoning, which would otherwise be homogenized after ~94 years.

Nevertheless this maximum duration should be considered with great caution as previous diffusivity data of P in olivine were recently challenged by Nelson et al. (2024), suggesting that P homogenization would have required several thousands of years. Similarly, we did not observe Fe zoning in olivines, but only preserved Fe zoning in clinopyroxene (figures 3b, 3c, and 4c) in CB samples, which we use here to further constrain residence times at EAV.

Models of Fe-Mg interdiffusion in BC8 (the most primitive CB) olivines that are < 100 µm (figures 2 and 4) within a melt of BC8 interstitial glass composition at ~1050 °C (figures 9 and 11) give total olivine homogenization times between ~72 days and ~4.7 years depending on crystallographic axis (99.99% of diffusive re-equilibration; figure 14).

Diffusion along the fastest axis ([001]) likely governs Fe-Mg homogenization. Because the
absence of Fe-Mg zonation in these olivines is ubiquitous, we can consider that the minimum time required to obtain 100% diffusive re-equilibration of CB olivines to be of 72 days (figure 14). Models of Fe-Mg interdiffusion in BC8 clinopyroxenes no larger than 200 µm (figures 2 and 4) within a melt of BC8 interstitial glass composition at ~1000 °C (figure 11) give total homogenization times between ~15 kyr and ~61 kyr (99.99% of diffusive re-equilibration; figure 14). Here, we consider that Fe-Mg zonations in clinopyroxene can be preserved up to 50% diffusive re-equilibration. Although Müller et al. (2013) only provided a diffusion coefficient along the [001] axis for Fe-Mg interdiffusion in clinopyroxene, it is the fastest diffusive axis and thus provides a conservative estimate of the upper limit of EAV residence times at ~6070 years (figure 14). It is also possible that the system existed for significantly longer if it is periodically subjected to replenishment by new melt that drives chemical zoning within crystals. In that case, ~6070 years would be the age of the last replenishment.

In both olivine and clinopyroxene, increased temperature and/or decreased grain size will significantly accelerate Fe-Mg interdiffusion (figure 14). In the range of values explored here, grain size has a stronger impact than temperature (figure 14). If we consider BC8 as representative of CBs, the total lifetimes of mushy magmatic reservoirs at EAV should be >72 days, and <6070 years, significantly longer than the residence times of magmas feeding lava flows (Field et al., 2012a). This plutonic activity could be one of the processes accommodating extension at EAV, as was proposed for the Dabbahu rift segment (Field et al., 2012b; Wright et al., 2012).

Rift break-up at Erta Ale: Continental or oceanic crust?

The question remains as to whether EAV is at the final stages of continental break-up. If we consider that break-up is achieved when the crust becomes oceanic rather than continental (e.g., DMM source, lack of continental lithosphere, crustal thickness, juvenile crust...), our
results have implications for typical arguments in this debate. Hutchison et al. (2018) have suggested that the lack of differentiated magmas at EAV is characteristic of recent break-up achievement. Our finding of Si-rich magmas in the active plumbing system of EAV might therefore challenge this interpretation. Nevertheless, although acidic magmas are rare in oceanic crust, they do exist at both fast- and slow-spreading centers, mainly as felsic intrusions in the plutonic sections (e.g., France et al., 2009; Nguyen et al., 2018; Chen et al., 2019). Hydrous partial melting of hydrothermally altered mafic crust, and protracted differentiation processes are both proposed to form such Si-rich magmas (e.g., Grimes et al., 2013; Erdmann et al., 2015; Erdmann & Koepke, 2016). Similarly to what we observe herein at Erta Ale Range and EAV, contamination of MORB by those Si-rich magmas has been reported at oceanic centers (e.g., Coogan et al., 2003; Koepke et al., 2011; France et al., 2014; Fischer et al., 2016). Following this comparison with oceanic spreading centers, we therefore conclude that the identification of protracted differentiation within the EAV plumbing system is not discriminating to discuss the advancement of the break-up process. The absence of a continental crustal signature and evidence of contamination by hydrothermally altered low-δ¹⁸O material have been cited as indicating proto-oceanic crust formation at Erta Ale Range (Barrat et al., 1998; Hutchison et al., 2018), but such features are also observed in continental settings when magma fluxes are high enough to shield the plumbing system from the continental crust (e.g., Troch et al., 2020), and we thus believe that such an observation is not discriminating either to decipher on the advancement of the break-up process.

More significantly, geophysical data and related geodynamic modelling suggest that the continental lithospheric mantle is absent at the rift axis, and that decompression melting in the absence of a strong plume-related thermal anomaly governs the melting process (Rychert et al., 2012). Nevertheless, those data were challenged by present day magma composition that is rather consistent with hot and deep melting below a 60-80 km-thick mantle lithosphere
870 (Ferguson et al., 2013). In any cases magma source at Erta Ale Range and EAV is not similar
to MORB (depleted convecting upper mantle), neither in terms of trace element signature
871
(figure 6 and 7), nor isotopically (Barrat et al., 1998). The crustal thickness at Erta Ale Range
872
(~15 km; Hammond et al., 2011; Keir et al., 2013) is also atypical of oceanic spreading
873
centers. Furthermore, Moore et al. (2019) suggested that the shallow depth of magma storage
874
at Erta Ale Range indicates complete break-up, but only ultrafast oceanic spreading centers
875
have magmatic reservoirs at ~1.5–2 km depth where the shallow axial reservoirs have been
876
reported by geophysical studies at EAV. In contrast, the spreading rate at Erta Ale Range is
877
more similar to slow or ultraslow oceanic spreading centers (McClusky et al., 2010) where
878
magmatic reservoirs are expected at 3–5 km depth (Morgan & Chen, 1993). Although this is
879
similar to our inferred depths of the shallow reservoirs at EAV, we rather propose that the
880
EAV plumbing system is transcrustal, extending to ~12 km depth.
881

882
Finally, despite EAV and Erta Ale Range magmas no longer interacting with the
continental crust, the crust is better classified as transitional and highly injected, indicating
883
that Erta Ale is within the ongoing break-up stage, but has not surpassed this stage because
884
the crust cannot be defined as oceanic.
885

886

887 CONCLUSIONS

888 We report trachytic and rhyolitic compositions recorded in cognate block samples at Erta Ale
Volcano, the first description of such plutonic-type samples there. They formed by a
889
combination of protracted fractional crystallization of a parental basaltic magma and reactive
890
porous flow in a mushy shallow magmatic reservoir. These samples represent the endogenous
891
growth of EAV, which has rarely been sampled. Oxygen isotopic data indicate that EAV
892
magmas are isolated from any high-δ¹⁸O continental crust during their evolution in the
893
plumbing system. Nevertheless, they interact with a shallow, high-temperature,
894
hydrothermally altered low-δ¹⁸O component. Similar data and interpretations at the scale of
the Erta Ale Range reflect that EAV has a rifting style characterized by a complex transcrustal
magmatic plumbing system, allowing for both local extreme melt evolution and magmatic
accommodation of extension. We propose a detailed model of this plumbing system
architecture. Despite the well-developed juvenile igneous crust at EAV, a transcrustal
plumbing system that spans ~15-km-thick crust and includes shallow rhyolitic reservoirs is
dissimilar to any known oceanic spreading center plumbing system. Our results therefore
highlight that, despite high magma fluxes along the Erta Ale continental rift segment and at
EAV, continental break-up cannot be considered completed in terms of igneous architecture,
but rather transitional between continental and oceanic crust.

905

## DATA AVAILABILITY STATEMENT

The data underlying this article are available in the OTELO Research Data Repository, at
https://dx.doi.org/10.24396/ORDAR-145, and are available in the article online
supplementary material.

910

## ACKNOWLEDGEMENTS

We thank the 3 anonymous reviewers for their constructive comments, as well as the associate
editors Madeleine Humphreys and Julia Hammer, and the editor in chief Georg Zellmer for
their editorial handling, which greatly improved the quality of this manuscript. The work
presented herein was made possible through the financial support of the ANR project n° 20-
CE49-0012 and to the Université de Bretagne Occidentale. We thank Jacques Durieux and
Bernard Marty for providing the lava lake samples, Chantal Peiffert assistance during the LA
ICP-MS analyses, Jessica Langlade for her help with EPMA analyses, and Nordine Bouden
for his assistance during SIMS analyses. Jakub Szlachetko, Baptiste Mary, and Manuel
Munoz are thanked for their assistance during µXANES measurements, and help during data processing. Roman Botcharnikov is thanked for sharing glass standards. Jean-Pierre Oldra is thanked for thin section preparation and sample preparation. Valentine Charvet's guidance and training in chemical map processing using the thermodynamic model was greatly appreciated.

We acknowledge Raphaël Pik for discussions that enriched our understanding and interpretation of the data presented herein, and Robert Dennen for formatting and proofreading the manuscript.

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1347 C.S., 2005. 4.4 billion years of crustal maturation: oxygen isotope ratios of magmatic

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1352 Watts, E.J., Gernon, T.M., Taylor, R.N., Keir, D., Pagli, C., 2023. Magmatic evolution during

1353 proto-oceanic rifting at Alu, Dalafilla and Borale Volcanoes (Afar) determined by trace
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1355 https://doi.org/10.1016/j.lithos.2023.107311

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1359 https://doi.org/10.1016/j.jvolgeores.2020.107094

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1362 White, W.M., 1997. Chapter 9: Stable Isotope Geochemistry.

1363 Wilke, M., Farges, F., Petit, P.-E., Brown, G.E., Martin, F., 2001. Oxidation state and

1364 coordination of Fe in minerals: An Fe K- XANES spectroscopic study. American

1365 Mineralogist 86, 714-730. https://doi.org/10.2138/am-2001-5-612

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1369 land. Nature Geosci 5, 242-250. https://doi.org/10.1038/ngeo1428

1370 Xu, W., Rivalta, E., Li, X., 2017. Magmatic architecture within a rift segment: Articulate

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1376 Zhang, W.-Q., Liu, C.-Z., Dick, H.J.B., 2021. Evidence for Multi-stage Melt Transport in the

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1379 https://doi.org/10.1093/petrology/egaa082

1380 Zhao, Z.-F., Zheng, Y.-F., 2003. Calculation of oxygen isotope fractionation in magmatic

1381 rocks. Chemical Geology 193, 59-80. https://doi.org/10.1016/S0009-2541(02)00226-7

1382

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1391

1392
1393 **Table legend**

1394

<table><caption>Table 1. Sample names, types, locations (coordinates), and emplacement dates (when known).</caption><thead><tr><th>Sample</th><th>Type</th><th>Latitude</th><th>Longitude</th><th>Comment</th></tr></thead><tbody><tr><td>BC1</td><td>Cognate Block</td><td>13.61475</td><td>40.67166667</td><td></td></tr><tr><td>BC2</td><td>Cognate Block</td><td>13.61619444</td><td>40.67197222</td><td></td></tr><tr><td>BC3</td><td>Cognate Block</td><td>13.61619444</td><td>40.67197222</td><td></td></tr><tr><td>BC4</td><td>Cognate Block</td><td>13.61619444</td><td>40.67197222</td><td></td></tr><tr><td>BC5</td><td>Cognate Block</td><td>13.61705556</td><td>40.67277778</td><td></td></tr><tr><td>BC6</td><td>Cognate Block</td><td>13.61475</td><td>40.67355556</td><td></td></tr><tr><td>BC7</td><td>Cognate Block</td><td>13.60941667</td><td>40.67741667</td><td></td></tr><tr><td>BC8</td><td>Cognate Block</td><td>13.60955556</td><td>40.67766667</td><td></td></tr><tr><td>BC9</td><td>Cognate Block</td><td>13.60894444</td><td>40.667</td><td></td></tr><tr><td>BC10</td><td>Cognate Block</td><td>13.60894444</td><td>40.667</td><td></td></tr><tr><td>BC11</td><td>Cognate Block</td><td>13.60894444</td><td>40.667</td><td></td></tr><tr><td>BC12</td><td>Cognate Block</td><td>13.60894444</td><td>40.667</td><td></td></tr><tr><td>BC15</td><td>Cognate Block</td><td>13.60866667</td><td>40.67422222</td><td></td></tr><tr><td>BC13</td><td>Lava flow</td><td>13.60891667</td><td>40.66763889</td><td>1973 eruption</td></tr><tr><td>BC14</td><td>Lava flow</td><td>13.60736111</td><td>40.6735</td><td>1973 eruption</td></tr><tr><td>BC16</td><td>Lava flow</td><td>13.60302778</td><td>40.67319444</td><td>Spatter cone<br>Outside caldera; Overlays</td></tr><tr><td>BC17</td><td>Lava flow</td><td>13.60425</td><td>40.71744444</td><td>BC18</td></tr><tr><td>BC18</td><td>Lava flow</td><td>13.60425</td><td>40.71744444</td><td>Outside caldera</td></tr><tr><td>BC21</td><td>Lava flow</td><td>13.60591667</td><td>40.66983333</td><td>Overlayed by BC13 & BC14</td></tr><tr><td>ERTA</td><td>Pele's hair</td><td></td><td></td><td>2003 field campaign</td></tr><tr><td>LL1</td><td>Lava lake</td><td></td><td></td><td>Durieux 2001</td></tr><tr><td>LL2011</td><td>Lava lake</td><td></td><td></td><td>Marty 2011</td></tr></tbody></table>

1396

1397

1398

1399

1400
1401 **Figures legends**

1402

1403 Fig. 1. (a) Geological map of Erta Ale volcano (1:100,000 scale; from Barberi & Varet, 1970) highlighting its northern and southern calderas. (b) A 2024 CNES satellite image of the same area (courtesy of Google Earth) indicating the locations of samples BC17 and BC18.
1404
1405 (c) Enlarged view of the northern EAV caldera showing its north and south pit craters and sampling locations. During the 2003 field campaign, the permanent lava lake was in the southern pit crater. (d) Regional location map showing the rift segments of focused magmatic intrusion (black filled ovals) after Rooney et al. (2014). Small inset map shows the plate boundaries. MER = Main Ethiopian Rift, TGD = Tendaho-Goba'ad discontinuity, Ar. = Arabian Plate, So. = Somalian Plate, and Nu. = Nubian plate. Background topography derived from NASA/Jet Propulsion Laboratory (JPL) Shuttle Radar Topography Mission (SRTM) database. (e) and (f) Circular inset pictures documenting cognate blocks (CB) size and emplacement that indicate that they were ejected from Erta Ale volcano.

1415

1416 Fig. 2. Thin section (plane-polarized light, ppl; cross-polarized light, xpl; or reflected light, as indicated) and backscattered electron (BSE) microphotographs of lava flow (LF) and cognate block (CB) samples. (a) BC8 (CB, ppl); interstitial matrix glass (gl) with apatite (Ap). (b) BC18 (LF, xpl); aggregated phenocrysts of plagioclase and olivine in a microlithic matrix. (c) BC13 (LF, xpl); a plagioclase with a spongy core and zoned rim. (d) BC8 (CB, xpl); plagioclase (Pl) inclusion within an olivine (Ol). (e) BC8 (CPB, ppl); melt inclusions (MI) in olivine. (f) BC8 (CB, reflected light), containing chalcopyrite (Ccp). (g) BC11 (CB, ppl); doleritic texture. (h) BC4 (CB, ppl); amphibole (Amp) overgrowth around a clinopyroxene (Cpx) core. (i) BC8 (CB, xpl); an inherited spongy plagioclase with zoned rims and irregular boundaries in a gabbroic texture. (j) BC4 (CB, BSE); exsolution textures
1426 between ilmenite and titanomagnetite. (k) BC8 (CB, BSE); a melt inclusion in olivine, with
1427 oxide (ox) and apatite having crystallized from the melt inclusion. (l) BC8 (CB, BSE); a
1428 reaction rim between olivine and the trachytic matrix glass to form orthopyroxene (Opx).

1429
1430 Fig. 3. Chemical compositions of main mineral phases in lava flows (bottom; LF) and
1431 cognate blocks (top; CB). (a) plagioclase (An content). (b) olivine (Fo content). (c) pyroxene
1432 (Mg#). Open symbols indicate analyses of mineral rims. Sample are organized in the y axis by
1433 order of plagioclase An minimum content, from the most An poor at the top to the most An-
1434 rich at the bottom. This order of presentation is kept in panel b and c, and correspond to the
1435 color gradient of the samples, between yellow for the An rich and red for the An poor.

1436
1437 Fig. 4. Chemical maps of (a) Al₂O₃, (b) CaO, (c) FeO, and (d) TiO₂ (wt.%) in sample BC8
1438 and (e) interpretive phase map. For clarity, parts of the CaO map with >25 wt.% CaO have
1439 been removed, implying apatite are in the same color as the holes (black) on the CaO map.
1440 Mapping data were processed using XMapTools 3.4 (Lanari et al. 2014; 2019).

1441
1442 Fig. 5. Major element bulk-rock compositions of lava flow samples, cognate blocks, in-situ
1443 melt inclusions (corrected from post-entrapment processes; see Supplementary data 1), and
1444 interstitial glasses. Erta Ale Range literature data points displaying compositions suggesting
1445 the involvement of magma mixing are circled in black (see section Evidence for hybridization
1446 of magmas en route to the surface). FeOᵢ = FeO + (Fe₂O₃/1.11). Literature data are from
1447 Barberi et al. (1974), Bizouard et al. (1980), Barrat et al. (1998), Field et al. (2012), Hagos
1448 et al. (2016), and Watts et al. (2020).

1449
Fig. 6. (a, c) Chondrite-normalized rare earth element (REE) abundances in (a) bulk-rock lava flows and cognate blocks and (c) melt inclusions and interstitial glasses. (b, d) Primitive-mantle-normalized trace element abundances in (b) bulk-rock lava flows and cognate blocks and (d) melt inclusions and interstitial glasses. Chondrite and primitive mantle values are from Sun & McDonough (1989). Literature data are from Barrat et al. (1998), Field et al. (2012), and Hagos et al. (2016).

Fig. 7. (a) La<sub>N</sub>/Yb<sub>N</sub> vs. La, (b) MgO vs. La, and (c) Zr vs. La diagrams for lava flows, cognate blocks, melt inclusions, and interstitial glasses. Literature data are from Barrat et al. (1998), Field et al. (2012), and Hagos et al. (2016). Fractional crystallization (FC) and assimilation-fractional-crystallization models (AFC) are shown on the La<sub>N</sub>/Yb<sub>N</sub> vs. La graph in solid and dashed lines, respectively. Details of these models are described in subsection 'Thermodynamic calibration of the differentiation sequence' and in Supplementary data 6.

Fig. 8. (a) Oxygen isotopic compositions vs. SiO<sub>2</sub> contents in interstitial glasses and melt inclusions, after correcting from IMF and matrix effect (see Supplementary data 2a & 2c). (b) Oxygen isotopic compositions vs. Fo in olivine, after correcting from IMF and matrix effect (see Supplementary data 2a & 2c). (c) Calculated δ<sup>18</sup>O<sub>melt</sub> (‰) vs. melt MgO content (wt.‰). δ<sup>18</sup>O<sub>melt</sub> values are either directly those measured in glasses and melt inclusions or are calculated olivine using a fractionation factor (Δ<sub>olivine-melt</sub> = δ<sup>18</sup>O<sub>olivine</sub> - δ<sup>18</sup>O<sub>melt</sub> determined experimentally by Zhao & Zheng (2003)). MgO* is the melt MgO content, either directly measured by EPMA in glasses and melt inclusions (SiO<sub>2</sub> and MgO contents of melt inclusions were recalculated with Petrolog3 (see Supplementary data 1)) or calculated for the olivines.
We determined the composition of the melt in equilibrium with the olivine composition based on the thermodynamic MELTS model presented in subsection 'Thermodynamic calibration of
1475 **the differentiation sequence**'. Details of the fractional crystallization, equilibrium crystallization, and assimilation-fractional crystallization models are given in subsection 1476
1477 **'Evidence for interactions with hydrothermally altered reservoir margins: Insights from oxygen isotopes'** and Supplementary data 7. Analytical uncertainties are smaller than the symbol sizes. CB, cognate block; LF, lava flow. Erta Ale Range and Volcano literature values from Barrat et al. (1998) are derived from whole-rock δ¹⁸O values assuming that they provide a close approximation to δ¹⁸O_melt. MORB, continental crust and hydrothermally-altered rocks ranges are from Eiler (2001). 1478
1481
1482
1483

1484 Fig. 9. (a) Temperature estimations using five different thermometers (n = number of analyses used in the calculation) and comparison with other thermometry studies and direct lava lake measurements from the literature. (b) Oxygen fugacity estimates using four different oxybarometers and comparison with literature data. (c) Pressure estimates obtained using the machine learning-based thermobarometer of Weber & Blundy (2024). Literature data sources: a, Field et al. (2012); b, de Moor et al. (2013); c, Giggenbach & Le Guern (1976); d, direct lava lake measurements (Le Guern et al., 1979; Burgi et al., 2002; Oppenheimer & Yirgu, 2002; Oppenheimer et al., 2004; Sawyer et al., 2008; Spampinato et al., 2008); e, Bizouard et al. (1980). 1485
1486
1487
1488
1489
1490
1491
1492
1493

1494 Fig. 10. Thermodynamic rhyolite-MELTS fractional crystallization model that best fits both natural Erta Ale Range and EAV data. (a) The evolution of the mineral assemblage (plagioclase, clinopyroxene, olivine, Fe-Ti oxide, and apatite, reported in cumulative normalized proportions of solid phases only) and the amount of liquid remaining during cooling. (b) SiO₂ and (c) TiO₂ contents in the liquid vs. liquid MgO content during the modelled fractional crystallization sequence compared to natural data from figure 5. The 1495
1496
1497
1498
1499
mixing model was obtained by drawing a line between the lava lake composition and the interstitial glass composition. In all panels, the dashed vertical line indicates the limit between steps 1 and 2 of the model (see subsection 'Thermodynamic calibration of the differentiation sequence'). Details of the thermodynamic calculations are presented in Supplementary data 2.

Fig. 11. Temperature map of cognate block BC8 calculated from the chemical compositions of mineral phases (clinopyroxene, plagioclase, olivine, Fe-Ti oxides, apatite, orthopyroxene, glass) based on the thermodynamic rhyolite-MELTS model results. Details of the thermodynamic calculations and map translation are presented in supplementary data 2.

Fig. 12. Schematic model of the plumbing system architecture at Erta Ale Volcano (modified after Cashman et al., 2017). Numbers indicate the various processes occurring. (1) Initial magma differentiation upon early crystallization of clinopyroxene and plagioclase at depth (~400 MPa). (2, 3) Protracted differentiation at shallower depths (~100 MPa), encompassing (2) the onset of olivine crystallization, increasing the TiO₂ content of the melt, followed by (3) Fe-Ti oxide crystallization at <4 wt.% MgO, sharply decreasing the TiO₂ content of melt. (4) Mixing between primitive and differentiated magmas (see subsection 'Evidence for hybridization of magmas en route to the surface'). (5) Thermal evolution of the mushy part of the plumbing system. (6) Shallow interactions with hydrothermally altered reservoir margins (see subsection 'Evidence for interactions with hydrothermally altered reservoir margins: Insights from oxygen isotopes'). (7) Depth estimates from pressure (see subsection 'Oxy-thermobarometry' and Fig. 9c) assume 100 MPa ≈ 3 km depth. The evolution of the MgO and water contents of the liquid during cooling as modeled in rhyolite-MELTS are shown at right (see subsection 'Thermodynamic calibration of the differentiation sequence'). The model
1525 results are annotated with the temperatures of first crystallization of the rock forming
1526 minerals (dashed gray lines), equilibrium temperatures determined using different
1527 thermometers (hatched boxes; see subsection 'Inferences from thermodynamic modeling and
1528 oxy-thermobarometry'), and the mean MgO compositions of the lava lake and of interstitial
1529 glasses and melt inclusions in CBs.

1530
1531 Fig. 13. Thermal and compositional evolution of lava flow and cognate block samples.
1532 Clinopyroxene (Mg#), olivine (Fo), plagioclase (An), Fe-Ti oxide (TiO2), interstitial glass
1533 (MgO), and bulk (MgO) compositions are used to infer the thermal evolution of magmas
1534 based on the LLD and CLD obtained by application of our thermodynamic rhyolite-MELTS
1535 fractional crystallization model. The bulk-rock (melt) composition represents the start of the
1536 thermal and compositional evolution of each sample. As suggested by barometer calculations,
1537 those compositions represent deep magmatic evolution. The samples are plotted vertically in
1538 order of their bulk MgO contents and horizontally by taking the rhyolite-MELTS temperature
1539 for each composition as the starting temperature for each sample. More evolved in-situ
1540 mineral, glass, melt inclusion compositions document the end of each sample's thermal and
1541 compositional evolution. Barometric calculations applied to interstitial glasses indicate that
1542 this later magmatic evolution mainly occurs in the shallow crust.

1543
1544 Fig. 14. Percentage of Fe-Mg re-equilibration (= [initial concentration – concentration at the
1545 given time step] / [initial concentration – equilibrium concentration]) in olivine or
1546 clinopyroxene cores modeled as a function of time for different temperatures, grain sizes, and
1547 along different crystallographic axes in olivine. The thickest curves indicate modeled
1548 conditions similar to those observed and estimated for CB sample BC8 (<100 µm olivines
1549 equilibrated at 1050 °C along the slowest diffusive axes; <200 µm clinopyroxenes
equilibrated at 1000 °C along the fastest diffusive axis). The gray shaded area highlights the range of acceptable timescales for sample BC8, bracketed by 100% olivine re-equilibration and 50% clinopyroxene re-equilibration. Diffusion coefficients for Fe-Mg interdiffusion in clinopyroxene are from Müller et al. (2013) along the [001] axis (fastest diffusive axis); those for Fe-Mg interdiffusion in olivine are from Dohmen & Chakraborty (2007) along the [001] (fastest diffusive axis) and [100]/[010] axes (slowest). Calculations were performed at ΔFMQ -0.5 and 100 MPa. We performed 1D numerical models using the concentration-dependent version of Fick's second law. The crystal-melt boundary was considered open to the melt (an infinite Fe-Mg reservoir), so rim compositions were held constant at Fo₅₅ for olivine and Mg#50 for clinopyroxene (representative of the most evolved compositions in sample BC8).

The clinopyroxene core composition was set at Mg#80 (the most primitive composition observed in BC8 and, because BC8 olivines are fully re-equilibrated, we used a conservative value of Fo₈₅ for olivine cores (the more primitive composition measured in all our samples)).

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<figure><img src="image_2.png" /><figcaption>Fig. 1.</figcaption></figure>

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<figure><img src="image_3.png" /><figcaption>Fig. 2.</figcaption></figure>
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<figure><img src="image_4.png" /></figure>

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<figure><img src="image_5.png" /><figcaption>1615 Fig. 4.</figcaption></figure>
<figure><img src="image_6.png" /><figcaption>Fig. 5.</figcaption></figure>
<figure><img src="image_7.png" /><figcaption>1620 Fig. 6.</figcaption></figure>

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<figure><img src="image_8.png" /><figcaption>1653 Fig. 8.</figcaption></figure>

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<figure><img src="image_9.png" /><figcaption>1672 Fig. 9.</figcaption></figure>

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<figure><img src="image_10.png" /><figcaption>1682 Fig. 10.</figcaption></figure>
<figure><img src="image_11.png" /><figcaption>1685 Fig. 11.</figcaption></figure>
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<figure><img src="image_12.png" /></figure>

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<figure><img src="image_14.png" /><figcaption>1712 Fig. 13.</figcaption></figure>

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<figure><img src="image_15.png" /></figure>
<figure><img src="image_16.png" /></figure>

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