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Geochimica et
Cosmochimica Acta

JOURNAL OF THE GEOCHEMICAL SOCIETY
AND THE METEOROLOGICAL SOCIETY

High CO₂ in the mantle source of ocean island basanites

Francesco Maria Lo Forte, Federica Schiavi, Estelle F. Rose-Koga, Silvio G. Rotolo, Maximilien Verdier-Paoletti, Alessandro Aiuppa, Vittorio Zanon

PII: S0016-7037(24)00039-5

DOI: https://doi.org/10.1016/j.gca.2024.01.016

Reference: GCA 13298

To appear in: *Geochimica et Cosmochimica Acta*

Received Date: 28 August 2023

Accepted Date: 17 January 2024

Please cite this article as: Maria Lo Forte, F., Schiavi, F., Rose-Koga, E.F., Rotolo, S.G., Verdier-Paoletti, M., Aiuppa, A., Zanon, V., High CO₂ in the mantle source of ocean island basanites, *Geochimica et Cosmochimica Acta* (2024), doi: https://doi.org/10.1016/j.gca.2024.01.016

This is a PDF file of an article that has undergone enhancements after acceptance, such as the addition of a cover page and metadata, and formatting for readability, but it is not yet the definitive version of record. This version will undergo additional copyediting, typesetting and review before it is published in its final form, but we are providing this version to give early visibility of the article. Please note that, during the production process, errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.
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High CO₂ in the mantle source of ocean island basanites

Francesco Maria Lo Forte a, Federica Schiavi b, Estelle F. Rose-Koga c, Silvio G. Rotolo a,d,
Maximilien Verdier-Paoletti c, Alessandro Aiuppa a, and Vittorio Zanon f

corresponding author: Francesco Maria Lo Forte

corresponding author mail address: francescomaria.loforte@unipa.it

**Abstract**

Some of the most CO₂-rich magmas on Earth are erupted by intraplate ocean island volcanoes. Here, we characterise olivine-hosted melt inclusions from recent (< 10 ky) basanitic tephra erupted by Fogo, the only active volcano of the Cape Verde Archipelago in the eastern Atlantic Ocean. We determine H₂O, S, Cl, F in glassy melt inclusions and recalculate the total (glass+shrinkage bubbles) CO₂ budget by three independent methodologies. We find that the Fogo parental basanite, entrapped as melt inclusion in forsterite-rich (Fo₈₀-₈₅) olivines, contains up to ~2.1 wt% CO₂, 3-47% of which is partitioned in the shrinkage bubbles. This CO₂ content is among the highest ever measured in melt inclusions in OIBs. In combination with ~2 wt.% H₂O content, our data constrain an entrapment pressure range for the most CO₂-rich melt inclusion of 648-1430 MPa, with a most

<footnote>a Dipartimento di Scienze della Terra e del Mare, Università degli Studi di Palermo, Palermo, 90123, Italy.</footnote>

<footnote>b Université Clermont Auvergne, CNRS, IRD, OPGC, Laboratoire Magmas et Volcans, F-63000 Clermont-Ferrand, France</footnote>

<footnote>c Institut des Sciences de la Terre d'Orléans (ISTO) UMR 7327, CNRS-Université d'Orléans-BRGM, 1A rue de la Férollerie, 45071 Orléans, Cedex 2, France.</footnote>

<footnote>d Istituto Nazionale di Geofisica e Vulcanologia, Sezione di Palermo, Via Ugo La Malfa 153,90146, Palermo, Italy.</footnote>

<footnote>e Muséum National d'Histoire Naturelle de Paris (MNHN), Institut de Minéralogie, de Physique des Matériaux, 61 rue Buffon CP52 75005 Paris, France</footnote>

<footnote>f Instituto de Investigação em Vulcanologia e Avaliação de Riscos, Universidade dos Açores, Rua Mãe de Deus, Ponta Delgada, 9501-801, Portugal.</footnote>
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conservative estimate at 773-1020 MPa. Our results therefore suggest the parental Fogo melt is stored in the lithospheric mantle at minimum depths of ~27 to ~36 km, and then injected into a vertically stacked magma ponding system. Overall, our results corroborate previous indications for a CO₂-rich nature of alkaline ocean island volcanism. We propose that the Fogo basanitic melt forms by low degrees of melting (F=0.06-0.07) of a carbon-enriched mantle source, containing up to 355-414 ppm C. If global OIB melts are dominantly as carbon-rich as our Fogo results suggest, then OIB volcanism may cumulatively outgas as high as ~16-21 Tg of carbon yearly, hence substantially contributing to the global deep carbon cycle.

**Keywords:** alkaline ocean islands, intraplate volcanism, Fogo volcano, melt inclusions, Cape Verde.

# 1. Introduction

Quantifying the initial amount of volatile elements (H₂O, CO₂, S, Cl, and F) transported by OIB magmas is essential for understanding the cycling of volatiles from Earth's deep interior (e.g., Hirschmann, 2006, 2018; Dasgupta and Hirschmann, 2010), and their role for mantle melting, and magma generation and extraction (e.g., Foley, 2011; Dasgupta, 2018). Although sustaining only ~10% of the annual magma output on Earth, OIB magmas can contain more than 5% volatiles, and therefore significantly contribute to the annual magmatic gas output (e.g., Cabral et al., 2014; Longpré et al., 2017; Boudoire et al., 2018; Taracsák et al., 2019; Burton et al., 2023), especially if compared to more voluminous - but less volatile rich - Mid Ocean Ridge Basalts (MORBs) (e.g., Hauri et al., 2019). Oceanic intra-plate magmas have some of the most CO₂-enriched magmas on Earth (Longpré et al., 2017; Taracsák et al., 2019), likely reflecting their derivation from low melting degree of a metasomatised (C-enriched) mantle source (Foley et al., 2009, 2011; Dasgupta et al., 2013; Foley and Fischer, 2017; Dasgupta, 2018; Aiuppa et al., 2021).

Glassy melt inclusions (MI), small droplets of silicate melt entrapped in crystals during their growth in magmas or during episodes of crystal fracturing, are - at least in principle - physically and chemically isolated systems shielded from the surrounding environment by the host crystal (e.g., Lowenstern 1995; Danyushevsky et al. 2002; Schiano 2003; Wallace 2005; Cannatelli et al. 2016; Wallace et al., 2021; Rose-Koga et al., 2021). As such, the entrapped melt should escape extensive volatile loss upon magma ascent/storage in the crust prior/during eruption. Upon quenching as glass, MI should allow recording the dissolved volatile contents at the P-T-X conditions of entrapment, and eventually the parental melt volatile contents if entrapment has occurred at/near source (and before volatile saturation is reached) (e.g., Métrich and Wallace, 2008; Edmonds and Wallace, 2017). However, the task is complicated by a range of processes that can potentially modify MI composition after it is entrapped in the crystal (e.g., Anderson 1974; Roedder 1979; Danyushevsky et al., 2002; Wallace et al., 2021), including post-entrapment crystallization at the MI-host crystal interface (e.g., Danyushevsky et al., 2000), diffusive exchanges of major, trace and volatile (H⁺) elements (e.g., Gaetani and Watson, 2000, 2002; Danyushevsky, 2001; Cottrell et al., 2002; Gaetani et al. 2012) and above all - shrinkage bubble formation (Anderson, 1974; Roedder, 1979; Maclennan, 2017). This latter process is especially relevant for volatile species, as it leads to partitioning of poorly soluble CO₂ from the melt inclusion glass into the shrinkage bubble. Shrinkage bubbles can contain up to >90% of the bulk MI CO₂ content (e.g., Mironov and Portnyagin, 2011; Hartley et al., 2014; Mironov et al., 2015; Moore et al., 2015; Esposito et al., 2016; Aster et al., 2016; Rasmussen et al., 2018, 2020), demonstrating the caveats of traditional approaches (Lowenstern, 2015) that focus on measuring volatiles in glass only. Still, accurately quantifying the shrinkage bubble CO₂ content is technically challenging (e.g., Naumov et al., 2003),
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and few examples are available of reconstructions of the total (bubble+glass) MI CO₂ budget in OIB (e.g., Taracsák et al., 2019; Tucker et al., 2019, Burton et al., 2023).

Some of the most alkali-rich OIB magmas are erupted in the Cape Verde Archipelago, in the central eastern Atlantic. This archipelago lies on the largest intraplate bathymetric anomaly in the Earth's oceans – the Cape Verde Rise - interpreted as the surface expression of a hotspot sustained by the ascent of a deeply-rooted mantle plume (Holm et al., 2006; French and Romanowicz, 2015). The presence of highly alkaline silica undersaturated rocks (nephelinites, melanephelinites, and melilitites), as well as carbonatites (Hoernle et al., 2002; Bonadiman et al., 2005), are evidence for a heterogeneous, oxidized (Moussallam et al., 2019), potentially C-enriched mantle source, as also indicated by the CO₂-rich signature of fumaroles (Aiuppa et al., 2020; Melián et al., 2021).

In this study we use primitive MI composition to confirm the carbon-rich signature of the Cape Verde magma source. We report on the major, trace, and volatile composition of olivine-hosted MI from basanitic tephra erupted at Fogo, the youngest active volcano of the Cape Verde Archipelago. Our goal is to determine the initial (parental melt) volatile contents, in the attempt to (i) test the carbon-enriched source hypothesis, (ii) refine our understanding of the architecture of the magma system, and (iii) more broadly, provide additional evidence for the carbon-enriched nature of alkali rich, OIB magmatism.

## 2. Volcanological Background

Fogo is the only active volcano in the Cape Verde Archipelago, in the central eastern Atlantic (Figure 1). Volcanic activity on the islands started ~26 Ma (Torres et al., 2010), and is dominated by silica-undersaturated alkaline melts (e.g., Gerlach et al., 1988), and rarer carbonatites (Hoernle et al., 2002; Mourão et al., 2012). Subaerial volcanic activity has been divided into four eruptive phases: (i) The Ribeira de Almada Group (~4.5 Ma, Foeken et al., 2009); (ii) The Monte Barro Group (of unknown age); (iii) The Monte Amarelo Group, whose activity preceded the lateral collapse of Monte Amarelo at circa 62 to 123 ka (Day et al., 1999; Foeken et al., 2009; Martínez-Moreno et al., 2018); and (iv) The Chã das Caldeira Group (Foeken et al., 2009) related to the recent growth of Pico do Fogo central volcano. The Pico do Fogo volcano erupted 28 times since the 1500s. Eruptions are sourced by a complex, vertically elongated magma plumbing system, with main magma storage/stagnation zones at ~24, ~13-19 and ~9-12 km depth (Hildner et al., 2011, 2012; Klügel et al., 2020; Lo Forte et al., 2023).
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<figure><img src="image_2.png" /><figcaption>**Figure 1.** Modified Satellite map of Fogo Island from Google Earth, showing sampling sites with their relative pictures. Insets show the location of the Cape Verde archipelago (top right), and Fogo Island (bottom right).</figcaption></figure>

**3. Materials and Methods**

**3.1. Sample preparation**

We report data for 18 MI hosted in 12 olivine phenocrysts from three different tephra samples representative of the last eruptive phase of Fogo volcano (see tables S1 and S2 of Supplementary Material S1, for further details). In detail, we analysed (i) 8 MI hosted in 5 olivines from the clinopyroxene- and olivine-rich tephra samples of São Jorge (SJ), of the early Holocene period of activity (~10 ka), and (ii) 10 MI hosted in 7 olivines from the recent basanitic tephra samples of the 1951 and 2014/15 eruptions. Samples were prepared according to the methodology described by Rose-Koga et al., (2021). Tephra samples were initially sieved to separate grain size populations in the 250–4000 µm range. For each sample, about 400 olivine phenocrysts were hand-picked from these grain size fractions. Olivine crystals were mounted on glass slides using Crystal Bond 509®, subsequently they were doubly polished with silicon carbide paper and alumina powders to attain a maximum thickness of 80–100 µm. Twenty phenocrysts, containing 26 MIs with size greater than 25 µm were selected for analysis.

**3.2. Analytical techniques**
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Electron microprobe analysis (EMPA) was used to determine the major element compositions of both host olivines and melt inclusions, as well as Cl, and F concentrations in glasses. Volatile element concentrations (H₂O, CO₂, S, and Cl) in glasses were measured by nano-SIMS. The CO₂ concentration in shrinkage bubbles in MI was determined by Raman micro-spectroscopy. Finally, Laser Ablation-ICPMS was utilized for trace element analysis. The full compositional dataset is listed in tables S3, S4, S5, S5.1, and S5.2 of Supplementary Material S1.

### 3.2.1. Electron microprobe analysis

Electron microprobe analyses were performed with a Cameca SxFiveTactis with a LaB6 source equipped with five wavelength-dispersive spectrometers at the Laboratoire Magmas et Volcans (LMV), in Clermont-Ferrand, France. Major elements and volatiles were analysed in the same analytical session, with an accelerating voltage of 15 kV, an 8 nA beam current, and a 20 µm defocused beam for major elements. Chlorine, sulphur, and fluorine analyses were conducted using an accelerating voltage of 15 kV, a 40 nA beam current and a 20 µm defocused beam, following the acquisition program outlined in the Cameca Peak Sight software (see Rose-Koga et al., 2020 for more details). The olivines were analysed with an accelerating voltage of 15 kV and 15 nA beam current. The total concentration of volatiles was determined by summing the counts from multiple iterations. Total counting times were 100 s for F and Cl peaks. The detection limits for F and Cl were 300 and 150 ppm, respectively. KE12 LMV, CFA47 (Spilliaert et al., 2006), and ALV-98I R23 (Hekin and Walker, 1987) glasses and the San Carlos olivine were used as standards. Typical errors on the samples (1σ error of mean) are less than 10% for all elements analysed both in MI and in olivines.

### 3.2.2. Nano-SIMS analysis

High spatial resolution measurements of volatiles in minor and trace amounts in MI were conducted using the CAMECA nanoSIMS50 of the Muséum National d'Histoire Naturelle de Paris. All measurements were carried using a focussed Cs⁺ primary ion beam accelerated to 16 keV in high vacuum conditions (< 8x10⁻¹⁰ mbar). Prior to acquisition, a 5 × 5 µm² region was pre-sputtered with a 280 pA primary beam for 2 minutes to remove the Au coat and ensure a clean surface. The current was then reduced to 20 pA and the inner 3 × 3 µm² were rastered over a 32 × 32 pixels matrix with a 1 ms/px dwell time during 100 cycles (~2 min). A numerical blanking was applied to 32 × 32 px matrix to consider only the innermost 1 × 1 µm² region, this in order to limit the contribution of potential contamination. A cloud of electron was generated above the sample's surface to avoid the accumulation of positive charges induced by the implantation of the Cs⁺ beam. ¹²C, ¹⁶O¹H, ²⁸Si, ³²S, and ³⁵Cl were acquired simultaneously in multicollection mode on 5 electron multipliers (Ems). Each measurement was corrected for the 44 ns of the detector's deadtime. We used a 20 × 140 µm entrance slit (#3) with a 200 µm diameter aperture slit (#2) and a 10% energy slit to optimize the Mass-Resolving Power (MRP) without compromising our transmission. This setting enabled a MRP > 8000 which ensured the separation of ¹⁶O¹H and its isobaric interference ¹⁷O. The standards used for the calibration of the different elements were: two glasses NIST 610 and 612; a Suprasil crystal; two basaltic glasses MRN-G1 and IND-G2 (Shimizu et al., 2017) and three internal standards STR 10-11-13 (Créon et al., 2018). All of this enables to achieve a statistically robust linear regression spanning a range of concentrations from wt% (H₂O) to hundreds of ppm (CO₂, S, Cl) (Figure S1 in the Supplementary Material S2). The nano-SIMS technique is typically associated with higher analytical precision (< 5%) and lower detection limits (< 10 ppm typically) for Cl (Rose-Koga et al., 2021) relative to electron probe (Rose-Koga et al., 2020). The Suprasil standards is a synthetic crystal known for its purity. It has been extensively used alongside Herasil crystal to assess the blank values for H₂O and CO₂ (e.g., Shimizu et al., 2021). By setting their concentrations
* Sciencesconf
at 0 in our calibrations we systematically subtract these blank contributions. Errors were estimated
as the quadratic sum of the measured 1σ Poisson ratio error and the 1σ prediction/confidence
interval local estimation and were subsequently reported as 2σ errors. For each measurement
corrected by our determined linear regression we can extract an error associated with the estimation
through the 95% confidence interval. This error is then quadratically added to the error on the
measurement itself (poissonian error on ratios).

**3.2.3. Raman Spectroscopy**

Raman spectra were collected using an inVia confocal Raman micro-spectrometer manufactured by Renishaw and equipped with a 532.1 ± 0.3 nm diode-pulsed solid state laser (200 mW output power), a Peltier-cooled CCD detector of 1040 × 256 pixels, and a Leica DM 2500 M optical microscope with a motorised XYZ stage, at the LMV. Scattered light was collected via a back-scattered geometry. Laser power was about 10 mW on the sample surface. A 100x microscope objective, a grating of 2400 grooves mm⁻¹, and a 20-µm slit aperture (high confocality setting) were used, which resulted in spectral resolution better than 0.4 cm⁻¹ and in vertical and horizontal spatial resolutions of approximately 2-3 and 1 µm, respectively, near the surface of the sample. Calibration of peak positions was performed at room temperature, based on the 520.5 cm⁻¹ peak of silicon and the 1331.5 cm⁻¹ peak of a diamond standard; two Neon emission bands of 568.98163 and 576.44188 nm wavelengths were used to check the linearity of the spectrometer. For CO₂ measurements in MI bubbles, the spectra were recorded from 725 to 1880 cm⁻¹ using the WiRE™ 4.4 software. Then, spectra were acquired in the 60-1320 cm⁻¹ wavenumber range to detect presence of mineral phases (e.g., carbonates, sulphates) or fluid SO₂. Finally, spectra were collected in the 2800-3900 cm⁻¹ region to detect H₂O and/or OH⁻ molecules. Acquisition time for a single CO₂ analysis was fixed at 60 s. Spectra were initially treated for removal of cosmic rays, and then an appropriate polynomial baseline correction was applied before fitting a combined Gaussian-Lorentzian curve to the CO₂ Fermi diad (e.g., Moore et al., 2015; Mironov et al., 2020; Schiavi et al., 2020). The analytical uncertainty associated with CO₂ density measurements is < 0.05 g/cm³ (Boudoire et al., 2023). Examples of Raman spectra acquired on shrinkage bubbles in melt inclusion are shown in Figure 2. In all shrinkage bubbles, only CO₂ was detected as fluid phase (Figure 2a).

**3.2.4. Laser Ablation-ICPMS**

Trace-element measurements on melt inclusions were carried out using a laser ablation system (193 nm Excimer Resonetics M-50E) associated with an inductively coupled plasma mass spectrometer (Agilent 7500 cs LA-ICP-MS at the LMV). Analysis was performed following procedures outlined in the previous studies (e.g., Rose-Koga et al., 2012). We used a pulse energy of about 3 mJ, a spot diameter between 15 and 33 µm, and a laser pulse frequency of 2–3 Hz, depending in the inclusion size to keep a fluence at sample surface of about 4 J/cm⁻². The background was measured for 30–40 s before ablation and analysis time was approximately 100 s. Data reduction was performed using the GLITTER software (Van Achterbergh et al., 2001). This technique uses CaO (measured by electron probe) as an internal standard. Reproducibility and accuracy of the analysis were constrained by systematic analysis of two standards BCR2-G and SRM NIST 612 at the beginning, in the middle and at the end of the session. Typical errors on the samples (1σ error of mean) are less than 10% for all trace elements analysed.
<figure><img src="image_3.png" /><figcaption>**Figure 2.** (a) Raman spectrum of melt inclusion shrinkage bubble (sample Fog 43-8-1). The bubble is made of pure CO₂, as revealed by the two strong peaks at 1284 and 1387 cm⁻¹, identified as Fermi</figcaption></figure>
diad (Kawakami et al., 2003). The separation between the two Raman bands of CO₂ is a reliable densimeter for the CO₂ fluid. (b) Raman spectra collected from the shrinkage bubble of Fog 43-8-1 (grey spectrum) and Fog 43-6-2 (orange and light blue spectra, acquired during tomographic analysis) MI samples. In the bubble, minerals such as Fe-magnesite and nahcolite were identified, as evidenced by peaks at 318, 737, and 1092 cm⁻¹ for Fe-magnesite, (Mg,Fe)CO₃ (Boulard et al., 2012) and the peaks at 685, 1045, and 1266 cm⁻¹ for nahcolite, NaHCO₃ (Fall et al., 2007). Both spectra also exhibit the characteristic olivine peaks at 822 e 854 cm⁻¹.

# 4. Results

## 4.1. Whole-rock and mineral chemistry

The studied tephra samples are nepheline-normative (Ne= 17.6–21.2 wt.%) basanites (SiO₂ = 41.73–43.78 wt.%; K₂O + Na₂O = 6.69-7.68 wt.%) (table 1 of Supplementary Data S1), with a common mineral assemblage represented by clinopyroxene, olivine, plagioclase, and Fe-Ti oxides. These tephra samples are part of the same suite analysed (for major elements and fluid inclusions) by Lo Forte et al. (2023). The olivine grains have euhedral shapes and exhibit a narrow range of compositions (forsterite mol% Fo₈₀.₈₅) (table S3 of Supplementary Material S1). The olivine phenocrysts of sample SJ (Fog 43) are slightly more forsteritic (Fo₈₃.₈₅) than the 1951 (Fog 8) and 2014/15 (Fog 11) olivines (Fo₈₀.₈₄). The same SJ olivines exhibit higher CaO (0.34–0.40 wt%) and NiO (0.13–0.15 wt%) contents than the 1951 and 2014/2015 olivines (CaO, 0.31-0.39 wt.%; NiO, 0.03-0.15 wt.%). Mineral chemistry results are detailed in table S3 of Supplementary Material S1.

## 4.2. Melt inclusions: size and morphology.

Olivines contain abundant MI, but only 13 % of them were large enough (> 25 µm) for trace element analyses. All studied inclusions are perfectly glassy, and elliptical in shape (Figure 3), with the long axis (a) ranging between 17.0 µm and 94.0 µm (mean, 49.6 µm) and the short axis (b) between 11.3 µm and 83.1 µm (mean, 40.7 µm). Different criteria are used in the literature to determine the size of the third semi-axis c (e.g., axis a > b =c, Hauri et al., 2018; Shaw et al., 2008, 2010; Wanless et al., 2014, 2015; geometric mean, Ni et al., 2017); direct measurements can be used to test different assumptions (Mironov et al., 2020). Here we use the criteria of Moore et al. (2015) that assumes axis (c) equals to the arithmetic mean of axes (a) and (b) of the ellipsoid. The semi-axis (c) is therefore estimated at between 14.2 µm and 88.5 µm (mean, 37.8 µm). Shrinkage bubbles were observed inside all the MI (table S7 of Supplementary Material S1). However, in 31 % of the selected MI the bubbles were lost during the polishing process as the MI surface was exposed for analysis. The MI in which bubbles were lost are not considered further (except samples Fog 43-3-1 for which we report glass CO₂ contents only). There is no textural evidence of MI decrepitation. Roughly half of the shrinkage bubbles contain mineral phases such as Fe-magnesite, nahcolite, and sulphates (Figure 2b).
<figure><img src="image_4.png" /><figcaption>**Figure 3.** Measurements of the major (a) and minor (b) axes of the MI (Fog 43-6-3) for the determination of the melt volume and measurement of the diameter (d) of the shrinkage bubble for the calculation of its volume.</figcaption></figure>

**4.3. Melt inclusion major element compositions and post-entrapment correction (PEC)**

Melt inclusion glasses are foiditic to basanitic and compositionally overlap with bulk rocks (Fig. 4a). They typically contain 39.89-45.95 wt.% SiO₂, 3.15-5.85 wt.% MgO, 3.13-5.32 wt.% Na₂O, and 8.07-10.74 wt.% FeOt (Fig. 4). A variety of processes may alter the melt composition after entrapment (see review in Wallace et al., 2021), including: (i) post-entrapment crystallization (PEC) of olivine at the MI-host mineral interface during slow cooling (Kent et al., 2008); (ii) volatile loss from the melt to the shrinkage bubbles (Wallace et al., 2015) or due to inclusion decrepitation (e.g., MacClennan, 2017); (iii) hydrogen ('water') diffusive loss (Hauri, 2002; Massare et al. 2002; Portnyagin et al. 2008, 2019; Chen et al. 2011; Gaetani et al. 2012); (iv) iron, magnesium diffusive exchange between the inclusion and its host mineral (e.g., Danyushevsky et al., 2000; Cottrell et al., 2002; Gaetani and Watson, 2000, 2002). We correct the composition of olivine-hosted MI for PEC using Melt inclusion modification corrections (MIMiC) (Rasmussen et al., 2020). The code employs the (KD Fe-Mg) model of Toplis (2005) and the olivine-melt thermometer of Putirka et al. (2007) to calculate the Fe-Mg olivine-melt partitioning (at redox conditions estimated by Moussallam et al., 2019) and to estimate magma temperature, respectively. Melt inclusion compositions, both pre- and post- MIMiC correction, are provided in tables S5 and S6 of Supplementary Material S1. PEC correction requires the addition of 0.03 to 9.28 % olivine to restore melt inclusion compositions to equilibrium with their hosts. The average PEC correction is 3.85 % for São Jorge (SJ) MI and 2.17 % for the two recent eruptions. Following PEC correction, melt inclusion compositions range 39.87-45.74 wt.% SiO₂, 4.00-7.36 wt.% MgO, 3.02-5.20 wt.% Na₂O, and 8.36-11.57 wt.% FeOt. On the total alkali-silica diagram (Figure 4a), the PEC-corrected MI span the foidite to basanite compositional fields, with the most SiO₂-rich MI (41.74-45.74 wt.%) overlapping the range of the whole-rocks (compositions given in Lo Forte et al., 2023) and of bulk rocks from the literature (Hildner et al., 2011, 2012; Mata et al., 2017). These MI have MgO contents varying from 4.00 to 5.76 wt.%, TiO₂ from 2.44 to 3.06 wt.%, and P₂O₅ from 0.84 to 1.36 wt.%. The CaO/Al₂O₃ ratios vary between 0.88 and 0.97 for SJ MI, and between 0.44 and 0.84 for the 1951 and 2014/15 MI (Figure 4b).
We model fractional crystallisation (FC) (Figures 4b-d) to determine the extent of crystallisation at the time of entrapment, and to test if the melt follows the liquid line of descent (LLD) prior to be entrapped, or if instead points to some Fe-Mg diffusive exchange during crystallisation (Figure 4b, c). We start our simulations with a hypothetic parental melt composition corresponding to the primitive lava Fogo 16 from Hildner et al. (2012) (MgO = 9.79 wt.%; FeO = 12.03 wt.%). We assume an initial H₂O content of 2.0 wt.% (see below) and an oxygen fugacity of 1.5 log units above the Fayalite-Magnetite-Quartz buffer (ΔFMQ +1.5; Moussallam et al., 2019). Starting from this parental melt composition and an initial temperature of 1150 °C, we simulate the cooling-crystallisation trajectory using Petrolog3 (Danyushevsky and Plechov, 2011), using the melt-mineral equilibrium models of Danyushevsky, (2001) for olivine, clinopyroxene, and plagioclase, and Ariskin and Barmina, (1999) for magnetite. Results (Fig. 4b,c,d) show that the modelled LLD accurately reproduces the whole-rock compositional trend. After PEC correction, MI also match more closely the LLD and the bulk rocks trend (Figure 4b,c,d), implying diffusive Fe loss was minor (if any). No correction was therefore applied. All raw and corrected MI compositions are listed in Tables S5 and S6 of Supplementary Material S1.
<figure><img src="image_5.png" /></figure>

**Figure 4** Total alkali-silica diagram (a) and major element compositions of melt inclusions (b, c, d) together with whole-rock compositions from Hildner et al. (2011, 2012), Lo Forte et al. (2023) and Mata et al. (2017). The solid lines are liquid lines of descent (LLD), calculated from a primitive parent melt (SiO₂ 42.14 wt.%, MgO 9.79 wt.%, FeO 12.03 wt.%, Al₂O₃ 12.41 wt.%, CaO 12.96 wt.%) using Petrolog3 (Danyushevsky and Plechov, 2011). The LLD are calculated assuming initial H₂O content of 2 wt.% and oxygen fugacity of ΔFMQ +1.5 (Moussallam et al., 2019). We used the melt-mineral equilibrium models of Danyushevsky (2001) for olivine, clinopyroxene and plagioclase, and of Ariskin and Barmina (1999) for magnetite. The orange LLD was calculated for FC > 55%, to fit the more evolved MI (d). The grey area represents the range of forsterite component (Fo₈₀₋₈₅) in the analysed olivines hosting MI (a).
## 4.4. Trace element composition

Table S5.2 of Supplementary Material S1 reports the concentrations of a selected suite of trace elements in the studied MI. The primitive mantle normalised trace element patterns (Fig. 5) (this study) match those of the bulk-rocks (Hildner et al., 2011, 2012; Mata et al., 2017). No negative Hf anomaly is observed, and the Zr/Hf ratios (averaging at 39) are comparable to the typical primitive mantle value (~36; Palme and O'Neil, 2003). The LILE (Large Ion Lithophile Elements)-enriched composition of Fogo melts is additionally corroborated by elevated Ba/La (15.0-20.3), Rb/Sr (0.05-0.07) and Rb/La (0.98-1.54) ratios (Weaver, 1991).

<figure><img src="image_6.png" /></figure>

**Figure 5.** Primitive mantle-normalised (Sun and McDonough, 1989) trace element patterns in MI samples (orange patterns; average values from all samples are used). The trace element patterns are intermediate between the compositional ranges of enriched mantle sources (EM-I and EM-II; data from Willbold and Stracke, 2006) and the HIMU mantle source (Willbold and Stracke, 2006). Note the significant Nb and Ta enrichment relative to the light REE. The red line identifies the compositions of Mid-Oceanic Ridge Basalts (MORB; Sun and McDonough, 1989).

## 4.5. Volatiles

All the melt inclusion glasses are volatile-rich and contain 1.97 to 2.79 wt.% H₂O and 9002 to 16044 ppm CO₂ in SJ MI (Fig. 6a). MI glasses from recent (1951-2014/15) eruptions are slightly less CO₂-rich than the SJ MI glasses and contain 1.20 to 2.92 wt.% H₂O and 1239 to 9102 ppm CO₂ (Fig. 6; Table S5.1 Supplementary Material S1). CO₂ in MI glasses broadly decreases with increasing degree of melt evolution, e.g., with decreasing CaO/Al₂O₃ (Fig. 6b) and MgO (Fig. 6d) and increasing FeO/t/MgO (Fig. 6c). The SJ MI glasses exhibit higher S concentrations (2676-7794 ppm) than the 1951 and 2014/15 MI (1146 to 5509 ppm) (values measured by nano-SIMS, Tables 5 and S5.1 Supplementary Material 1). Chlorine concentrations, measured by nano-SIMS, range from 947 to 1358 ppm in SJ MI and from 956 to 1415 ppm for the recent eruptions. The measured
fluorine concentrations are 1285-2216 ppm in SJ MI glasses and 1617-2249 ppm in the 1951-2014/15 MI glasses. For each element, a general least-squares linear regression was calculated, along with both 95% confidence and prediction intervals. This process allowed us to extract an error associated with the estimation through the 95% confidence interval for each measurement corrected by the determined linear regression. Subsequently, this error was quadratically added to the measurement itself. The absolute and relative errors were most pronounced for very low concentrations outside the range covered by our set of standards. As we extrapolated to higher concentrations, absolute errors increased. However, for measured concentrations (on average, >2000 ppm), the relative error rose less drastically compared to lower concentrations. For further details, please refer to Table S9 in Supplementary Material S1 and Figure S2 in Supplementary Material S2. The PEC-corrected concentrations vary as follows: 1.9-2.7 wt.% (SJ MI) and 1.2-2.8 wt.% (1951 and 2014/15 eruptions) for H₂O (Fig. 6a); 2621-7718 ppm (SJ) and 1120-5250 ppm (1951-2014/15) for S; 917-1328 ppm (SJ) and 934-1415 ppm (1951-2014/15) for Cl (Fig. 7). As typical of alkaline magmas in general (Webster et al., 2018), the Fogo MI glasses are especially Cl- and F-enriched (up to 1415 and to 2249 ppm) relative to other OIB localities (Fig. 7), being only matched by the El Hierro basanite MI (Taracsák et al., 2019). The Fogo MI have F/Cl ratios in the 1.2 to 2.4 range, i.e., higher than F/Cl values of most arc magmas, exception made for F/Cl reaching 2.5 in the Trans Mexican arc (Vigouroux et al., 2008) and in the Central American arc (Sadofsky et al., 2008; Fig. 7). This F-rich signature is a distinctive characteristic of OIB magmas in general (Kendrick et al., 2014).

<figure><img src="image_7.png" /><figcaption>**Figure 6.** CO₂ concentrations vs. H₂O (a), CaO/Al₂O₃ (b) and FeOt/MgO (c) ratios, and MgO (d). The SJ MI exhibit higher H₂O and CO₂ contents than those from the more recent eruptions in 1951 and 2014-15. The orange symbols represent data related to the total CO₂ content in the MI (glass + bubble), corresponding to measurements obtained through Raman-based total CO₂ estimates (see section 5.1 for further details).</figcaption></figure>
<figure><img src="image_8.png" /></figure>

**Figure 7.** In the F vs. Cl plot, our data are compared to MI of different OIB localities and from volcanoes of different geodynamic contexts, such as Ascension (Webster and Rebbert, 2001), Etna (Métrich et al., 1993, 2004), La Soufriere (Heath et al., 1998), St. Helens (Blundy and Cashman, 2005), Parícutín (Rowe et al., 2011), Trans Mexican belt (Vigouroux et al., 2008), Central American belt (Sadofsky et al., 2008), Vanuatu (Métrich et al., 2011), Eifel (Harms and Schmincke, 2000), Aeolian islands (Di Martino et al., 2011), Masaya (Pérez et al., 2020), and Kamchatka volcanoes (Portnyagin et al., 2011). The F/Cl ratios of our MI (1.16-2.41) fall in the field of OIB volcanoes and partially overlap the composition of El Hierro MI.
# 5. Discussion

## 5.1. Reconstructing the total (glass+bubble) CO₂ content in MI

Our results show that the Fogo MI glasses contain up to ~1.6 wt.% CO₂ (Fig. 6), plus ~2 wt.% H₂O and ~0.48 wt.% S, on average. However, all MI also contain CO₂-rich bubbles (Fig. 2). There is a positive relationship between total MI volumes and bubbles volumes (Fig. 8a), indicating that all MI contain similarly low bubble volumetric fractions (0.2-8.6 %). We conclude that the contemporaneous entrapment of melt and fluid inclusions is unlikely (Rasmussen et al., 2020), and therefore all fluid bubbles found are shrinkage bubbles. Hence, as recently demonstrated elsewhere (Taracsák et al., 2019), reconstructing the total MI CO₂ budget requires a procedure in which the CO₂ masses contained in (i) shrinkage bubbles and (ii) glass are both fully taken into account.

The CO₂ mass contained in the glass is initially calculated by multiplying glass mass in the MI (glass density × glass volume) by CO₂ concentration in the glass (cfr. 4.5). Glass density is derived using MIMiC (Rasmussen et al., 2020), which uses the relationships of Lesher and Spera (2015), while glass volume is estimated following the procedure of Moore et al. (2015).

Next, we quantify shrinkage bubble CO₂ using three different, independent methodologies: (i) the Raman method (Esposito et al. 2011; Hartley et al. 2014; Moore et al. 2015; Rasmussen et al. 2018); (ii) the Simulated-volume method (Rasmussen et al., 2020); and (iii) the Observed-volume method (e.g., Shaw et al., 2010; Wanless et al., 2014; Hauri et al., 2018; Tucker et al., 2019). Our aim is to use the level of agreement/disagreement of these different methods as a proxy for the level of accuracy/inaccuracy of such bubble CO₂ estimates (Fig. 8).

In the Raman method (Hartley et al., 2014; Moore et al., 2015; Moore and Bodnar, 2019; Rasmussen et al., 2020), the bubble CO₂ mass is derived as the product of bubble CO₂ density × bubble volume. Raman indicates a pure-CO₂ gas phase composition for the shrinkage bubble (Figure 2 a); we hence apply the densimeter by Wang et al. (2011) for the calculation of the bubble CO₂ density.

The other two CO₂ estimates (Simulated and Observed) are derived using the MiMiC code (Rasmussen et al., 2020). The Simulated-volume method relies on estimating bubble volumes and CO₂ contents from modelling the growth of bubbles during the cooling and quenching of the melt inclusions (rather than directly measuring bubble volumes). This bubble growth model assumes that exchange of CO₂ between melt and bubble terminates at a CO₂ closure temperature ($T_c$) for diffusion (generally higher than the glass-transition temperature) followed by further bubble expansion with no CO₂ loss to bubble (Rasmussen et al., 2020).

Finally, the Observed-volume method (e.g., Shaw et al. 2010; Wanless et al. 2014; Hauri et al. 2018; Tucker et al. 2019) uses directly measured bubble volumes and entails determining the CO₂ mass within the vapor bubble by applying the modified Redlich-Kwong (MRK) equation of state. The method relies on the assumption that vapor-melt equilibrium is maintained throughout the entire bubble quenching/growth history. The temperature of vapor-melt equilibrium is determined based on the calculated glass-transition temperature, which however can vary significantly by several hundred degrees Celsius, depending on factors such as melt viscosity and cooling rate (Giordano et al., 2005).

All three CO₂ restoration methods have their caveats and limitations which can potentially lead to overestimation/underestimation of the total MI CO₂. Use of the Raman restoration method (e.g., Esposito et al., 2011; Moore et al., 2015), in particular, is challenged by difficulties in (i) correctly measuring melt inclusion geometry, and hence glass vs. bubble volumes (Moore et al., 2015) and
(ii) quantifying the relative volumetric fractions occupied by minerals and fluid in the bubble (Schiavi et al., 2020). Determination of melt inclusion geometry (glass and bubble volumes) in principle requires accurate 3D scanning microtomographic analysis, which is rarely available. Our glass/bubble volumes (Table S7 of Supplementary Material S1) are estimated from 2D analysis, and on the assumption that the length of the third semi-axis c (normal to the polished plane) equals to the arithmetic mean of axes a and b (Moore et al., 2015); this assumption should, at least in principle, introduce a relatively minor (5%) error (Wallace et al., 2015). More problematic is to account for the carbon mass sequestered in mineral phases inside the bubble. Recent studies have found that mineral phases such as carbonates, sulphides, sulphates and native sulphur can precipitate inside bubbles during cooling (e.g., Kamenetsky et al., 2002; Hartley et al., 2014; Moore et al., 2015; Esposito et al., 2016; Robidoux et al., 2018; Guzmics et al., 2019; Tucker et al., 2019; Schiavi et al., 2020; Venugopal et al., 2020; Buso et al., 2022; Feignon et al., 2022). Precipitation of these minerals can potentially sequester large fractions of the bubble-carried CO2, hence ignoring their presence can result in significant (up to 40%) underestimation of the total MI CO2 (Schiavi et al., 2020). We detected the presence of carbonate phases, such as Fe-magnesite and nahcolite (Supplementary Material S2 Figure S2), and sulphates, in roughly 50 % of the Raman spectra. The proportions of the different phases are highly variable. Raman microtomography revealed significant volumes of carbonates crystallised on the walls of the MI Fog 43-6-2 bubble, amounting to ~35 vol% according to an approximate average estimate based on few 2D sections extracted from the 3D map (Supplementary Material S2 Figure S2). This corresponds to ~25% of the CO2 sequestered in the bubble (compared with 9% reported in Table S7 of the Supplementary Material S1). We conclude that the total CO2 content can be underestimated (by about 15-20 % for MI Fog 43-6-2) if carbon stored in carbonates inside the bubble is neglected. However, considering the overall positive correlation observed between glass MI CO2 and shrinkage bubble density (Fig. 8b), we conclude that precipitation of carbonates is likely to be marginal in all samples except Fog 43-8-1 and 43-6-2 (that coherently exhibit the most intense Raman peaks associated with carbonates).

The second CO₂ restoration methodology used (Simulated-volume method) overcomes some of the limitations above, as it does not rely on measured bubble volumes, but rather on estimated bubble volumes based on modelling bubble growth during MI cooling and quenching. The detailed sensitivity study of Rasmussen et al. (2020) demonstrates that the amount of cooling experienced by MI, from their entrapment temperature (T_c) to their CO₂ closure temperature (T_c), is the key parameter controlling the extent of CO₂ loss to bubbles. According to Rasmussen et al. (2020), this cooling extent (T_c - T_e) depends on MI dimension (measured texturally) and cooling rate, for which we unfortunately have no independent estimate. In such conditions, MIMiC adopts a calculated cooling rate derived by solving the Dodson equation for the magnesium closure temperature (T_c(MgO)); this procedure can potentially introduce a maximum error of 50% on the restored CO₂ concentrations (when the used T_c(MgO) is in excess of by one order of magnitude).

Finally, the Observed-volume method calculates the CO₂ mass in the bubble by applying an equation of state and performing a mass balance. In this approach, the observed bubble volume, measured under the microscope, used in the calculation assumes that melt-bubble equilibrium is maintained throughout the entire bubble growth history. However, if equilibrium is not maintained, the calculation based on the observed volume may not be accurate. Additionally, this method assumes the glass transition temperature (T_g) as the final melt-bubble equilibrium temperature, but precisely assessing T_g is complicated by its strong dependence on factors such as cooling rate, melt composition, melt viscosity, and volatile content (Rasmussen et al., 2020).

Ultimately, since all restoration methods have limits and caveats, a multi-method approach is advisable (Rasmussen et al., 2020). The outputs of the three restoration methods used here are compared in Figures 8c and 8d. We find that the Raman spectroscopy measurements and simulated-volume approach yield CO₂ concentrations that agree within 9.7 % on average (range: 0.4-22.6 %).
The agreement between these two independent methodologies provides confidence in our restored total MI CO₂ values. The total MI CO₂ contents determined with the observed-volume method also positively correlate with those inferred from Raman analysis (Fig. 8d), but typically plot above the 1:1 line, indicating that the observed-volume approach may systematically over-estimate the MI CO₂ budget, as already suggested (Moore et al., 2018; Rasmussen et al., 2020). Rasmussen et al. (2020) also found that the concentrations reconstructed from both Raman measurements and the simulated-volume method are quantitatively consistent with those obtained by re-homogenising MI at high temperature and pressure – this latter procedure being considered as that yielding the most accurate CO₂ reconstructions. Ultimately, while we are well aware of the many existing potential sources of uncertainty, we conclude that our calculations bring nonetheless robust evidences that (i) in some cases of OIB MI, a large fraction (up to 50 %) of the MI CO₂ is contained in the bubbles (Mironov et al., 2015), and (ii) the total (glass+bubble) CO₂ content in the Fogo parental melt is certainly > 1 wt. %, and potentially as high as ~2 wt. % (the highest value obtained by the Raman and Observed-volume methods). For the sake of simplicity, in the discussion below we refer to the Raman-based bubble (and hence total) CO₂ estimates, but we note that these overlap with those derived from the bubble growth model using MiMIC by a factor of <10%.

<figure><img src="image_9.png" /><figcaption>**Figure 8.** (a) Plot of bubble volumes as a function of MI volumes; isovolumetric lines corresponding to bubble/MI volume ratios of 0.2 %, 1.0 %, 2.0 %, 5.0 %, and 8.0 % are also shown. (b) The plot shows a positive correlation between the CO₂ calculated in the glass and the density of the shrinkage bubble, except for the two samples (Fog 43-8-1 and 43-6-2) which exhibit intense Raman peaks associated with carbonates, specifically Fe-magnesite and nahcolite crystals on the bubble wall (Fig. 2). As shown in (b), the overlooked presence of carbonates in the bubble leads to an underestimation of the total CO₂ content. Excluding these two outliers, the samples show a good</figcaption></figure>
correlation (R² = 0.802). Plot of the total CO₂ concentration in MI, calculated using the Raman and Simulated methods (c) and the Raman and Observed methods (d). In plot (c), there is a strong correlation (R²=0.891), with CO₂ concentrations from Raman and Simulated-volume methods showing a 1:1.1 relationship. In contrast, in plot (d), the total CO₂ calculated using the Raman and Observed-volume methods is not strongly correlated (R²=0.806), with a ratio of 1:1.6. The dashed lines correspond to a ratio 1:1.

## 5.2. A CO₂-rich parental melt

Finally, glass CO₂ mass and shrinkage bubble CO₂ mass are summed to obtain total CO₂ mass in MI, which is ultimately converted in total CO₂ concentration in the MI. All calculations, for each MI analysed, are given in Table 1 and in Table S7 of Supplementary Material S1. Using the Raman method results described above, we find that bubbles contain 281-5546 ppm CO₂ (for the SJ MI) and 857-2920 ppm CO₂ (for the 1951 and 2014/15 MI). The total CO₂ contents thus range from 1309 to 21502 ppm for SJ MI, and from 1084 to 9102 ppm for the 1951 and 2014/15 eruptions (Fig. 6). We thus estimate that 3 to 47 % of the total MI CO₂ content is contained in the bubbles.

<table><caption>**Table 1.** Summary of the total CO₂ content (bubble + glass) calculated in the analysed melt inclusions using three different methods (Raman, Simulated, and Observed).</caption><thead><tr><th>Sample</th><th>CO₂ wt% in MI<br>(Raman Method)</th><th>CO₂ wt% in MI<br>(Simulated Method)</th><th>CO₂ wt% in MI<br>(Observed Method)</th></tr></thead><tbody><tr><td>Fog 43-3-1</td><td>1.25</td><td>1.43</td><td>/</td></tr><tr><td>43-5-1</td><td>2.15</td><td>1.95</td><td>2.75</td></tr><tr><td>43-6-1</td><td>1.49</td><td>1.48</td><td>2.31</td></tr><tr><td>43-6-2</td><td>1.23</td><td>1.42</td><td>2.15</td></tr><tr><td>43-6-3</td><td>1.16</td><td>1.18</td><td>1.74</td></tr><tr><td>43-6-4</td><td>/</td><td>/</td><td>/</td></tr><tr><td>43-7-1</td><td>0.13</td><td>/</td><td>/</td></tr><tr><td>43-8-1</td><td>1.01</td><td>1.40</td><td>/</td></tr><tr><td>Fog 8-1-1</td><td>0.40</td><td>0.54</td><td>0.77</td></tr></tbody></table>
<table><tr><td>8-4-1</td><td>0.89</td><td>1.03</td><td>1.41</td></tr><tr><td>8-7-1</td><td>0.74</td><td>0.92</td><td>0.79</td></tr><tr><td>8-8-1</td><td>0.91</td><td>1.16</td><td>1.69</td></tr><tr><td>8-11-2</td><td>0.62</td><td>0.47</td><td>0.97</td></tr><tr><td>8-11-3</td><td>/</td><td>/</td><td>/</td></tr><tr><td>8-11-4</td><td>0.11</td><td>/</td><td>/</td></tr><tr><td>8-11-5</td><td>/</td><td>/</td><td>/</td></tr><tr><td>Fog 11-2-1</td><td>0.21</td><td>0.22</td><td>0.46</td></tr><tr><td>11-3-1</td><td>0.82</td><td>0.99</td><td>1.08</td></tr></table>

Our results show that the basanitic melts from Fogo are exceptionally CO₂-rich. Their volatile-rich nature is additionally demonstrated by high halogen contents (Fig. 7). The estimated CO₂ content of the parental melt (2.15 wt.%) is up to a factor 10 higher than for other OIB localities (e.g., Erebus in Antarctica, Moussallam et al., 2014; Fernandina and Santiago in the Galapagos, Koleszar et al., 2009; Piton de la Fournaise in Reunion Island, Di Muro et al., 2014; Kilauea in the Hawaii, Sides et al., 2014; Pico in the Azores, Métrich et al., 2014) and Iceland (e.g., Bardabunga, Bali et al., 2018, and Laki, Hartley, 2014) (Fig. 9a). In addition to confirming that bubbles do matter if the CO₂ MI budget is to be reconstructed (Moore et al., 2015; Rasmussen et al., 2020; Buso et al., 2022), our results also corroborate recent work that shows that alkali rich OIB magmas can be especially carbon-rich, as observed at El Hierro (CO₂ up to 4 wt.%; Longpré et al., 2017; Moussallam et al., 2019; Taracsák et al., 2019) and La Palma (3-5 wt.%; e.g., Burton et al., 2023) in the Canary Archipelago, and at Piton de La Fournaise in La Reunion Island (3.5 ± 1.4 wt.%, Boudoire et al., 2018) (Fig. 9a) (note that in these studies shrinkage bubble CO₂ has been accounted for with either observations or modelling). The fact that alkali rich OIB volcanoes erupted magmas richer in CO₂ than alkali poor OIBs has been documented at a global scale (Sun and Dasgupta, 2023), although we see no evidence in our MI of the extremely high (9-10 wt.%) CO₂ contents inferred by these authors for the Cape Verde archipelago parental melts. Rather, the ~2 wt. % CO₂ content we estimate for the Fogo parental melt is at the lower range of the measured/inferred (Hudgins et al., 2015; Aiuppa et al., 2021) CO₂-enriched signatures of alkaline mafic melts from continental rift context (e.g., in the East African rift system; Foley et al., 2012). This link between high CO₂ compositions and melt enrichment in alkalis (and other incompatible elements; Aiuppa et al., 2021) is further explored below (Fig. 10).
<figure><img src="image_10.png" /><figcaption>**Figure 9.** (a) H₂O and CO₂ concentration in studied MI and in other OIB occurrences. Isobars were calculated using an average composition of our MI samples at temperature of 1150 °C and an oxygen fugacity of ΔFMQ +1.5 (Moussallam et al., 2019). Degassing paths, in a closed system, were calculated using the volatile saturation models by Moretti et al. (2003) and MagmaSat (Ghiorso and Gualda, 2015), with the same input melt composition and temperature and volatile contents of 21502 ppm and 2.0 wt.% for CO₂ and H₂O respectively. (b) In the CO₂ content vs. Pressure graph, different saturation models are compared. The solid lines represent the outputs of the different models, instead the dashed lines represent the extrapolations of these models.</figcaption></figure>

**5.3. Clues from CO₂ vs. trace-element correlations**

A potentially useful approach to test the accuracy of our reconstructed MI CO₂ content is to normalize them to the measured Nb and Ba contents in the same MI glasses (Fig. 10). These trace elements are thought to have similar bulk mineral/melt distribution coefficients to that of carbon (e.g., Rosenthal et al., 2015; Hirschmann, 2018). Hence, the CO₂/Nb and CO₂/Ba ratios should not be fractionated during mantle melting and fractional crystallisation of mafic melts. If a mantle-sourced melt has not degassed CO₂ during crustal emplacement, it is thus expected to preserve a record of its source mantle CO₂/(Nb, Ba) ratio (e.g., Saal et al., 2002; Michael and Graham, 2015; Shimizu et al., 2016; Le Voyer et al., 2017; Hauri et al., 2018; Tucker et al., 2019). Following these premises, the expected range of mantle CO₂/Nb (Fig. 10a) and CO₂/Ba (Fig. 10b) have been inferred (see grey filled diagonal bands in Figure 10) from analysis of MORB that, owing to their volatile-depleted nature, are thought not to have reached volatile saturation before eruption (undegassed MORB; Michael and Graham, 2015) or entrapped as MI (e.g., Saal et al., 2002; Le
Voyer et al., 2017; Moussallam et al., 2023). OIB magmas, in contrast, owing to their volatile enriched nature, do systematically reach saturation during ascent (Dixon and Stolper, 1995) and, even more significantly, form post-entrapment shrinkage bubbles that can sequester a large fraction of the MI CO₂ upon cooling and decompression (Hartley et al., 2014; Moore et al., 2015; Wallace et al., 2015; Tucker et al., 2019; Schiavi et al., 2020). As noted above, most MI results from the literature refer to glass measurements only, in which no correction for CO₂ in the bubbles has been attempted (with exceptions; Taracsák et al., 2019; Tucker et al., 2019); unsurprisingly, these plot below the mantle field in Figure 10, implying their CO₂/trace element ratios underestimate the source mantle ratios, owing to extensive CO₂ loss to bubbles, pre- and post-entrapment (e.g., Hartley et al., 2014).

Considering the highest reconstructed total CO₂ contents only, our MI exhibit a relatively narrow range of CO₂/Nb (277-283) and CO₂/Ba (26-22) (Fig. 10). This is at the lower limit (CO₂/Nb), or below (CO₂/Ba), the interval of representative ratios proposed for the upper mantle, based on the analysis of MORB MI and/or undegassed MORB glasses (Fig. 10). It is important to recall that, in their recent global compilation, Sun and Dasgupta (2023) have proposed that the mantle source of OIB volcanism is characterised by higher CO₂/Nb (~1850) and CO₂/Ba (~226) ratios than the depleted mantle source feeding MORB volcanism. Overall, Fig. 10 indicates that, in spite of our applied correction for CO₂ in shrinkage bubbles, our reconstructed total CO₂ concentrations likely still underestimate the real parental melt CO₂ contents, potentially because of early (before MI formation) CO₂ separation from melt into the vapour phase (e.g., Rosenthal et al., 2015; Le Voyer et al., 2017; Hirschmann, 2018; Aiuppa et al., 2021).

<figure><img src="image_11.png" /><figcaption>**Figure 10.** CO₂/trace element ratios. (a) In the Nb versus CO₂ plot, most of the literature “degassed” MI and our MI with the CO₂ content measured in the glass only (white diamond with</figcaption></figure>
light blue border) plot below the CO₂-Nb mantle array (represented by the grey area), as defined by the un-degassed MI and glasses of the MORB (characteristic upper mantle CO₂/Nb ratios from refs. Saal et al., 2002; Cartigny et al., 2008; Le Voyer et al., 2017; Hirschmann, 2018). Contrarily, our MI with the total CO₂ content (orange symbol, CO₂ glass+CO₂ bubble) fall in the mantle array field. (b) same as (a) but using Ba. Our MI with the total CO₂ content fall within the CO₂/Ba mantle range, 74-120 (Saal et al., 2002; Cartigny et al., 2008; Le Voyer et al., 2017; Hirschmann, 2018). In contrast, literature “degassed” MI and our MI with CO₂ measured only in the glass plot below the CO₂/Ba mantle array.

**5.4. High CO₂ implies deep magma storage**

The important implication of the high CO₂ contents of the studied MI is that magma must be stored deep in the Fogo plumbing system prior to eruption (Figs. 9b, 11). To demonstrate this, we use the inferred H₂O-CO₂ MI concentrations (Fig. 9a), in combination with available saturation models (Fig. 9b), to estimate the MI entrapment pressures (assuming MI are volatile saturated at the time of their entrapment). However, the conversion of volatile contents into pressures requires selection of one of the several saturation models existing in the literature (e.g., MagmaSat, Ghiorso and Gualda, 2015; Iacono-Marziano et al., 2012; D-Compress, Burgisser et al., 2015; Moretti et al., 2003; VolatileCalc, Newman and Lowenstern, 2002; Jiménez-Mejías et al., 2021; Shishkina et al., 2014), all of which have been calibrated over different compositional ranges and at relatively low pressures, and hence for lower CO₂ concentrations than the concentrations considered here). In addition to being calibrated at relatively low pressures (< 800 MPa, the model of Eguchi and Dasgupta, 2018 being the only exception), these different volatile saturation models predict CO₂ solubilities in Fogo melts that are highly variable, and typically span a factor 2 to 4 (or more) from the least and the most conservative estimate (Fig. 8b). This implies that, for a given MI CO₂ content, a range of MI entrapment pressures can be derived, with the VolatileCalc and Iacono-Marziano et al. (2012) models typically returning the lowest entrapment pressures, and the Jiménez-Mejías et al. (2021) model the highest (Fig. 9b). It is important to remind that some of these models have been calibrated for tholeiitic compositions, and their results are therefore shown for illustrative purposes only. The MagmaSat model outputs some intermediate entrapment pressure conditions in all cases (Fig. 9b).

We focus on the range (2095-21502 ppm) of total CO₂ contents, which converts into a minimum-maximum range of entrapment pressures of 100-290 MPa (217 MPa with MagmaSat) for the most CO₂-depleted inclusion (Fog 11-2-1; 0.21 wt. %) and to 648-1430 MPa (773 MPa with MagmaSat) for the most CO₂-rich inclusion (Fog 43-5-1; 2.1 wt. %). The upper range of MI entrapment pressures is higher than the pressures obtained previously at Fogo using clinopyroxene-liquid barometry (300-750 MPa) and fluid inclusion microthermometry (200-740 MPa) (Hildner et al. 2011, 2012; Klügel et al., 2020; Lo Forte et al., 2023). Ultimately, these pressures indicate MI may have been entrapped during early olivine crystallisation at depths of 26 to 35 km, well within the lithospheric mantle (extending from ~12 to ~60 km depth underneath Fogo Island; Liu et al., 2021).

We then use the composition (major elements and volatile contents) of this deeply trapped Fog 43-5-1 MI as the starting condition for model simulations of volatile degassing upon decompression. We run two model simulations of degassing using the saturation codes of Moretti et al. (2003) and MagmaSat (Ghiorso and Gualda, 2015). Both codes have been used previously to model magma degassing in a variety of volcanic systems erupting magmas of different composition (Aiuppa et al., 2007, 2010, 2017, 2022; Edmonds et al., 2010; Marini et al., 2011; Oppenheimer et al., 2011; Pino et al., 2011; Moretti 2013a, 2013b; de Moor et al., 2016; Bamber et al., 2022; Halldórsson et al., 2022; Romero et al., 2023; Sharpe et al., 2022; Boulanger and France, 2023; Liu et al., 2023; Mitchell et al., 2023). With both volatile saturation codes, we model dissolved volatile contents in
melt along a closed-system decompression path, starting from 1200 MPa pressure (final pressure, 0.1 MPa), in isothermal conditions (T = 1150 °C), and at ΔFMQ +1.5. Results of our model calculations, illustrated in Figure 9a, indicate that the modelled melt volatile compositions reproduce well the range of H₂O-CO₂ contents measured in our MI (Fig. 9a), providing confidence in our choice of volatile saturation models and model output. For comparison, we also show in Figure 9 the modelled degassing path calculated for (i) isothermal decompression in open-system conditions (orange curve) and (ii) non-isothermal (temperature range, 1128-985 °C) closed-system degassing with crystallisation taken into account (green curve). The latter model trend was obtained by linking together (a) MI saturation pressures (from H₂O-CO₂ contents) with (b) MI major element compositions, and hence (by matching against the modelled LLD; Fig. 4) melt temperature and extent of crystallisation. The open system run predicts some extreme CO₂ depletions at low pressure (that are not observed in the MI dataset; see inset in Fig. 9a); while the non-isothermal closed system+crystallisation model trend predicts some larger (> 2 wt. %) H₂O enrichments at low pressure that potentially explain the composition of some more evolved El Hierro MI.

Our results (Fig. 9) suggest magma storage conditions at Fogo to be significantly deeper than previously thought. Previous work, based on fluid inclusion microthermometry (Hildner et al., 2011, 2012; Klügel et al., 2020; Lo Forte et al., 2023) and on clinopyroxene-liquid geothermobarometry (Hildner et al., 2011, 2012; Klügel et al., 2020), have revealed the existence of a complex, vertically stacked magma storage system, with a main zone at 17–25 ± 5 km depth (Fig. 11) overlain, in the 9–12 ± 2 km depth range, by a sequence of levels of temporary magma stagnation, where differentiation eventually takes place prior/during eruption (generating phonotephritic melts, such as those erupted during the 1951, 1995 and 2014-2015 eruptions). Our volatile data, especially those of the CO₂-richest – and hence more deeply trapped – MI (2.15 wt.%, Fog 43-5-1), extend the range of magma storage conditions downward. The entrapment (saturation) pressure range, inferred for this MI, is between 648 and 1430 MPa (i.e. about 20 to 45 km depth), depending on the saturation model used (see Fig. 9b). If we concentrate on the models that account for the increase of CO₂ solubility with melt alkalinity (and silica undersaturation), the inferred MI entrapment pressures are 773 (MagmaSat) and 1020 MPa (Moretti et al., 2003), respectively. With the mantle-crust reference model previously utilised by Lo Forte et al. (2023), these pressures correspond to entrapment depths of ~27 km (MagmaSat) and ~36 km (Moretti et al., 2003). The shrinkage bubble in this MI contains a small amount of minerals (cfr. 5.1), implying its Raman-derived CO₂ bubble density (and hence bubble CO₂ content) is not severely affected by mineral precipitation. Although it is well possible that the measured CO₂ content underestimates the real CO₂ content of the parental melt, because of pre-entrapment CO₂ loss to bubbles (cfr. 5.2), we interpret such MI saturation pressures as indicative of a deep magma storage at 31.5 ± 4.3 km depth, at (or nearby) a main geological boundary bounded by the Elastic effective Thickness (Te; located at ~30 km depth; Pim et al., 2008).

Ultimately, our data prove that pre-eruptive magma storage/ascent underneath Fogo occurs/starts within/from a deep magma storage zone at 31.5 ± 4.3 km, in the lithospheric mantle. This is right above a cluster of deep earthquakes (38-43 km depth; Leva et al. 2019) that may correspond to fracturing events during magma accumulation zone replenishment events with more deeply sourced melts, rising from the convecting mantle source (at > 60 km). Such deep pre-eruptive magma storage conditions are similar to those of other intraplate oceanic islands, e.g., El Hierro (Taracsák et al., 2019), Piton de la Fournaise (Boudoire et al., 2018), La Palma (Dayton et al., 2023), and the Azores islands (e.g., Zanon et al., 2023).
<figure><img src="image_12.png" /></figure>
**Figure 11.** Schematic cross-section (not to scale) showing the entire magma plumbing system of Fogo volcano. The model is based on the framework proposed by Lo Forte et al. (2023). Our data provide new constrain on the location of the roots of the volcano at 31.5 ± 4.3 km depth, which feed the upward magma storage zone at 13-25 km depth, as well as the magma stagnation zone located at 9-12 km depth. The blue stars represent the deep earthquakes recorded by Leva et al. (2019), in proximity of the deep magma accumulation zone. The right-hand side boxes present the previous barometry results. Consistent with Liu et al. (2021), the lithosphere-asthenosphere boundary (LAB) is located at 60 km depth. The three isotopic mantle components, which sign the magmatism of Cape Verde, were obtained from prior literature works (e.g., Gerlach et al., 1988; Doucelance et al., 2003; Escrig et al., 2005; Martins et al., 2010; Jackson et al., 2018a).

5.5. High carbon in the source mantle, and implications for the global carbon cycle

We use our parental melt CO₂ content inferred above (2.15 wt. %) to estimate the carbon content in the mantle source feeding volcanism at Fogo. To this aim, we use a modal batch-melting equation (Shaw, 1970):

$$C_{pm} = \frac{C_0}{[F + D(1 - F)]} \approx C_0/F \quad (1)$$

in which C<sub>pm</sub> stands for the CO₂ in mantle-derived parental melts (2.15 wt. % in our case), C₀ is the initial abundance of CO₂ in the source mantle, D is the CO₂ solid-melt partition coefficient (assumed at 0.00055; Rosenthal et al., 2015), and F is the melt fraction (0 to 1). This partial melting degree F in the Fogo mantle source is estimated at 0.06-0.07 by using MI major element composition to derive the final stage mantle-melt equilibration temperature (~1597 °C) and pressure (~4.8 GPa), and hence melting degree, according to the methodology of Sun and Dasgupta (2020, 2023). With these numbers, the 2.15 wt. % parental melt CO₂ content converts into 355-414 ppm C in the Fogo mantle source.

We caution our estimated carbon range should be considered as a conservative lower limit, since our MI potentially entrap an already partially degassed melt (see above). Also, the model is intrinsically simplistic as it assumes a homogeneous (with fixed C content) mantle, while we know the mantle is in fact spatially heterogenous. In spite of these limitations, we conclude our results unambiguously prove the CO₂-rich nature of alkali-rich mafic melts feeding intraplate volcanism at Cape Verde. Note that, given the very low D for CO₂ (~0; Rosenthal et al., 2015), the use of a fractional meting equation (in place of batch melting) would not lead to any obvious difference.

Our results, combined with recent observations at other OIB localities globally (e.g., Anderson and Poland, 2017; Longprè et al., 2017; Boudoire et al., 2018; Hauri et al., 2018; Miller et al., 2019; Taracsák et al., 2019; Tucker et al., 2019), suggest that a C enriched OIB mantle source is likely to be the rule rather than the exception. Our estimated Fogo mantle content (355-414 ppm C) is well within the OIB mantle source range recently derived by Sun and Dasgupta (2023) (330-400 ppm), but well above the maximal C content of 164±55 ppm inferred by Rosenthal et al. (2015) for the mantle source of intraplate basalts.

This carbon-enriched nature of global OIB, and of their mantle sources, has important consequences for our understanding of the mass/volume of deep carbon reservoirs, and of the deep carbon cycle - the transfer of carbon from the deep reservoir(s) (mantle+core) to the exosphere (ocean+atmosphere). First, our results contribute to the still open debate on the carbon content of the Bulk Silicate Earth, for which available estimates remain divergent, from ~110 ppm C (Hirschmann, 2018) to 330-450 ppm C (Marty et al., 2020) or even up to ~500 ppm C (Marty, 2012). The high C content of the OIB source mantle manifestly leans toward estimates based upon
Ar (Marty, 2012) or C/N ratio (Marty et al., 2020) terrestrial inventories that require the deep convecting mantle to be C-rich, possibly >10 more enriched than the Depleted Mantle (DMM) that sustains MORB volcanism (~20-56 ppm C; Voyer et al., 2017; Hauri et al., 2019). We need to bear in mind, however, that radiogenic isotopes have clearly proven the presence of a variety of recycled (slab-related) crustal materials in the OIB mantle (e.g., Hofmann and White, 1982; Hofmann et al., 1997; Hoernle et al., 2002; Cabral et al., 2013; Ellam, 2021), and that it is therefore well possible that a substantial fraction of the high C in OIB is in fact derived from a recycled crustal carbon source (Cabral et al., 2014). Carbon isotope results for OIB magmas would be vital to test this hypothesis but are unfortunately available for only a handful set of OIB. Where available, however, results suggest a heavy (crustal) carbon signature in the mantle (Sandoval-Velasquez et al., 2021). This implies that high carbon OIB may in fact reflect local infiltration of C-enriched metasomatic fluids/melts resulting from melting of carbonated lithologies (e.g., carbon-rich sediments and/or carbonated oceanic lithosphere) entrained into the mantle during past subduction events (Dasgupta, 2018). Such a situation may indeed explain the enriched nature of Fogo magmas.

Secondly, our results with those recently obtained elsewhere (Boudoire et al., 2018; Taracsák et al., 2019) are critical to refining estimates of the OIB contribution to deep carbon flux to the atmosphere. Previous work (Dasgupta and Hirschmann, 2010; Hauri et al., 2019) has found carbon outgassing from our planet interior to be primarily controlled by the volumetrically dominant MORB volcanism, and by less voluminous but more C-rich calcalkaline volcanism. However, MORB magmas are CO₂-poor (typically <<0.1 wt.%; see Fig. 10), while calcalkaline magma CO₂ contents lie somewhere between 0.1 and 1 wt % CO₂ (Plank and Manning, 2019). If the majority of OIB is as C-rich as Fogo, OIB volcanism is expected to contribute disproportionally (relative to other forms of volcanism) to the deep carbon flux to the atmosphere. Previous inventories (Dasgupta and Hirschmann, 2010) have found the carbon flux through OIB volcanism to be highly difficult to estimate, reflecting the often degassed (C-depleted) nature of OIB-carried MI (especially when no correction for shrinkage bubbles CO₂ is applied), and because of their extreme heterogeneity (in terms of depleted vs. enriched compositions). Because of these complexities, Dasgupta and Hirschmann (2010) conservatively derived their OIB carbon flux (1.2-30 Tg/yr) by assuming it to be ~10–50% of the MORB C flux. If for simplicity we assume global OIB to be dominantly alkaline in nature and to contain 2.15 wt. % CO₂ as our Fogo parental melt, then with a cumulative OIB magma extrusive output of 0.3-0.4 km³/yr (Crisp, 1984) the global OIB carbon flux would be 16-21 Tg/yr, and > 100 Tg/yr if degassing of intrusive magma (1.5-2 km³/yr; Crisp, 1984) is also taken into account.

These initial calculations are subject to large uncertainties. For example, tholeiitic OIB (e.g., Hawaii) or near-ridge hot-spot related OIB (e.g., Iceland) are less CO₂-poor than alkali rich OIB (Fig. 10), and the fact that all alkali rich OIB are as C-rich as Fogo (or El Hierro; Taracsák et al., 2019) is virtually untested. Ultimately, the volumetric proportions of enriched and non-enriched OIB would need to be constrained with a reasonable level of detail. Thus, our calculations here tentatively indicate that the global OIB carbon flux might be significantly higher than previously thought.

## 6. Conclusions

Olivine-hosted MI from recent (≤10 ka) Fogo basanitic tephra samples exhibit some of the highest (glass + bubbles) CO₂ contents (up to 2.15 wt. %) on record. The Fogo MI are also found to additionally contain up to 2.73 wt. % H₂O, 8205 ppm S, 1459 ppm Cl and 2249 ppm F. Our results are based on the use of three independent methodologies to correct for CO₂ contained in shrinkage bubbles, whose general agreement corroborates our reconstructed total CO₂ concentrations. The inferred MI CO₂/Nb (277-283) and CO₂/Ba (26-22) ratios are lower than the characteristic mantle
ratios, suggesting our MI – although exceptionally CO₂-rich – do still underestimate the real CO₂ contents in primary melts (potentially containing up to 8-9 wt % CO₂; Sun and Dasgupta, 2023) due to pre-entrapment CO₂ loss to bubbles. This notwithstanding, our results suggest that the basanitic magmas sustaining Fogo activity may originate from a metasomatized, carbon-rich mantle source containing a minimum of 355-414 ppm C, in agreement with recent estimates for global OIB mantle sources. Prior to eruption, these parental basanitic magmas would remain stored in the lithospheric mantle in the ~27 to ~36 km depth range, deeper than previously inferred (Hildner et al., 2011, 2012; Klügel et al., 2020; Lo Forte et al., 2023).

Our results demonstrate the importance of measuring shrinkage bubble CO₂ to correctly reconstructing the architecture of magmatic plumbing systems, and confirm the exceptional carbon-rich signature of alkali rich OIB magmas and their carbon-rich mantle sources. In spite of their relatively low eruptive volumes, OIB eruptions may contribute disproportionally to global volcanic carbon outgassing from the mantle into the Earth's atmosphere. We tentatively estimate the global OIB carbon flux to potentially be as high as 17-23 Tg/yr.

## Acknowledgements

This study was funded by the Deep Carbon Observatory (Alfred P. Sloan Foundation, USA), the Ministero dell'Università e Ricerca (Italy, Grants n. 2017LMNLAW and 2022HA8XCS), and Fundação para a Ciência e Tecnologia, Portugal (MAGAT project, ref. CIRCNA/OCT/0016/2019). LMV and IVAR are acknowledged for providing Francesco Maria Lo Forte access to laboratories. Dr. Jean-Luc Devidal (LMV) is acknowledged for assistance during electron probe analyses. Guilherme Bettencourt (IVAR) helped during sampling fieldwork at Fogo. The authors wish to acknowledge constructive comments from the Editor (Jeffrey G. Catalano), the Associate Editor (Adreas Stracke), M. Portnyagin, and one anonymous reviewer.

## Data Availability

Data are available through Mendeley Data at https://doi:10.17632/gr2dy9grdn.1.

## Appendix A. Supplementary Material

The Supplementary Material comprises an excel file labelled as Supplementary Material S1, containing multiple sheets with detailed data information: Whole rock; Sample ubication; Host olivine analysis; Scoria analysis; Uncorrected MIs analysis; Corrected MIs analysis; Total MI CO₂ contents; MI Saturation Pressures. Additionally, there is a .pdf file labelled as Supplementary Material S2, which includes the calibration of nano-SIMS standards (Figure S1) and the Raman tomography of MI Fog 43-6-2 (Figure S2).

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**Declaration of interests**

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

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