Magma ascent and degassing processes of the 2011 and 2017-18 eruptions of Shinmoedake in Kirishima volcano group, Japan, based on petrological characteristics and volatile content of magmas

doi:10.21203/rs.3.rs-1449052/v1

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Magma ascent and degassing processes of the 2011 and 2017-18 eruptions of Shinmoedake in Kirishima volcano group, Japan, based on petrological characteristics and volatile content of magmas

https://doi.org/10.21203/rs.3.rs-1449052/v1

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published 23 May, 2023

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The eruption activity of Shinmoedake in the Kirishima volcanic group of Japan resumed in 2017–18 after a quiet period that lasted between 2011 and 2017. Subplinian eruptions preceded lava effusion in 2011; however, no subplinian eruption occurred in 2017–18. Petrological studies and melt inclusion analyses were conducted to investigate the ascent and degassing of the magma to understand the cause of the different eruption styles. Chemical analysis of the melt inclusions from the 2011 eruption indicates that mafic magma with high volatile content (6.2 wt% H₂O, 1.4–0.25 wt% CO₂)ascended into the shallow felsic magma (3.7–1.9 wt% H₂O, 0.048–0.025 wt% CO₂) at depths of 4–5 km. Calculations indicate that the mafic magmas were of lower density (1717–1835 kg m^-3) than the felsic magma (2264–2496 kg m^-3) at 125 MPa and that the two magmas were mixed. The 2011 mixed magma with high volatile content (4.0 wt% H₂O, 0.14–0.70 wt% CO₂)had a bubble volume of approximately 50 vol% at 50 MPa, which is likely to have caused the subplinian eruption. The whole-rock and chemical compositions of the plagioclase-, clinopyroxene-, and orthopyroxene phenocryst cores were similar in 2018 and 2011, suggesting that the 2018 magma was a remnant of the 2011 magma. Chemical analyses of the groundmass from 2018 and the MELTS calculation indicate that the magma approached chemical equilibrium over the period 2011 to 2018. Melt inclusion analyses and volcanic gas observation gave a lower bulk volatile content for the 2018 magma (3.0–2.1 wt% H₂O, 0.10–0.087 wt% CO₂) compared to the 2011 magma. Comparison of the degassed-magma volumes estimated from the S and Cl contents of the melt inclusions, SO₂ flux and volcanic gas composition, and erupted-magma volume indicates that excess degassing has been occurring in the magma due to convection since March 2011, which may have decreased the volatile content of the magma. The relatively low volatile content meant that the 2018 magma could not erupt explosively and lava was instead erupted via effusion.

Kirishima volcano

Shinmoedake

Eruption

Petrology

Magma mixing

Magma ascent

Degassing

Melt inclusions

Volatile content

Bubble volume

The behavior of volatiles in magma chambers is a highly important factor that both influences and controls eruptions. A vapor phase of more than 5 vol% that accumulated in the magma chamber of Mt. Pinatubo by 1991 may have been associated with the Plinian eruption that occurred shortly afterward (Gerlach et al. 1996). However, the effective degassing of magma in the chamber of an approximately 1 km³ dacitic-lava dome at Mt. Unzen in 1991–95 resulted in an effusive eruption (Ohba et al. 2008). These results indicate that eruption styles depend on the volatile content and degassing processes of a magma chamber. It is therefore important that the volatile behavior and degassing processes of magmas are studied in order to better understand the mechanisms surrounding volcanic eruption. Petrological studies investigating volcanic rocks allow elucidation of the physical and chemical conditions of a magma chamber, including factors such as temperature, chemical composition, and the properties of an end-member magma if mixing occurred. Melt-inclusion analysis is a powerful method for estimating the volatile content of the melt in a magma prior to eruption [e.g., Anderson 1973; Anderson et al. 1989; Johnson et al. 1994; Lowenstern 2003]. Because melt inclusions preserve the chemical compositions of a melt at the time of phenocryst growth, the volatile content of the melt in each end-member magma resulting from the magma-mixing process can be estimated. Moreover, combining melt-inclusion analysis with volcanic gas observation can be used to estimate the volume of a degassed magma.

This study is an investigation of the magma ascent and degassing processes during a series of eruptions from 2011–18 at Shinmoedake in the Kirishima volcanic group. The 2011 eruptions were marked by a series of phreatomagmatic- and subplinian eruptions followed by vulcanian explosions and the effusion of lava within the summit crater (Nakada et al. 2013). Continuous degassing activity with intermittent vulcanian explosions and continuous ash emissions were also observed from January to September 2012 (Mori and Kato 2013). The 2017–18 eruptions involved vulcanian explosions and lava effusion in the summit crater during March 2018 that were followed by small ash eruptions in October 2017; however, no subplinian eruptions occurred. Such differences provide a useful opportunity for us to investigate the relationship between eruption style and the physical and chemical conditions of a magma. Petrological studies and melt-inclusion analyses of the Shinmoedake eruptions were therefore conducted to understand the physical and chemical conditions in the magma chambers and the amount of degassed magma present during the different active periods. Geological, geophysical, and geochemical observations were thus combined to model the magma ascent and degassing processes during the 2011 and 2017–18 Shinmoedake eruptions.

Overview Of The Kirishima Volcano Group And The 2011 And 2017–18 Shinmoedake Eruptions

The Kirishima volcanic group is located in southern Kyushu, Japan (Figure. 1). The group extends over an area of about 20 × 30 km and contains more than 20 small volcanoes (Imura and Kobayashi 2001). Historically, the group can be divided into older and younger volcanoes; activity started in the older volcanoes at around 1.2 Ma and continued until several hundred thousand years ago, while the younger volcanoes first became active around 300 ka with activity continuing to the present time. The composition of the rocks in the younger Kirishima volcanoes ranges from olivine basalt to pyroxene andesite and hornblende-bearing pyroxene dacite. Shinmoedake, one of the younger volcanoes, first became active at approximately 18 ka. Since 742, all eruptions have occurred at the Shinmoedake and Ohachi volcanoes in the southeastern part of the Kirishima volcanic group, except for Ioyama (Figure. 1). The eruptions at Ohachi in 788 and 1235 and Shinmoedake over the period 1716–17 were explosive and ejected pyroclastic flows that caused significant disasters. Some minor and phreatic eruptions also occurred at Shinmoedake in 1959, 1991, and 2008 (Geshi et al. 2010).

The 2011 series of Shinmoedake eruptions started with phreatomagmatic eruptions on 19 January 2011 (Kato and Yamasato 2013; Miyabuchi et al. 2013; Nakada et al. 2013). Three subplinian eruptions during the period 26–27 January (the “P1, P2, and P3 events”) produced a total tephra mass of 1.7–2.8 × 10¹⁰ kg (Maeno et al., 2014). These eruptions were followed by lava effusion within the summit crater (28 January to 1 February) and vulcanian explosions. A dense-rock equivalent (DRE) of approximately 15 × 10⁶ m³ of effused lava was produced (Kozono et al. 2013; Nakada et al. 2013). Observations of the volcano by the Global Positioning System (GPS) indicated that it was undergoing continuous inflation from December 2009 (Kawamoto et al. 2011), while significant deflation by approximately 14 × 10⁶ m³ was observed as a result of the subplinian eruptions and lava effusion (Imakiire and Nishimura 2012; Nakao et al. 2013). Vulcanian explosions and continuous ash emissions from the summit crater occurred intermittently from February–September (Nakada et al. 2013), and an estimated 4.2 × 10⁸ kg of tephra was erupted during seven events from 24 February to 7 September (Nishiki et al. 2013). Petrological study of ash samples from the eruptions during March–June indicated that more than half of the volume of the ash was composed of essential material (Oishi et al. 2013; Suzuki et al. 2013a). GPS observation indicated that inflation restarted on 2 February and continued until November 2011, at which point it ceased (Imakiire and Nishimura 2012; Nakao et al. 2013). The intense gas emission that started with the eruptive activity peaked at 490 kg s^− 1 SO₂ (42600 t d^− 1 SO₂) on 28 January (Mori and Kato 2013), after which it gradually decreased to reach approximately 1 kg s^− 1 SO₂ by the end of 2011 (Mori and Kato 2013).

The activity then resumed with small eruptions of ash on 11–16 October 2017 after a rest period of approximately 6 y. A small ash eruption that started on 1 March 2018 rapidly changed to ash eruptions with vulcanian explosions and lava effusion within the summit crater on 6 March 2018. The lava effusion continued until 9 March, and the vulcanian explosions and ash eruptions continued intermittently until 27 June 2018. One difference in the eruption styles observed in 2011 and the 2017–18 is that subplinian eruptions did not occur in 2017-18. Approximately 15 × 10⁶ m³ Dense-Rock Equivalent (DRE) of effused lava was produced during this period (Chiba et al. 2018; Oikawa et al. 2018). Continuous inflation of the volcano observed from 1 July 2017 to 1 March 2018 (JMA, 2019) was followed by deflation of approximately 7 × 10⁶ m³ (JMA, 2019b) during the vulcanian eruptions and the lava effusion of 1 − 10 March 2018. The inflation then restarted and continued until 31 January 2019. The intense gas emission that started in October 2017 and peaked at 390 kg s^− 1 SO₂ (34000 t d^− 1 SO₂) on 7 March 2018 (JMA, 2019) decreased gradually to reach approximately 1 kg s^− 1 by October 2018 (JMA, 2019).

Sampling And Analytical Methods

Samples

Eruptive products that were analyzed from the 1716–17 and 2011 eruptions of Shinmoedake volcano (Table 1) include gray and white colored pumices from the 1716-17 eruption (hereafter “1716–17GP” and “1716–17WP”; Table 1), pumices from the phreatomagmatic eruption on 19 January 2011 (hereafter “2011–1PP”), gray and white pumices from the subplinian eruptions during events P2 and P3 on 26–27 January 2011 (hereafter “2011–2SGP and 2011–2SWP” for event P2 and “2011–3SGP” for event P3), and a bomb and lapillus from the vulcanian explosions of February 2011 (“2011–4VB” and “2011–5 VL”). Although both gray and white pumices were ejected during the 2011 subplinian eruptions (Suzuki et al. 2013b; Tomiya et al. 2013), the majority was gray. Ash particles and lapilli from the vulcanian and ash eruptions during March–August 2011 were also examined in this study (“2011–6VA”, “2011–7VL”, “2011-8A”, and “2011–9A”; Table 1) along with eruptive products from the 2018 eruptions (Table 1); gray colored pumices and scoria from the vulcanian explosion on 6 and 9 March 2018 (hereafter “2018–1VGP” and “2018–1VS” for the 6 March products and “2018–2VGP” and “2018–2VS” for the 9 March products; Table 1), three ballistic blocks from the vulcanian eruptions (“2018–3BB1, BB2, and BB3”), two dark-gray and lithic lava fragments (“2018–4L1 and L2”), a gray and porous lava fragment (“2018-4L3”), and a banded and porous lava fragment (“2018-4L4, L5”). Scoria from the 1235 eruption of Ohachi volcano was also analyzed (“1235S” in Table 1). Detailed descriptions of the samples are summarized in Table 1 and their sampling locations are noted in Fig. 1.

Analytical methods

The major-element compositions of the 1235S, 2011-7VL, and 2018 eruptive products were determined using wavelength-dispersive X-ray fluorescence analysis (XRF; Table 2; Ishizuka et al. 2014). Glass beads, prepared by fusing 1:10 mixtures of 0.5 g subsamples and lithium tetraborate, were analyzed using a Panalytical Axios XRF spectrometer at the Geological Survey of Japan (GSJ), AIST (Advanced Industrial Science and Technology). The pumices, scoria, bomb, lapilli, ashes, and lava fragments from the 2011 and 2018 eruptions were mounted in epoxy resin, ground with sandpaper, and then polished with diamond powder (1 µm). The mode compositions and porosities of the samples were determined using backscattered electron (BSE) images by scanning electron microscopy (SEM; Additional file 1). Chemical analyses of the phenocrysts, melt-inclusions, groundmass microlites (< 0.1 mm), and the bulk composition of the groundmass were carried out using electron probe micro-analyzers (EPMA; JEOL JXA-8900 and JEOL JXA-8530F at GSJ, AIST). The experimental conditions used for the analysis of minerals and groundmass were identical to those reported by Saito et al. (2002) and Saito et al. (2001). Compositional profiles across the olivine, plagioclase, and pyroxene phenocrysts from core to rim were also determined by EPMA. The major-element contents and the minor S and Cl contents of the melt inclusions were determined by EPMA, with an accelerating voltage of 15 keV, a beam current of 10 nA, and a defocused beam with a diameter of 5 µm (Table 3). The S Kα radiation wavelength was used for S analysis to determine the sulfide and sulfate proportions in the inclusions (Wallace and Carmichael 1994). The H₂O content of the inclusions was estimated by EPMA, for which it was assumed that the H₂O content was equal to the difference between 100% and the sum of the other measured contents. The error of ± 1 wt% that was used for this estimation is based on repeated analyses of standard glass samples (Saito et al., 2018).

The H₂O and CO₂ contents of the melt inclusions in the eruptive products from 1235, 1716-17, 2011, and 2018 were measured by secondary ion mass spectrometry (SIMS) (Tables 1 and 3). The mounted samples were polished with Al₂O₃ powder (1 µm) and coated with gold prior to the SIMS analyses with a Cameca ims-1270 SIMS (installed at GSJ, AIST). A detailed description of the SIMS analyses performed can be found in Saito et al. (2010) and Saito et al. (2018). A Cameca nanoSIMS 50L (installed at GSJ, AIST) was used to measure the H₂O and CO₂ contents in the melt inclusions from the 2018 eruptions. In both analyses, Cs⁺ ions were used as the primary beam and negatively charged secondary ¹H, ¹²C, and ³⁰Si ions were collected (Hauri et al. 2002). A limited area measuring 10 × 10 µm on the sample surface was analyzed to avoid the crater-edge effect (Saito et al. 2010). A normal-incidence electron gun was used for charge compensation of the sample based on the methods by Kita et al. (2004). SIMS calibration lines were made for H₂O and CO₂ on each measurement day using the reference glasses. Repeated analyses of the same melt inclusions indicated an analytical error of not more than ± 0.2 wt % for H₂O and ± 0.0028 wt % for CO₂ (Saito et al., 2010).

Whole-rock and mode compositions

The products of the 2011 eruptions (2011-2SGP, 2011-3SGP, 2011-4VB, 2011-5VL and 2011-7VL) were found to have andesite composition (SiO₂ = 57–58 wt.%), which is similar to that of 1716-17GP (Table 2; Fig. 2). 2011-2SWP has a dacite composition (SiO₂ = 63 wt.%), which is similar to 1716-17WP. Suzuki et al. (2013b) concluded that most of the ejecta from the 2011 eruptions are mixed magma products (SiO₂ = 57–58 wt.%, 960–980 ºC) of felsic mgama (silicic andesite having SiO₂ = 62–63 wt.% and 870 ºC) and mafic magma (basaltic andesite having SiO₂ = 55 wt.% and 1030 ºC). The products of the 2011 eruptions (2011-2SGP, 2011-3SGP, 2011-4VB, 2011-5VL, and 2011-7VL) with andesite composition are therefore mixed products and 2011-2SWP, with a dacite composition, was derived from the silicic andesite magma chamber.

The products of the 2018 eruptions have SiO₂ content of 58–59 wt.% and K₂O content of 1.6–1.7 wt.%, which is similar or a little more evolved than 2011-2SGP and 2011-3SGP. However, the porous gray lava samples from the 2018 eruptions (2018-4L3 and 2018-4L5) (Table 2; Fig. 2) have similar chemical compositions to 2011-2SWP. This result indicates that two different magmas that were similar to those from 2011 and 1716–17 were erupted in 2018; one an andesitic magma with SiO₂ = 57–59 wt.% and K₂O = 1.5–1.7 wt% and the other dacite with SiO₂ = 63 wt.% and K₂O = 2.4–2.5 wt%. This suggests that the same magma system was present in 1716–17 and 2018. On the other hand, the major-element composition of 1235S is basaltic andesite and is similar to lava from the 7 ka eruption at Takachihonomine (Table 2). 2011-2SGP and 2011-3SGP are composed of 20–21 vol% plagioclase phenocrysts, 1–2 vol% clinopyroxene- and orthopyroxene phenocrysts, < 1 vol% olivine phenocrysts, < 1 vol% Fe-Ti oxides, and 72–74 vol% groundmass (Additional file 1). 2011-2SWP has a mode composition similar to the subplinian gray pumice. 2011-4VB, 2011-5VL, 2011-6VA, and 2011-7VL also have mode compositions that are similar to those of the subplinian gray pumices from January 2011, but their groundmass contents (66–71 vol%) are slightly lower. The porosities of products from the subplinian eruptions (42–54 vol% of 2011 levels at -2 to -4) are a little higher than those of the vulcanian bomb, ash, and lapilli (2–40 vol% of 2011 levels at -4 to -7; Additional file 1). 2018-1VGP and 2018-2VGP are composed of 20–22 vol% plagioclase phenocrysts, 8 vol% clinopyroxene and orthopyroxene phenocrysts, < 1 vol% Fe-Ti oxides, and 70–72 vol% groundmass (Additional file 1). 2018-1VS and 2018-2VS have slightly less plagioclase (16–17 vol%) and a slightly higher pyroxene content (11 vol% clinopyroxene and orthopyroxene phenocrysts) than the vulcanian gray pumices, although they have similar groundmass contents (72–73 vol%). The groundmass contents of 2018-4L1, 2018-4L2 (67–69 vol%), and 2018-4L3 (54 vol%) are lower than those of the vulcanian gray pumices and scoria. 2018-4L3 has the highest phenocryst content of 46 vol% (and the lowest groundmass content of 54 vol%). Considering its andesite composition (SiO₂ = 62.65 wt.%; Table 2), it is likely a silicic andesite end member (SiO₂ = 62–63 wt% and 43 vol% phenocryst content) estimated by Suzuki et al. (2013b). The pumices, scoria, and lava from the 2018 eruptives contain no olivine phenocrysts, except for 2018-2VS with olivine phenocrysts of less than 1 vol% (Additional file 1). The rare olivine phenocrysts in 2018-2VS have a reaction rim composed of orthopyroxene (Additional file 2l). These observations suggest that olivine could not stabilize in the 2018 magma chamber.

Chemical compositions of minerals and groundmasses

The cores of the plagioclase phenocrysts (hereafter “plagioclase-phenocryst cores”) in the 2011 eruption products have variable composition (An_{42 − 94}) with major An_{84 − 94} and An_{50 − 62} populations (Fig. 3a). The two populations indicate that mafic magma has been mixed with high-An plagioclases, and felsic magma with low-An plagioclases. Significant variation was also observed in the rims of the plagioclase-phenocrysts (hereafter “plagioclase-phenocryst rims”) with two peaks at around An_{50 − 60} and An_{66 − 76} (Fig. 3b). In contrast, the plagioclase cores (< 0.1 mm) in the groundmass (hereafter “groundmass-plagioclase cores”) show variation of An_{40 − 90} with only one peak observed (Fig. 3c). The small variation indicates that the groundmass is a mixed product. In comparison, the significant variation observed in the plagioclase-phenocryst rims suggests that the plagioclase phenocrysts did not regrow after mixing, implying that the eruption occurred soon after mixing occurred. Similar variations are observed in the plagioclase-phenocryst cores (An_{44 − 91}) and rims (An_{50 − 90}), and in the groundmass-plagioclase cores (An_{57 − 88}) of the 2018 products (Fig. 3a-c). This similarity indicates that the magmas that were erupted in 2011 and 2018 formed via mixing of the same end-member magmas. Of the 222 plagioclase phenocrysts from 2011, 10% (2011-1PP to 2011-7VL) have rims reaching more than 0.1 mm in length (Additional file 2), while 28% of the 85 plagioclase phenocrysts in the 2018 products have such wide rims. The increase in the fraction of plagioclases with wide rims suggests that the magma chamber was able to crystallize over this period. Clinopyroxene-phenocrysts from the 2011 products have cores of Wo_{34 − 45}En₄₀₋₄₇Fs₁₃₋₂₀ and Mg#70–77 (Fig. 3g), and rims of Wo_{33 − 45}En₄₀₋₅₀Fs₁₃₋₁₉ and Mg#70–80 (Fig. 3h), while the groundmass-clinopyroxene cores show Wo_{15 − 44}En₃₄₋₆₂Fs₁₃₋₃₆ and significant variation in Mg# (Mg#60–76) (Fig. 3i). Similar variations are observed in the clinopyroxenes from the 2018 eruptions (phenocryst cores of Wo_{39 − 46}En₃₈₋₄₅Fs₁₄₋₁₈ and Mg# 69–76, phenocryst rims of Wo_{38 − 46}En₃₇₋₄₆Fs₁₄₋₁₈ and Mg# 68–76, and groundmass-clinopyroxene cores of Wo_{12 − 43}En₄₀₋₆₄ Fs_{12 − 28} and Mg# 67–79; Fig. 3g-i). The similar chemical composition of the clinopyroxene-phenocryst cores from 2011 and 2018 indicates that these clinopyroxenes were from the same source.

The orthopyroxene-phenocrysts in the 2011 products have cores of Wo_1 − 6En_{61 − 77}Fs₂₁₋₃₆ and Mg#63–79, and rims of Wo_2 − 5En_{62 − 74}Fs₂₃₋₃₈ and Mg#62–76 (Fig. 3j-k), with the groundmass-orthopyroxene cores showing similar composition (Wo_2 − 11En₅₇₋₇₄Fs₂₂₋₃₅ and Mg# 60–76), (Fig. 4l). The Mg# is bimodally distributed in the orthopyroxene-phenocryst rims (Mg#64–68 and Mg#72–74; Fig. 3e). Most of the groundmass-orthopyroxene cores show a peak at Mg#70–74. The high-Mg# orthopyroxene-phenocryst rims therefore appear to have grown simultaneously with the formation of the groundmass-orthopyroxene cores. The low-Mg# rims on some of the orthopyroxene phenocrysts indicates that the period between mixing and eruption was too short to allow high-Mg# rims to grow on all phenocrysts. Orthopyroxene phenocrysts from the 2018 eruptions have cores of Wo_2 − 4En_{61 − 75}Fs₂₂₋₃₇ and Mg# 62–77 and rims of Wo_2 − 5En_{61 − 74}Fs₂₃₋₃₆ and Mg# 63–77 with groundmass-orthopyroxene cores of Wo_2 − 8En_{63 − 76}Fs₂₀₋₃₄ and Mg# 65–79 (Fig. 3j-k). The orthopyroxene-phenocryst cores show bimodal Mg# distribution, similar to those in the 2011 products (Fig. 3j), indicating that the magmas that were erupted in 2011 and 2018 formed by mixing of the same end-member magmas. Although the orthopyroxene-phenocryst rims have bimodal Mg# distribution in the same manner as the 2011 samples (Mg#64–68 and Mg#72–74; Fig. 4e), more high-Mg# (Mg#72–74) orthopyroxene-phenocryst rims formed as compared to 2011. This discrepancy suggests that the high-Mg# orthopyroxene-phenocryst rims from the 2018 samples were able to grow in the magma chamber over the period 2011–2018.

All clinopyroxene phenocrysts in the 2011 products observed in this study have rims of less than 0.05 mm (Additional file 2). On the other hand, 13% of the 68 clinopyroxene phenocrysts from 2018 have rims reaching more than 0.05 mm in length, and the maximum width of the clinopyroxene-phenocryst rims from the 2018 eruptions is 0.16 mm. The fraction of orthopyroxene phenocryst with rims larger than 0.05 mm increased from 3% in 2011 (130) to 21% in 2018 (82). The maximum width of the orthopyroxene-phenocryst rims also increased from 0.07 mm in 2011 to 0.16 mm in 2018. The increases in the numbers of clinopyroxenes and orthopyroxenes with wide rims and the maximum width of the rims from 2011 to 2018 suggest that the magma chamber crystallized over time, as did the plagioclase-phenocryst rims.

The olivine phenocrysts in the 2011 products have core compositions of Fo_{74 − 80} (Fig. 3d). The olivine-phenocryst rims and groundmass-olivine cores from 2011 vary more (Fo_{60 − 76} and Fo_{52 − 78}, respectively) than the olivine-phenocryst cores in terms of composition (Fig. 3e and 3f). The olivine phenocrysts have rims that show normal zoning with widths of less than 0.06 mm (Additional file 2), while olivine phenocrysts are rarely observed in the 2018 products (Additional file 1) and the phenocrysts that are present (2) have large orthopyroxene rims (Additional file 2l). No groundmass-olivines were found in the 2018 products. These observations indicate that the physical and chemical conditions during the formation of the 2018 magma were not stable enough for olivines to form.

The groundmasses of the eruption products from January 2011 (2011-1PP, 2011-2SGP, and 2011-3SGP) have andesite compositions: 61–62 wt% SiO_2, and 2.1–2.2 wt% K₂O (normalized to volatile-free basis; Additional file 3; Fig. 4), except for 2011-2SWP, which has a rhyolitic composition (71 wt% SiO_2, and 3.8 wt% K₂O; Additional file 3; Fig. 4). The groundmasses of the products from February 2011 (2011-4VB and 2011-5VL) have slightly SiO₂- and K₂O-rich compositions (63–64 wt% SiO₂ and 2.4 wt% K₂O) compared to the subplinian gray pumices. The groundmasses in the eruption products from March-August 2011 (2011-6VA to 2011-9A) have richer SiO₂ and K₂O compositions (62–65 wt% SiO₂ and 2.0–2.8 wt% K₂O), except 2011-8A (GG-type). The groundmass of 2011-8A (GG-type) has a highly evolved composition (68 wt% SiO₂ and 3.5 wt% K₂O), which is similar to that of 2011-2SWP. These results indicate that the melt in the magma chamber gradually transformed to a high-SiO₂ and -K₂O composition between January and August 2011. The groundmasses of 2018-1VGP, 2018-1VS, 2018-2VGP, 2018-2VS, and 2018-4 L1 and L2 (dark-gray and lithic lava) have a dacite composition: 64–66 wt% SiO_2, and 2.5–2.9 wt% K₂O (Additional file 5; Fig. 4). These compositions are similar to or a little evolved compared to the eruptive products from March–August 2011. This similarity suggests that the melt in the 2018 magma was a remnant of the melt in the 2011 magma. On the other hand, the groundmass of 2018-4L3 (gray and porous lava) has a rhyolite composition (75 wt% SiO_2, and 4.5 wt% K₂O; Additional file 3; Fig. 4), which is the highest SiO₂ and K₂O contents among the groundmasses observed in the 2011–2018 products.

Melt inclusions

Four melt inclusions from 1235S, 16 inclusions from 1716-17GP, 79 inclusions in the 2011 eruptive products (2011-1PP to 2011-7VL), and 24 inclusions from 2018-1VGP were analyzed in this study (Table 1; Table 3; Additional file 4). Most inclusions are glassy, but several contain minerals and/or bubbles, and are found in the olivine, clinopyroxene, orthopyroxene, plagioclase, and ilmenite phenocrysts (Table 3; Additional file 4). Most inclusions are elliptical or quadrilateral in shape and range from 0.01 to 0.15 mm in diameter (Fig. 5). BSE images of the melt inclusions in the olivines show overgrowth of 1–2 µm. The chemical compositions of these inclusions were corrected for post-entrapment crystallization by adding the host-olivine component until an olivine-melt equilibrium was reached, assuming that K_D = 0.30 (Roedder and Emslie 1970; Saito et al. 2010). Olivine overgrowth of 0–5.3 wt% was estimated via this method (Table 3; Additional file 4).

One inclusion found in an olivine from 1235S has andesite composition (59 wt% SiO₂ and 1 wt% K₂O) while three inclusions in the clinopyroxenes from 1235S have dacite composition (63–68 wt% SiO₂ and 1–2 wt% K₂O) (Table 3; Fig. 6). Sixteen inclusions in 1716-17GP have dacite to rhyolite compositions (62–76 wt% SiO₂ and 2–5 wt% K₂O; Table 3; Fig. 6; Additional file 4). Inclusions from the 2011 eruptions can be divided into two types: andesite inclusions in olivine phenocrysts (34 inclusions) with 53–62 wt% SiO₂ and 1–3 wt% K₂O, and dacite-rhyolite inclusions in the plagioclase, clinopyroxene, and orthopyroxene phenocrysts (8 inclusions with 72–75 wt% SiO₂ and 4–5 wt% K₂O in the plagioclase phenocrysts, 18 inclusions with 68–76 wt% SiO₂ and 3–7 wt% K₂O in the clinopyroxene phenocrysts, and 17 inclusions with 68–78 wt% SiO₂ and 3–5 wt% K₂O in the orthopyroxene phenocrysts). The inclusions from the 2018 eruptions have andesite to rhyolite compositions; 6 inclusions in the plagioclase phenocrysts have 72–73 wt% SiO₂ and 4 wt% K₂O, 6 inclusions in the clinopyroxene phenocrysts have 71–76 wt% SiO₂ and 4–5 wt% K₂O, and 12 inclusions in the orthopyroxene phenocrysts have 60–75 wt% SiO₂ and 2–4 wt% K₂O. The chemical composition of the andesite and rhyolite inclusions did not differ greatly (Fig. 4). Less-evolved chemical compositions of < 60 wt% SiO₂ and < 2 wt% K₂O in the 2011 samples were not observed in the inclusions from 2018, which have intermediate compositions rather than the andesite and dacite-rhyolite inclusions observed in the 2011 eruptives.

Significant variation is observed in the H₂O, CO₂, and S contents of the andesite inclusions within the olivine phenocrysts, which range from 1.4 to 7.0 wt% for H₂O (by SIMS), from 0.001 to 0.054 wt% for CO₂, and from 0.019 to 0.170 wt% S (Fig. 8; Table 3; Additional file 4). These large variations are uncorrelated with the SiO₂ and K₂O contents of the inclusions, indicating that the differences were not the result of fractional crystallization. Furthermore, inclusions with high H₂O content (> 4 wt%) have high S content (0.148–0.170 wt%; Fig. 8; Table 3). These results indicate that the observed variation may have been caused by the exsolution of H₂O and S from the andesite melt. On the other hand, relatively small variation was observed in the Cl content of the inclusions (0.053–0.092 wt%), except for 2 inclusions in 11-4V-p8i1 and 11-1P-p1I1 that have high Cl contents (0.165 wt% and 0.121 wt%, respectively). These high Cl content inclusions are accompanied by high SiO₂ and K₂O contents (61.80 wt% SiO₂ and 2.91 wt% K₂O and 61.16 wt% SiO₂ and 2.84 wt% K₂O, respectively), suggesting that the increase in the Cl contents was due to crystallization.

Large variation is also observed in the H₂O, CO₂, and Cl contents of the dacite-rhyolite inclusions from the 2011 eruptions, which range from 1.4 to 4.8 wt% for H₂O (by SIMS), 0.002 to 0.048 wt% for CO₂, and 0.040 to 0.166 wt% for Cl (Fig. 8; Table 3; Additional file 4). These large variations are uncorrelated with the SiO₂ and K₂O contents in the same inclusions, indicating that the differences were not caused by fractional crystallization but were rather the result of degassing in the dacite-rhyolite melt. Most of the inclusions have S contents of less than 0.02 wt%, although four of the samples: 11-1P-p7i1, 11-4V-p6i1, 11-4V-p7i1, and 11-4V-p12i1 have inclusion S contents of 0.034–0.061 wt% S. These inclusions with relatively high S contents and low SiO₂ contents (67.85–70.35 wt%; Table 3; Additional file 4) suggest that the high S contents were the result of mixing high S content mafic melt with rhyolite melt.

The andesite-rhyolite inclusions in the 2018 eruptions also show large variations in terms of H₂O, S and Cl contents, which range from 0.5 to 5.0 wt% for H₂O (by SIMS), 0.003 to 0.138 wt% for S, and 0.065 to 0.166 wt% for Cl (Fig. 7; Table 3; Additional file 4). The CO₂ contents in these inclusions are less than 0.027 wt%. The maximum H₂O, CO₂, and S contents in the andesite inclusions from the 2018 eruptions (5.0 wt% H₂O, 0.021 wt% CO₂, and 0.138 wt% S) are lower than those in the andesite inclusions from the 2011 eruptions (7.0 wt% H₂O, 0.054 wt% CO₂, and 0.0170 wt% S) (Fig. 7; Table 3; Additional file 4), suggesting that the andesite melt in the 2018 magma was volatile-poor compared to the 2011 magma. The large variations in the H₂O, S, and Cl contents of the andesite-rhyolite inclusions from 2018 are roughly correlated with the SiO₂ contents of the inclusions; those with lower SiO₂ content have higher H₂O, S, and Cl contents (Fig. 7). The results indicate that these variations were likely caused when mafic melt with high H₂O, S, and Cl contents mixed with felsic melt of low H₂O, S, and Cl contents.

Physical And Chemical Conditions In The Magma From 2011–18

Temperature of the magma

Two-pyroxene thermometry (Putirka 2008) was applied to the compositions of the borders of intergrown clinopyroxene and orthopyroxene phenocrysts from the 1716–17, 2011, and 2018 eruptions. A temperature of 948 ºC was obtained for the intergrown phenocrysts from the 1716–17 eruptions and the intergrown phenocrysts from all eruptions in 2011 (2011-1 PP to 2011-9A) gave similar temperatures (908–945 ºC; Additional file 5), except 2011-8A, for which the intergrown phenocryst yielded a higher temperature estimate (1055 ± 10 ºC). 2011-2SWP, with a dacite whole-rock composition (Table 2), showed the lowest temperature of 908 ºC in the 2011 eruptives. Two-pyroxene thermometry indicated that the temperature of the 2018 eruptives (914–944 ºC) are similar to those of the 2011 eruptives. 2018-4L3, with a whole-rock dacite composition, showed the lowest temperature of the 2018 eruptives at 914 ºC, which is similar to the temperature of 2011-2SWP (908 ºC) with dacite composition.

FeTi oxide thermometry (Anderson et al. 1993) was also applied to the border compositions of intergrown magnetite and ilmenite phenocrysts from the 1716–17, 2011, and 2018 eruptions (Additional files 5 and 6). The compositions of the intergrown phenocrysts in 1716-17GP gave a temperature of 940 ºC (Additional file 5) and an oxygen fugacity (log fO₂) of -9.9 log units (2.0 log units above the FMQ buffer; Additional file 6), while the compositions of the intergrown phenocrysts from the 2011 eruptions (2011-1 PP to 2011-5VL) gave temperatures of 891–977 ºC (Additional file 5) and oxygen fugacities (log fO₂) of -12.1 to -9.7 log units (0.6–1.3 log units above the FMQ buffer; Additional file 6). The temperature estimated for the 2011 magma by FeTi oxide thermometry is similar to that obtained by two-pyroxene thermometry, except for the slightly higher temperature obtained for 2011-4VB (977 ºC) by FeTi oxide thermometry. In addition, the temperatures of 2018-1VGP, 2018-1VS, 2018-2VGP, 2018-2S, 2018-4L1, and 2018-4L2 (944–985 ºC) estimated by FeTi oxide thermometry are similar to those obtained by two-pyroxene thermometry (914–944 ºC). The oxygen fugacities of these eruptives (-10.5 to -9.9 log units and 1.2–1.4 log units above the FMQ buffer) are identical to those of the 2011 eruptives (-10.7 to -9.9 log units), except for 2011-1PP. 2018-4L3 showed the lowest temperature (868 ºC) and the lowest oxygen fugacity (-11.2 log units) of the 2018 products.

The results of two-pyroxene thermometry and FeTi oxide thermometry applied to the 1716–17, 2011, and 2018 eruptions indicate that the temperature and oxidation state of the magma chamber did not change from 1716–17 to 2018. The lowest temperature 868 ºC, which was obtained for 2018-4L3 with a dacite chemical composition (Table 2), is identical to the temperature estimated for the felsic magma end member (870 ºC for a silicic andesite magma with 62–63 wt% SiO₂) by Suzuki et al. (2013b). 2018-4L3 was therefore derived from the low-temperature silicic part of the magma chamber. 2011-2SWP, with a dacite chemical composition (Table 2), shows a relatively low temperature (908 ºC), suggesting that it originated in the low-temperature silicic part of the magma chamber and was heated by high-temperature magma before the eruption.

We then applied the olivine-liquid thermometer (Putirka 2008) to the chemical compositions of the melt and the host phenocrysts of 22 olivine-hosted melt inclusions from the 2011 products in order to estimate the temperature of the mafic end member magma that included olivine phenocrysts (Table 3; Additional file 4). Temperatures of 940–1041°C (average 990°C with standard deviation of ± 25°C) were obtained, which is similar or slightly lower than the estimate for basaltic andesite magma obtained by Suzuki et al. (2013b). We also applied plagioclase-liquid thermometers (Putirka 2008) to the chemical compositions of the melt and the host phenocrysts of 10 plagioclase-hosted melt inclusions, clinopyroxene-liquid thermobarometers (Putirka 2008) to the chemical compositions of the melt and the host phenocrysts of 20 clinopyroxene-hosted melt inclusions, and orthopyroxene-liquid thermobarometers (Putirka 2008) to the chemical compositions of the melt and the host phenocrysts of 14 orthopyroxene-hosted melt inclusions (Table 3; Additional file 4). The temperature estimates for the plagioclase-, clinopyroxene-, and orthopyroxene-hosted melt inclusions in the 2011 products (938 ± 23 ºC, 951 ± 18 ºC, and 978 ± 36 ºC, respectively) are similar to those obtained by the two-pyroxene thermometry (912–945 ºC) and FeTi oxide thermometry (891–977 ºC). In addition, the temperature estimates for the plagioclase-, clinopyroxene-, and orthopyroxene-hosted melt inclusions in the 2018 products (973 ± 15 ºC, 955 ± 33 ºC, and 979 ± 29 ºC, respectively) are similar to those obtained by two-pyroxene thermometry (914–944 ºC) and FeTi oxide thermometry (944–985 ºC).

Magma pressure

The large variability observed in the H₂O and CO₂ contents of the andesite inclusions in the olivine phenocrysts and the dacite-rhyolite inclusions in the clinopyroxene- and orthopyroxene-phenocrysts from 2011 (Fig. 8) is not related to the SiO₂ content (Fig. 7), suggesting that magma degassing occurred with a decrease in the pressure. Gas saturation pressures were calculated from the H₂O and CO₂ content of the inclusions by using the melt-H₂O-CO₂ solubility models proposed by Newman and Lowenstern (2002) and Ghiorso and Gualda (2015) (Saito et al. 2018). Gas saturation pressures of 62–486 MPa (average of 232 MPa with standard deviation of ± 132 MPa, n = 22) were obtained for the andesite inclusions in the olivine phenocrysts via the solubility model of Newman and Lowenstern (2002). The pressure range corresponds to a depth of 2–20 km under a lithostatic pressure gradient. Gas saturation pressures of 127–225 MPa (average of 182 MPa with standard deviation of ± 36 MPa, n = 9) were calculated by the solubility model of Ghiorso and Gualda (2015), which is in the range of the estimates made using the solubility model of Newman and Lowenstern (2002). The pressure range corresponds to a depth of 5–9 km under a lithostatic pressure gradient. These results indicate that mafic magma ascended from a depth of at least 9 km to 2 km as the olivine crystallization occurred.

Large variability was also observed in the H₂O and CO₂ contents of the dacite-rhyolite inclusions in the clinopyroxene- and orthopyroxene-phenocrysts from the 2011 products, which is not related to the SiO₂ content (Fig. 7), suggesting that magma degassing occurred alongside a decrease in pressure. Gas saturation pressures were calculated using the H₂O and CO₂ content obtained for the inclusions by the two melt-H₂O-CO₂ solubility models. Gas saturation pressures of 36–199 MPa (average of 120 MPa with standard deviation of ± 50 MPa, n = 10) were obtained via the solubility model of Newman and Lowenstern (2002). The average pressure (120 MPa) corresponds to a depth of 5 km under a lithostatic pressure gradient. Gas saturation pressures of 126–186 MPa (average of 152 MPa with standard deviation of ± 25 MPa, n = 5) were calculated using the solubility model by Ghiorso and Gualda (2015), which is within the range obtained in the above estimates. These results indicate the presence of a felsic magma at a pressure of 120–152 MPa (corresponding to a depth of 5–6 km). The average gas saturation pressures obtained by the two solubility models (120 ± 50 MPa and 152 ± 25 MPa) are in good agreement with the pressure estimated for silicic andesite magma (125 MPa) by Suzuki et al. (2013b).

Gas saturation pressures were similarly calculated using the H₂O and CO₂ content of the andesite-rhyolite inclusions in the 2018 products and the above two melt-H₂O-CO₂ solubility models. Gas saturation pressures of 74–205 MPa (average of 118 MPa with standard deviation of ± 45 MPa, n = 7) were obtained using the solubility model by Newman and Lowenstern (2002), except for the extremely low gas saturation pressures observed in 18-1V-p2i1 and 18-1Vp4i1. The pressure range corresponds to a depth of 3–8 km under a lithostatic pressure gradient. The gas saturation pressures for the inclusions calculated by the solubility model by Ghiorso and Gualda (2015), were generally in the range 114–194 MPa (except for 18-1V-p2i1), with an average of 144 MPa and standard deviation of ± 44 MPa, n = 3, which is within the range of the above estimates. These estimates indicate that the 2018 magma was at depths of 3–8 km as the pyroxenes crystallized within the magma. The maximum estimate (8 km) is also similar to the source depth of the crustal deformation (7 km bsl) from July 2017 to March 2018 reported by JMA (2019b). The average gas saturation pressures estimated by the two solubility models (118 ± 45 MPa and 144 ± 44 MPa) are the same as those estimated from the dacite-rhyolite inclusions in the 2011 products (120 ± 50 MPa and 152 ± 25 MPa ), indicating that the felsic magma was located at the same depth from 2011 to 2018.

Gas saturation pressures were also calculated using the H₂O and CO₂ content of the dacite-rhyolite inclusions in 1716-17GP and the andesite-dacite inclusions in 1235S using the two melt-H₂O-CO₂ solubility models. Gas saturation pressures of 71–106 MPa (average of 83 MPa with standard deviation of ± 20 MPa, n = 3) were obtained for the dacite-rhyolite inclusions in 1716-17GP via the solubility model by Newman and Lowenstern (2002) while a pressure of 95 MPa was calculated for one plagioclase-hosted inclusion by the solubility model of Ghiorso and Gualda (2015). Extremely low gas saturation pressures were obtained for KP-p4i1. These results suggest that the 1716–17 magma was stored at a depth of 3–4 km. Gas saturation pressures of 89–136 MPa (average of 119 MPa with standard deviation of ± 34 MPa, n = 3) and 117–122 MPa (average of 120 MPa with standard deviation of ± 2 MPa, n = 3) were obtained for the andesite-dacite inclusions in 1235S using the solubility model of Newman and Lowenstern (2002). The results suggest that the 1235S magma was stored at a depth of 4–5 km. These results suggest that the magmas were generally stored at a depth of approximately 4 km before the eruptions.

Bulk volatile content of magma

Suzuki et al. (2013b) concluded that a mafic magma, a basaltic-andesite magma with a temperature of 1030 ºC and 9 vol% phenocrysts of olivine and plagioclase, was mixed with a felsic magma, a silicic-andesitic magma with a temperature of 870 ºC and 43 vol% phenocrysts of plagioclase, pyroxene, and FeTi-oxide before the 2011 eruptions. The andesite inclusions in the olivine phenocrysts were likely derived from the mafic magma while the dacite-rhyolite inclusions in the pyroxene and plagioclase phenocrysts were sourced in the felsic magma. The H₂O, S, and Cl contents of a melt in the 2011 mixed magma can be calculated from the volatile contents in the andesite and rhyolite inclusions, assuming that the melt was produced by the mixing of a melt from the mafic magma with a melt from the felsic magma (Table 4). The following assumptions were made; (1) inclusion 11-2SG-p9i1, with the highest H₂O content of all the andesite inclusions from 2011-2SGP, can represent the mafic melt, (2) inclusion 11-2SG-p13i1, with the highest H₂O content of all the rhyolite inclusions from 2011-2SGP, can represent the felsic melt, and (3) the mixed melt will have a SiO₂ content of 62 wt%, considering the SiO₂ content of 2011-2 SGP (Additional file 5). This calculation gave a H₂O content of 5.9 wt%, a S content of 0.107 wt%, and a Cl content of 0.066 wt% for the 2011 mixed melt. Based on the content in a phenocryst from 2011-2SGP (28 vol%), a bulk content of 4.0 wt% H₂O, 0.072 wt% S, and 0.044 wt% Cl (Table 4) was obtained. Both the andesite inclusions in the olivine phenocrysts and the dacite-rhyolite inclusions in the pyroxene and plagioclase phenocrysts of the 2011 products have lower CO₂/H₂O mass ratios than those of the volcanic gas (Fig. 8). The disagreement in the CO₂/H₂O mass ratios of inclusions and volcanic gas is most likely due to the super-saturation of CO₂ at the time of inclusion entrapment. This indicates that merely measuring the melt inclusions might lead to underestimation of the total volatile content in the magmas, especially less-dissolved volatile species such as CO₂ (Papale, 2005). Because the volcanic gas observed in 2011 was emitted from a mixed magma, we calculated the bulk CO₂ content of the magma from the bulk H₂O and S contents and the mass CO₂/S and CO₂/H₂O ratios of the volcanic gas (Table 4). A bulk CO₂ content of 0.70 wt% was calculated from the bulk S content (0.072 wt%) and the mass CO₂ and S ratio in the volcanic gas (9.8; A1 magma in Table 4). A bulk CO₂ content of 0.14 wt% was thus calculated using the bulk H₂O content (4.0 wt%) and the mass CO₂ and H₂O ratio in the volcanic gas (0.035; A2 magma in Table 4).

Similar to the method used for the 2011 mixed magma, the H₂O, S, and Cl contents of a melt in the 2018 magma were calculated from the volatile contents in the andesite and rhyolite inclusions, assuming that the mixed melt in the 2018 magma was composed of a mafic melt and a felsic melt (Table 4). The following assumptions were made; (1) inclusion 18-1V-p5i1, with the highest H₂O content of all the inclusions in 2018-1VGP, can represent the mafic melt, (2) two felsic melts with low or high H₂O contents were considered because the rhyolite inclusions are particularly varied in terms of H₂O content (0.5–3.7 wt%); inclusion 18-1V-p1i1, with the highest H₂O content of all the rhyolite inclusions in the 2018-1VGP represents felsic melt 1, and inclusion 18-1V-p2i1, with the lowest H₂O content among all the rhyolite inclusions in 2018-1VGP represents felsic melt 2, and (3) a mixed melt has a SiO₂ content of 65 wt%, considering the SiO₂ content of 2018-1VGP (Additional file 3). Mixing the 2018 mafic melt with 2018 felsic melt 1 gave a H₂O content of 4.5 wt%, S content of 0.094 wt%, and Cl content of 0.091 wt% for 2018 mixed melt 1. Mixing the 2018 mafic melt with 2018 felsic melt 2 gave a H₂O content of 2.1 wt%, S content of 0.082 wt%, and Cl content of 0.011 wt% for 2018 mixed melt 2. Based on the content of a phenocryst in 2018-1VGP (28 vol%), we obtained a B1 magma with a bulk content of 3.0 wt% H₂O, 0.063 wt% S, and 0.060 wt% Cl and a B2 magma with a bulk content of 2.1 wt% H₂O, 0.054 wt% S, and 0.054 wt% Cl (Table 4). Assuming super-saturation for CO₂ at the time of inclusion entrapment, we also calculated the bulk CO₂ content of the 2018 magma from the bulk S contents and the mass CO₂/S ratio in the volcanic gas. A bulk CO₂ content of 0.10 wt% was obtained from the bulk S content (0.063 wt%) and the mass ratio of CO₂ and S in the volcanic gas in October 2017 (1.6; B1 magma in Table 4). A bulk CO₂ content of 0.087 wt% was calculated from the bulk S content (0.054 wt%) and the same mass ratio of CO₂ and S in the volcanic gas (B2 magma in Table 4).

In addition, the bulk H₂O, S, and Cl contents (6.2 wt% H₂O, 0.14 wt% S, and 0.050 wt% Cl) were estimated for the 2011 mafic magma from those of the 2011 mafic melt and phenocryst content (8.9 vol%) of the magma estimated by Suzuki et al. (2013b). The CO₂ content of the 2011 mixed melt was calculated at 1.0 wt% from the bulk CO₂ content of A1 magma and its phenocryst content (28 vol%). Assuming that the CO₂ content of the 2011 felsic melt was equal to that of inclusion 11-2SG-p13i1 (0.048 wt%), which is the maximum content in the rhyolite inclusions of the 2011 products, mass balance calculation gave a CO₂ content of 1.5 wt% for the mafic melt from the 2011 mixed melt and the mixing ratio (mafic:felsic = 0.67:0.33). The bulk CO₂ content of the 2011 mafic magma (M1 magma in Table 4) was calculated at 1.4 wt% considering the phenocryst content (8.9 vol%). In the same way, in terms of the A2 magma, this calculation gave a CO₂ content of 0.28 wt% for the mafic melt in the 2011 mixed melt comprising the A2 magma and the mixing ratio, resulting in a bulk CO₂ content of 0.25 wt% for the 2011 mafic magma (M2 magma in Table 4).

We could now estimate the bulk H₂O, CO₂, S, and Cl contents of the 2011 felsic magma, assuming that the H₂O, CO₂, S, and Cl contents of inclusion 11-2SG-p13i1 represent those of the felsic melt and no super-saturation of CO₂ occurred at the time the inclusions were entrapped. The bulk H₂O, CO₂, S, and Cl contents of felsic magma depend on the phenocryst content. Assuming that the felsic magma contained no phenocryst, a bulk content of 3.7 wt% H₂O, 0.048 wt% CO₂, 0.007 wt% S, and 0.087 wt% Cl were obtained ((F1 magma in Table 4). On the other hand, if the phenocryst content was 43 vol%, as estimated for silicic-andesite magma by Suzuki et al. (2013b), a bulk content of 1.9 wt% H₂O, 0.025 wt% CO₂, 0.004 wt% S, and 0.045 wt% Cl is obtained (F2 magma in Table 4).

Chemical composition of melt in the magma

The chemical compositions of the groundmass in the 2018 products are slightly evolved as compared to the 2011 products (Fig. 4) in spite of the similarities in the whole-rock chemical composition and chemical compositions of the phenocrysts from the two masses. MELTS calculation (Gualda et al. 2012) was used to investigate the difference with the whole-rock composition of 2011-2SGP (Table 2) and a temperature of 928°C, pressures of 5–500 MPa, a FMQ + 2 buffer of fO₂, and the bulk H₂O and CO₂ contents of the A1 and A2 magmas (4.0 wt% H₂O and 0.70 wt% CO₂ or 0.14 wt% CO₂; Additional file 8). The MELTS calculations indicate the presence of crystallized plagioclase, clino- and ortho-pyroxene, and FeTi-oxide phenocrysts without any olivine phenocrysts (Additional file 8). The gas saturation pressures for the dacite-rhyolite inclusions in the 2011 products (36–199 MPa) that were obtained in the previous section suggest that the pressure in the 2011 felsic magma chamber was 50–200 MPa. Melts in the A1 and A2 magmas have SiO₂ contents ranging from 65.37–71.48 wt% with 2.28–3.64 wt% K₂O at pressures of 50–200 MPa (Fig. 4). These chemical compositions are slightly SiO₂- and K₂O-rich compared to the groundmass in the 2011 products (Fig. 4). Tomiya et al. (2013) proposed that the injection of mafic magma into the chamber occurred several weeks to several days before the 2011 subplinian eruptions. Therefore, the difference between the chemical composition of the melts in the A1 and A2 magmas and the groundmass occurred because the 2011 mixed magma had not reached chemical equilibrium due to the short time period between mixing and eruption.

The MELTS calculation was also applied to the whole-rock composition of 2018-1VGP (Table 2) with a temperature of 944°C, pressures of 5–500 MPa, a FMQ + 2 buffer of fO2, and the bulk H₂O and CO₂ contents of the B1 and B2 magmas (3.0 wt% H₂O and 0.10 wt% CO₂ for B1 magma, 2.1 wt% H₂O and 0.087 wt% CO₂ for B2 magma; Additional file 8). The MELTS calculation indicated crystallization of the plagioclase, clino- and ortho-pyroxene, and FeTi-oxide phenocrysts (Additional file 8) with melts containing SiO₂ contents of 66.31–70.76 wt% and 2.43–3.40 wt% K₂O at pressures of 50–200 MPa. These chemical compositions are similar to the results for the A1 and A2 magmas and are similar or slightly SiO₂- and K₂O-rich compared to the groundmass in the 2018 products, except for that obtained for FeO (Fig. 4). The similar chemical compositions obtained for the groundmass in the 2018 products and the melts obtained by the MELTS calculation suggest that the 2018 magma was closer to chemical equilibrium. This suggests that the mixed magma from 2011 remained in the magma chamber and that the chemical reactions between the minerals and melt produced the 2018 magma. The rarity of olivine phenocrysts and lack of existence for groundmass olivines in the 2018 products support this hypothesis because the results of the MELTS calculation indicate that olivines could not crystallize in the A1, A2, B1, and B2 magmas. The olivine phenocrysts that originated in the 2011 mafic magma, which were injected into the 2011 felsic magma, were probably dissolved in the mixed magma over time. The increase in the fractions of plagioclase phenocrysts with wide rims of > 0.1 mm, and clinopyroxene and orthopyroxene phenocrysts with wide rims of > 0.05 mm from the 2011 magma to the 2018 magma (mentioned previously), also supports this hypothesis.

In order to evaluate this hypothesis, compositional profiles from the core to the rim of 13 orthopyroxene phenocrysts from 2011-1PP to 2011-7VL and 12 orthopyroxene phenocrysts from 2018-1VGP and 2018-1VS were measured, all of which show rims with reverse zoning under EPMA (Additional file 2). The compositional profiles of the orthopyroxene phenocrysts in the 2011 products show a large change in the Mg# at a distance of 0.02–0.03 mm from the rims, with widths of less than 0.005 mm. The Mg# profiles were calculated via Mg-Fe diffusion (Saunders et al. 2012) using a residence time of 1, 10, and 100 years at a temperature of 928 ºC. The observed Mg# profiles were similar to those obtained for 1 year using the calculation (Additional file 2). On the other hand, the compositional profiles of the orthopyroxene phenocrysts in the 2018 products show a large change in the Mg# at a distance of 0.04–0.10 mm from the rims. The widths of the Mg# range from 0.010 to 0.028 mm, which is larger than those obtained for the 2011 products, suggesting that Mg and Fe diffusion continued over this time period. The observed Mg# profiles appear to be closer to those obtained for 10 years by calculation (Additional file 2). Therefore, the observed compositional profiles of the core to rim of orthopyroxene phenocrysts in the 2011 and 2018 products are consistent with the above hypothesis, although more detailed analysis is needed in the future.

Bubble volume and magma density

The density contrast between magma and crust can control the ascent of magma; magma can become trapped within a magma chamber if the density contrast becomes negligible (e.g., Walker 1989). In addition, the bubble volume of magma can control its eruption style; magmas with high bubble volume cause explosive eruptions, while those with poor in bubble volume lead to effusive eruptions. We therefore need to obtain the bubble volume and the density of the magma to investigate the magma ascent and eruption processes. Both bubble volume and magma density are highly dependent on the bulk volatile content of a magma. Assuming that the gas bubbles that formed by the exsolution of H₂O and CO₂ from the melt did not separate from the magma during its ascent, the bubble volume and densities of the mafic and felsic magmas of the 2011 eruptions, the 2011 mixed magma, and the 2018 magma at different depths were calculated using the bulk H₂O and CO₂ contents of the magma (Saito et al., 2018). Details of the calculation methods are given in Additional file 8 and Additional file 4 of Saito et al. (2018).

The conditions used for calculation of the 2011 mafic magmas (M1 and M2 magmas) were as follows; whole-rock composition and temperature of 1030°C of basaltic-andesite magma estimated by Suzuki et al. (2013b), an oxygen fugacity controlled by the NNO buffer, and the bulk H₂O and CO₂ contents of the M1 and M2 magmas (Table 4). The major-element compositions of the mafic melts at pressures of 50–500 MPa are similar to those of the andesite inclusions in the olivine phenocryst (Fig. 6). The H₂O and CO₂ contents of the mafic melts at pressures of 50–500 MPa are also similar to the distribution of the andesite inclusions (Figs. 7c, 7d and 8a). These similarities indicate that the calculation results are close to the actual conditions. The calculation indicates that the mafic magmas (M1 and M2) had bubble volumes of 12.7–19.7 vol% and densities of 1987–2144 kg m^− 3 at a pressure of 200 MPa, indicating that mafic magma contained abundant bubbles and was of low density before it was injected into the felsic magma.

The conditions used for calculation of the 2011 felsic magmas (F1 and F2 magmas) were; a whole-rock composition of 2011-2SWP, the temperature of 870°C for silicic-andesite magma estimated by Suzuki et al. (2013b), an oxygen fugacity controlled by 2 log units above the FMQ buffer, and the bulk H₂O and CO₂ contents of the F1 and F2 magmas (Table 4). The major-element compositions of the felsic melts at pressures of 50–200 MPa are similar to those of the dacite-rhyolite inclusions in the 2011 products (Fig. 6). The H₂O and CO₂ contents of the felsic melts at pressures of 50–200 MPa are also similar to the distribution in the dacite-rhyolite inclusions (Figs. 7c, 7d and 8a). The calculation indicates that the felsic magmas (F1 and F2) had bubble volumes of < 8.4 vol% and densities of 2264–2496 kg m^− 3 at a pressure of 125 MPa (Additional file 8), which is the pressure of the felsic magma chamber estimated using the gas saturation pressures for dacite-rhyolite inclusions in the 2011 products. Because the mafic magmas had lower densities of 1717–1835 kg m^− 3 than the felsic magma at the same pressure, the density contrast may have caused the two magmas to mingle and mix.

These calculations also provide information about the bubble volume and density of the 2011 mixed magma. Assuming that the chemical equilibrium obtained for the mixed magma is correct, the 2011 mixed magma (A1 and A2 magmas) had a bubble volume of 3.2–9.2 vol% and a density of 2280–2401 kg m^− 3 at a pressure of 200 MPa, and 48.1–50.4 vol% and 1348–1402 kg m^− 3 at a pressure of 50 MPa. However, the difference between the chemical composition of the groundmass and that obtained by the MELTS calculation for the 2011 products reveals that the mixed magma had not attained chemical equilibrium upon eruption. Consequently, the above estimation for the A1 and A2 magmas might be an extreme example. Therefore, we calculated the bubble volume and density of the mixed magmas M1 (or M2) and F1 (or F2) by using the mass ratio for each magma in the mixing process (M1 or M2 : F1 or F2 = 0.65:0.35; Additional file 9). This case defines the other extreme in which the two magmas simply “mingled” without any chemical reaction. The bubble volume and density of the mixed M1 + F1 magma are 13.8–57.1 vol% and 1162–2118 kg m^− 3 at pressure ranges of 50–200 MPa while those of the M2 + F2 magma are 13.8–53.5 vol% and 1162–2147 kg m^− 3, respectively, at the same pressure ranges (Fig. 9; Additional file 9). These calculations indicate that the 2011 mixed magma had a high-volume content of approximately 50 vol% at a low pressure of 50 MPa, whether chemical equilibrium was attained in the magma or not.

We also calculated the bubble volume and density of the 2018 magma, assuming that chemical equilibrium was attained in this magma. The 2018 magma (B1 and B2 magmas) had a bubble volume of 0.5–1.1 vol% and a density of 2483–2549 kg m^− 3 at a pressure of 200 MPa, and 24.4–37.4 vol% and 1657–1983 kg m^− 3 at 50 MPa. The bubble volume of the 2018 magma (24.4–37.4 vol%) at a pressure of 50 MPa is smaller than those calculated for the M1 + F1 and M2 + F2 magmas (53.5–57.1 vol%) and the A1 and A2 magmas (48.1–50.4 vol%) at the same pressure (Fig. 9). The lack of subplinian eruptions in 2018 is therefore considered to have been because of the low bubble volume in the shallow part of the magma at this time.

Degassed-magma Volume

The mass of degassed magma during volcanic activity is one of the important factors used in investigating the degassing process. We calculated the mass of degassed magma based on the estimated bulk volatile content of the 2011 and 2018 magmas, the measured SO₂ flux, and the chemical composition of the magmatic gas emitted from the summit crater, using the equation:

M_VE = ( C_MM – C_DM ) × M_DM,

where M_DM is the mass of degassed magma (kg), C_MM is the bulk volatile content of the mixed magma (kg kg^− 1), C_DM is the bulk volatile content of the degassed magma (kg kg^− 1), and M_VE is the mass of the volatile material emitted from the summit crater (kg). The bulk volatile content of 0.072 wt% S and 0.044 wt% Cl (Table 4) that was obtained for the A1 (or A2) magma (C_MM) was used to calculate the mass of degassed magma in 2011–12. The volatile content of degassed magma (C_DM) was estimated from the S and Cl contents of the groundmass in the 2011 products (Additional file 3), assuming phenocryst contents of 28 vol% and no S or Cl contents in the phenocrysts. The mass of emitted sulfur was calculated from the SO₂ flux and the mole ratio of SO₂/H₂S (Additional file 10). The mole ratio was 8 on 15 March 2011, before decreasing to 0.8–3.3 after April (Shinohara 2013). Mole ratios of 8 and 2 were therefore assumed for these periods, respectively. The H₂O/SO₂ and CO₂/SO₂ ratios were assumed to be 560 and 8 (Shinohara 2013) and a Cl/S ratio of 0.235 was based on the ratio in the ash-leachate (Vinet N Person. Communication). The mass of degassed magma (kg) was then converted to the degassed-magma volume (x10⁶ m³) for comparison with the erupted-magma volume and crust deformation, assuming a magma density of 2500 kg m^− 3. The mass of the degassed magma in 2017–18 were also calculated for two cases in which the magma had (1) the bulk S content of the B1 magma (0.063 wt% S) and (2) the bulk S content of the B2 magma (0.054 wt% S) before degassing (Table 4). The bulk S content of the magma after degassing was calculated from the minimum S content of the groundmass (0.002 wt% of 2018-4L1 in Additional file 3) and the phenocryst content (28 vol%) of the 2018 products. The mole ratio of SO₂ and H₂S in the volcanic gas was assumed to be 2.7 based on observation of the volcanic gas (Additional file 10). A detailed description of the calculation used for the degassed magmas is given in Table 5 and Additional file 10. If the average, minimum, and maximum SO₂ flux values are observed over a single day, the maximum value is at most two times larger and the minimum value 0.3 times larger than the average (Additional file 10). This indicates that an error of 30–200% for the degassed-magma volume is unavoidable.

The degassed-magma volume in 2011–12 is summarized for four periods, depending on the eruption style (Table 5). The degassed-magma volume on 26–27 January (4–5 × 10⁶ m³) is smaller than the volume of magma that was erupted (7–11 × 10⁶ m³ ; Maeno et al. 2014) during the subplinian eruptions on 26–27 January. However, the estimation that was used to obtain the volume of degassed magma is likely to have been underestimated, because no SO₂ flux measurement was made during the most active periods of eruption. The volume of degassed magma on 28 January–1 February (24–32 × 10⁶ m³) was 1.6–2.1 times larger than that of erupted magma in the same period (15 × 10⁶ m³; Kozono et al. 2013; Nakada et al. 2013). Considering the error in the SO₂ flux measurement mentioned above, the degassed-magma volume is likely consistent with erupted-magma volume. This indicates that the SO₂ gas emission on 28 January–1 February could be explained by degassing of the lava effused in the summit crater. On the other hand, the degassed-magma volume on 2 February–7 September (13–18 × 10⁶ m³) is more than 65 times greater than that of the erupted magma in the same period (less than 0.2 × 10⁶ m³; Nishiki et al. 2013). Furthermore, during the period 8 September 2011–26 September 2012, approximately 2–3 × 10⁶ m³ of magma was degassed although none was erupted. The excess degassing during the period 2 February 2011–26 September 2012 indicates that the magmatic gas was derived from the non-erupted magma located in a deeper part of the chamber. Because the magma chamber is located at a depth of approximately 5 km (Suzuki et al. 2013b), such degassing may be due to magma convection in a conduit (Ohwada et al. 2013; Shinohara 2013). This excess degassing may have decreased the bulk volatile content of the 2011 magma in the chamber over time.

GPS observation indicates that deflation reaching 14 × 10⁶ m³ occurred between 25 January and 1 February, which could be explained by the discharge of magma from the chamber to the surface (Nakada et al. 2013). The crustal deflation changed to inflation on 25 February, and this continued until December 2011. GPS observation indicates that the total volume of the inflation was 11 × 10⁶ m³ (Table 5). The inflation could be explained by continuous replenishment of the magma chamber from a deep source (Nakada et al. 2013; Nakao et al. 2013; Suzuki et al. 2013b). The mafic-magma proportion of the degassed magma on 2 February 2011– 26 September 2012 could be 10–14 × 10⁶ m³ using a mixing ratio of 0.65 for mafic to mixed magma (Suzuki et al. 2013b) and a degassed-magma volume of 15–21 × 10⁶ m³ (Table 5). This estimate is similar to the inflation of the chamber observed by GPS, suggesting that mafic magma was injected to the chamber from deeper in the volcano during February–December 2011, causing inflation of the crust, gas emission, and small eruptions.

The degassed-magma volume in 2017–18 is summarized for five periods, depending on the kind of eruption that occurred (Table 5). The degassed-magma volume estimated for 11 October 2017–28 February 2018 (14 × 10⁶ m³) is much larger than the erupted-magma during this period. The degassed-magma volume for 1 March–9 March (29 × 10⁶ m³) is 1.9–2.2 times greater than the magma erupted in the same period (15 × 10⁶ m³; Chiba et al. 2018). Considering the error in the SO₂ flux measurement, the degassed-magma volume on 1 March–9 March (29 × 10⁶ m³) is similar to the volume of erupted magma, indicating that all the SO₂ gas emitted in the summit crater was derived from the lava erupted during this period. On the other hand, the degassed-magma volume for 10 March–27 June (5 × 10⁶ m³) is 17–20 times larger than erupted-magma volume in the same period (0.3 × 10⁶ m³; Oikawa et al. 2018). Furthermore, 3 × 10⁶ m³ of magma was degassed from 28 June–13 October 2018, although no magma was erupted. Similar to 2011–12, a large degassed magma volume that was greater than the erupted magma volume was observed in a period of vulcanian eruptions occurred alongside ash and gas emissions, suggesting that the degassing occurred by convection in a conduit (Ohwada et al. 2013; Shinohara 2013). GPS observation indicates that the deflation of the crust after 1 March reverted to inflation on 10 March and that the inflation continued until January 2019, with a total volume of 10 × 10⁶ m³ (Table 5). As discussed for 2011–12, the magma chamber must have undergone replenishment from a deep source during this period, causing inflation of the crust, gas emission, and small eruptions.

Ascent And Degassing Processes Of The Magma

Based on the above estimation of the physical and chemical conditions in the magmas and the degassed-magmas, the magma ascent- and degassing processes were estimated for the 2011 and 2017–18 Shinmoedake eruptions. The subplinian eruptions in January 2011 were the result of mixed mafic- and felsic magma. The mafic magma had basaltic-andesite composition while the felsic magma had silicic-andesite composition, and the proportion of basaltic-andesitic magma in the mixed magma ranged from 0.6 to 0.7 (Suzuki et al. 2013b). The felsic magma must have been stored at a pressure of 120–152 MPa (equivalent to a depth of 5–6 km), considering the gas saturation pressures for the dacite-rhyolite inclusions in the 2011 products that were obtained using the two solubility models (120 ± 50 MPa and 152 ± 25 MPa). The felsic magma had relatively lower bulk volatile content (1.9–3.7 wt% H₂O, 0.025–0.048 wt% CO₂, 0.004–0.007 wt% S, and 0.045–0.087 wt%; Table 4), bubble volumes of < 9.3 vol%, and a bulk density of 2241–2496 kg m^− 3 at 125 MPa (Fig. 9). The density structure beneath Shinmoedake is estimated to be 2000–2500 kg m^− 3 at depths of less than 0 km bsl and approximately 2500 kg m^− 3 at depths of 0–1 km bsl (JMA, 2013). Assuming that the density of the crust is 2500 kg m^− 3 at depths of more than 0 km bsl, the felsic magma may not be able to ascend to shallower depths because of the small contrast between the density of the felsic magma and the crust. Considering that the major-element composition of 1716–17WP is similar to that of 2011-2SWP (Table 2; Fig. 2), which was likely derived from the felsic magma, the felsic magma can be assumed to have remained in the chamber following the 1716–17 eruptions.

The mafic magma with bulk volatile content of 6.2 wt% H₂O and 0.25–1.4 wt% of CO₂ may have ascended from a depth of 19 km (at a pressure of 486 MPa) based on the gas saturation pressures of the andesite inclusions in the olivines from the 2011 products ((i) in Fig. 9). The mafic magma had bubble volumes of 27.6–33.0 vol% and densities of 1717–1835 kg m^− 3 at 125 MPa (M1 and M2 in Additional file 8). The lower density of the mafic magma than the felsic magma (2241–2496 kg m^− 3; Fig. 9: Additional file 8) suggests that it may have injected into the felsic magma chamber, promoting mixing of the two magmas ((ii) in Fig. 9). Comparison between the results of the MELTS calculation and the chemical composition of the groundmass in the 2011 products indicates that the mixed magma did not attain chemical equilibrium before eruption. Densities of 1875–2022 kg m^− 3 for the M1 + F1 and M2 + F2 magmas and 2050–2147 kg m^− 3 for the A1 and A2 magmas were obtained at a pressure of 125 MPa. Both densities are lower than that of the surrounding felsic magma at the same pressure (2241–2496 kg m^− 3; Fig. 9), suggesting that the mixed magma may have continued ascending ((iii) in Fig. 9). Furthermore, the ascending mixed magma had bubble volumes of 51.3–57.1 vol% (M1 + F1 and M2 + F2 magmas; density of 1186–1222 kg m-³) or 48.1–50.4 vol% (A1 and A2 magmas; density of 1348–1402 kg m-³) at a pressure of 50 MPa. Such bubble-rich magma could have caused the subplinian eruptions. We should mention here that the above densities and the bubble volumes were calculated with the assumption that gas bubbles that formed from the exsolution of H₂O and CO₂ from the melt do not separate from the magma during its ascent. If the bubbles separated from the magma and rose up to an upper part of the chamber, the magma erupted at the subplinian eruptions could have a higher bubble volume than the above estimates. The chemical composition of the groundmass in March–August 2011 seems more evolved than that in January–February, and becomes close to that predicted by MELTS for a melt in the mixed magma, indicating that the chemical reaction partially proceeded in the magma during the period March–August. On the other hand, the volume of degassed-magma from 2 February 2011 to September 2012 was more than 75 times greater than that of the eruptive products during the same period. This suggests that the magma in the chamber underwent degassing due to the convection of magma in a conduit. This excess degassing may have decreased the bulk volatile content in the 2011 magma.

The similar whole-rock composition and variations in the plagioclase-, clinopyroxene-, and orthopyroxene-phenocryst cores of the 2011 and 2018 magmas indicates that the 2011 mixed magma remained following eruption. Similar gas saturation pressures (with averages of 118 ± 45 MPa and 144 ± 44 MPa by the two solubility models) were obtained from the H₂O and CO₂ contents of the andesite-rhyolite inclusions in the 2018 products, indicating that the depth of the magma chamber did not change from 2011 to 2018. However, the rarity of the olivine phenocrysts, wide rims of plagioclase, clinopyroxene and orthopyroxene phenocrysts in the 2018 products and the comparison between the MELTS calculation and the groundmass in the 2018 products indicates that that chemical reaction of the magma proceeded to reach equilibrium. The estimated bulk volatile contents for the 2018 magma (3.0 wt% H₂O and 0.10 wt% CO₂ for B1 magma and 2.1 wt% H₂O and 0.087 wt% CO₂ for B2 magma; Table 4) are lower than those obtained for the 2011 magma. The decrease in the volatile content may have been caused by the excess degassing of the 2011 magma in the magma chamber following the eruptions. The density of the 2018 magma was 2374–2498 kg m^− 3 at a pressure of 125 MPa, which is similar to that of the surrounding felsic magma (2241–2496 kg m^− 3; Fig. 9). Considering the small contrast between the density of the 2018 magma and the surrounding felsic magma, the 2018 magma may not have ascended via buoyancy. Instead, the injection of new magma into the bottom of the 2018 magma chamber, inferred from the crust inflation from June 2017 to March 2018 (JMA 2019b), may have forced the 2018 magma into shallower depths, resulting in the 2018 eruptions. We speculate that the new magma had relatively low bulk volatile content because if the magma had high volatile content like the 2011 mafic magma, it could have turned over to mix with the 2018 magma before eruption. The lower bubble volume of the 2018 magma (24.4–37.4 vol% at 50 MPa) is likely to have led to effusive rather than explosive eruption. After the lava effusion in March 2018, the volume of degassed magma present was greater than the volume of erupted magma ( more than 27-fold), suggesting the degassing of magma by convection in a conduit. GPS observation indicates a total volume of 10 × 10⁶ m³ for the inflation that occurred from 10 March to 31 January 2019 (Table 5), suggesting that new magma injected into the chamber from greater depths caused the degassing activity.

(a) Petrological observation and melt-inclusion analysis of products from the 2011 and 2018 eruptions at Shinmoedake in the Kirishima volcanic group were conducted to investigate the magma ascent and degassing processes. Whole-rock composition and chemical compositions of plagioclase-, clinopyroxene-, and orthopyroxene-phenocryst cores from the 2018 products were found similar to those in the 2011 products, suggesting that the 2018 magma was a remnant of the 2011 magma. Comparison of the mode composition and chemical composition of the groundmass in the 2018 products with the results obtained using MELTS indicates that the magma approached chemical equilibrium between 2011 and 2018, suggesting that the changes observed could be explained by the magma approaching chemical equilibrium.

(b) Melt-inclusion analysis reveals that the inclusions obtained for the 2011 eruptions can be divided into two types; andesite inclusions in olivine phenocrysts and dacite-rhyolite inclusions in plagioclase, clinopyroxene, and orthopyroxene phenocrysts. Large variations were observed in the volatile content of the andesite and dacite-rhyolite inclusions, which are uncorrelated with the SiO₂ or K₂O contents in the inclusions, indicating degassing of the andesite and dacite-rhyolite melt. The inclusions from the 2018 eruptions have andesite to rhyolite compositions with no gaps in the chemical compositions. Large variations are also observed in the H₂O, S, and Cl contents of the andesite-rhyolite inclusions from the 2018 eruptions, which are roughly correlated with the SiO₂ contents of the inclusions, indicating that mafic melt with high H₂O, S, and Cl contents was mixed with felsic melt of low H₂O, S, and Cl contents.

(c) Two-pyroxene thermometry (Putirka 2008) was applied to the compositions of the borders of intergrown clinopyroxene and orthopyroxene phenocrysts from the 1716–17, 2011, and 2018 eruptions, with results indicating that the temperature of the 2018 magma was similar to those of both the 2011 and the 1716–17 magmas. The application of FeTi oxide thermometry to the compositions at the borders of the intergrown magnetite and ilmenite phenocrysts from the 2018 products gave temperature and oxygen fugacity estimates similar to those obtained for the 2011 and 1716–17 magmas. These results indicate that the temperature and oxidation state of the magma chamber did not change from 1716–17 to 2018.

(d) Gas saturation pressures of the andesite inclusions in the 2011 products calculated using the H₂O and CO₂ content range from 62 to 486 MPa (with averages of 232±132 MPa and 182±36 MPa from the two solubility models). The gas saturation pressures of the dacite-rhyolite inclusions in the 2011 products range from 62 to 199 MPa and the averages obtained using the two solubility models (120±50 MPa and 152±25 MPa) are in good agreement with an estimate (125 MPa) for silicic-andesite magma made by Suzuki et al. (2013b). The gas saturation pressures of the andesite-rhyolite inclusions in the 2018 products range from 74 to 205 MPa (averages of 118±45 MPa and 144±44 MPa by the two solubility models), which is similar to the estimation made for the 2011 felsic magma.

(e) The bulk volatile contents of the magmas associated with the 2011 and 2018 eruptions were estimated based on the volatile content of the melt inclusions and the chemical composition of volcanic gas; the 2011 mixed magma showed 4 wt% H₂O and 0.70–0.14 wt% CO₂, 2011 mafic magma showed 6.2 wt% H₂O and 1.4–0.25 wt% CO₂, 2011 felsic magma showed 3.7-1.9 wt% H₂O and 0.048–0.025 wt% CO₂, and 2018 magma showed 3.0–2.1 wt% H₂O and 0.10–0.087 wt% CO₂.

(f) The bubble volume and density of the mafic and felsic magmas from the 2011 eruptions, 2011 mixed magma, and 2018 magma were calculated at pressure rages of 5–500 MPa using the bulk volatile content and whole-rock composition of the magmas. The calculation shows that the mafic magmas were of lower density than the felsic magma, and that the mafic magma rose into the felsic magma within the magma chamber, promoting mixing. The 2011 magma had a particularly high bubble volume just before eruption (about 50 vol% at a pressure of 50 MPa), which caused the subplinian eruption. The 2018 magma had a lower bubble volume (24.4–37.4 vol% at a pressure of 50 MPa) and thus failed to cause subplinian eruptions.

(g) The mass of degassed magma present was calculated based on the estimated volatile content of the magma, the measured SO₂ flux, and the chemical composition of the magmatic gas emitted from the summit crater. The degassed-magma volume from 2 February 2011 to September 2012 was more than 75 times as large as that of the eruptive products during the same period, suggesting that the degassing of magma in the chamber was due to magma convection in a conduit. This excess degassing decreased the bulk volatile content of the 2011 magma.

(h) The magma ascent and degassing processes of the 2011 and 2017–18 Shinmoedake eruptions are also proposed. Mafic magma with a high volatile content of 6.2 wt% H₂O and 1.4–0.25 wt% CO₂ ascended from at a depth of 19 km before the 2011 eruptions. The mafic magma, with high bubble volume and low density, was injected into and mixed with the felsic magma located at a depth of 5–6 km. The bubble volume of the mixed magma increased to approximately 50 vol% at a depth of 2 km, followed by fragmentation, and resulted in subplinian eruptions on 26–27 January 2011. Degassing of the magma in the chamber due to magma convection in a conduit started in March 2011, together with small eruptions. This excess degassing may have decreased the bulk volatile content of the magma. The mixed magma then proceeded towards chemical equilibrium. The 2018 magma, with a low volatile content of 2.1–3.0 wt% H₂O and 0.087–0.13 wt% CO₂, ascended again from the magma chamber, probably due to the injection of new magma into the bottom of the chamber, which likely started in 2017. This magma, with relatively low bubble content, could not cause a subplinian eruption and erupted instead via effusion in March 2018.

AIST: Advanced Industrial Science and Technology

BSE: Backscattered electron

bsl: below sea level

CD: Volatile content of the degassed magma

CM: Volatile content of the magma

DRE: Dense rock equivalent

EPMA: Electron probe micro analyzer

GSJ: Geological Survey of Japan

MD: Mass of degassed magma

MV: Mass of volatile material emitted from the crater

SEM: Scanning Electron Microscopy

SIMS: Secondary Ion Mass Spectrometry

XRF: X-ray fluorescence analysis

Ethics approval and consent to participate

Not applicable.

Consent for publication

Not applicable.

Availability of data and materials

The datasets used and/or analysed during the current study are available from the corresponding author on reasonable request.

Competing interests

The authors declare no competing interests.

Funding

Not applicable.

Authors' contributions

GS carried out the SEM, EPMA, and SIMS analyses of the eruptive products and the Melts calculation and drafted the manuscript.

TO perform the geological survey of the 2018 eruption at Shinmoedake, Kirishima volcano group and collected the samples.

OI carried out the whole-rock analyses of the eruptive products.

All authors read and approved the final manuscript.

Acknowledgements

We thank Dr. Nicolas Vinet for providing data on the Cl/S ratio of ash-leachate from the Shinmoedake eruption. We thank Drs. Nobuo Geshi, Yoshihiro Ishizuka, Ryuta Furukawa, and Masayuki Oishi for providing the rock samples used in this research and their helpful comments. We thank Dr. Isoji Miyagi for helping SIMS measurements. We would like to thank Editage (www.editage.jp) for English language editing.

Authors' information

All authors are researchers at the Geological Survey of Japan, AIST.

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Magma ascent and degassing processes of the 2011 and 2017-18 eruptions of Shinmoedake in Kirishima volcano group, Japan, based on petrological characteristics and volatile content of magmas

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Abstract

Figures

Introduction

Overview Of The Kirishima Volcano Group And The 2011 And 2017–18 Shinmoedake Eruptions

Sampling And Analytical Methods

Samples

Analytical methods

Results

Whole-rock and mode compositions

Chemical compositions of minerals and groundmasses

Melt inclusions

Physical And Chemical Conditions In The Magma From 2011–18

Temperature of the magma

Magma pressure

Bulk volatile content of magma

Chemical composition of melt in the magma

Bubble volume and magma density

Degassed-magma Volume

Ascent And Degassing Processes Of The Magma

Summary And Conclusions

Abbreviations

Declarations

References

Tables

Supplementary Files

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