# Sulfur isotope and trace element systematics in arc magmas: seeing through the degassing via a melt inclusion study of Kyushu Island volcanoes, Japan

Masataka Kawaguchi, Kenneth T. Koga, Estelle F. Rose-Koga, Kenji Shimizu, Takayuki Ushikubo, Akira Yoshiasa

► To cite this version:

Masataka Kawaguchi, Kenneth T. Koga, Estelle F. Rose-Koga, Kenji Shimizu, Takayuki Ushikubo, et al. Sulfur isotope and trace element systematics in arc magmas: seeing through the degassing via a melt inclusion study of Kyushu Island volcanoes, Japan. Journal of Petrology, 2022, 63 (7), 10.1093/petrology/egac061. hal-03852651

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1

2 Sulfur isotope and trace element systematics in arc magmas: seeing
3 through the degassing via a melt inclusion study of Kyushu Island
4 volcanoes, Japan

5

6 Masataka Kawaguchi*,

7 1. Université Clermont Auvergne, CNRS, IRD, OPGC, Laboratoire Magmas et Volcans, F-

8 63000 Clermont-Ferrand, France, masataka.kawaguchi@etu.uca.fr

9 2. Graduate School of Science and Technology, Kumamoto University, 2-39-1 Kurokami,

10 Chuo-ku, Kumamoto-shi, Kumamoto 860-8555, Japan, 181d9061@st.kumamoto-u.ac.jp

11 Kenneth T. Koga, Université Clermont Auvergne, CNRS, IRD, OPGC, Laboratoire Magmas et

12 Volcans, F-63000 Clermont-Ferrand, France, ken.koga@uca.fr

13 Estelle Rose-Koga, Université Clermont Auvergne, CNRS, IRD, OPGC, Laboratoire Magmas

14 et Volcans, F-63000 Clermont-Ferrand, France, estelle.koga@uca.fr

15 Kenji Shimizu, Kochi Institute for Core Sample Research, Japan Agency for Marine-Earth

16 Science and Technology (JAMSTEC), 200 Monobe-otsu Nankoku, Kochi 783-8502, Japan,

17 shimmy@jamstec.go.jp

18 Takayuki Ushikubo, Kochi Institute for Core Sample Research, Japan Agency for Marine-Earth

19 Science and Technology (JAMSTEC), 200 Monobe-otsu Nankoku, Kochi 783-8502, Japan,

20 ushikubot@jamstec.go.jp

21 Akira Yoshiasa, Graduate School of Science and Technology, Kumamoto University, Kurokami

22 2-36-1, Kumamoto 860-8555, Japan, yoshiasa@kumamoto-u.ac.jp

23

24 * Indicate the corresponding author

25
26 **Abstract**

27 Sulfur is a minor element in magmas but one of the major volatile elements released in 28 volcanic systems, from the magma to the fluid phase upon ascent. Not only are sulfur 29 gases potentially toxic for humans and plants, they are also involved in causing drastic 30 climate changes after major volcanic eruptions. Therefore, studies are carried out by the 31 geoscience community to assess the magmatic sulfur flux by looking at the sulfur content 32 and isotopes in erupted products, with the ultimate aim of improving understanding of 33 the sulfur cycle in subduction zones. Kyushu island in Japan hosts 25 volcanoes among 34 which 11 are active and represent a natural hazard for the local population. It is perhaps 35 the most suitable site for the study of the sulfur cycle for its availability of recent volcanic 36 deposits and its many, highly monitored volcanoes. We investigated sulfur and sulfur 37 isotope compositions of the magma source of Kyushu Island arc using olivine-hosted melt 38 inclusions in mafic tephras and lavas, from 8 volcanoes (9 Holocene samples) going from 39 Northern Kyushu with Oninomi, Yufu, Kuju, and Aso, to Southern Kyushu volcanoes 40 such as Kirishima-Ohachidake, Kirishima-Shinmoedake, Sumiyoshi-ike, and 41 Kaimondake, and one back arc volcano, Fukue-Onidake. We measured major, trace and 42 volatile elements and S isotopes (δ³⁴S) in melt inclusions. Magma composition recorded 43 in the inclusions ranges from basalt to andesite (SiO₂ ranging from 40.3 to 60.7 wt. %). 44 For each edifice, we identified the least degassed and least differentiated compositions 45 based on volatile and trace element systematics, and selected the melt inclusions closest 46 to their primitive melts. Comparing these primitive magmas, Sr/Y underlines a 47 compositional dichotomy between volcanoes from northern (Sr/Y>20) and southern 48 Kyushu (Sr/Y<20), separated by a non-volcanic area corresponding to the subduction of 49 the Kyushu-Palau ridge. The δ³⁴S in melt inclusions range from -0.32±0.79 ‰ to 50 +9.43±0.47 ‰ (2σ) and trace the source of the magma from the different volcanoes, rather 51 than degassing or crustal fractionation processes. δ³⁴S is not fractionated by the nature of 52 the fluid (aqueous or melt) metasomatizing the mantle wedge, therefore it is not the first 53 order factor controlling the sulfur isotope variations. Instead, this study illustrates the 54 need for a heavy δ³⁴S component, likely sulfate from seawater, contained in the agent that 55 metasomatized the mantle beneath the arc. If such an observation is confirmed in other 56 subduction zones, sulfur isotopes in melt inclusions may be an effective way to trace
57 seawater input into the mantle beneath arcs.

58

59 **Keywords**

60 volatile elements, trace elements, slab input, seawater, subduction.

61
## Introduction

The behavior of sulfur (and of its isotopes) in magma is complex because of its multiple valency and its partitioning into melt, gas, fluids and S-bearing minerals. Because sulfur can become saturated in the magmatic system and crystallize as sulfides, its geochemical behavior is expected to differ significantly from lithophile trace elements, such as lanthanide series rare earth elements (REE), or high field strength elements (Nb, Ta, Hf: HFSE). Yet, assessments of the sulfur budget and its evolution in magma are important to volcanology and mantle geochemistry, especially since SO₂ in volcanic gas is considered as an indicator of magmatic flux beneath a volcano (e.g. Symonds et al. 1994; Métrich and Mandeville, 2010). Furthermore, the sulfur isotope composition of arc magma sources is challenging to determine precisely, partly because significant sulfur isotope fractionation is overprinted in erupted products (e.g. Mandeville et al., 2009). It appears nearly impossible to determine the source sulfur isotope composition through subaerial eruption products. Instead, the systematic study of sulfur isotope composition in olivine-hosted magmatic inclusions is much more revealing, as the olivine can prevent diffusive sulfur-loss through the host crystal (e.g. Bucholz et al., 2013). There have also been successful attempts recently to find primitive sulfur isotopic compositions in melt inclusions from Lassen volcano and the 2018 Kilauea flow (Muth and Wallace 2021; Lerner et al. 2021). However, other olivine inclusions have not been as successful in sampling unmodified magmas; the melt inclusion studies of the Lesser Antilles, Canary Islands and Kamchatka have recorded indications of degassing or sulfide fractionation (Bouvier et al., 2008; Beaudry et al., 2018; Gurenko et al., 2018; 2021, respectively), with δ³⁴S values ranging from -9.0 to +7.0 ‰, -5.9 to +3.2 ‰ and -1.6 to +12.3 ‰, respectively. Clearly, further efforts to interpret sulfur systematics in different arc systems are needed to help improve our understanding of the sulfur cycle in subduction zones.

Kyushu Island is one of the four main islands of Japan and is known for its high concentration of active volcanoes; 12 volcanoes are considered active along the 240 km segment of Kyushu. This makes it an ideal location to sample Holocene eruptions along an arc section. Despite this ideal setting, there are no systematic studies reporting melt inclusion compositions along this arc or characterization of mantle variations below these arc volcanoes. Furthermore, sulfur fluxes of the volcanoes are well-monitored, and for
93

example, Aso volcano is known for its persistent sulfur emissions (approximately 600 tons per day; Carn et al. 2017; Japan Meteorological Agency, 2020). It is therefore an ideal arc to carry out a systematic investigation of sulfur along the arc segment. This study reports the variation in sulfur isotopes recorded in this volcanic segment, and its relation to the geodynamic setting of this arc.

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# Geological setting

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## Tectonic setting

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The modern volcanic activity of Kyushu Island is related to the subduction of the Philippine Sea plate (Nakada and Kamata, 1991; Miyoshi et al., 2008ab, 2010; Shibata et al., 2014). This is composed of two segments of different ages, divided by the Kyushu-Palau ridge (Fig. 1a): a young, hot northern segment (Shikoku Basin, 15–26 Ma; Okino et al., 1994), and an old, cold southern segment (West Philippine Basin, 30–54 Ma; Shibata et al., 1977; Deschamps and Lallemand, 2002; Ishizuka et al., 2018). The Kyushu-Palau ridge is a remnant of Eocene–Oligocene arc that split away from the Izu-Bonin-Mariana arc, due to spreading in the Shikoku Basin (e.g., Seno and Maruyama, 1984). Recent seismic and geomagnetic studies indicate the presence of the subducting Kyushu-Palau ridge beneath central Kyushu, precisely beneath the non-volcanic region between Aso and Kirishima volcanoes (e.g. Park et al., 2009; Fig. 1b). The slab has been subducting aseismically, down to a depth of 430 km beneath Kyushu Island (Zhao et al., 2012; Fig. 1b).

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P- and S-wave seismic tomography images identify a low-velocity zone along the subducting slab down to a depth of 300–400 km, but with no contrast between the two segments (Zhao et al., 2012; Liu and Zhao, 2016). On the other hand, the receiver function tomography shows dehydration at much shallower depths, with hydrated oceanic crust present to a depth of 70 km beneath Southern Kyushu, but only reaching 50 km beneath Northern Kyushu (Abe et al., 2013). This suggests extensive dehydration beneath the fore-arc region in the north. Lastly, numerical simulations place the location of the low-velocity zone beneath the volcanic front to fore-arc region in Northern Kyushu and beneath the volcanic front to the back-arc side in the case of central to Southern Kyushu. This north-south dichotomy is due to the contrasting thermal state of the slab because of
124 its age variation (Iwamori, 2007).

125 Low-velocity and high-Poisson's ratio anomalies are identified in the upper mantle
126 beneath the back-arc region (Huang et al., 2013; Asamori and Zhao, 2015). This
127 anomalous region extends from the Tsushima Strait in the north (between the Korean
128 peninsula and Kyushu Island) to the southern edge of Kyushu Island. This anomaly
129 extends beneath the back-arc to a depth of more than 200 km, reaching below the slab. It
130 potentially indicates the existence of mantle upwelling, induced by a slab window. This
131 mantle upwelling, originating beneath the Philippine Sea plate, is possibly caused by a
132 deep dehydration reaction of the Pacific slab located below (at depths of 600-670 km),
133 and suggests the presence of convective circulation above the stagnant Pacific slab, with
134 the mantle flowing across the Philippine Sea plate via a window or around its leading
135 edge. Such "big mantle wedge" convection is thought to be the cause of the continental
136 rift systems and intraplate volcanism in East Asia (Tatsumi et al., 1990; Zhao et al., 2009;
137 Zhao and Tian, 2013).

138

139 Volcanism

140 The modern volcanism in Kyushu Island is grouped into three main volcanic
141 provinces based on geographical distribution and geochemical signature: Northern
142 Kyushu, Southern Kyushu, and the back-arc region. The volcanic activity in the frontal-
143 arc of Northern Kyushu is characterized by hot-subduction zone magmatism, such as the
144 eruption of adakitic-magmas from Yufu-Tsurumi volcano (YF in Fig. 1b; Sugimoto et
145 al., 2006; Shibata et al., 2014) and the calc-alkaline high-magnesium andesites (6-3 Ma;
146 Nakada and Kamata, 1991; Kakubuchi et al., 1995). Also in the north, high-alumina
147 island-arc basalts are found in Kuju (KJ) and Aso (AS) volcanoes from 2-0 Ma (Nakada
148 and Kamata, 1991; Mahony et al., 2011). The temporal evolution of the source mantle,
149 showing progressive increase of the slab component, potentially explains the variations
150 in major and trace element signatures of the area (Nakada and Kamata, 1991). Oninomi
151 (ON) basalt is an exception with an adakite and OIB-like affinity (i.e. high Sr/Y no HFSE
152 depletion), and an eruption age of 16 ka (Ohta et al., 1992; Shitaoka et al., 2014).

153 In Southern Kyushu, the modern activity is characterized by calc-alkaline magma
154 for both the back-arc (7.6 – 0.4 Ma; Nagao et al., 1999) and the volcanic front, where
Kirishima (KR) Sumiyoshi-ike (SM), Sakurajima (SK) and Kaimondake (KM) volcanoes are located (Mahony et al., 2011). Alkali basalts erupted at ~1 Ma were reported in the back-arc side (Uto and Uchiumi, 1997; Kita et al., 2012) and at 0.08 Ma (Aojiki cone adjacent to Sumiyoshi-ike volcano, Kita et al., 2012). In the back-arc region, alkali to tholeiitic basalts have erupted episodically from ~15 Ma (Mahony et al., 2011). Onidake volcano on Fukue Island (FK, Fig. 1b; referred to as Fukue hereafter) is a typical active back-arc volcano in this region. Both alkali basalts and calc-alkaline andesites erupted at similar ages, in a zone between the frontal-arc and back-arc regions, such as Unzen volcano (UZ, Fig. 1b) and the adjacent area, (Sugimoto et al., 2005; Kita et al., 2001).

Magmatic H₂O concentrations are estimated to be 3.6 – 6.0 wt. % for Northern Kyushu volcanoes, calculated from the observed high-An plagioclase and whole rock compositions (Zellmer et al., 2012). This indicates that the primary basaltic magma was potentially water-rich, although source lithology and slab-derived fluid/melt characteristics are poorly constrained. Miyoshi et al. (2008ab; 2010) determined boron concentrations in basaltic rocks and identified an increase in the B/Nb ratio from Yufu toward Kirishima, that they interpreted as an increase in slab-derived fluid contribution. Similarly, partial melting of the subducted young Philippine Sea plate has been suggested based on the systematics of its adakitic signature and Sr-Nd-Pb isotope ratios of rocks in Northern Kyushu (Sugimoto et al., 2006; Shibata et al., 2014).

## Samples and Methods

### Sample description

Nine olivine-bearing rocks were selected from 49 samples, covering major active volcanoes of Kyushu Island (Fig. 1ab and Table 1). We focused on lapilli-sized samples to have a higher chance of finding quenched volcanic products (thus avoiding daughter mineral crystallization within melt inclusions) and to minimize the effect of the diffusive volatile-loss (Danyushevsky et al., 2002; Lloyd et al., 2013). An exception was made for Yufu volcano, where we collected a block of lava from a pyroclastic flow deposit (YF1L) since no olivine-bearing tephra was found. The inclusions of Kuju (KJ5S), Aso (KSS), Kirishima-Ohachidake (KROHTH), Kirishima-Shinmoedake (KRSM11), Sumiyoshi-ike (SM1S), Kaimondake (KM1S12a) and Fukue (FKONON) were glassy while those of
YF1L were completely crystallized. In the case of Oninomi (ON2S), the inclusions hosted in the core of the olivines were mostly rounded and crystallized, while glass found at the rim of skeletal olivines tended to be clear, tubular and often open to the outside of the host crystal. Although the olivines of FKONON were also skeletal, the melt inclusions were fully enclosed. Complete descriptions of the appearance and size of the samples are given in Supplementary material Table S1 and document S2, following the recommended practices for melt inclusion studies (Rose-Koga et al., 2021).

Whole rock sample descriptions are from the literature and brief descriptions are summarized here: YF1L is a mafic enclave found in a Nonokusa pyroclastic-flow deposit on the southern slope of the edifice (Ohta et al., 1990; Okuno et al., 1999), which is interpreted as being equivalent to the mafic endmember of the Yufu-Tsurumi volcanic group (Ohta and Aoki, 1991). The olivine-bearing mafic blocks are typically vesiculated augite-hornblende-plagioclase basalt to basaltic andesite. ON2S is a coarse black scoria forming the edifice of Oninomi scoria cone. Its composition ranges from basalt to basaltic andesite containing plagioclase, olivine, augite, occasional orthopyroxene, destabilized hornblende, and Ti-Fe oxides. It also contains abundant granitic xenoliths (Ohta et al., 1992). KJ5S is the Hiijidake scoria, a brown scoria fall deposit, erupted from the Hiijidake cone of Kuju volcano. It is a basaltic andesite containing plagioclase, clinopyroxene, orthopyroxene, olivine (Kamata, 1997; Nagaoka and Okuno, 2014; Kawanabe et al., 2015). KSS is the Kishimadake scoria, a coarse black scoria fall deposit, erupted from the Kishimadake cone of Aso volcano (Miyabuchi and Watanabe, 1997), which is the same sample as in Kawaguchi et al. (2021) and identified as the mafic endmember of Holocene Aso eruptions. KROHTH is the Takaharu scoria, a black vesiculated basaltic andesite scoria fall deposit, erupted from Ohachidake cone, from the Kirishima volcanic group, which corresponds to the fall unit ThT of Tsutsui et al. (2007). The related lava unit contains plagioclase, clinopyroxene, orthopyroxene and olivine. KRSM11 is a mixed andesitic pumice erupted in 2011 from Shinmoedake cone, from the Kirishima volcanic group, and corresponds to Unit 2 of Miyabuchi et al. (2013), whose bulk composition is andesitic, containing phenocrysts of plagioclase, clinopyroxene, orthopyroxene, olivine, magnetite and ilmenite. SM1S is a fine to coarse black scoria produced by a phreatomagmatic eruption of Sumiyoshi-ike volcano (Moriwaki et al., 1986). It is a basaltic scoria with olivine, plagioclase and clinopyroxene phenocrysts and glassy
groundmass with microlites. KM1S12a is a lithic-rich dense to vesiculated scoria containing plagioclase, clinopyroxene and olivine phenocrysts, and corresponds to the fall unit Km12a of Fujino and Kobayashi (1997), erupted from Kaimondake volcano (Kawanabe and Sakaguchi, 2005); FKONON is a vesiculated coarse scoria, and corresponds to the Onidake scoria fall unit of Nagaoka and Furuyama (2004), erupted from the Onidake cone, from Fukue volcanic group.

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## Melt inclusion preparation

Olivine-hosted melt inclusions of these nine samples were prepared for chemical analysis. Loose olivine crystals (250 to 630 µm) were hand-picked from crushed and sieved samples. The dimensions of the inclusions and vapor bubbles were measured under a microscope (Table S1). Crystallized inclusions of YF1L and ON2S were homogenized using a Vernadsky-type microscope heating stage following the method described by Le Voyer et al. (2010). A detailed description of the heating procedure is given in Supplementary material document S1. Fully enclosed inclusions were exposed using a silicon carbide mat for coarse polishing, and corundum mats (3 µm and 1 µm mats) were used to complete the polishing. To avoid possible carbon contamination, diamond-based paste was avoided. Polished olivine crystals were subsequently mounted in indium for analysis, and a final polish with a 0.25 µm alumina paste was performed to remove scratches on the exposed surface.

## Analytical methods

Volatile element concentrations (H₂O, CO₂, S, Cl, and F) in the inclusions were measured by SIMS (Cameca IMS-1280HR of Kochi Institute of Core Sample Research, JAMSTEC, Japan) following the procedure described by Shimizu *et al.* (2017), comparable to that of Rose-Koga et al. (2020) for F, S and Cl. In brief, we used a 20 keV Cs⁺ primary ion beam of 300 – 500 pA defocused to be 10 – 15 µm in diameter. Secondary ions were accelerated at 10 kV. Then a 10 keV electron beam with a diameter of ~100 µm was applied for electrostatic charge compensation over the area of the Cs beam spot. The field aperture was set to the size corresponding to 5 x 5 µm on the sample surface. Mass resolving power was set at ~6000 to separate mass interferences (for example, to separate ³⁴S¹H interference from ³⁵Cl requires 5120 MRP; Burdo and
Morrison 1971). Negative secondary ions of ¹²C, ¹⁶OH, ¹⁹F, ³⁰Si, ³¹P, ³²S, and ³⁵Cl and the mass position of 11.9 amu were measured by an axial electron multiplier using the peak switching method. Each analysis consisted of 10 cycles, and the total measurement time was ~6 min. Typical relative standard deviations (1σ) obtained from repeated analysis of a secondary basaltic glass standard from the East Pacific Rise, EPR-G3 (Shimizu et al., 2017; Supp. Table S9), were 1.4, 3.2, 0.9, 2.5 and 1.5% for H₂O, CO₂, S, Cl and F, respectively, and the measured values of EPR-G3 usually agreed with the known value. In the event that the secondary standard value deviated significantly, the calibration curves were revised.

Sulfur isotopes in the inclusions were measured by SIMS at 1) Kochi Institute of Core Sample Research, JAMSTEC, Japan (Cameca IMS-1280HR) and 2) Centre de Recherches Pétrographiques et Géochimiques (CRPG)-CNRS-Nancy, France (Cameca IMS-1270). At Kochi, analyses were conducted following the procedure described by Shimizu et al. (2019). To summarize, we used a 20 keV Cs⁺ ion beam of ~0.5 nA defocused to be ~10 µm in diameter. A 10 keV electron beam of ~100 µm diameter was used to flood the sample surface for electrostatic charge compensation. Secondary ions were accelerated to 10 kV. The field aperture was set at a size corresponding to 15 × 15 µm² on the sample surface. Negative secondary ions of ³²S and ³⁴S were measured in multi-detection mode with an FC and axial EM, respectively. Mass resolving power was set at ~2200 and ~5000 for FC and EM detectors, respectively, to separate mass interferences of ³¹P¹H from ³²S and of ³³S¹H from ³⁴S. Each measurement consisted of 20 s pre-sputtering, 120 s auto-centering of ³²S to the field and contrast apertures, and 5 s × 50 cycles for measurements. The total measurement time for each analysis was ~7 minutes. At CRPG, the set-up was similar except we used a stronger primary beam of 3 nA defocused to ~15 µm in diameter. Mass resolving power was set to 4500. Negative secondary ions of ³²S and ³⁴S were measured in the multi-detection mode with an FC and axial EM, respectively. Each measurement consisted of 60 s pre-sputtering and 5 s × 40 cycles for measurements, giving a total analysis time of ~5 minutes for each spot. Before each measurement, we adjusted the center for ³²S. In both cases, the standard basaltic glass, EPR-G3 (δ³²S<sub>VCDT</sub>=-0.09±0.55‰, 2SD; Shimizu et al., 2019), was systematically measured during the analytical session to monitor EM drift and to define the instrumental mass fractionation factor (IMF) for ³⁴S: IMF = (³⁴S/³²S)measured/( ³⁴S/³²S)reference). At
CRPG, other standard glasses were analyzed with EPR-G3, and we did not find any matrix-dependent variation for IMF, unlike in the recent report of Taracsak et al. (2021) (further discussion in Supplementary document S1). Sulfur isotopes in the samples are expressed in δ³⁴S relative to the Vienna Canyon Diablo Troilite (VCDT, ³⁴S/³²S = 1/22.6436; Ding et al., 2001; δ³⁴S = [(³⁴S/³²S)unknown/IMF/( ³⁴S /³²S)VCDT - 1] × 1000, in per mil). The analytical error (2σ) obtained from repeated measurements of the EPR-G3 standard during the analytical session was less than 1.2 ‰ for both SIMS laboratories, and the internal precision (2σ) is also reported for individual analysis, with an average of 0.8 ‰ (Supplementary material Table S4).

Major element compositions of melt inclusion and host-olivine phenocryst were determined using a CAMECA SX100 electron microprobe at the Laboratoire Magmas et Volcans, Clermont-Ferrand, France (LMV). The analytical settings for silicate glass were 15 kV acceleration voltage, 4 nA beam current, with counting times at 30 s for Ti, at 20 s for Si, Al, Fe, Mg, Na and K, and at 10 s for Mn, Ca, and P. For olivine, we used a 15 nA beam current with counting time at 30 s for Ca, at 20 s for Si and Al, at 15 s for Ti and Cr, and at 10 s for Fe, Mn, Mg, and Ni. Beam diameter was set at 5 or 10 µm depending on the size of the inclusion, and focused for olivine measurements. Relative analytical uncertainties (1σ) obtained from repeated measurements of the basaltic glass standard (VG-2, Juan de Fuca ridge basalt, Jarosewich et al., 1980; Dixon et al., 1991) were less than 0.6% for SiO₂, 1% for Al₂O₃, FeO and CaO, 3.5% for TiO₂ and Na₂O, 10% for MnO, 1.5% for MgO, 4.5% for K₂O, 17% for P₂O₅ (Supplementary Table S3).

In addition, the S, Cl and F concentrations of 12 melt inclusions were also determined, using a 40 nA beam current with counting times of 50 s for S and Cl, and 200 s for F. For each analysis, an average of 5 measurements was taken. The standard deviations (1σ) for S, Cl, and F were less than 143 ppm, 95 ppm, and 397 ppm, respectively. Further details of the analytical methods, accuracy and verification, are found in Rose-Koga et al. (2020).

Sulfur speciation of selected inclusions was determined by measuring a relative sulfur Kα peak position in a melt inclusion sample compared to the standard anhydrite mineral using the EPMA (Carroll and Rutherford, 1988), with an assumption that S²⁻ and S⁶⁺ are the only two relevant species in silicate melt (e.g. Jugo et al., 2010). We took an average of three spots per melt inclusion, and the standard error (1σ) was approximately
±10% S<sup>6+</sup>/S<sup>Total</sup> for samples with S concentration > 1000 ppm, which is the case for 84% of the melt inclusions.

Trace element concentrations in melt inclusions were determined using a laser ablation system (193 nm Excimer Resonetics M-50E laser ablation system) coupled with an inductively coupled plasma mass spectrometer, Agilent 7500 cs (LA-ICP-MS) at LMV following the classical procedure described by Rose-Koga et al. (2012, 2017). Analyses were conducted with a laser pulse frequency of 1–3 Hz, a spot diameter between 15 and 33 µm depending on the MI size, and a fluence of 2.8 mJ/cm². Twenty-six masses were collected: ⁷Li, ⁴⁵Sc, ⁵¹V, ⁶⁵Cu, ⁸⁵Rb, ⁸⁸Sr, ⁸⁹Y, ⁹⁰Zr, ⁹³Nb, ¹³⁷Ba, ¹³⁹La, ¹⁴⁰Ce, ¹⁴¹Pr, ¹⁴⁶Nd, ¹⁴⁷Sm, ¹⁵³Eu, ¹⁵⁷Gd, ¹⁶³Dy, ¹⁶⁶Er, ¹⁷²Yb, ¹⁷⁵Lu, ¹⁷⁸Hf, ¹⁸¹Ta, ²⁰⁸Pb, ²³²Th, ²³⁸U. The internal reference mass was ⁴⁴Ca, with the CaO concentration taken from the electron microprobe analysis. A typical signal acquisition started by collecting a background for 30s, which was followed by the laser firing for 70 s or less depending on the thickness of the inclusion. Data reduction was performed by the GLITTER software (GEMOC, Macquarie University, Australia). BCR-2G, NIST 610 and NIST 612 glasses (Gao et al., 2002; Gagnon et al. 2008) were also analyzed to monitor the reproducibility and accuracy of the analyses. Typical relative standard errors on samples (1σ error of mean; σ/√n, where n is the number of cycles) were < 5% for V, Rb, Sr, Y, Zr, Ba, La, and Ce, < 8% for Sc, Nb, Pr, and Nd, < 15% for Cu, Sm, Eu, Gd, Dy, Er, Hf, Pb, Th, and U, < 25% for Li, Yb, Lu, and Ta.

## 333 Correction for post-entrapment modification

### 334 Post-entrapment crystallization

Major, volatile and trace element concentrations of melt inclusions were corrected for post-entrapment crystallization (PEC; e.g., Danyushevsky et al., 2000) by incremental addition of equilibrium olivine (0.1 wt. % step) into the melt, until it reached equilibrium with the host olivine (Toplis, 2005). We adopted a constant Fe<sup>2+</sup>/Fe<sup>Total</sup> ratio calculated from the empirical equation of Kilinc et al. (1983) from the averages of measured SKα peak shifts. For example, a single fO<sub>2</sub> value was assumed to represent a rock sample (Jugo et al., 2010). Olivine-melt equilibrium temperatures were calculated using the olivine-saturated melt thermometer of Sugawara (2000), corrected for the effect of H<sub>2</sub>O on the
olivine liquidus temperature (Médard and Grove, 2008). The correction for post-entrapment Fe-Mg exchange (also called “diffusive Fe-loss”; Danyushevsky et al., 2000) was not necessary (see Supplementary material document S1), because the FeO concentration in a melt inclusion plots along the compositional trend defined by the whole rock data. Concentration of volatile and trace elements was adjusted by assuming their perfect incompatibility in the host olivine (Table S2). The corrected concentrations are used throughout this paper (Table S8). Raw data are reported in Supplementary Tables S3, S4 and S5). It should be noted that PEC added less than 6 % of olivine to melt, for 95% of the sample.

## Vapor-bubble formation

In general, more than 40–90% of the initially dissolved CO₂ in the melt at the time of entrapment is lost to shrinkage bubbles, when vapor bubble growth occurs within an inclusion (e.g. Wallace et al., 2015). Therefore, the total CO₂ concentration in a melt inclusion was determined as (dissolved + exsolved),

$$CO_2^{\text{Total}} = CO_2^{\text{Melt}} + \phi \frac{\rho_v}{\rho_m} \times 10^6$$

where CO₂Melt is the dissolved CO₂ concentration in the melt, φ is volume fraction of the bubble in the melt inclusion, ρv is the volumetric CO₂ vapor density, and ρm is the melt density, calculated from the glass composition. The φ was determined based on measurements of the two visible axes, a and b, under the microscope (Table S1) assuming the length of the third axis as c = (a + b) / 2 (e.g. Tucker et al., 2019). The ρv value was estimated using RhyoliteMELTS ver. 1.2.0 (Ghiorso and Sack, 1995; Gualda et al., 2012; Ghiorso and Gualda, 2015), based on CO₂-H₂O solubility at the glass transition temperature, Tglass, and an equation of state (EOS) for the fluid (Duan and Zhang, 2006). The Tglass value was calculated from the temperature dependence of the viscosity at the given hydrous glass composition and a constant viscosity of 10¹² Pa·s (Giordano et al., 2005, 2008). The total CO₂ concentrations were first calculated using the measured glass composition and the equation above, and they were then corrected for PEC in the same way as for the other volatile elements. This EOS calculates ρv with an assumption that the bubble and melt are in equilibrium until the bubble ceases to expand. Since carbon (C) diffusion is temperature dependent, if the diffusive closure temperature is effectively higher than Tglass, the bubble would expand without C exsolution, resulting in an
overestimation of $\rho_v$ and of the total CO$_2$ concentration (e.g. Pichavant et al., 2013; Bali et al., 2018). If this non-equilibrium expansion has occurred, the estimated total CO$_2$ concentration represents a maximum value.

We conducted an additional test based on the method of Moore et al. (2015), which calculates the expected vapor bubble volume in a melt inclusion to assess the extent of the thermal contraction of the melt and host olivine due to post-entrapment cooling. This calculation provides a mechanical upper limit of the bubble size formed by post-entrapment cooling. For example, the test concludes that a melt inclusion contained a pre-entrapment bubble if a measured bubble volume fraction exceeds the expected value. In this case we did not apply the “CO$_2$ bubble correction”. This is because the calculation of total CO$_2$ is strongly dependent on the bubble volume fraction. Therefore, if a pre-entrapment bubble had been present, because the observed bubble volume would have been larger than the expected shrinkage bubble, it would result in an overestimation of the total CO$_2$.

The major source of error of the total CO$_2$ in the inclusions is the uncertainty of the melt inclusion volume, as the volume is estimated assuming an ellipsoid with measurements of only two axes. A recent numerical simulation reports that the assumption taking an arithmetic mean for the third ellipsoid axis yields a relative error (1σ) of 37% to 48% (Tucker et al., 2019). We conservatively assigned a 50% (1σ) relative uncertainty to inclusion volumes. The uncertainties (1σ) on the total CO$_2$ concentration therefore ranged from 10 to 50% after accounting for the volume error (Table S6).

# Results

## Major element compositions

The melt inclusions in this study varied from basalt to andesite with SiO$_2$ ranging from 40.3 to 60.7 wt. % (Table S2, Fig. 2). Fo content [Mg / (Fe + Mg) × 100 in mole] of host olivine varied from 65.2 to 83.1 (Table S7). The composition of melt inclusions formed tight clusters for each tephra sample, except Aso and Oninomi. The melt inclusions of Fukue, Yufu and Sumiyoshi-ike were classified as alkali-basalt to picrobasalt and basanite. Those of Kuju, Kirishima-Ohachidake, Kirishima-Shinmoedake and Kaimondake were sub-alkaline basaltic andesites. Melt inclusions of Aso exhibited a wide range of SiO$_2$ and were classified as transitional basalt and sub-alkaline basaltic
andesite to andesite (Kawaguchi et al., 2021). The melt inclusions of Oninomi varied from basanite to basalt, except for one andesite inclusion. Basanite magmas are rare around Kyushu Island especially at the volcanic front (Iwamori, 1991, 1992). Similar basanite rocks are found in the northern part of the Izu-Bonin volcanic arc (Oshika et al., 2014) and in primitive rocks from the central Chugoku district, SW Japan (Iwamori, 1991, 1992). The subalkaline melt inclusions from Northern Kyushu were higher in total alkali (Na₂O + K₂O) compared to those from the Southern Kyushu for the same SiO₂ concentration.

## Trace element compositions

The melt inclusions in this study were enriched in large-ion lithophile elements (LILE) over high field strength elements (HFSE) resulting in high LILE/HFSE values (Fig. 3), characteristic of island arc basalts (IAB; e.g., Elliott, 2003). Positive K, Pb and Sr anomalies were present. Positive Sr anomalies in samples from the Northern Kyushu (Yufu, Kuju, Aso) were stronger than those from Southern Kyushu (Kirishima, Sumiyoshi-ike, Kaimondake). The trace element patterns of Fukue and Oninomi displayed no Nb-Ta negative anomalies and high Nb concentration (Nb ≥ 23 ppm), as found for the whole rock data (e.g., Ohta et al., 1992; Miyoshi et al., 2008). Melt inclusions from Oninomi had higher light rare earth element concentrations (LREE), and lower HREE compared to those from Fukue (Fig. 3). Two Oninomi melt inclusions were 4 to 25 times richer in Ba than the other melt inclusions.

425

## Volatile element and sulfur isotope compositions

H₂O and CO₂ concentrations in melt inclusions were corrected for PEC and post-entrapment vapor bubble growth, and the corrected values were used throughout this paper. H₂O concentrations varied between 0.01 and 6.06 wt. % (Table S2), and there was no systematic difference between Northern and Southern Kyushu. H₂O concentrations within the same tephra sample varied less than 1 or 2 wt. % (Fig. 4). CO₂ concentrations ranged from a few ppm to 1.2 wt.% (Fig. 4). S, Cl and F concentrations varied from 142 to 3800 ppm, from 89 to 1440 ppm, and from 55 to 1170 ppm, respectively (Table S2). The high volatile contents indicate successful sampling of less-degassed, near-primitive volatile in magmas.
Chlorine concentrations in the melt inclusions varied from 89 to 1440 ppm but with less than 7% of the melt inclusions with Cl <500ppm (8 out of 117; Table S2). Most Cl concentrations were higher than those of MORB (max ~500 ppm Cl; e.g. Le Voyer et al., 2015), indicating the addition of Cl to the sub-arc mantle from the subducting slab (e.g., Straub and Layne, 2003). Similarly, F concentrations in melt inclusions (55 to 1170 ppm) were higher than those of MORB (71 to 576 ppm; e.g. Le Voyer et al., 2015), indicating a variable degree of F addition from the subducted slab. These elevated Cl and F values are characteristic of arc basalts (e.g., Rose-Koga et al., 2014; Van den Bleeken and Koga, 2015).

Measured δ³⁴S values ranged from -0.3 ± 0.8 ‰ to +9.4 ± 0.5 ‰ (Table S4), and were significantly higher than those of MORB (-0.91±0.50‰, Kanehira et al., 1973; Sakai et al., 1984; Peters et al., 2010; Labidi et al., 2012, 2014), and within the range reported for bulk S isotope data of Japanese arc volcanic rocks (from -0.2 to +18.3‰, Ueda and Sakai, 1984). Most samples have δ³⁴S > +3.4‰ (Table S4), with Fukue and Oninomi showing relatively lower δ³⁴S values (<+2.5‰ with one exception at +3.63‰; Table S4).

452

## 453 Discussions

The discussion is organized into six sections. The first three sections concern the determination of the primitive magmas of each volcano, for geochemical characteristics of major, trace, and volatile elements as well as sulfur isotopes: entitled “Major and trace element composition of the primitive magma”, “Volatile element concentration in the primitive magma”, and “S isotope composition of the primitive magma”. Then, the following two sections discuss the relationship between the geochemical characteristics of the primitive magmas and the geodynamic characteristics of the Kyushu Arc, which are entitled “The magma groups of Kyushu”, and “Geodynamic significance of the source regions”. Finally, the last section, “Sources of sulfur in the subduction zone”, outlines the contributions of sulfur isotope signatures to the arc.

Identification of the primitive magma makes it possible to characterize the magma source mantle, which is a mixture of mantle and slab-derived components in this case of arc volcanism. However, arc lavas are commonly modified through interaction with the
crust before eruption. An olivine-hosted melt inclusion is considered as being a droplet of magma trapped early in the growing crystal, and is generally assumed to have undergone negligible modification through crystal fractionation in the crust as olivine is typically the first liquidus phase in basalt. However, melt inclusions may not be completely free from modifications prior to entrapment. Therefore, it is necessary to exclude geochemical signals relating to any pre-entrapment crustal processes such as crystal fractionation, magma mixing, wall-rock assimilation, and pre-entrapment volatile element degassing, before interpreting the chemical variations of melt inclusion compositions. Once such an analysis has been done, the remaining variation represents the melting and metasomatic processes in the source region, and its relationship to the geodynamic setting.

### Major and trace element composition of the primitive magma

Combinations of incompatible and compatible elements and their ratios are useful indicators for discriminating the effects of fractional crystallization, magma mixing and partial melting of source (e.g. Fourcade and Allegre, 1981; Schiano et al., 2010). Following the method of Schiano et al. (2010), we determine for each volcano, if the variation in melt inclusion composition is due to mixing, partial melting or crystal fractionation (Supplementary material document S1). In detail, some of the inclusions showed insignificant compositional variation within a sample (Fig. 5a), and those compositions were placed at the mafic end of whole rock variation (Fig. S4). In other cases, the inclusions showed variations of Ba/Nb ratio, which is a trace element ratio of two highly incompatible elements that does not vary during mantle partial melting and crystal fractionation in magma (Fig.5b). Overall, each sample appears to form either a distinct cluster (e.g. Kirishima, Kuju), or a trend (e.g. Kaimondake, Yufu) when plotted using different sets of major and trace elements (Fig. 5). From this initial analysis, we concluded that a distinct primitive magma (or magmas) exists for each sample, and we determined a primitive melt composition for each volcano (Table 2).

All inclusion samples, except for those of Fukue, show significant variations in trace element ratios, such as Rb/Nd and Th/Nd. This indicates that magma mixing and/or partial melting is the dominant process rather than fractional crystallization (Fig.S3) since these incompatible element ratios do not fractionate with early crystallization of phases
such as olivine, plagioclase, and pyroxene (see inset, Fig. S3). Inspecting the geochemical variations for individual volcanoes (Fig. S4), we found that the magmas of Yufu, Aso, Sumiyoshi-ike and Kaimondake showed geochemical systematics indicating magma mixing and/or crustal assimilation. On the contrary, fractionation of elements during partial melting was identified for the magmas of Kuju, Kirishima-OH and Kirishima-SM. Lastly, the variations of the final group, Fukue and Oninomi, were best explained by crystal fractionation (Fig. S3). Because the major and trace element data of Fukue formed a tight cluster, we chose the least evolved melt based on trace elements (i.e. lowest concentrations), as being representative of the primitive melt of Fukue (A-I, Fig. S3).

These primitive melt compositions have the lowest SiO₂ contents of the magma series for each volcano, with high MgO, and low incompatible element concentration (Table 2). Trace element ratios are expected to reflect the composition of mantle derived magma, but the Fo contents of the host-olivines were more evolved (from 73 to 83) than the expected mantle value of 89. This suggests that our “primitive magmas” have potentially suffered from pre-entrapment fractional crystallization and/or pre-entrapment magma mixing. Alternatively, low Fo content can be acquired through Fe-Mg exchange with an evolved magma while maintaining the trace element abundance at the time of entrapment (e.g. Cottrell et al., 2002; Rose-Koga et al., 2021). In this case, primitive trace element and isotope compositions are conserved in the inclusion.

Melt inclusions of Oninomi show a wide range of SiO₂ concentration from 42.7 to 58.6 wt. %, with the variation being greater than that reported for the bulk rocks (51.0 – 54.5 wt. % SiO₂, Ohta et al., 1992). While 10 out of the 11 inclusions have less than 48.5 wt. % SiO₂, one melt inclusion has an andesitic composition (U-I, 58.6 wt. % SiO₂, Table S2). This andesitic inclusion is tubular and glassy, trapped in the rim of an olivine, indicating a late-stage entrapment during the crystallization of a cooler magma (Table S1 and Supplementary document S2). We exclude this inclusion during the rest of the discussion, because of its texture, and the reported evidence of crustal assimilation in Oninomi basaltic lava, such as the presence of granitic xenoliths and disequilibrium quartz phenocrysts (Ohta et al., 1992), and a high ⁸⁷Sr/⁸⁶Sr ratio (0.704821±0.000011, Kita et al., 2001; Shibata et al. 2014). Similarly, the anomalous enrichment of Ba, and LREE found in the two melt inclusions (L-I and DZ-I) is again inferred to be indicative of crustal assimilation (Fig. 3). The remaining inclusions show nearly identical trace
element patterns (Fig. 3), suggesting a similar origin. We chose the least evolved melt in this group as the primitive melt composition to represent Oninomi volcano (Z-I).

The melt inclusions from Aso volcanoes (KSS), especially the ones from the eruptions of the last 3700 years, have been described previously, and showed compositional variations and phenocryst textures consistent with magma mixing (Kawaguchi et al., 2021). Typically, evolved melt inclusions were trapped in reversely zoned olivines, with lower Fo content in the core (< 72) than the rim (68-78), while primitive melts were trapped in normally zoned olivine. The data in this study corroborate such a conclusion. As shown in Fig.S4, the melt inclusions and reported bulk rock data of Aso volcano indicate magma mixing (Fig. S4). For Aso (KSS), the least evolved melt of the mafic endmember is taken as being closest to the primitive composition (K-II).

There is no previous trace element data reported for Kaimondake volcano, and this limited our “test” to find definitive conclusions on the systematics of its magma composition. Kaimondake illustrates a weak concave curve on Rb versus Rb/V and Rb versus Rb/Sr plots, indicating magma mixing between a mafic magma and a more evolved magma (Fig.S4). Based on this, the composition of a melt inclusion from the mafic endmember was chosen as the primitive magma composition (J-I).

Melt inclusions of Sumiyoshi-ike (SM1S) show a variation requiring three component mixing (Supplementary material document S1). First, significant variations in K₂O (or K₂O/TiO₂), Rb and Ba/Nb are seen for a given SiO₂ (Fig. 5, and Fig. S5), indicating the presence of at least two distinct parental magmas since such incompatible element variations cannot be generated by fractional crystallization or partial melting of a single source. Furthermore, the plot of Ba/Nb against K₂O/TiO₂ (or Rb/Sr) shows that a third component is required to explain the compositional variation of the inclusions (Fig. S5ab). One component appears to be granite, added via crustal assimilation, as already identified by Pb isotope systematics (Brown et al., 2020). The other two endmembers are the low K₂O/TiO₂ with the lowest Ba/Nb endmember (S-II), and the low K₂O/TiO₂ with the highest Ba/Nb (AA-I). These are considered to be the two endmembers of the primitive magmas.

For Yufu volcano (YF1L), a negative linear correlation between SiO₂ and incompatible element concentrations, such as K₂O and Rb, indicates the mixing of two magmas. This correlation is independent of the mixing reported from the bulk rock
chemical variations between a basaltic and a dacitic magma (Fig. 5 and Fig. S4, Ohta et al., 1990; Ohta and Aoki, 1991). Furthermore, this negative correlation cannot be generated by crustal assimilation, since the partial melt of a crust is SiO₂-rich melt (e.g. granite) with high incompatible element concentrations (e.g., DePaolo, 1981). Melt inclusions of Yufu show a significant variation in Ba/Nb, which correlates inversely with SiO₂ concentration. The lack of correlation with K₂O/TiO₂ or CaO/Al₂O₃ ratios (not shown) indicates the Ba/Nb variations are not due to the assimilation of plagioclase or of crustal rocks. Therefore, the geochemical systematics of Yufu melt inclusions indicate mixing, and the two endmembers have two distinct primitive melt compositions: high and low Ba/Nb (we selected BQ-III and BD-I, respectively).

Magmas of Kirishima (KROHTH, KRSM11) and Kuju (KJ5S) show geochemical systematics consistent with a variation of degrees of partial melting from a single source (Fig.S4). For Kirishima volcano, this interpretation is consistent with whole rock studies showing the existence of two parental magmas generated by different degrees of partial melting of a similar source (Inoue, 1985; Miyamoto, 1997). The SiO₂ variations are so restricted that we decided to rely on trace elements for the choice of the most primitive compositions. We chose the melt inclusions with the consistently lowest trace element concentrations (i.e. highest degree of partial melting) representing the primitive magma compositions for Kuju (AD-V) and Kirishima (Ohachidake, G-I and Shinmoedake, E-II).

## Volatile element concentrations in the primitive magma

Volatile concentrations, especially the concentration of H₂O and CO₂, in primitive arc magmas are critical geochemical parameters that can constrain models of magma genesis. While it has been studied extensively, the current consensus is that it would be highly improbable to be able to measure the primitive H₂O and CO₂ content directly from geological materials. In fact, because of the high-diffusivity of H₂O in host-olivines, measured H₂O concentrations of melt inclusions likely equilibrated at crustal magma chamber depths (Portnyagin et al., 2008; Chen et al. 2011; Gaetani et al., 2012; Ferriss et al., 2018). Nevertheless, the H₂O concentrations trapped in melt inclusions are higher than those of erupted lava, and it is thus considered to provide a minimum estimate. On the contrary, the other volatile elements, such as CO₂, S, Cl and F, do not diffuse out
through the host-olivine within the lifetime of a magma (e.g. Bucholz et al., 2013) and they retain the values at the time of entrapment. It should be noted that if CO₂ degassing starts before entrapment of the melt in the host olivine, the primitive CO₂ composition cannot be trapped in a melt inclusion.

In this study, several methods were combined to estimate the minimum H₂O content in the primitive magmas of each edifice and the values are reported in Table 2. First, the melt inclusions from Kuju, Kirishima, and Kaimondake displayed H₂O values from 2.8 to 6.1 wt. % (Table S2). These would commonly qualify as “high-H₂O” values for the arc magma (e.g. see compilation in Cooper et al., 2012). Alternatively, the H₂O content in magma can be estimated using a plagioclase-hygrometer (Lange et al., 2009; Waters and Lange, 2015) in combination with an olivine-liquid thermometer (Sugawara, 2000; Médard et al., 2008) with a typical uncertainty of 23 °C and 0.6 wt. % H₂O (Sakuyama et al., 2014). It should be noted that the hygrometer records the water content in the magma at the time of plagioclase crystallization, which may happen after melt entrapment in olivine. Again, this is a minimum H₂O estimate for the primitive magma. For example, for the case of Aso, the plagioclase hygrometer gives an H₂O concentration of 4.7 wt.%, while measured H₂O concentration in melt inclusions is approximately 3.0 wt. % (Table S2, and Kawaguchi et al., 2021). In this case, 4.7 wt. % is taken as the minimum primitive H₂O content.

Another method for estimating H₂O content in magma is to adopt the maximum H₂O/K₂O ratio of the melt inclusions (Fig. 6a). As these elements are incompatible to silicate minerals in basalt, the ratio does not vary with crystallization. The highest ratio should indicate the minimum H₂O loss, if a magmatic suite shares the same parental magma. In the case of mixed magmas, H₂O/K₂O of the mafic endmember likely defines the least degassed ratio. This K₂O normalization also eliminates the effect of post-entrapment crystallization. Specifically, we looked for the highest H₂O/K₂O values of the melt inclusions in this study and the whole rock H₂O/K₂O from previous studies, in which H₂O content from the plagioclase hygrometer is combined with measured K₂O concentration (Table 3; Zellmer et al., 2012; 2014; Kawaguchi et al., 2021). The H₂O concentration is then calculated by multiplying the highest H₂O/K₂O ratio by the K₂O concentration of the selected primitive melt (Table 2). This approach is especially effective for the samples from Oninomi, Yufu, Aso, and Kirishima samples, as H₂O in
the inclusions was re-equilibrated to low values. It should be noted that the highest H₂O/K₂O ratios among the melt inclusions for each edifice correspond to those of the selected primitive ones identified in the previous section (Table S2), confirming the least degassed characteristic of our chosen primitive magma compositions. The exceptions were for Kuju, Yufu, and Sumiyoshi-ike, for which the highest H₂O/K₂O is not that of the selected inclusions representing primitive magma compositions. For Kuju, the difference in H₂O/K₂O between the highest melt inclusion and the selected primitive endmember was within measurement uncertainty, so we chose the value of the endmember (KJ5S-AD-V). Yufu and Sumiyoshi-ike melt inclusions both have two parental magmas (see previous section) possibly indicating two sources with different H₂O concentrations; H₂O/K₂O of the magma cannot be assumed to be constant throughout mixing. In addition, judging from the low H₂O concentrations, these melt inclusions lost a significant amount of their initial water. While we used the same approach to present the highest H₂O concentrations, those of Yufu and Sumiyoshi-ike remain truly minimum estimates.

We used the maximum CO₂ concentration of each sample as the minimum estimate for the primitive value, using a similar reasoning as for H₂O. We considered the bubbles present in melt inclusions as an exsolved fluid phase, and “added the CO₂ back” to the melt using the equation of state method (e.g., Tucker et al., 2019). The CO₂ content in the melt inclusions is reported in Table S2, and the comparison before and after the correction is displayed on Figure 4. The highest CO₂ content is found in the alkali basalt of Fukue, reaching 1.2 wt. %. Taking these corrected values, nearly all melt inclusions were formed at depths exceeding 10 km (Figure 4). The highest entrapment pressure determined using the volatile content of an inclusion is 12 kb (Fukue), and 7.4 kb (Aso) for the inclusions from the volcanic front.

Cl and F concentrations in primitive magmas are likely representative of primary concentrations, since these elements exsolve at shallow pressure and depth (~100 MPa for Cl and ~10 MPa for F; Spilliaert et al., 2006). In addition, the expected minimum entrapment depth inferred from H₂O and CO₂ concentrations in magma is mostly greater than 100 MPa, so that Cl and F could not have exsolved before entrapment. Therefore, the Cl and F concentrations of the selected melt inclusions (Table 2) are considered as being those of the primitive melt compositions. The primitive F and Cl concentrations
range from 250 to 1160 ppm and 580 to 1270 ppm, respectively. These values yield F/Nd ratios that are systematically higher by 27 to 45 than a typical MORB mantle value of 21±5 (Workman et al., 2006) or 20 (NMORB Shimizu et al. 2016). Similarly, Cl/K varies from 0.074 to 0.43, higher than the canonical mantle value of 0.0075 ± 0.0025 (Saal et al. 2002). Such high ratios are consistent with all other arc magmas and attest to these magmas incorporating a volatile component from the slab (e.g. Van den Bleeken and Koga, 2015).

S concentrations in primitive magmas are assessed from melt inclusion compositions, taking into account potential modifications by degassing and sulfide fractionation. S is a volatile element that is incorporated into the gas phase, and is also a major element constituting the sulfides that can crystallize in magmas. In silicate magmas, sulfur solubility depends on redox conditions and therefore on the relative proportions of sulfide and sulfate ions in the melt (e.g. Carrol and Webster, 1994; Jugo et al., 2005). Observed high dissolved S concentrations are characteristic of oxidized- and hydrous-basalts found in subduction zones (Roggensack, 2001; Wallace, 2005; Webster et al., 2010, Narvaez et al., 2018). The melt inclusions from our samples (except Kirishima-Ohachidake, Fukue and Oninomi) plot above the S concentration at the sulfide saturation (SCSS) curve (Fig. 7) and thus appear to be supersaturated in sulfide. The SCSS is calculated based on the saturation of FeS in which S²⁻ is the dominant S species. However, no sulfide was found in these samples, suggesting that the majority of S is in the form of S⁶⁺, preventing the formation of FeS. The remaining S is in the form of S²⁻, and the concentration must be lower than the SCSS. This explains why these samples are in fact undersaturated. There are a few exceptions (Y-I and BC-I of Sumiyoshi-ike) in which sulfide globules were present, suggesting that the SCSS is reached for some compositions. Lastly, S concentrations of all our melt inclusions are under-saturated with respect to anhydrite saturation (SCAS), indicating sulfate fractionation had not occurred (Li and Ripley, 2009; calculated SCAS between 5270 to 8560 ppm for our samples). Considering that high S concentrations were found in the inclusions, and that inclusions were trapped deep in the crust, we conclude that S depletion due to degassing is minimal, except for Yufu and Sumiyoshi-ike.

Fukue and Oninomi are different, and occur at the line of sulfide saturation (Fig. 7). This is consistent with the observations of sulfide globules commonly found in the
whole rock samples. For these samples, clearly it is not possible to determine the primitive S concentrations. The magmas are sulfide (S²⁻) saturated at the surface (Fig. 7) and as solubility of S decreases with pressure (e.g. Mavrogenes and O'Neill, 1999; Hart and Gaetani, 2006), then the magmas are also sulfide saturated at the depth of melting, assuming that the redox state of the sub-arc mantle remained reduced. Alternatively, because the sub-arc mantle is oxidized (e.g. Parkinson and Arculus, 1999; Kelley and Cottrell, 2009), then the magma remains undersaturated and the melt inclusions potentially record the primitive S compositions. In the latter case, the magmas of Fukue and Oninomi became reduced at a late stage, just prior to eruption, to explain the presence of sulfides. Kirishima-Ohachidake plots on the SCSS, but no sulfides were observed in the melt inclusions (Fig. 7). The sample has an elevated proportion of S⁶⁺ (0.5-0.9, Table S2), explaining the absence of sulfide (based on visual inspection of hand samples). Because the sample has the lowest S of all the Kyushu melt inclusions, and the concentration of H₂O, CO₂, and S are lower than those of the sample from the near-by Shinmoedake edifice, the melt inclusions of Kirishima-Ohachidake have potentially experienced substantial degassing prior to entrapment. We report the sulfur concentration of the previously selected G-I melt inclusions (800 ppm) as the probable primitive value, with a footnote (Table 2).

Sulfur may leave the magma through partitioning into the fluid during degassing (i.e. degassing of H₂O and CO₂). This S-loss via degassing results in a positive correlation of H₂O with S as reported in natural and experimental samples (e.g. Spilliaert et al. 2006; Lesne et al. 2011). Such correlations were found among the inclusions from Kirishima SM, Kaimondake, Kuju, but not for Sumiyoshi-ike, Kirishima-OH, and Yufu, whose samples showed sulfide saturation (Fig. 6a). The correlation of H₂O and S could alternatively derive from mixing, which is confirmed by correlations between the non-degassing elements F and S (Fig. 6b). It is possible to conclude that S variation is in fact due to processes other than degassing, and that the S concentration of the mafic endmember could therefore represent the primitive composition (e.g. Kaimondake). The compositional variations of Kuju and Kirishima are considered to represent a melting process. In this case, the highest degree melt represents the least fractionation from the source, and the lowest trace element concentrations (Table 2). Such primitive magmas globally coincide with the highest concentrations of S and H₂O in the melt inclusions
721 from a sample.

722 For Yufu, the S concentrations are paradoxical. The Yufu samples plot above the 723 SCSS line (Fig. 7; for example, up to 3793), suggesting a high proportion of S<sup>6+</sup> in the 724 magmas. However, the measured S speciation gives negligible S<sup>6+</sup> abundances in the melt 725 inclusions, implying reduced fO<sub>2</sub> conditions (Table S3). Furthermore, if the melt 726 inclusion is really reduced, sulfide should precipitate and the S concentration should be 727 at the SCSS, but no sulfides were found in Yufu melt inclusions. We argue that this 728 paradox is a consequence of the heating procedure specific to the sample. As Yufu melt 729 inclusions were partially crystallized, they were re-homogenized with the heating stage 730 (e.g. Le Voyer et al., 2010). During this process, a purified He flux was flushed in and 731 the oxygen fugacity was kept between 10<sup>-10</sup> and 10<sup>-9</sup> atm, to avoid oxidation of the host 732 mineral and to ensure an efficient quench. These extremely reduced conditions affected 733 the S speciation of the melt inclusion, even if the sample were to have originally been 734 oxidized. In fact, independent analyses from amphibole compositions of basaltic mafic 735 enclaves found in Yufu indicate oxidized conditions (ΔNNO+1 ± 0.4; Ohta et al., 1990; 736 Ridolfi et al., 2010). Thus, we conclude that the Yufu magma was originally oxidized and 737 the high-S concentrations of its melt inclusions reflect the value at the time of entrapment. 738

738 Mineral phases such as carbonates, sulfides, sulfates and native sulfur were reported 739 inside the bubbles upon cooling (Kamenetsky et al., 2002, Moore et al., 2015; Robidoux 740 et al., 2018). Here we did not study the inside of the inclusion bubbles. If such S-bearing 741 phases were present on the wall of the vapor bubble, primary S concentrations would be 742 significantly underestimated. Although it is impossible to examine the exposed vapor 743 bubble *a posteriori* after polishing, the bubbles in the melt inclusions of this study did not 744 contain visible solid precipitates (inspection under an optical microscope, Supplementary 745 material document S2). Without this quantification, we assumed our oxidized melt 746 inclusions to be undersaturated, and to not crystallize either anhydrites or sulfides. 747

748

### *S* isotope composition of the primitive magma

749 Identifying the source of magmatic volatiles through S-isotope characterization 750 requires careful evaluation of all potential modifications to the initial isotopic signature. 751 In general, modification and variation of sulfur isotope compositions is caused by two 752 processes: mixing of sulfur derived from isotopically distinct sources, and isotopic
fractionation due to changes in sulfur speciation. For the latter case, a significant equilibrium sulfur isotope fractionation factor ($\alpha_{gas-melt}$) was calculated for the degassing of a magma ($\alpha_{gas-melt}$ = 0.996 to 0.998 at the degassing conditions of the Mazama magma, e.g. Mandeville et al., 2009), leading to $\delta^{34}$S$_{melt}$ being heavier than equilibrium $\delta^{34}$S$_{gas}$ at oxidized conditions (typically higher than at ~ΔFMQ+1; Marini, 2011; Fiege et al., 2015). The implication was that open-system degassing potentially reaches an even higher $\delta^{34}$S$_{melt}$ (Marini et al., 1994; example of Krakatau 1883 eruption, Mandeville et al., 1998). For the other case, where modification of S-isotope composition is due to mixing, previous studies have interpreted variations in S isotope composition of arc lavas as resulting from the addition of a slab-derived seawater sulfur component to the sub-arc mantle (Sasaki and Ishihara, 1979; Alt et al., 1993, 2012; Evans et al., 2014). However, these studies pointed out that the fractionation of S isotopes during the transformation mechanisms of seawater $SO_4^{2-}$ to dissolved S in primitive arc magma is poorly constrained. The variation observed in the present sulfur isotope data (Fig. 8) is caused by mixing and is unlikely to have been influenced by degassing, because the melt inclusions of this study trapped magmas before degassing, as inferred from the high CO$_2$ contents (note, the bubble corrected values), with the exception of the Kirishima-OH sample.

The sulfur fractionation model showed that the variation in $\delta^{34}$S is maximal during open-system degassing (Marini et al., 2011; Supplementary document S1). In detail, the model predicts that even if there was extensive degassing and all the melt inclusions were affected by open-system degassing of up to 50% of the initial S, the $\delta^{34}$S of the melt inclusions would be changed by a maximum of 2 ‰, which is the 2σ error of our $\delta^{34}$S measurements. The $\delta^{34}$S measurement is insensitive to less than 50% degassing due to its uncertainty in this study. In the case of extremely degassed whole rocks, a study has reported $\delta^{34}$S of +18.0 ‰ for S contents down to 9 ppm (Aso Nakadake; Ueda and Saki, 1984). Therefore, systematic degassing of 50% is unlikely among the inclusion data here, since S concentrations of Kyushu melt inclusions are in the range of those reported for other primitive arc basalts in the world (e.g., as high as ~3000 ppm in Cascades, Wallace, 2005). Furthermore, there is no along-arc correlation between $\delta^{34}$S and Cu (sulfide fractionation indicator, Beaudry et al. 2018), while local correlations within 3 ‰ are present, we conclude that the observed isotopic variation is not controlled by sulfide
fractionation. The δ³⁴S values of our primitive melt inclusions therefore reflect the values of the metasomatized mantle below the volcanoes (Table 2).

A correlation between δ³⁴S and S concentrations would be expected if magma degassing were to be significant. There is no such correlation among melt inclusions from each individual Kyushu volcano (Fig. 8a). The inclusions Yufu are the only samples showing a weak negative correlation (Fig. 8a) and the values of other volcanoes are clustered around a single value. There is an overall weak negative correlation for samples from several of the volcanic edifices. However, this correlation is less likely to be due to degassing as it would require a common S concentration in the parental magmas and exactly the same fractionation mechanism during degassing of individual edifices spread across the island. Sulfur depletion through degassing can be illustrated by normalizing the concentration with a non-degassing, incompatible element such as K₂O, (Fig. 8b). The negative correlations with δ³⁴S are slightly more apparent for Yufu, and Sumiyoshi-ike. Given that trace element concentrations of these samples were controlled by mixing rather than degassing (Supplementary material document S1), we would argue that the negative trend is also due to mixing. Therefore, the δ³⁴S values of the mixing endmembers potentially represent the magma source values, considering their mafic character and high S concentration. Both endmembers are reported as the potential primitive magmas of Yufu and Sumiyoshi-ike (Table 2). For Oninomi and Fukue, δ³⁴S values are systematically lower than those of other oxidized magmas. In fact, the S concentration of these samples was controlled by sulfide saturation, so the low δ³⁴S values are potentially explained by fractionation during sulfide precipitation, which leaves the remaining melt slightly lighter than the oxidized magmas (e.g. Marini et al., 2011).

## The three magma groups of Kyushu

Magma sources are identified using trace element ratios of highly incompatible elements. Firstly, an inspection of the trace element spidergrams shows that Fukue and Oninomi magmas are different to the others due to their absence of Nb depletion, otherwise common in arc magmas (Fig. 3). While the differences between Northern and Southern Kyushu may not be obvious from the spidergram, they can be distinguished on ratio-ratio plots. The plot of Ba/Zr against Nb/Zr illustrates the contrast between the arc
816 magmas and the mantle array derived from MORB samples (Fig. 9a, Ba/Nb ~ 7, note also
817 model mantle compositions are indicated). Because the contribution of slab-derived
818 material to the magma source increases the Ba/Zr ratio (Fig. 9a; e.g. Leeman et al. 2005),
819 arc magmas show elevated Ba compared to Nb. Fukue and Oninomi magmas plot close,
820 and almost parallel, to the MORB mantle array, while other Kyushu magmas plot above
821 it and reflect the addition of slab-derived material, rich in Ba. Furthermore, Northern
822 Kyushu magmas (open symbols in Fig. 9a) show, in general, higher Ba/Zr (and Ba/Nb),
823 than Southern Kyushu, with the exception of one endmember of Yufu.

824 The distinct groups are also clearly visible in Fig. 9b (Sr/Y against Th/Nb). The
825 MORB melt inclusion compositions (small light blue triangles) cluster near the values of
826 the model depleted MORB mantle, with Sr/Y = 2.3 and Th/Nb = 0.053 (Workman and
827 Hart, 2005). For peridotite partial melting, Sr/Y in MORB melt inclusions show limited
828 variation, whereas the Kyushu melt inclusions show a variation of more than a factor of
829 5, with those of the North (open symbols) plotting at higher Sr/Y values (>30) than those
830 of the South (<30) and MORB. There is also a variation in Th/Nb, a ratio of highly
831 incompatible elements (Fig. 9b). This variation is caused by the input of a mobile agent
832 from the slab to the arc magma source (i.e. metasomatism), and the resulting variation is
833 characteristic of arc magmatism (see light pink solid circle representing arc olivine-hosted
834 melt inclusions). The Th/Nb of the slab agent is potentially characterized both by its
835 physical nature (aqueous flux, silicate melt, or supercritical fluid) and by the source
836 composition (sediment, altered oceanic crust). For example, Nb is strongly partitioned
837 into rutile, D<sub>Nb</sub><sup>rutile/melt</sup> ranging 14 to 550 (Foley et al. 2000; Klemme et al., 2005), and
838 the Th abundance can be controlled by allanite and apatite (Hermann, 2002). As these
839 minerals are commonly found in the sedimentary and magmatic part of the slab, the slab-
840 issued flux can be characterized by these phases (Klimm et al, 2008). Furthermore, the
841 value of Th/Nb in the altered oceanic crust is expected to be similar to that of MORB,
842 and variations in mantle fertility only causes a range in Th/Nb from 0.05 to 0.14 (Fig.
843 9b). On the contrary, subducting sediments can vary significantly, from 0.10 to 1.5, with
844 an average of 0.77 (Plank and Langmuir, 1998). As HFSE elements are poorly soluble in
aqueous fluid, it is generally inferred that only silicate melt, especially that deriving from subducting sediment, is the most effective slab agent that can change Th/Nb ratio.

The high Sr/Y values are often interpreted as being adakite signatures, as they are the result of partial melting of a slab containing residual garnet (e.g. Defant and Drummond, 1990; Martin et al. 2005). In this scenario, melting, instead of dehydration, is the process required to fractionate Sr and Y. High Sr/Y values in Northern Kyushu are comparable to those from other arc systems, for example the Southern Cascades (1 to ~10% slab melt addition into source mantle, Walowski et al., 2016). Alternatively, high Sr/Y ratios are potentially derived from magma interacting with garnet and/or amphibole in the lower arc crust during magma ascent (e.g. Davidson et al., 2007). Because the melt inclusions of this study have experienced little crystal fractionation and the primitive chemical components were carefully selected (Supplementary material document S1), the Sr/Y ratios of our inclusions are unlikely to reflect the extensive lower crust interaction processes required to change the Sr/Y ratios.

Putting together all of the characteristics of Fig. 9, Northern Kyushu magmas are higher in Sr/Y and Ba/Zr than Southern Kyushu ones, and the volcanic front magmas (Northern and Southern Kyushu) are higher in Th/Nb and lower in Nb/Zr than MORB. There are therefore three sources of the Holocene mafic magmas. Among those, Oninomi and Fukue are the closest to the mantle array (Fig. 9a), with the lowest Ba/Zr enrichments of about 1.5 and 1.8, respectively. This very modest slab contribution, particularly for Fukue, is explained by its back arc location, where the distance to the slab surface is over 400 km (Fig. 1). A low Ba/Th enrichment is also seen in Akashima volcano from the same Fukue volcanic group (Kuritani et al., 2017). While Ba/Th and Ba/Nb are close to the mantle values, Sr/Y is significantly different from the depleted mantle value, both for Fukue and Oninomi. Because the DMM can yield a maximum Sr/Y of ~10 (determined using [Sr]DMM /[Y]DMM *D<sub>Y</sub><sup>mantle/melt</sup> /D<sub>Sr</sub><sup>mantle/melt</sup>) during mantle partial melting (Workman and Hart, 2005), and the Sr/Y values of Fukue and Oninomi are >20, they must reflect the source composition.

The melt inclusions from the other edifices plot above the mantle array (Fig. 9a).
For example, fluid derived from the slab increased the Ba/Zr of melt inclusions (Fig. 9a) by a factor of 2.8 (Kaimon) to 9 (Aso). Th/Nb varies by up to a factor of 10 (1.0 for Kirishima Shinmoedake) compared to that of the MORB-cluster (~ 0.1). The slab contribution is the most significant for the two Kirishima samples, Aso and Yufu (Fig. 9a), and is interpreted as being mainly due to silicate melt addition causing high Th/Nb and Sr/Y (Fig. 9ab). On the contrary, for Sumiyoshi-ike and Kaimondake the contribution is weaker and likely due to aqueous fluids.

## Geodynamic significance of the source regions

The presence of these three magma groups indicates that there are at least three sources responsible for magma genesis below Kyushu Island. Compositional variation of arc primitive magmas is derived from variations in the degree of melting and the source composition of the metasomatized mantle, whose composition is controlled by the quantity and composition of flux added from the slab (e.g. Stolper and Newman, 1994; Tatsumi and Eggins, 1995). Such variation is observed on a scale of a whole arc (1000s of km) to within a single volcano (e.g. Le Voyer et al., 2010). The studied volcanic area covers a length of 250 km, and the subducting plate underneath has a discontinuity in the plate, separating a young northern plate from an older southern one (Shikoku Basin and West Philippine Basin, respectively, Fig. 1).

First, we examine the potential heterogeneity of the mantle below Kyushu volcanoes prior to metasomatism from the slab. This is particularly important for the Northern Kyushu volcanoes, for which basalts were previously described as “OIB-like” for their enriched trace element characteristics (e.g., Nakada and Kamata, 1991). In fact, Fukue and Oninomi magmas are alkali basalts, as commonly found in oceanic islands, with slightly elevated Th/Nb values compared to MORB-DM (Fig. 9b), and the Th/Nb ratio is closer to the value of the bulk silicate earth, which is the enriched mantle (McDonough and Sun, 1995; Palme and O'Neill, 2003; Lyubetskaya and Korenaga, 2007). Similar systematics can be seen on a Ba/Zr - Nb/Zr plot (Fig. 9a). These diagrams also show that arc magmas in general have a significant addition of Ba and Th compared to Nb. The addition of a component rich in Ba and Th to a typical upper mantle (i.e.
depleted mantle) potentially creates similar characteristics to an “enriched” mantle and detecting the difference between the two is not easy from these diagrams. In detail, given the ranges of Th/Nb and Ba/Nb of enriched and depleted mantle are much smaller than the variation of these ratios observed in arcs, we conclude that the initial mantle heterogeneity does not significantly impact the trace element characteristics of arc magmas. Therefore, it does not matter if we consider the mantle as being initially enriched or depleted before the slab agent addition. For the following discussion, we consider depleted mantle (Workman and Hart, 2005; Salters and Stracke, 2004) as the reference mantle composition below the arc volcanoes prior to metasomatism.

The dichotomy of Sr/Y values separating North and South Kyushu is thus explained by slab contribution (Fig. 9b and Fig. 10). Fractionation of Sr/Y is due to the presence of garnet as a residual phase, and the residual garnet is potentially present in the mantle, metasomatized mantle, and slab (i.e. sediment and AOC). Following the consensus about the thermal structure of the arc mantle (e.g., Syracuse et al. 2010), the hottest part of the mantle below a volcano is located about midway between the bottom of the Moho and the top of the slab, which is approximately 65-75 km (Fig. 1). This is not a depth at which melt would form in garnet peridotite (e.g. Walter, 1998). Considering that there is no strong upward movement of the solid mantle below a volcano, the maximum melting is expected to be attained at the level of this hottest area of the mantle, and this is likely the condition of equilibration between mantle and melt. The garnet signature cannot be produced in the mantle beneath the volcanic front. The residual garnet signature must therefore come from the slab either from subducted sediments or AOC (Defant and Drummond, 1990; Martin et al. 2005), or the metasomatized mantle near the slab-mantle interface (e.g. Lara and Dasgupta, 2020). Furthermore, Sr/Y is more strongly fractionated by the garnet-melt pair, than by garnet-aqueous fluid (Martin et al., 2005). On this basis, the high Sr/Y magmas in Northern Kyushu are more consistent with the presence of a silicate melt near the slab-mantle interface with the presence of residual garnet in sediment and/or AOC.

There remains the question, does the difference between the north and south simply
represent the presence/absence of silicate melt or does it represent variable contributions from different lithologies of the slab (i.e. AOC and sediments)? The answer to this question is non-unique, as the exact compositions of the subducting material are undeterminable, as are the exact pressure-temperature conditions at which the dehydration/melting reactions take place. Nevertheless, there exists a wealth of databases which allow an educated guess to be made about the composition of the slab material (e.g. Plank and Langmuir, 1998; Plank, 2014), and geodynamic models of pressure-temperature paths along the subduction zones (Syracuse et al., 2010). Here, we used the model package, Arc Basalt Simulator 5 (ABS 5, Kimura 2017), to test the role of the fluxes from sediments and AOC, as well as the presence/absence of silicate melt (Supplementary document S1 for detail). ABS 5 conveniently includes sediment and AOC compositions from around Japan. Furthermore, it includes the models of pressure-temperature trajectory, pseudosection diagrams for related lithologies, a composite slab with six lithologies, and chromatographic fluid/melt modification adjustments. In Fig. 11a,b, the endmember fluids/melts derived from a sediment (Nankai sediment, ABS 5) and AOC (Shikoku basin, ABS 5) are plotted, together with depleted mantle. As expected, the Th/Nb ratio of the fluxes derived from AOC is close to the value of the depleted mantle, and those derived from the sediments have significantly higher Th/Nb. Because the composition of the slab derived agent controls the ratios of highly incompatible elements, both fluid/melt phases from sediment and AOC must modify the source mantle to cover the range of variation observed (see mixing curve in Fig. 11b). However, because of the contrast in the proportion of water in the fluid (~90%) and melt (~10%), it requires a lower aqueous fluid flux to attain the equivalent extent of wet mantle melting. In other words, to reach the same amount of melted mantle, the proportion of slab melt needed will be greater than that of fluid. On Fig. 11, this tendency is displayed by the position of thick dashed lines, which are the result of the flux melting model with degrees of melting ranging from 5-20%. For example, the arc magmas derived from slab melt (blue and orange lines) systematically plot at higher Sr/Y than those derived from fluid (Fig. 11b). This explains why slab derived melt modifies the mantle more than fluid at a given
961 melting percentage.

962 The compositions of the mantle fluxing agent illustrate the first-order influence on
963 the trace element characteristics of arc magma, as has been known and discussed for many
964 years (e.g., Tatsumi, 1989; Kessel et al., 2005). As shown in Fig. 11, one approach is to
965 use a detailed forward model to predict the flux from the slab. Alternatively, ABS 5 also
966 includes a function of the inverse fit, by iteratively searching for the best fit solution by
967 changing conditions of fluid release along the slab. Given the multiple parameters, it was
968 fairly easy to create an inverse model that yielded results showing a contrast of slab
969 surface temperatures between North and South Kyushu. However, we concluded that the
970 fit-results of the inverse calculation of ABS 5 do not strengthen the conclusion already
971 shown with Fig. 11, so we did not discuss this further. Instead, we tested a mixing model
972 using three components: the depleted mantle, a sediment flux, and an AOC flux (Table
973 S8), and calculated the proportion of a mixture satisfying the observed incompatible ratios
974 (Th/Nb, Rb/Ba, Ba/Nb, Sr/Y, Sr/Nd, Nb/Zr). The results of the calculation also confirm
975 the conclusion drawn from Figs. 9, 10, and 11. Based on the proportion of the mixing
976 model (Fig. 12), Northern Kyushu volcanoes show stronger slab input than the Southern
977 volcanoes, suggesting a higher amount of flux due to higher temperatures. Sediment -
978 AOC contributions are also variable but do not correlate with the geographical
979 distribution of volcanoes. Lastly, we caution that this trend is just based on the Holocene
980 mafic melt inclusions that we could find, and the existence of the north-south dichotomy
981 needs to be tested by extending the sampling to other mafic eruptions of different ages.

982 The Northern Kyushu group, Yufu, Kuju, and Aso, are associated with silicate melt
983 input from the slab, and the southern Kyushu group, Kirishima, Sumiyoshi-ike,
984 Kaimondake, are associated with aqueous fluids (Fig. 11b). However, the amount of flux
985 from sediment compared to the amount of flux from AOC creates different sub-groups.
986 Aso, Kirishima and one endmember of Yufu (BQ-III) have higher Th/Nb values than
987 Kuju, Kaimondake, Sumiyoshi-ike, and the other endmember of Yufu (BD-I). This
988 difference is explained by a higher proportion of sediment flux than that of AOC (Fig.
989 11ab). Aso, Kirishima, and Yufu (BQ-III) plot closer to the mixing curves of sediment-
depleted mantle (Fig. 11ab). Curiously, this variation of sediment-AOC contributions does not correlate with the geographical north-south separation, or with the depth to the slab surface, or the dip of subduction. It should be noted that Yufu volcano melt inclusion magmas are a mixture of both melts, that with more sediment-derived flux and that with less. This variability within a single volcano indicates that primitive magmas arriving at the base of the crust can be heterogeneous for different reasons. We also note that the degrees of dehydration of the slab in the forearc region along Kyushu are variable, with patchy regions of high and low water contents in the forearc mantle wedge, as identified by a seismic tomography study (Abe et al. 2013). This indicates that availability of “water” in the slab below the arc volcanic front may vary with respect to spatial and perhaps temporal distributions, causing non-systematic contributions of sediment and AOC fluxes.

Magmas from Fukue and Oninomi are significantly different from those of the main volcanic front. They appear to record significantly different mantle processes. Fukue magma is slightly higher in Ba/Th and Th/Nb than the depleted mantle alignment (Fig. 9a). Fukue volcano is located in the back-arc region, where slab surface depth is approximately 400 km. At such a depth, the slab is mostly dehydrated and unlikely to contribute significant aqueous flux that could lead to melting. Instead, magma genesis at Fukue is caused by mantle upwelling, potentially responding to the extension of the back-arc region, because Fukue is located broadly at the northern extension of Okinawa trough. If the degree of partial melting is low, it explains the LREE enriched pattern of the spidergram (Fig. 3). Alternatively, the magma genesis could be related to a deep fluid released from the older Pacific plate stagnating at the mantle transition zone, extending over a much broader area than the young Philippine Sea plate (Kuritani et al., 2017). This model is constrained by isotopic composition of Fukue magmas, which are similar to deeply recycled sediment (EM2), and provides the mechanism for broader magmatic activities extending to China. A recent seismic tomography study appears to indicate the that Philippine Sea plate is present beneath Fukue (Zhao et al., 2012; Liu and Zhao, 2016), and if the contribution of the Pacific plate is critical, such mantle upwellings must flow
1019 around the upper Philippine sea plate.

1020 Lastly, Oninomi volcano is unique in character, quite different from the near-by
1021 volcanoes (e.g. Yufu, Kuju, and Tsurumi-dake, not studied). The spidergram of Oninomi
1022 magma shows no arc-magma-like character (Fig. 3), and it is somewhat similar to that of
1023 Fukue. The Ba/Nb and Th/Nb values are close to those of depleted mantle, suggesting it
1024 records little input from the slab. It seems that a small degree of partial melting of a
1025 slightly metasomatized mantle would be sufficient to explain the magma composition. As
1026 the volcano is located in the middle of a 40 km wide graben, active for the last 6 Ma
1027 (Kamata and Kodama, 1994), then if the extension is accompanied by lithospheric
1028 thinning, corresponding upward decompressing in the mantle below could cause adiabatic
1029 melting (McKenzie and Bickle, 1988). It should be noted that mantle melting in response
1030 to a small change in crustal load has been observed and demonstrated in Iceland, where
1031 deglaciation has caused magmatism (Slater et al., 1998). We consider that this upward
1032 decompression melting in response to the extension is the most plausible mechanism for
1033 magma genesis at Oninomi.

1034

1035 Sources of sulfur in the subduction zone

1036 The sulfur isotope compositions of melt inclusions from the Kyushu volcanoes in
1037 this study reflect the compositional variation of magma sources rather than degassing
1038 process. Now we need to look at what could be the main contribution to sulfur isotope
1039 variation in the subduction zone magmas? In our data set, δ³⁴S weakly correlates with
1040 Th/Nb, which is sensitive to slab input (Fig. 13a). As Th/Nb indicates the composition of
1041 the metasomatized magma source, and the depleted mantle is at -0.91 ‰ for δ³⁴S, and
1042 0.06 for Th/Nb, the variation is reflecting the slab input. Comparing with other melt
1043 inclusion measurements from Antilles and Kamchatka (Bouvier et al., 2008; Gurenko et
1044 al., 2018; 2021), our data plot in a similar range, showing global systematics. Fig. 13b
1045 illustrates a lack of systematics of δ³⁴S and Cl/F, which is strongly influenced by the
1046 presence of fluid and hydroxyl minerals (amphibole and mica). It is inferred that an
1047 aqueous fluid flux from lithologies with residual hydroxyl minerals yields higher Cl/F
1048 than a slab melt flux (e.g. Van den Bleeken and Koga, 2015). For example, Cl/F of a cold
subduction volcano is 3.0±2.0 (e.g. melt inclusions from Iwate volcano; Rose-Koga et al., 2014), and melt inclusions from typical fresh MORB are expected to have values around 0.4±0.4 (Le Voyer et al., 2015, 2019; Wanless et al., 2014). Kyushu magmas cluster around Cl/F = 2.1±0.8. We conclude from this systematics that the nature of the fluid (aqueous or melt) is not the main factor controlling the variation in sulfur isotopes. We have also examined other trace element indicators that are considered sensitive to fractionation processes rather than source composition. For example, V (either concentration or ratios normalized to other trace elements) does not correlate with δ³⁴S. This suggests that variation in sulfur isotopes may not be due to changes in oxidation state of the fluid carrying the sulfur. Similarly, δ³⁴S does not correlate with Sr/Y (residual garnet), F/Nd (residual hydroxyl minerals or breakdown), or La/Sm (degree of melting). The variation in δ³⁴S in Kyushu magma must be related to the compositional variation of the slab component, which is transported into the mantle wedge by a flux of mobile agents, likely a solute rich melt.

Positive correlations between δ³⁴S and trace element ratios are also found, for example with Rb/Ba, and Pb/Ce, in addition to Th/Nb (Fig. 14). Such recurring correlations underlines that the variation in δ³⁴S of these Kyushu magmas is again related to the slab flux. Careful inspection of these correlations identifies two distinct contributions in addition to the depleted mantle. In Fig. 14a, the positive trend for the samples from the main volcanic front (excluding Fukue and Oninomi) can be explained by mixing between a first endmember with a low δ³⁴S and a Rb/Ba lower than that of depleted mantle (approximately δ³⁴S < +3 ‰, Rb/Ba < 0.03), and a second endmember with a high δ³⁴S and high Rb/Ba. Along this positive trend, the Rb/Ba variations of the slab flux come from the different source composition, combined with Rb/Ba fractionations due to the presence of specific minerals, for example (i) the presence of residual phyllosilicates, which have an affinity for Rb and Ba, and can fractionate them from the slab flux, (ii) Ba is preferentially incorporated into feldspar relative to a silica rich melt. As for the sources from the slab, first, the AOC is expected to inherit the Rb/Ba ratio of MORB, because only slight addition of Rb is expected during MORB alteration (0.09; Staudigel, 2003), and this value is similar to that of the depleted mantle (0.089; Workman and Hart, 2005 and 0.073; Salters and Stracke, 2004). During subduction of this AOC there is no significant amount of phyllosilicates in the metamorphosed basalt
and gabbro, therefore neither endmember can be derived from the AOC. We note that
altered basalts sampled at the ocean floor often have high Rb/Ba (e.g. Shu et al., 2017;
Hickey-Vargas et al., 2018) but such basalts are volumetrically insignificant compared to
the total mafic material subducted. Second, the sediment flux component used in Fig. 11
gives Rb/Ba as 0.007 and 0.009 for fluid and melt respectively. Even considering oceanic
manganese nodules and carbonates, they all have Rb/Ba <0.01 (e.g. Li and Schoomaker,
2005). So, based on the Rb/Ba ratio, the sediments would be the best candidate for the
endmember with low Rb/Ba and low δ³⁴S. Subduction zone sediments are reported to
have a range of negative δ³⁴S values (-40 to +10 ‰; e.g. Sasaki and Ishihara, 1979; Lew,
1981; Alt and Burdett, 1992; Canfield and Farquhar, 2009; Alt and Shanks, 2011; Peters
et al., 2010).

For the high Rb/Ba endmember, we considered other specific slab components. For
example, the Rb/Ba ratio of shale is high (0.24) and seawater has a Rb/Ba of 8 (Li and
Schoomaker, 2005). The corresponding δ³⁴S values range from -16 to -38 ‰ (Sasaki and
Ishihara, 1979; Cruse and Lyons, 2004) for shale, and +21 ‰ for seawater (Rees et al.,
1978). Seawater sulfur is sulfate, and thus seawater derived material rich in sulfate tends
to give strong positive sulfur isotope values. The systematics of δ³⁴S and Rb/Ba suggest
a slab component of seawater itself (or rocks impregnated by seawater sulfate). Forearc
serpentinites also show variable but high Rb/Ba, because of varying interactions with
seawater (Savov et al. 2005; Deschamps et al., 2010; Albers et al. 2020). Similarly,
feldspar-bearing sandstone, as well as granite from south-west Japan, is also reported to
have high Rb/Ba (Shinjoe 1997; Kiminami et al. 2009). These rocks are magnetite-
bearing (and often sulfide-bearing) and sulfate is not expected to be stable. Only brown
clasts of serpentinites are reported to have an oxidized state due to extensive interaction
with seawater expected to have occurred near the seafloor (Albers et al. 2020). Because
of this, it is not clear if serpentinite is an adequate sink for high δ³⁴S, commonly related
to sulfate. Alternatively, serpentinites in the subducting slab can potentially carry sulfides
with positive δ³⁴S, due to extensive seawater interaction. Such serpentinites are expected
to release sulfate-bearing fluids with δ³⁴S = 14 ‰, at the time of serpentinite dehydration
to chlorite harzburgite (Alt et al. 2013). It is not clear if these ultramafic rocks are the
definitive carrier satisfying both the high δ³⁴S and high Rb/Ba, but it appears that they are
certainly the potential carriers of a seawater sulfur isotope signature.
Considering the discussion above and our current knowledge, we conclude that seawater is a plausible vector responsible for causing the high δ³⁴S and high Rb/Ba geochemical signature in Kyushu arc magmas. Direct injection of seawater into the source region of arc magmas is a possible scenario as there are reports indicating the presence of seawater in the subduction zone eclogite (seawater fluid inclusions; e.g. Sumino et al., 2010; Kawamoto et al., 2013). The model of the deep injection of seawater fluids contradicts with observations of mantle-like sulfur isotopes in the most primitive deep arc cumulates (Lee et al., 2018), where the δ³⁴S in a suite from Sierra Nevada cumulates correlates with an index of magma evolution indicating that high δ³⁴S is a result of crustal assimilation and crystal fractionation process (Lee et al., 2018). Consequently, this suggests an efficient release of sulfate from the slab before the arc magmatic front, leaving the reduced sulfide with lighter δ³⁴S as the dominant sulfur species transported to the deep mantle (Lee et al., 2018). Similar conclusion about the negative δ³⁴S of the slab component was presented by Li et al. 2020 based on the measurements of sulfides in metamorphic veins. On the contrary, Walters et al. 2019 concluded that the slab component had a positive δ³⁴S based on their analysis of sulfides in subducted ophiolites. Lastly, the analysis of subarc xenoliths indicates the presence of positive δ³⁴S primitive magma (Bénard et al. 2018). We think this disparity illustrates the complexity of sulfur behavior in magma. It should be noted that sulfur in the Kyushu inclusions are dominantly sulfate, while the cumulates in Lee et al.'s study (2018) and the slab component of Li et al. (2020) suggest the presence of reduced magma and crystallization at reduced condition dominated by sulfide. Such a contrast may mean that there is a range of redox states of arc magmas controlling the sulfur isotopic signature.

The ratio of Pb/Ce correlates with δ³⁴S. This correlation is consistent with the interpretation based on the systematics of Fig. 14a (δ³⁴S vs. Rb/Ba). The flux from the slab gives a Pb/Ce value ranging from 0.13 ~ 0.45 for an AOC-derived flux (ABS 5 model in our study), of 1.1 to 4.0 for a sediment-derived flux, and 1.7 for seawater (Li and Schoomaker 2005). Furthermore, no negative δ³⁴S component is needed to describe the systematics of the Kyushu magmas, so long as a strongly negative sulfur isotope component is subducted. For example, sulfide in sediments and oceanic crust is as low as - 40 ‰ (Peters et al., 2010; Alt and Shanks, 2011). The lack of a negative component suggests that the dominant sulfur mobility from the slab is carried out by sulfate rather
than sulfide, and the negative δ³⁴S of sulfide is perhaps not incorporated into the slab flux.
While our findings may apply only to the local sulfur cycle in Kyushu volcanoes, it shows
potential isotopic filters operating at the subduction zone that are primarily controlled by
initial sulfur speciation. Lastly, some negative δ³⁴S values are reported in melt inclusions
and embayments from the Lesser Antilles, where there are known to be large sediment
inputs into the mantle wedge (Bouvier et al., 2008), and a decreasing δ³⁴S with decreasing
S concentration for melt inclusions from Kamchatka (Gurenko, 2021); they are
interpreted as the result of sulfur fractionation in the crust (i.e. sulfide precipitation and/or
degassing). The presence or absence of negative δ³⁴S slab flux is thus critical in
understanding S cycle systematics in the subduction zone system.

## 1156 Conclusions

- Olivine-hosted magmatic melt inclusions along the Kyushu arc in this study are
  from mafic tephra deposits. We identified the melt inclusion with the most
  primitive magma composition for each volcano.

- The highest water content was 6.1 wt. %, in agreement with the high values of
  other arcs, while the CO₂ in the inclusions was lost. The minimum entrapment
  pressures determined by the volatile content of inclusions reaches the base of the
  arc crust, indicating deep entrapment of the inclusion.

- We found three primitive magma groups: North and South Kyushu, and a group
  with little slab input (Oninomi and Fukue). The northern group systematically
  shows higher Sr/Y values than the southern group. This can be explained by the
  slab agent being hotter and recording a stronger slab melt signature with residual
  garnet.

- Magma at Oninomi volcano is likely produced by a small degree of mantle melt
  with nearly no input from the slab. The crustal extension in the area may have
  triggered such a small degree of melting.

- Fukue is located in the back-arc and its magma with small slab input is consistent
  with its geodynamic setting.

- The sulfur isotope composition of the magma source was determined by carefully
  examining the trace element systematics and checking that the melt inclusions
1176 were not influenced by any degassing processes. Careful analysis of such results
1177 reveals that sulfur isotopes from the Kyushu arc magmas are characterized by a
1178 mixture of the depleted mantle, slab sediment, and a third component, which could
1179 potentially be seawater.

1180

1181 **Credit authorship contribution statement**

1182 Masataka Kawaguchi field, measurements and calculation. Kenneth Koga
1183 conceptualization, field, measurements, supervision, calculation and writing. Estelle
1184 Rose-Koga conceptualization, field, supervision and writing. Kenji Shimizu S isotope
1185 measurements, review and editing. Takayuki Ushikubo S isotope measurements, review
1186 and editing. Akira Yoshiasa field, supervision.

1187

1188 **Funding**

1189 M.K. was supported by a Fellowship from the French Embassy in Tokyo, Japan
1190 and Campus France #947957L. E.R.-K. and K.K. acknowledge funding by Région
1191 Aubergne Rhône-Alpes program SCUSI (Soutien aux Coopéra- tions Universitaires et
1192 Scientifiques INternationals, #17.009169.02).

1193 **Acknowledgements**

1194 This work is the main part of the PhD of Masataka Kawaguchi. KK and ER-K thank
1195 A. Gurenko for his technical help during our Nancy SIMS session, Nordine Bouden,
1196 Johan Villeneuve, and Michel Champenois of the SIMS Team for their warm welcome
1197 and professionalism. MK and ER-K are indebted to Dr. Jean-Luc Devidal for his support
1198 and assistance during the long EPMA and LA-ICPMS sessions at LMV. The authors
1199 deeply thank Dr. Yasuhisa Tajina, Dr. Hideto Naruo and Dr. Toshiaki Hasenaka for their
1200 help on the field, some outcrops could not have been found without them, and precious
1201 on-spot discussions clarified the geological context. MK acknowledges Prof. Mitsuru
1202 Okuno and Tomohisa Tamura for providing articles and information on the outcrops and
1203 MK thank PhD colleague Guillaume Georgeais, for his help on the field during sampling
1204 in Japan. All authors thank Fran van Wyk de Vries for English editing. Constructive
1205 reviews from two anonymous reviewers improved the clarity of our arguments and data
1206 presentation. Finally, we want to acknowledge the editorial work of Madeleine
1207 Humphreys and Georg Zellmer. MK was supported by a Fellowship from the French
1208 Embassy in Tokyo, Japan and Campus France. ER-K and KK acknowledge funding by
1209 Région Auvergne Rhône-Alpes program SCUSI (Soutien aux Coopérations
1210 Universitaires et Scientifiques Internationales). This is Laboratory of Excellence
1211 ClerVolc contribution number 540.

1212
1213

**Figure captions**

1214
1215

1216 **Fig. 1** Map of the Japanese Island arc and Kyushu sample locations. a) Map shows the tectonic
1217 setting of Japanese Islands (after Miyoshi et al., 2008 and Mahony et al., 2011). Gray solid lines
1218 represent the volcanic front. Major ridge and plateau of the Philippine Sea plate are indicated by
1219 gray dashed lines. The formation ages of the Philippine Sea plate are from Okino et al. (1994,
1220 1998) and Deschamps and Lallemand (2002), and that of the Pacific plate is from Müller et al.
1221 (2008). Location of the study area is indicated by a red square. b) Map of Kyushu Island showing
1222 target volcanoes. Open and solid triangles indicate the locations of Quaternary volcanoes, with
1223 active volcanoes shown as solid red triangles. Gray solid lines and dotted lines denote the iso-
1224 depth contours to the upper boundary of the seismic (Nakamura and Kaneshiro, 2000; Baba et al.,
1225 2002; Nakajima and Hasegawa, 2007; Hirose et al., 2008; Saito, 2017) and aseismic parts of the
1226 subducting plate (Zhao et al., 2012), respectively. The blue field and red dotted line mark the
1227 inferred subducting Kyushu-Palau ridge and slab fracture, respectively (Park et al., 2009). FK:
1228 Fukue, UZ: Unzen, TS: Tsurumi, ON: Oninomi, YF: Yufu, KJ: Kuju, AS: Aso, KR: Kirishima,
1229 SM: Sumiyoshi-ike, SK: Sakurajima, KM: Kaimondake.

1230

1231 **Fig. 2.** Total alkali ($Na_2O + K_2O$) vs. $SiO_2$ diagram for olivine-hosted melt inclusions. Dashed
1232 lines represent boundaries of classification of volcanic rocks (Le Maitre et al., 2002). Black solid
1233 line indicates the boundary between alkalic and sub-alkalic volcanic rocks (Miyashiro, 1978). All
1234 oxide concentrations have been corrected for post-entrapment crystallization and normalized to
1235 100 wt. % as volatile-free. Reported whole rock compositions of Kyushu Island volcanic rocks
1236 are shown as small solid circles for comparison: pink, volcanic rocks along the volcanic front;
1237 grey, other volcanic rocks of Kyushu. WR data: Sawamura and Matsui, 1957; Shinno, 1966;
1238 Nakamura, 1971; Nakada, 1986; Nakada and Kamata, 1988; Kakubuchi and Matsumoto, 1990;
1239 Ohta et al., 1990, 1992, Ohta and Aoki, 1991; Iwamori et al., 1991, 1992; Nagao et al., 1992;
1240 Kakubuchi et al., 1995; Imura and Kobayashi, 2001; Kita et al., 2001, 2012; Uto et al., 2004;
Kawanabe and Sakaguchi, 2005; Sugimoto et al., 2005, 2006; Miyoshi et al., 2005, 2008a, b, 2010, 2011; Sakuyama et al., 2009, 2014; Hoang et al., 2013; Tajima et al., 2013, 2014; Shibata et al., 2013, 2014; Kawanabe et al., 2015; Yamasaki et al., 2016; Kuritani et al., 2017; Brown et al., 2020; Nche et al., 2021.

**Fig. 3.** Primitive mantle-normalized trace element compositions of melt inclusions. The evolved melt inclusions of KSS (higher than 52 wt.% SiO₂ on hydrous basis) and assimilated inclusions of ON2S are shown in different colors, as they may be influenced by secondary processes. Composition of the primitive mantle is taken from Sun and McDonough (1989). The order of elements after Pearce et al. (2005).

**Fig. 4.** H₂O and CO₂ concentrations in melt inclusions. The concentrations have been corrected for PEC and for gas bubbles. Gas bubble-uncorrected CO₂ concentrations measured in melt inclusion are also shown in light gray. Dashed contour lines represent solubility at 1, 3 and 5 kbars (from inside to outside) at 1100 °C, calculated for a typical basaltic melt of 49 wt.% SiO₂ and 1100 °C using VolatileCalc (black lines, Newman and Lowenstern, 2002) and for our K-rich melt inclusion (FKONON-O-I) using MagmaSat (red lines, Ghiorso and Gualda, 2015). Error bar is standard error (1σ).

**Fig. 5** Variations of melt inclusion compositions. (a) K₂O vs. SiO₂ shows an absence of overall correlation of Kyushu magmas. For individual volcanoes, Aso, Kaimondake and Kirishinma-Shinmoedake (SM) magmas show a positive correlation and Yufu shows a weak negative correlation. (b) Ba/Nb vs. SiO₂ shows a scatter of melt inclusion compositions, while showing tight clusters for each volcano. The scatter suggests the absence of a common geochemical process along the arc. Some individual clusters show a small variation in Ba/Nb, a ratio of two highly incompatible elements, indicating the absence of mixing, either of the source, magma, or crust. On the contrary, Yufu and Sumiyoshi-ike magmas show a negative slope with significant variation of Ba/Nb, indicating the presence of mixing. Note that the evolved melt inclusions of Aso are excluded for clarity (higher than 52 wt.% SiO₂ on hydrous basis).

**Fig. 6.** K₂O-normalized volatile element abundances in the melt inclusions. (a) H₂O and S show a positive correlation (except for Oninomi, Sumiyoshi-ike, and Yufu). Such a correlation is often interpreted as a result of magmatic degassing. However, (b) F vs. S also shows a strong correlation (i.e. Kaimon) indicating that the correlation is not caused by H₂O-loss through degassing because
1275 F is only weakly soluble in magmatic fluids. Note, the horizontal axis is on a log-scale to illustrate
1276 the wide range of S/K₂O values.

1277
1278 **Fig. 7** Comparison of measured sulfur concentrations in primitive melt inclusions with the sulfur
1279 content at sulfide saturation (SCSS). Concentrations for all the other melt inclusions are shown
1280 as grey symbols. Symbols are the same as in Figure 2. Dashed black line indicates the 1:1 line.
1281 SCSS was calculated according to Fortin et al. (2015) using measured H₂O concentration and the
1282 trapping temperature (Ttrap of Table S7) of melt inclusions at 0.1 GPa. For primitive melt
1283 inclusions, we used minimum H₂O concentrations estimated from the highest H₂O/K₂O ratio
1284 (Table 3). We took the average measured H₂O concentration of the melt inclusions from each
1285 edifice and assigned this value to the melt inclusions of the same edifice for which H₂O data was
1286 missing. Decrease of SCSS caused by pressure increments does not change our conclusions, since
1287 it does not significantly decrease the S concentration. For example, an increase in pressure from
1288 0.1 GPa to 0.5 GPa decreases the S concentration in the primitive melt inclusions by a maximum
1289 of 300 ppm.

1290
1291 **Fig. 8.** Sulfur isotope composition plotted against sulfur abundance. (a) δ³⁴S vs. S plot to illustrate
1292 isotope fractionation due to degassing. Here, there is no obvious trend, except perhaps for Yufu.
1293 (b) δ³⁴S vs. S/K₂O plot may illustrate the degassing process better by compensating concentration
1294 increase due to magma differentiation. Even here, Yufu, Kaimondake and Sumiyoshi-ike show a
1295 weak negative trend. Error bars represent 2 standard errors, corresponding to approximately 95
1296 % confidence interval.

1297
1298 **Fig. 9** Slab-derived material contribution to the magma source. a) Ba/Zr vs. Nb/Zr plot. Small blue
1299 triangles are melt inclusion and glass MORB data (Shaw et al. 2010; Wanless et al., 2012; 2014;
1300 Le Voyer et al. 2015; 2017; 2019; Shimizu et al. 2019), and they define a trend with a slope of
1301 Ba/Nb of 7, shown with a blue line (also called the "MORB-OIB mantle array; Leeman et al.,
1302 1990). Small pink circles are arc melt inclusions (Straub and Layne, 2003; Churikova et al., 2007;
1303 Elburg et al., 2007; Portnyagin et al., 2007; Sadofsky et al., 2008; Vigouroux et al., 2008;
1304 Wysoczanski et al., 2006; Bouvier et al., 2008, 2010a,b; Sorbadere et al., 2011, 2013; Rose-Koga
1305 et al., 2014; Walowski et al., 2016; Narvaez et al., 2018; Barker et al., 2020). Blue circles with a
1306 cross indicate the models of the depleted mantle (Workman and Hart 2005; Salters and Stracke
1307 2004). Purple circles with a cross indicate the model bulk silicate Earth (McDonough and Sun
1308 1995; Palme and O'Neill, 2003; Lyubetskaya and Korenaga, 2007). All the melt inclusions from
Kyushu volcanoes plot above the blue line, in the arc magma field. b) Sr/Y vs. Th/Nb plot. Melt inclusions from volcanoes of Northern Kyushu (open symbols) have Sr/Y values higher than those from Southern Kyushu volcanoes (filled symbols), with a Sr/Y value of 20 separating the two groups (see text for explanation).

**Fig. 10** Sr/Y ratio and Y concentration of melt inclusions. As in Fig. 9, pink and blue filled symbols indicate olivine hosted melt inclusion compositions for arc magmas, and MORB, respectively. Symbols filled in white are samples from Northern Kyushu, and filled with color indicates samples from Southern Kyushu. Sr/Y clearly separates the north from the south.

**Fig. 11** Mixing and melting trajectories of subduction zone endmembers are overlaid on Fig. 9ab. Blue cloud is MORB melt inclusion data, and pink cloud is arc basaltic melt inclusion data. Green square indicates the depleted mantle (Workman and Hart, 2005), and filled circles are slab derived fluxes: blue, silicate melt from AOC; light blue, aqueous fluid from AOC; orange, sediment melt; brick-red, sediment fluid (exact compositions of these fluids are in Table S8). Black lines are mixing curves between a flux and the mantle. Metasomatized mantle plots along this curve. In the case of three component mixing between the mantle and other fluxes, the metasomatized mantle plots in the region between two mixing curves. Thick colored lines indicate the trajectory of a melting model (modified after the flux model of Van den Bleeken and Koga 2015). Tick marks indicate 5% intervals of the degree of melting ranging from 5 to 20 %. The larger degree of melting approaches the mixing curve. It should be noted that variation caused by melting does not deviate significantly from the mixing lines. (a) the figure illustrates a clear mantle array, and the input from the slab is the main cause of deviation from the array. The melt flux is more efficient than the fluid flux at raising Ba/Zr. (b) the figure indicates high Sr/Y, a signature of residual garnet. Again, the melt flux is more efficient at producing high Sr/Y magmas than the fluid flux.

**Fig. 12.** Results of mixing calculations for the primitive end members. Note that trace element characteristics are dominantly controlled by the depleted mantle. Small additions of fluxes from the slab are sufficient to satisfy the model. It should be noted that this calculation is based on the incompatible element ratios. Specifically, the optimal solution for mixing was confirmed by element ratios rather than concentration match. Thus the mixing fraction is independent of the element fractionation by partial melting. Northern Kyushu samples are more strongly influenced by slab signature, i.e. a slightly smaller fraction of the deleted mantle.
1343

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