# Fluid Fluxing and Accumulation Drive Decadal and Short-Lived Explosive Basaltic Andesite Eruptions Preceded by Limited Volcanic Unrest

S Utami, F Costa, Ph Lesage, Patrick Allard, H Humaida

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S Utami, F Costa, Ph Lesage, Patrick Allard, H Humaida. Fluid Fluxing and Accumulation Drive Decadal and Short-Lived Explosive Basaltic Andesite Eruptions Preceded by Limited Volcanic Unrest. Journal of Petrology, 2021, 62 (11), 10.1093/petrology/egab086. hal-03893281

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# Fluid Fluxing and Accumulation Drive Decadal and Short-Lived Explosive Basaltic Andesite Eruptions Preceded by Limited Volcanic Unrest

S. B. Utami¹,*, F. Costa¹,², Ph. Lesage³, P. Allard⁴, and H. Humaida⁵

¹Asian School of the Environment, Nanyang Technological University 50 Nanyang Avenue, N2-01B-27 Singapore, 639798; ²Earth Observatory of Singapore, Nanyang Technological University 50 Nanyang Avenue, N2-01A-15 Singapore, 639798; ³ISTerre, Université Savoie Mont Blanc-Université Grenoble Alpes-Université, Gustave Eiffel, CNRS, IRD, 38000, Grenoble, France; ⁴Institut de Physique du Globe de Paris, Université de Paris, CNRS UMR71541, Rue Jussieu, 75005, Paris, France; and ⁵Jl. Cendana No.15, Semaki, Kec. Umbulharjo, Kota Yogyakarta, Daerah Istimewa Yogyakarta 55166, Indonesia

*Corresponding author: E-mail: sri.budhi.utami@gmail.com

Received 26 February 2021; Revised 19 August 2021; Accepted 6 October 2021

## Abstract

Some volcanoes are known for repeatedly producing explosive but short-lived eruptions (< half a day) every decade or so. These eruptions are often preceded by limited unrest signals and short run-up times to eruption (a few hours to months), and thus they are difficult to anticipate. Some well-documented examples are the 1990 and 2014 eruptions of Kelud volcano in Indonesia, or the 2015 Calbuco eruption in Chile. Here we interrogate the rock record and obtain insights into the processes and pre-eruptive conditions that led to the 1990 Kelud eruption, which we integrate with monitoring data (seismicity, lake temperature and hydro-acoustics, sulfur emissions) towards a conceptual model for this type of events. Mineral-melt geothermobarometers indicate that the basaltic andesite magma carried a crystal cargo from as deep as 15–19 km, and reached volatile saturation at 4–9 km with 2–4 wt.% water in the melt. The textures and compositional zoning of orthopyroxene and plagioclase do not support intrusion of more primitive magma as the driver for the eruption, and we instead propose that pre-eruptive fluid accumulation and high-temperature fluid fluxing from depth (likely dominated by CO₂) played a major role in priming the eruption to occur. Such pre-eruptive gas accumulation is also supported by mass balance calculation of the emitted excess SO₂ gas. Mg-Fe diffusion profiles in reversely zoned pyroxenes constrain timescales of weeks to months before eruption for fluid addition to the reservoir, and such events may be recorded in the monitoring signals, especially in the change of hydroacoustics and water lake temperature, and probably in the seismic swarms. We propose that fluid exsolution and accumulation in the shallow reservoir plays a crucial role in modulating and triggering short-lived explosive eruptions with brief unrest at Kelud and probably other volcanoes worldwide.

**Key words:** volcano monitoring; magma storage conditions; diffusion timescales; Kelud; Kelut

## INTRODUCTION

Explosive eruptions can generate lethal hazards, and thus special attention is needed to monitor and properly interpret the signs of unrest in anticipation of the eruption (Jolly and de la Cruz, 2015). Yet for many eruptions, dedicated volcano monitoring efforts are not enough to anticipate their explosivity, when the eruption will occur, and how it will progress with time. This is partly because the relationship between the magmatic processes controlling the erup-

tions and associated unrest patterns are poorly understood. Explosive eruptions can often be preceded by a wide range of unrest patterns, from abrupt to gradual to progressive changes in precursory signals over times les varying from a few days to years before eruption, for example in Merapi 2006 vs. 2010 (Indonesia) (Ratdomopurbo et al., 2013; Surono et al., 2012). Some explosive eruptions are preceded by exceptionally short run-up times (a few hours to a few weeks), which makes it extremely challenging for civil protection authorities
<figure><img src="image_2.png" /><figcaption>**Fig. 1.** Map of 1990–2014 volcanic deposits at Kelud volcano showing the distribution of the 1990 deposits from tephra fall and pyroclastic density currents (PDC) as documented by Bourdier *et al.* (1997). These are overlain by pyroclastic density flows from the 2014 eruption (defined as Stage 1 and 3). Regions of flattened trees were digitized from Maeno *et al.* (2017). The Lirang seismic station was the single station with interpretable seismic data in 1990 (Lesage & Surono, 1995). Inset map shows the location of Kelud in the island of Java, Indonesia.</figcaption></figure>

to advise communities, given the limited time to activate mitigation plans.

Kelud volcano, located in East Java (Indonesia) is part of the Sunda arc volcanoes, and a good example of a volcanic system that has produced two explosive eruptions that were preceded by short unrest in 1990 and 2014. For example, before the event in 2014, volcano-tectonic earthquakes abruptly spiked starting from 3 February, dissipating until 10 February, before a final spike in seismicity occurred a couple of hours before the eruption at 22:50 on 13 February (e.g. Cassidy *et al.*, 2019; Hidayati *et al.*, 2018; Maeno *et al.*, 2016). Such a pattern is observed at other subduction zone volcanoes in the world (Passarelli and Brodsky, 2012, and references therein). Calbuco volcano in Chile, which erupted explosively in 2015, had very short run up times of seismicity a few hours before its first explosive phase that produced a volcanic cloud column >20 km tall (Arzilli *et al.*, 2019; Morgado *et al.*, 2019a, 2019b; Pardini *et al.*, 2018).

The link between magmatic processes, unrest signals and the explosive dynamics of these eruptions can be elucidated by combining various types of information: the whole-rock, mineral compositions, and textures combined with timescale indicators (e.g. Albert *et al.*, 2016; Kahl *et al.*, 2013; Rasmussen *et al.*, 2018); thermobarometric constraints and geophysical models (Chiarabba *et al.*, 2004; Patanè *et al.*, 2003; Widiyantoro *et al.*, 2018); petrological estimates of magma degassing compared to satellite measurements of volatile emissions (Caricchi *et al.*, 2014; Jay *et al.*, 2014); petrological indicators of magma volume, timescales and processes compared to unrest signals from monitoring records (Cheng and Costa, 2019; Kahl *et al.*, 2011). Such integrative approaches have been applied to investigate the processes and mechanisms that lead to different eruptions at a variety of volcanoes (e.g. Costa *et al.*, 2020), including Mt Etna (Kahl *et al.*, 2013), Eyjafjallajökull (Pankhurst *et al.*, 2018), Mt. St. Helens

(Saunders *et al.*, 2012), Ruapehu (Kilgour *et al.*, 2014), Shishaldin (Rasmussen *et al.*, 2018), Piton de la Fournaise (Albert *et al.*, 2019), or Kilauea. For many of these examples renewed mafic magma injection is proposed to be the main driving mechanism for the eruption.

Here we provide new constraints on the pre-eruptive conditions and the driving mechanism of the 1990 explosive eruption of Kelud, based on detailed investigations of its eruption deposits, minerals, and glass. We also obtain timescales for the magmatic processes which we combine with the instrumentally recorded unrest signals to propose a conceptual model for the eruption. We show that fluid accumulation in the upper parts of the reservoir during years of repose and fluxing of deep fluids that are likely CO₂-dominated may play a key role in promoting and driving the recurrent, short-lived explosive eruptions at Kelud. Such fluid accumulation and fluxing may be recognized in the monitoring signals recorded in a lake water present before the eruption. Our findings may thus be of importance towards a better understanding of future eruptions at Kelud and of analogous events other volcanoes worldwide.

## GEOLOGICAL SETTING

### Kelud volcano and lake

Kelud is a basaltic andesite stratovolcanic complex located in East Java (Indonesia) as part of the Sunda Arc subduction zone volcanoes, and has been classified as semi-plugged as there is no significant degassing during repose periods (Whelley *et al.*, 2015) (Fig. 1). The current edifice consists of two highly excavated craters (Kawah Dargo being the main one) with a westward opening, surrounded by four crypto-domes (Sumbing, Lirang, Gajah Mungkur, and Umbuk) (Wirakusumah, 1991). The location and age of the nearest
crypto-domes indicate a clockwise shift of eruptive centers over time, from the earliest crypto-dome (Gunung Lirang, 238 ky) to the current crater (Global Volcanism Program, 2013; Wirakusumah, 1991). More distal parasitic domes, such as the Gunung Umbuk show evidence for flank volcanic activity at ~39 kya (Wirakusumah, 1991). The valleys around the volcano are highly incised and steeply dipping, showing that the edifice is being actively eroded. Kelud hosts one of the few neutral pH volcanic lakes in the world, whose volume averaged ~2.5 × 10⁶ m³ prior to the 1990 eruption (Bourdier et al., 1997). Bernard and Mazot (2004) identified two hydrothermal systems feeding the pre-1990 crater lake: a deep Li-B-enriched alkali-chloride hydrothermal system, and a shallow Ca-Mg sulfate dominated hydrothermal system. The neutral pH indicates that an impermeable barrier exists somewhere between the magma reservoir and the shallow conduit of Kelud. After the devastating 1919 eruption (5160 casualties), a drainage tunnel system was constructed to drain the lake level and reduce the hazards from lahars during future eruptions. This drainage system was further deepened after the 1966 eruption (Caudron et al., 2012).

## Historical eruptions and hazards from Kelud

Kelud has erupted around 34 times since 1000 CE and all events were highly explosive except for two dome extrusions in 1920 and 2007 (Global Volcanism Program, 2013). Most eruptions generated a range of hazards, including lahars, landslides, ash fall, ballistics, and pyroclastic density currents (Global Volcanism Program, 2013). Lahars generated by the mixing of eruptive deposits with lake water have caused most of the fatalities, killing 5160 people in 1919, and up to 10 000 people in 1586 (Global Volcanism Program, 2013). Ashfall has produced widespread damage to the surrounding villages and cities (e.g. Blake et al., 2015; Bourdier et al., 1997; Pardyanto, 1990; Williams et al., 2020), with roof collapse due to ash accumulation killing 32 people in 1990. Ash also forms unconsolidated deposits that become washed down the flanks and the steep-sided valleys by heavy rains, forming cold lahars (Thouret et al., 1998). Pyroclastic density currents usually remain confined to inside the crater or proximal areas (Bourdier et al., 1997; Maeno et al., 2016). The last eruption in 2014 generated a 17-km high volcanic column, promoting ash circulation within the troposphere for several months (Kristiansen et al., 2015; Vernier et al., 2016) and causing extensive damage to infrastructure networks across multiple cities in Java (Blake et al., 2015). Despite its suddenness and widespread impact more than 160 000 people were evacuated in less than two hours, thanks to the high level of preparedness and coordination by the Centre for Volcanology and Geological Hazard Mitigation (CVGHM) (Andreastuti et al., 2016).

## The 1990 eruption

This eruption ejected a total volcanic volume of 0.13 km³ dense rock equivalent (DRE) (Bourdier et al., 1997), and released a total of ~170 kt of SO₂ into the troposphere according to satellite data (Global Volcanism Program, 2013). The unrest period leading up to the February 1990 eruption of Kelud was characterized by changes of the crater lake, hydroacoustics anomalies, and seismic swarms. According to Vandemeulebrouck et al. (2000) the unrest might have started in April 1989 with an increase in baseline of low frequency hydroacoustics signals, probably due to the onset of gas rising up and early boiling of the lake water. Four main stages can be recognized in the sequence of events that led to the 1990 eruption of Kelud (Table S1). All reported times refer to local time (GMT + 7):

1. Main unrest phase (November 1989–January 1990): this was marked by an increased number of daily earthquakes and the appearance of two seismic swarms in December 1989, and then January 1990. Lesage & Surono (1995) attributed these swarms to fracturing of a brittle plug blocking the volcano conduit. Changes were also observed in the chemistry, pH, and color of the lake (Badrudin, 1994; Kadarsetia et al., 2006). The lake bottom and surface temperature increased from 30–60°C and from 30–48°C, respectively (Pardyanto, 1990). Vandemeulebrouck et al. (2000) interpreted the increasing hydroacoustic signals in the high-frequency band as due to increased gas bubbling in a convective column, and signals in the ultrasonic frequency bands as related to pressure variations in the hydrothermal system.

2. Temporary quiescence (January–9 February 1990): this lasted about three weeks and both seismic and hydroacoustic signals decreased to their background levels.

3. Escalation (9–10 February 1990): renewed unrest suddenly started after midnight on 9 February with >400 low magnitude earthquakes recorded until 9:00 am on 10 February, the day of the eruption (Lesage & Surono, 1995). In response, the Volcanological Survey of Indonesia ordered an evacuation at 10:00 am (Pardyanto, 1990).

4. The eruption (10 February 1990): The eruption started at 11:41 am, with seven phreatomagmatic explosions in one hour. From 12:35 pm, it then rapidly developed as a sub-Plinian magmatic phase with a 10-km high column and lasted until 5:00 pm and its collapse generated pyroclastic flows (Global Volcanism Program 1990; Pardyanto 1990; Lesage & Surono 1995). Tephra of lapilli size fell on Margomulyo Volcano Observatory, about 7 km west of the crater, until 5:00 pm, and ash continued to fall during the night. After this major phase, smaller discrete explosive outbursts occurred on the 11 and 12 February that produced incandescent tephra and lightning (Global Volcanism Program, 1990). The main phase of the eruption ended around March 1990 (Global Volcanism Program, 2013), having completely emptied the crater lake and generated syn-eruptive hot lahars.

5. Heavy tropical rainfalls subsequently remobilized the fresh pyroclastic deposits to produce 33 cold lahar episodes until 28 March (Thouret et al., 1998). After the eruption, a new magmatic plug had solidified in the excavated crater and a new lake gradually reformed (Bourdier et al., 1997). According to the Volcano Survey of Indonesia, the eruption ended at the end of March 1990 (Global Volcanism Program, 1990).

# METHODOLOGY

## Sampling strategy

The studied rock samples were taken from a vertical outcrop in the southwestern part of Kelud's main crater that provides access to deposits from the 1951 to 1990 eruptions (Fig. 2). The studied outcrop was protected by dense vegetation before the 2014 eruption, and afterwards, by trees blown down by that eruption (Maeno et al., 2016).

The 1951 deposit is made of beds of ash and blocks, gently dipping downward away from the main crater, and contains both unaltered white pumices (> 64 mm), oxidized red pumices, and dark ones partially covered by soil and organic matter. Entrained lithic clasts (< 64 mm) and charcoal are minor components. A paleosol
<figure><img src="image_3.png" /><figcaption>**Fig. 2.** (a) Photograph of the main 1990 outcrop sampled for this study, overlying the previous deposits from 1966 and 1951 eruption. (b) schematic diagram of the 1990 stratigraphic units showing our sample distribution and names (adapted from Bourdier et al., 1997). All samples except for PL2 derive from the main eruption stage in the pumice flow and fall deposits, while PL2 (in bold) representing the late eruption stage is found embedded in the pumice fall deposit. Only the pumice fall (P) and pumice flow (PF) units are shown in the outcrop photograph.</figcaption></figure>

horizon separates the 1951 eruption deposit from the overlying 1966 eruption deposit. The latter deposit consists of massive to weakly bedded ash and block layers, containing bomb-sized pumice blocks and lithic clasts of comparable sizes. The 1966 pumices are light grey to white porphyritic blocks. Another paleosol horizon separates products from the 1966 and 1990 eruptions. The 1990 deposit gently dips upward away from the crater and consists of a mixture of ash and poorly to moderately-sorted grey to white pumice blocks, together with entrained charcoal and lithics. Pumice clasts range from sub-rounded lapilli to bomb-sized pumices blocks (up to 15 cm), embedded in a mixture of ash and smaller pumice fragments. Lithic ejecta are sub-rounded to sub-angular porphyritic andesites. We interpreted this lens as the *pumice flow layer* previously described by Bourdier et al. (1997). This layer is overlain by an upward dipping, clast-dominated bed containing sub-rounded to rounded lapilli- to bomb-sized pumice clasts, as well as lithics. We interpret this upper bed as the *pumice fall layer* described by Bourdier et al. (1997) in their composite stratigraphy.

We found two types of pumices within the 1990 deposits: (1) a volumetrically prevalent, light and vesicular pumice in both the pumice flow and fall deposits, and (2) a volumetrically minor, relatively dense but still vesicular pumice in the pumice fall deposit only. Combining our observations with the composite stratigraphy of Bourdier et al. (1997), we consider the prevalent pumice component as representative for the main eruption stage, whereas the minor pumice component represents the late eruption stage. Thereafter, we refer to these samples as ‘main-stage’ and ‘late-stage’. We studied six pumice samples from the main-stage and one from the late-stage. Our samples were collected from the outcrop described above during a field campaign in October 2017 and were complemented by samples previously collected by Humaida (2013) that we interpreted to be main-stage pumices. Lithic clasts are fragments of either past lava

flows excavated during the eruption or fragments of the plug that blocked the conduit (Lesage & Surono, 1995). The 1990 deposit stratigraphy closely resembles that of the subsequent 2014 eruption, compiled by Maeno et al. (2019) and Goode et al. (2018). We did not find scoriaceous pumice clasts in the 1990 deposit, but Bourdier et al. (1997) found scoria deposits in other outcrops, which were included in their composite stratigraphic column. Scoria was also produced during the 2014 eruption and lapilli-sized clasts were found around the crater floor.

## Analytical methodology

### Bulk sample characterization

Point counting was conducted on standard polished petrographic thin sections to determine the areal proportions of minerals and groundmass (> 1400 point per thin section), using a Pelcon mechanical point counting stage, and the JMicrovision software on optical images of the thin sections. Whole-rock major and trace element compositions were analyzed using X-Ray Fluorescence (XRF) with fused lithium methaborate/tetraborate discs—and inductively coupled plasma mass spectrometer (ICP-MS) at the Activation Laboratory in Ancaster, Ontario (Canada) following the approach outlined in Norrish and Hutton (1969). There was no alteration on the samples, and thus no additional steps were taken other than cutting the pumices into small 50 × 40 mm billets needed for whole rock analyses. All other preparation steps, including crushing and powdering, was done at Activation Laboratory.

### Textures, mineral and glass chemistry

Minerals and textures of interest were characterized and imaged using the JEOL JSM-7800F field emission scanning electron microscope (SEM) at the Asian School of the Environment, Nanyang
Technological University (NTU), Singapore. We used an accelerating voltage of 15 kV, 20 nA current to obtain back-scattered electron (BSE) images of various parts of the sample. Minerals and glasses were analyzed in both spot and traverse modes using a JEOL-JXA 8530F field emission EPMA at the Facility for Analysis Characterization Testing & Simulation (FACTS) at NTU. We determined the major element composition of plagioclase (Plag, n = 388), orthopyroxene (Opx, n = 1628), clinopyroxene (Cpx, n = 639), Fe-Ti oxides (magnetite - Mag, n = 53; ilmenite - Ilm, n = 35), and amphibole (Amph, n = 20). Quartz (Qz) was identified but not analyzed. Point and traverse analyses of Opx and Cpx were made with a focused beam of about 1 mm, accelerating voltage of 15 kV and a current of 20 nA. The relative errors for pyroxene analysis are 0.2–6% for Si, Al, Ca, Mg, Fe, and Mn; and 1–64% for Na, Ti, and K. For plagioclase, Mag, and Ilm analysis we used the same beam size and accelerating voltage, but with a 10 nA beam current. Relative errors for Fe-Ti mineral analysis are 0.2–7% for Fe, Ti, and Mg; 2–24% for Mn, Al and Si; and 8–99% for Cr. For Plag analysis, the relative errors are 0.4–4% for Si, Al, Na, Ca, and Fe, whereas for Mg and K, the relative errors are 1–55%. Amph was analyzed with a beam of 1 to 5 mm, a 15 kV accelerating voltage, and a 10 nA beam current. The relative errors are 0.4–6% for all major element oxide contents in Amph, except for Cr₂O₃, where they are 13–62%. The relative error for halogens in Amph are 32–56% for Cl and 19–58% for F. The EPMA standard reference materials for pyroxene and Amph analyses came from P & H (diopside - Si, Ca, Mg; corundum - Al, albite - Na, rutile - Ti, specularite - Fe, rhodonite - Mn, Chromium oxide - Cr, and orthoclase - K, fluorite - F, halite - Cl). The standard reference materials for Plag, came from P & H (albite - Na, Al, Si; wollastonite - Ca, orthoclase - K and olivine - Mg) and Astimex (olivine - Mg, Fe, plagioclase - Ca). The standard reference materials for Fe-Ti oxides came from Astimex (Almandine - Al, Si; titanium - Ti, diopside - Mg, chromium oxide - Cr; magnetite - Fe, rhodonite - Mn).

Chemical indicators used to characterize the minerals were calculated from standard structural formulae (Deer et al., 1992). Pyroxenes are characterized by Mg-number (Mg # = 100 × Mg / [Mg + Fe*] in mol, where Fe* is total iron as Fe²⁺), and their proportions of wollastonite (Wo = 100 × Ca / [Ca + Mg + Fe*] in mol), enstatite (En = 100 × Mg / [Ca + Mg + Fe*] in mol), and ferrosilite (Fs = 100 × Fe* / [Ca + Mg + Fe*] in mol). Plag composition is described with the anorthite content (An = 100 × Ca / [Ca + Na] in mol). For magnetite we computed the ülvospinel component, X<sub>usp</sub> = 3 × Ti / (Ti + Fe*), and for ilmenite, X<sub>Ilm</sub> = 2 × Ti / (Ti + Fe*), following Sauerzapf et al. (2008). For amphibole, we used the structural formulae of Leake et al. (1997) as included in Ridolfi et al. (2010).

Glass inclusions hosted in Cpx and interstitial glass were analyzed using a defocused beam of 10 mm, an accelerating voltage of 6 kV and a current of 5 nA. The following elements were analyzed: Si, Al, Ca, Na, K, Fe*, Mg, Mn, Ti, P, Cl, and S, with Na analyzed first to account for Na migration during analyses. Melt inclusion compositions were corrected for post-entrapment crystallization of the host mineral along the melt inclusion walls (e.g. Bucholz et al., 2013; Esposito et al., 2014; Gaetani & Watson, 2000). A hypothetical groundmass composition of each sample was obtained by subtracting the median major oxide compositions of the phenocrysts (Plag + Opx + Cpx + Mag), weighted to their modal abundances, from the whole-rock composition. For a few samples whose whole rock compositions was not available (samples 90–2 and 90–20), we used the average composition of all 1990 samples reported by Humaida

(2013). Such composition was used where appropriated to calculate some intensive variables of the magma.

## Geothermobarometry and magma storage conditions

### Mineral geothermobarometers

The pre-eruptive conditions of magma storage (temperature, T; pressure, P; oxygen fugacity, fO₂, and melt water content) were inferred using a combined approach of rhyolite-MELTS thermodynamic software (Gualda et al., 2012), nine mineral/melt geothermobarometers, oxybarometers, and a hygrometer, as described below. Reliable application of these geothermobarometers and volatile calculations require assessment of equilibrium. For mineral-mineral pairs, calculations were done using the core–core or rim–rim compositions of touching grains, where available. For mineral-melt pairs from the main-stage pumices, crystal core compositions were paired with the whole-rock compositions, and crystal rim compositions were paired with the measured interstitial glass composition. For late-stage samples, rim and microlite compositions were paired with the calculated groundmass composition rather than the measured interstitial glass, because we found that the latter was too evolved to be in equilibrium with the phenocrysts.

Pyroxene-melt equilibrium was tested using the Mg/Fe* distribution coefficient (K<sub>d</sub>) defined as:

$$K_{d\_crystal-melt} = \frac{X_{Fe}^{crystal} X_{Mg}^{melt}}{X_{Mg}^{crystal} X_{Fe}^{melt}} \quad (1)$$

where X is the mol fraction for the relevant mineral–melt and mineral–mineral pairs using K<sub>d</sub><sub>Cpx-melt</sub> = 0.28 ± 0.08 and K<sub>d</sub><sub>Opx-melt</sub> = 0.29 ± 0.06 (Putirka 2008). For Opx- and Cpx-melt geothermometers (equations 28a and 33 in Putirka (2008), respectively), we fixed the Fe²⁺/Fe<sub>total</sub> of the melts at the fO₂ reaction buffer constrained by Fe-Ti oxides equilibria as determined in late-stage pumices (Sauerzapf et al., 2008). We also added 3 wt.% water into the melt according to hygrometer results (see below). Varying the water content in the melt by ± 1 wt.% changes the calculated temperature by <15°C for pyroxene-melt geothermometers. The standard errors on temperatures estimated from the Cpx-melt and Opx-melt geothermometers are ± 42°C and ± 48°C (Putirka, 2008), respectively. For all pyroxene-based geothermometers, the pressure was fixed at 100 MPa. A pressure change of ± 100 MPa results in a temperature change of ± 5°C. No adjacent Opx-Cpx crystals appear to be in equilibrium using the K<sub>d</sub> criterion of Putirka (2008). Instead, we used the QUILF algorithm of Andersen et al. (1993) and we only report the temperatures for which the calculated and expected pyroxene compositions are similar within analytical uncertainty and give T errors <25°C.

For the amphibole-melt geothermobarometry, using the Mg-Fe* K<sub>d</sub><sub>Amph-melt</sub> of Putirka (2016) resulted in a very wide range of possible equilibrium melts. We thus used an additional equilibrium test based on the exchange of Si/Al (mol) (K<sub>d</sub><sub>Amph-melt</sub> = 0.96 ± 0.04; Li et al., 2021), from which we verified the Amph phenocrysts to be in equilibrium with the whole rock. We also applied the Amph geothermobarometry formulations of Ridolfi & Renzulli (2012) and Ridolfi et al. (2010) and we accepted all results that fulfilled the structural formulae of Ridolfi et al. (2010).

The fO₂ and temperature for the late-stage pumices were calculated from co-existing Mag-Ilm pairs and the oxybarometer calibrations of Sauerzapf et al. (2008), and Ghiorso & Evans (2008). We also used the Plag-melt hygrometer of Waters & Lange (2015)
to determine the pre-eruptive water content of the melt. For this, we considered the compositions of the crystal rims, and of the interstitial glass, together with the temperatures given by the Opx-melt geothermometer for the main-stage pumices, and by Fe-Ti oxides for the late-stage pumices. The calculated water content, temperature, and interstitial glass compositions were then combined with the H₂O-CO₂ saturation model of Ghiorso & Gualda (2015) in the MagmaSat application to estimate volatile saturation pressures. For some computations we considered the maximum CO₂ concentration (336 mg/kg) measured by Cassidy et al. (2016) in plagioclase-hosted melt inclusions of the 2014 products. We recognize that this may not be directly applicable to the 1990 eruption, and that is likely a minimal value for the melt, since no corrections was made for the potential CO₂ present in the bubble (e.g. Bucholz et al., 2013; Gaetani et al., 2012; Lloyd et al., 2013).

## Pyroxene diffusion modeling and calculations of timescales

The timescales of Fe-Mg re-equilibration in Opx and Cpx were obtained from our measured crystal zoning profiles and the diffusion model approaches described in Costa et al. (2008) and Flaherty et al. (2018). A total of 38 reversely zoned crystal profiles were used: 22 in the main-stage samples (Opx = 20, Cpx = 2), and 16 in the late-stage samples (Opx = 14, Cpx = 2). We used the zoning in Mg-number of rim-to-core profiles approximately perpendicular to the crystal c-axis (along *a* or *b* axis) since this is deemed to be the slowest diffusion direction and the least affected by 2D or 3D diffusion (cf. Krimer & Costa, 2017). For initial profile geometries, we used sharp step functions across two well-defined compositional plateaus that are close to the observed zoning profile shape and thus have less uncertainty (Costa & Morgan, 2010). We performed independent checks of the geometry and position of initial Mg-number distributions by comparing to Al₂O₃, which is a slower diffusing component and thus likely closer the initial profile that would be related to growth. As a boundary condition we assumed a constant composition at the crystal rim. We thus obtained a range of diffusion timescales around a best fit for each profile with the analytical uncertainty and diffusion coefficients determined following Flaherty et al. (2018).

For profiles in Opx crystals we used the Fe-Mg diffusion coefficient formulations of Dohmen et al. (2017) including the correction for composition, as well as the formulations of Ganguly & Tazzoli (1994) which give 2–3 times shorter diffusion timescales at the range of composition, temperature, and fO₂ values at Kelud. Fe-Mg zoning in Cpx crystals was modeled by using the diffusion coefficient of Müller et al. (2013), assuming isotropic diffusion. A mean rim temperature of T = 979 ± 48°C was used for modeling according to Opx rim-melt temperatures we have obtained (see below), with the fO₂ buffered at NNO +1 (log fO₂ = −9.60) according to the formulation of O'Neill & Pownceby (1993). For the late-stage Opx, we used a mean temperature of T = 972 ± 48°C calculated from the rim and calculated melt pairs, also with fO₂ at the same buffer (log fO₂ = −9.70). Data for modeling the Cpx crystals is from one sample (90–23) for which Cpx-melt temperatures average 1023 ± 42°C for the main-stage and 976 ± 42°C for the late-stage, all according to our geothermobarometry results presented in further sections.

## Calculation of exsolved fluid using SO₂ mass balance

We computed the mass of fluid exsolved from the melt by using the petrologic method (Devine et al., 1984). The highest sulphur

concentration in Cpx-hosted melt inclusions was taken to represent the pre-eruptive S content of the melt and the sulphur concentration of the interstitial glass to represent the degassed melt. The total amount of sulphur (as SO₂) outgassed during the 1990 eruption is given by the sulphur loss per unit mass of magma normalized to the residual melt proportions and multiplied by the DRE volume of the eruption. We compared our petrologic estimate of the outgassed amount of SO₂ with the total SO₂ emissions measured from space (0.17 Mt; TOMS, Global Volcanism Program, 2013) to evaluate the presence of an exsolved fluid phase prior to eruption (e.g. Costa et al., 2003; Costa et al., 2013; Jay et al., 2014).

# RESULTS

## Petrology, mineralogy, and geochemistry of main- and late-stage pumices

### Whole rock and glass chemistry

The whole rock composition of both, the main- and late-stage pumices is basaltic andesite (Table 2 and Fig. 3a), with 54–56 wt.% SiO₂, a total alkali silica (Na₂O + K₂O) of ~4 wt.%, and a K₂O content that straddles the boundary between the low- to mid-K series (Fig. 3b). Other major oxides show limited compositional variability (8–9 wt.% CaO, 2–4 wt.% MgO, 18–19 wt.% Al₂O₃, and 9–10 wt.% FeO*) (Table 2, Fig. 3c–f). Interstitial glass and melt inclusions from the main-stage samples are more evolved than the corresponding whole rock, ranging in composition from dacite to rhyolite (Fig. 3a). The late-stage pumices are slightly more evolved than the main-stage ones, and the interstitial glasses contain <1 wt.% of MgO, FeO* and CaO. The calculated groundmass melts obtained by subtracting phenocrysts from the whole rock are also dacitic (Table S3), and are compositionally intermediate between the whole rock and the measured interstitial glass or melt inclusions. Compositional trends in Harker diagrams of other major oxides (CaO, MgO, FeO, and Al₂O₃) versus SiO₂ are broadly linear and with a negative slope (Fig. 3c–f) which can be due to crystal fractionation and/or mixing between the crystals and melts of different origins (see Fig. S1).

Trace element concentrations in the whole rocks show a pattern typical of island arc magmas (Fig. S2). Primitive mantle-normalized trace element patterns reveal a relative enrichment in fluid mobile large ion lithophile elements (Cs, Rb, Ba, Sr), with a pronounced spike in Pb, and depletion in Nb. Chondrite-normalized rare earth elements (REE) concentrations show a relative enrichment in light REE ((La/Sm)ₙ = 1.48–1.73) and an almost flat pattern for the heavy REEs ((Dy/Yb)ₙ = 0.96–1.09). The main-stage pumice does not show significant Eu* anomalies (Eu* = Eu/(Sm × Gd)¹/² ~ 0.99 to 1.12), whereas the late-stage pumice has the highest positive Eu* anomaly (1.16). Differences in moderately volatile trace elements between main-stage and late-stage pumices involve a higher V/Cu ratio in the late-stage pumice (~9 vs. 3–6) indicating a shift in oxidation state of the late-stage erupted magma.

### Textural relations, mineralogy and mineral chemistry

The main-stage pumices are light gray to gray, with a porphyritic texture set in a vesicular, glass-rich, but microlite-bearing groundmass (Fig. 4a–j, Table 1). The phenocryst (crystals >100 mm) assemblage consists of Opx + Cpx + Plag + Mag + Amph and its pseudomorph (Amph*) ± olivine symplectite (Ol*) along with accessory apatite (Apa). The samples show a significant range of phenocryst content, from about 30 to 47 vol.%; plagioclase is the most abundant (< 14
<figure><img src="image_4.png" /><figcaption>**Fig. 3.** Normalized major element oxide compositions (versus SiO₂) of whole rocks and glasses from the 1990 Kelud eruption analyzed in this study. (a) total alkali content (Na₂O + K₂O); (b) K₂O; (c) CaO; (d) MgO; (e) Al₂O₃; (f) FeO. Greyscale domains represent sample compositions from other eruptions of Kelud (1919—Sjarifudin & Claproth 1989), 1951—Sjarifudin and Claproth 1989, this study; 1966—this study; Pre-1990—Bernard [unpub. data]; 2007—Humaida [2013], Jeffery et al. [2013]; 2014—Cassidy et al. [2016], Goode [2018]). Although we acknowledge that there is an apparent correlation between the compositional domains from other Kelud eruptions, this trend is the result of combining whole rocks and glasses together. Whereas the whole rock is consistently basaltic andesite and completely overlaps for all eruptions, the different glass compositions are more evolved and thus form a trend within each respective eruption.</figcaption></figure>

to 38 vol.%), followed by Opx (< 9 vol.%), Cpx (< 5 vol.%), Mag (< 5 vol.%), and Amph (< 0.2 vol.%). Phenocrysts of Plag, Opx or Cpx, are subhedral to euhedral or angular fragments of larger crystals.

Plag phenocrysts show a wide range of textures, including oscillatory zoning, sieved cores, and are typically broadly normally zoned with lower An at their rims than interiors (Fig. 4a). Unzoned Opx and Cpx phenocrysts (34% of the pyroxenes) are typically subhedral to euhedral and may host mineral and melt inclusions in their cores (Fig. 4b). Chemically zoned Opx phenocrysts (23% of the population) are typically reversely zoned, with a higher Mg-number rim mantling a rounded core (Fig. 4c). The compositional step from core to rim is sharp, even though some phenocrysts show more complex zoning at their rims (Fig. 4c). Normally zoned Opx phenocrysts (3%) are rare. Cpx phenocrysts are less commonly zoned; 18% show normal zoning in Mg-number, and about 1% reverse zoning (Fig. 4d). The cores of reversely zoned crystals are rounded or embayed, indicating a partial dissolution before rim regrowth. Mineral inclusions embedded in most phenocryst cores are commonly Fe-Ti oxides or apatite.

Glomerocrysts (13% of the total population) are made of Opx ± Cpx ± Plag ± Mag, with grain boundaries that can be rounded and sutured, or sub-rounded and partly exposed to the surrounding melt (Fig. 4f). We also found euhedral Cpx overgrowths over rounded and resorbed Opx phenocryst cores (3% of the total population), with inclusions of Mag outlining the core-overgrowths boundaries (Fig. 4g). Rare amphibole occurs as partially reacted phenocrysts (Fig. 4h). The pseudomorph reaction assemblage comprises anhedral fine-grained Cpx + Opx + Plag + Mag. Pseudomorphs after Amph breakdown consists of aggregates of Opx + Cpx + Plag + Mag in the original shape of the amphibole (Fig. 4i). Symplectites after olivine breakdown (Ol*, 1% of the total population) are made of Cpx + Opx + Mag (Fig. 4j). Microlites (< 100 mm) are mainly Plag, Cpx, Opx, Mag, and Apa, varying in habit from euhedral or skeletal to angular zoned fragments of larger crystals (Fig. 4e). Cpx microlites may have hourglass twinning, indicating rapid growth.

The late-stage pumices are light grey, with a porphyritic texture set in a vesicular and porous, glass-rich groundmass. They share the same rock-forming mineral assemblage as the main-stage pumices, but have higher phenocryst content (49 vol.%) and contain both quartz (Qz) and ilmenite. Opx, Cpx, and Plag phenocrysts can be
<table><caption>**Table 1:** Summary of sample description of 1990 samples used in our study. Asterisks (*) refer to the 1990 samples previously studied in Humaida (2013)</caption><thead><tr><th>Sample ID</th><th>KELUT90B*</th><th>1990-PL1</th><th>1990-BR</th><th>90-2*</th><th>90-20*</th><th>90-23*</th><th>1990-PL2</th></tr></thead><tbody><tr><td>Stage</td><td>Main</td><td>Main</td><td>Main</td><td>Main</td><td>Main</td><td>Main</td><td>Late</td></tr><tr><td>Latitude</td><td>-8.0</td><td>-7.9</td><td>-7.9</td><td>-7.9</td><td>-7.9</td><td>-7.9</td><td>-7.9</td></tr><tr><td>Longitude</td><td>112.0</td><td>112.3</td><td>112.3</td><td>112.3</td><td>112.3</td><td>112.3</td><td>112.3</td></tr><tr><td>Origin</td><td>Pumice fall</td><td>Pumice fall</td><td>Pumice flow</td><td>Pyroclastic flow*</td><td>Pyroclastic flow*</td><td>Pyroclastic flow*</td><td>Pumice fall</td></tr><tr><td>Hand specimen observations</td><td>light gray to gray porphyritic pumice</td><td>light gray to gray porphyritic pumice</td><td>light gray to gray porphyritic pumice</td><td>Thin section only</td><td>Thin section only</td><td>Thin section only</td><td>gray dense porphyritic pumice</td></tr><tr><td>Microscopic textures</td><td>Glomerocryst of plag + opx + cpx + mag. Overgrowth of cpx around rounded opx. Amph pseudomorphs and ol symplectites. Zoned and unzoned opx, cpx phenocrysts.</td><td>Glomerocryst of plag + opx + cpx + mag. Overgrowth of cpx around rounded opx. Contains partially reacted amph, amph pseudomorphs, and symplectite. Zoned and unzoned phenocrysts.</td><td>Glomerocryst of plag + opx + cpx + mag. Overgrowth of cpx around rounded opx. Ol symplectite. Zoned and unzoned phenocrysts. Amph pseudomorph.</td><td>Glomerocryst of plag + opx + cpx + mag. Amph pseudomorph. Overgrowth of cpx around rounded opx. Sieve-textured plag.</td><td>Glomerocryst of plag + opx + cpx + mag. Sieve textured plag. Zoned and unzoned phenocrysts.</td><td>Glomerocryst of plag + opx + cpx + mag. Overgrowth of cpx around rounded opx. Ol symplectite. Zoned and unzoned phenocrysts</td><td>Glomerocryst of plag + opx + cpx + mag. Overgrowth of cpx around rounded opx. All phenocrysts are zoned. Large (&gt;500um) amph pseudomorphs. Ol symplectite. Contains abundant qz bearing enclaves.</td></tr><tr><td>Matrix texture</td><td>vitric vesicular interstitial glass with variable vesicles. Rare microlites of plag + opx + cpx + mag</td><td>vitric vesicular interstitial glass with abundant fragments and microlites of (skeletal) plag + opx + cpx + mag. Opx microlites can be reversely zoned.</td><td>vitric vesicular interstitial glass with small vesicles. Rare microlites of plag + opx + cpx + mag</td><td>vitric vesicular interstitial glass</td><td>vitric vesicular interstitial glass</td><td>rare microlites of plag + opx + cpx + mag. Glass is partially devitrified.</td><td>vitric vesicular interstitial glass with thin, highly vesicular groundmass and thin (&lt;10um) vesicle walls. Microlites consist of normally zoned plag.</td></tr><tr><td>Other observations</td><td></td><td></td><td>One of oikocryst of opx surrounding plag + mag</td><td></td><td></td><td></td><td>Qz is anhedral and interstitial, crystallizing around phenocrysts.</td></tr><tr><td>Zoning type</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Plag</td><td>oscillatory, normal rim normal, reverse, multiple</td><td>oscillatory, normal rim normal, reverse, multiple</td><td>oscillatory, normal rim normal, reverse, multiple</td><td>No observations</td><td>oscillatory, normal rim normal, reverse, multiple</td><td>oscillatory, normal rim normal, reverse, multiple</td><td>oscillatory, normal rim normal, reverse, multiple</td></tr><tr><td>Opx/Cpx</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr></tbody></table>
<table><caption>**Table 1: Continued**</caption><thead><tr><th>Sample ID</th><th>KELUT90B*</th><th>1990-PL.1</th><th>1990-BR</th><th>90-2*</th><th>90-20*</th><th>90-23*</th><th>1990-PL.2</th></tr><tr><th>Zoned Minerals</th><th>Plag, Opx, Cpx</th><th>Plag, Opx, Cpx</th><th>Plag, Opx, Cpx</th><th>Plag, Opx, Cpx</th><th>Plag, Opx, Cpx</th><th>Plag, Opx, Cpx</th><th>Plag, Opx, Cpx</th></tr></thead><tbody><tr><td><strong>Modal Mineralogy # (vol.%)</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Glass</td><td>57.6</td><td>69.1</td><td>68.3</td><td>59.5</td><td>66.9</td><td>52.8</td><td>50.7</td></tr><tr><td>Plag</td><td>29.8</td><td>14.1</td><td>22.4</td><td>30.4</td><td>19.6</td><td>37.6</td><td>31.4</td></tr><tr><td>Opx</td><td>4.6</td><td>6.6</td><td>2.4</td><td>3.8</td><td>4.6</td><td>3.3</td><td>8.7</td></tr><tr><td>Cpx</td><td>5.2</td><td>5.4</td><td>3.9</td><td>4.2</td><td>5.3</td><td>3.9</td><td>3.3</td></tr><tr><td>Mag</td><td>2.9</td><td>4.8</td><td>3.0</td><td>1.9</td><td>3.6</td><td>2.4</td><td>4.3</td></tr><tr><td>Amph</td><td>&lt;0.1**</td><td>0.12</td><td>0.18**</td><td>0.12**</td><td></td><td></td><td>1.6**</td></tr><tr><td>Apa</td><td>&lt; 0.1</td><td>&lt; 0.1</td><td>&lt; 0.1</td><td>&lt; 0.1</td><td>&lt; 0.1</td><td>&lt; 0.1</td><td>&lt; 0.1</td></tr><tr><td>Ol</td><td>&lt; 0.1</td><td>&lt; 0.1</td><td>&lt; 0.1</td><td></td><td></td><td></td><td>&lt; 0.1</td></tr><tr><td>Ilm</td><td></td><td></td><td></td><td></td><td></td><td></td><td>&lt; 0.1</td></tr><tr><td>Total</td><td>100</td><td>100</td><td>100</td><td>100</td><td>100</td><td>100</td><td>98</td></tr><tr><td><strong>Crystal Content for crystals > 1 mm (vol.%)</strong></td><td>42.4</td><td>30.9</td><td>31.7</td><td>40.4</td><td>33.1</td><td>47.3</td><td>47.7</td></tr><tr><td><strong>Analyses Conducted</strong></td><td>XRF, SEM (WDS/EDS), EPMA</td><td>XRF, SEM (WDS/EDS), EPMA</td><td>XRF, SEM (WDS), EPMA</td><td>SEM (WDS)</td><td>SEM (WDS)</td><td>XRF, SEM (WDS/EDS), EPMA</td><td>XRF, SEM (WDS/EDS), EPMA</td></tr></tbody></table>

\#Plag = plagioclase; Opx = orthopyroxene; Cpx = clinopyroxene; Mag = magnetite; Amph = amphibole; Apa = apatite; Ol = olivine; Ilm = ilmenite

\*Samples from H. Humaida's unpublished doctoral thesis (Humaida, 2013)

\*\*Amph is pseudomorph

\*\*\*Ol has been turned to symplectite
<table><caption>**Table 2:** Major and trace element composition for 1990 whole rocks. KELUT90B was used as input for rhyolite-MELTS modeling</caption><thead><tr><th>Sample ID</th><th>KELUT-90B</th><th>KELUD-1990-PL1</th><th>KELUD-1990-BR</th><th>KELUD-1990-PL2</th></tr></thead><tbody><tr><td>Stage</td><td>Main</td><td>Main</td><td>Main</td><td>Late</td></tr><tr><td colspan="5"><em>Major elements (wt.%)</em></td></tr><tr><td>SiO<sub>2</sub></td><td>56.15</td><td>54.17</td><td>54.6</td><td>54.42</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>18.42</td><td>18.95</td><td>18.36</td><td>17.88</td></tr><tr><td>FeO*</td><td>8.74</td><td>9.35</td><td>9.65</td><td>9.26</td></tr><tr><td>MnO</td><td>0.2</td><td>0.2</td><td>0.21</td><td>0.2</td></tr><tr><td>MgO</td><td>3.65</td><td>3.66</td><td>3.81</td><td>3.81</td></tr><tr><td>CaO</td><td>9.1</td><td>9.05</td><td>9.2</td><td>9.08</td></tr><tr><td>Na<sub>2</sub>O</td><td>2.99</td><td>2.9</td><td>2.82</td><td>2.89</td></tr><tr><td>K<sub>2</sub>O</td><td>0.7</td><td>0.69</td><td>0.67</td><td>0.67</td></tr><tr><td>TiO<sub>2</sub></td><td>0.64</td><td>0.69</td><td>0.67</td><td>0.65</td></tr><tr><td>P<sub>2</sub>O<sub>5</sub></td><td>0.13</td><td>0.16</td><td>0.14</td><td>0.13</td></tr><tr><td>Total</td><td>100.7</td><td>99.8</td><td>100.4</td><td>99.0</td></tr><tr><td colspan="5"><em>Trace elements (mg/kg)</em></td></tr><tr><td>La</td><td>5.82</td><td>6.05</td><td>6.0</td><td>5.96</td></tr><tr><td>Ce</td><td>13.2</td><td>13.5</td><td>13.3</td><td>13.3</td></tr><tr><td>Pr</td><td>1.82</td><td>1.76</td><td>1.74</td><td>1.78</td></tr><tr><td>Nd</td><td>8.34</td><td>8.34</td><td>8.01</td><td>8.23</td></tr><tr><td>Sm</td><td>2.55</td><td>2.37</td><td>2.24</td><td>2.36</td></tr><tr><td>Eu</td><td>0.873</td><td>0.854</td><td>0.817</td><td>0.902</td></tr><tr><td>Gd</td><td>2.84</td><td>2.32</td><td>2.52</td><td>2.4</td></tr><tr><td>Tb</td><td>0.46</td><td>0.44</td><td>0.43</td><td>0.41</td></tr><tr><td>Dy</td><td>2.72</td><td>2.78</td><td>2.60</td><td>2.74</td></tr><tr><td>Ho</td><td>0.55</td><td>0.56</td><td>0.53</td><td>0.58</td></tr><tr><td>Er</td><td>1.72</td><td>1.67</td><td>1.64</td><td>1.7</td></tr><tr><td>Tm</td><td>0.287</td><td>0.251</td><td>0.248</td><td>0.256</td></tr><tr><td>Yb</td><td>1.89</td><td>1.71</td><td>1.71</td><td>1.7</td></tr><tr><td>Lu</td><td>0.29</td><td>0.27</td><td>0.25</td><td>0.27</td></tr><tr><td>Y</td><td>15.7</td><td>15.3</td><td>15.6</td><td>15.3</td></tr><tr><td>Zr</td><td>44</td><td>39</td><td>38</td><td>38</td></tr><tr><td>Nb</td><td>1.2</td><td>0.9</td><td>0.8</td><td>0.9</td></tr><tr><td>Hf</td><td>1.5</td><td>1.1</td><td>1.1</td><td>1.2</td></tr><tr><td>Ta</td><td>0.11</td><td>0.04</td><td>0.02</td><td>0.05</td></tr><tr><td>Th</td><td>0.93</td><td>0.86</td><td>0.83</td><td>0.85</td></tr><tr><td>U</td><td>0.38</td><td>0.34</td><td>0.33</td><td>0.35</td></tr><tr><td>Rb</td><td>13</td><td>12</td><td>12</td><td>12</td></tr><tr><td>Cs</td><td>1.0</td><td>0.8</td><td>0.8</td><td>0.8</td></tr><tr><td>Sr</td><td>577</td><td>605</td><td>580</td><td>575</td></tr><tr><td>Ba</td><td>484</td><td>479</td><td>463</td><td>465</td></tr><tr><td>Sc</td><td>20</td><td>19</td><td>20</td><td>19</td></tr><tr><td>V</td><td>183</td><td>179</td><td>180</td><td>179</td></tr><tr><td>Cr</td><td>30</td><td>&lt; 20</td><td>&lt; 20</td><td>80</td></tr><tr><td>Co</td><td>19</td><td>19</td><td>20</td><td>20</td></tr><tr><td>Cu</td><td>40</td><td>60</td><td>30</td><td>20</td></tr><tr><td>Pb</td><td>9</td><td>8</td><td>8</td><td>8</td></tr><tr><td>Zn</td><td>60</td><td>60</td><td>70</td><td>60</td></tr><tr><td>Sn</td><td>4</td><td>3</td><td>&lt; 1</td><td>1</td></tr><tr><td>W</td><td>1.9</td><td>2.3</td><td>2.5</td><td>0.5</td></tr><tr><td>Ga</td><td>16</td><td>15</td><td>16</td><td>16</td></tr><tr><td>Ge</td><td>1.4</td><td>1.3</td><td>1.3</td><td>1.3</td></tr><tr><td>Tl</td><td>0.49</td><td>&lt; 0.05</td><td>&lt; 0.05</td><td>&lt; 0.05</td></tr><tr><td>Sb</td><td>0.8</td><td>&lt; 0.2</td><td>0.3</td><td>0.4</td></tr><tr><td>Eu*</td><td>0.99</td><td>1.12</td><td>1.05</td><td>1.16</td></tr></tbody></table>

Continued
<table><caption>**Table 2: Continued**</caption><thead><tr><th>Sample ID</th><th>KELUT-90B</th><th>KELUD-1990-PL1</th><th>KELUD-1990-BR</th><th>KELUD-1990-PL2</th></tr></thead><tbody><tr><td>CI Chondrite normalized REE</td><td></td><td></td><td></td><td></td></tr><tr><td>(La/Yb)n</td><td>2.21</td><td>2.54</td><td>2.52</td><td>2.52</td></tr><tr><td>(La/Sm)n</td><td>1.48</td><td>1.65</td><td>1.73</td><td>1.63</td></tr><tr><td>(Dy/Yb)n</td><td>0.96</td><td>1.09</td><td>1.02</td><td>1.08</td></tr><tr><td>V/Cu</td><td>4.58</td><td>2.98</td><td>6.00</td><td>8.95</td></tr></tbody></table>

Eu* = Eu / (Sm × Gd)1/2, with all compositions are normalized to primitive mantle.

FeO* = total iron oxide, converted to Fe2+

See data on glass inclusions and calculated matrix in supplementary data.

chemically zoned. Plagioclase shows progressive, normally zoned rims (Fig. 5a). Opx phenocrysts are typically normally zoned (44% of the total) rather than reversely zoned (2%) (Fig. 5b). Phenocryst edges are serrated or show reaction rims, likely from partial melting of evolved interstitial glass. Glomerocrysts are common (39% of the total population surveyed in this study) and with the same mineralogy as those in the main-stage pumices (Fig. 5c). Cpx can be found as overgrowths on Opx (5%) and are more abundant than in main-stage pumices (Fig. 5d). The groundmass is highly porous with thin vesicle walls (< 10 mm) and rhyolitic interstitial glass (Fig. 5e). Microlites are mainly Plag plus some Cpx, Opx, Mag, and Apa (Fig. 5f). Fe-Ti oxides are Mag and Ilm, with Mag partially surrounded by frayed Ilm rims wherever Mag edges are exposed to melt (Fig. 5g). Quartz crystals commonly fill the interstices surrounding phenocrysts of Plag ± Opx ± Cpx ± Mag in enclaves, indicating that the magma cooled enough to evolve to high SiO₂ content and crystallize Qz (Fig. 5h). Amphibole is completely reacted to pseudomorphs (7%) that are considerably larger than the main-stage pseudomorph (> 500 mm) and appear as fine-grained Cpx + Opx + Plag + Mag assemblages in the shape of the original Amph crystal (Fig. 5i). Olivine symplectites (2% of the total population) are similar in morphology and mineralogy to those in main-stage pumices (Fig. 5j).

### Mineral compositions and traverses

The main-stage Plag ranges from An = 58 to 90 (Fig. 6a, Table 3), with cores showing a bimodal distribution at An58–75 and An75–90, whereas rims are mainly unimodal at An58–75. Plagioclase phenocrysts can be classified into two main textural types (Fig. 6): Type 1a Plag has a spongy, sieved core (~ An86) mantled by a zone at An = 70–90, which is in turn surrounded by a lower-An rim (58–70), some crystals can also have an intermediate sieved zone, which we call Type 1b. Type 2 crystals do not show significant zoning or textures aside from oscillatory zoning, with cores fluctuating between An60–75 and An75–90 and surrounded by lower An rim (An60–70) (Fig. 6d). Plagioclase microlites are zoned, with cores of An60–85 mantled by rims of An60–62 that are similar to Type 2 Plag phenocrysts. Compositional profiles in Type 1b crystals show one plateau that coincides with a light grey zone in BSE images (An80) and darker rim (An60). The outer zones of these crystals show multiple compositional fluctuations, from about An 62 to 75 (Fig. 6e–f). We have few Plag analyses from the late-stage pumice. These have similar phenocryst types and compositions as the main-stage pumices, but a major difference is the presence of a more evolved rim population at about An40. Such a low-An rim population is consistent with

the more evolved compositions of the interstitial melt in late-stage pumice.

Main-stage Opx phenocrysts display a bimodal distribution with a mode at Mg-number = 62–70 and another one at 70–74 (Fig. 7a, Table 4). The low Mg-number mode characterizes glomerocrysts, Opx cores with Cpx overgrowths, and most cores of both zoned and unzoned phenocryst. A Mg-number > 70 is common for rims of reversely zoned phenocrysts and microlites. Reversely zoned Opx phenocrysts (Fig. 7b) commonly display sharp compositional steps, with higher Mg-number in rims (72–74) than cores (66–70) (Fig. 7c). Opx crystals in late-stage pumice show a wide compositional spread (Mg-number = 58–72), and many are more evolved than Opx from the main-stage (Table 4). Most unzoned crystals have a core compositionally similar to main-stage Opx (Mg-number = 68–69), and some Opx phenocrysts show normal zoning (Fig. 7d), with Mg-number 64–68 in cores and 58–62 in rims (Fig. 7e). The composition of Opx in glomerocrysts (Mg-number = 59–72) overlap with those of phenocryst cores and microlites. Overgrowth Opx cores have similar compositions to main-stage overgrowth ones (Mg-number = 68–70). Opx in Amph breakdown textures overlap with most other compositions (Mg-number = 60–68).

Cpx phenocrysts in main-stage samples (Mg-number = 70–78, Wo38–45En40–48) show a narrower compositional range than Opx, and vary depending on the textural group (Fig. 8a and Table 4; for Wo and En distributions, see Fig. 5c). Phenocryst cores either overlap or have lower Mg-number (70–76) than rims (72–76), rarely Cpx also show reverse zoning (Fig. 8b–c). Glomerocryst Cpx cores overlap with phenocryst cores at Mg-number = 71–76. Cpx overgrowth and microlites record the largest compositional range in Mg-number (72–74), and in Wo and En content (Wo41–44En40–43) compared to the rest of the Cpx. Late-stage Cpx (Mg-number = 58–73, Wo38–48En31–48; Table 4), are more evolved than main-stage Cpx. Phenocrysts show normal zoning (Fig. 8d), with higher Mg-number in cores (= 70–76) than rims (= 68–70; Fig. 8e). Cpx in glomerocrysts display the widest compositional range (Mg-number = 62 to 79, Wo41–49En32–45). Cpx microlites and overgrowth (Mg-number = 72–74, Wo42–44En41–43) have higher Mg and Ca compared to the phenocrysts. Cpx found as pseudomorph (Mg-number = 69–74, Wo41–44En39–42) is likely produced from Amph breakdown.

The two partially reacted amphibole phenocrysts (Mg-number = 76–83) from the main-stage straddle the magnesio-hastingsite and tschermakite-pargasite end-members (Table 5) defined by Leake et al. (1997). Both phenocrysts have reaction rims consisting of an anhedral, fine-grained assemblage of Opx + Cpx + Plag +
<figure><img src="image_5.png" /><figcaption>**Fig. 4.** Plane-polarized (PPL) and BSE images of representative mineral textures pumices from the main eruptive stage. (a) Type 2 Plag phenocryst with a partially resorbed core mantled by a spongy-sieve textured growth zone with a normally-zoned rim. Light grey inclusions are apatite (Apa). (b) Cpx phenocryst with glassy inclusions (MI), Plag inclusions, and post-entrapment daughter minerals (PEM). (c) Euhedral reversely zoned Opx phenocryst with Apa and Mag inclusions. (d) Normally zoned Cpx and reversely zoned Opx. The normally zoned Cpx hosts many irregular melt inclusions. (e) Groundmass texture with plagioclase and pyroxene microlites surrounded by quenched interstitial glass. (f) Glomerocryst of Opx + Cpx + Mag showing both unsutured and sutured grain boundaries. (g) PPL image of Cpx overgrowth mantling rounded Opx core with Mag phenocrysts along boundaries. Orange dashed line shows the Opx-Cpx boundary (h) Partially reacted Amph grains with anhedral Mag phenocrysts attached. (i) Pseudomorph after Amph breakdown, filled with fine-grained assemblage of Opx + Cpx + Plag + Mag. (j) Symplectitic texture after olivine breakdown showing anhedral pyroxenes surrounded by stringy magnetite.</figcaption></figure>
<figure><img src="image_6.png" /><figcaption>**Fig. 5.** PPL and BSE images of representative mineral textures in the pumice from the late eruption stage. White and grey scale bars indicate 100 µm scale. (a) Normally zoned Plag phenocryst with glassy inclusions. (b) Multiply zoned Opx phenocrysts attached to partially reacted Mag glomerocryst, featuring reverse mantling growth zone followed by normal rim. (c) Glomerocryst comprising anhedral to subhedral Opx + Cpx + Plag + Mag. (d) PPL image of Cpx overgrowth mantling rounded Opx core. Orange dashed line marks the Opx-Cpx boundary. (e) Groundmass textures with Fe-Ti oxides, pyroxenes, and Plag microlites. (f) Magnified view of Opx + Apa + Plag + Cpx microlites; (g) Mag phenocryst with Ilm rims. (h) Quartz bearing enclaves, where Qz forms an interstitial phase around the pyroxene and plagioclase phenocrysts. (i) Pseudomorph of Amph, with Plag inclusions. (j) Symplectitic texture surrounded by Mag, Plag, and pyroxene phenocrysts.</figcaption></figure>
<figure><img src="image_7.png" /><figcaption>**Fig. 6.** (a) False color BSE images, histograms and representative traverses of plagioclase compositions from the main and late eruption stage, further divided into type 1–3 phenocrysts, microlites, and enclaves. (b) Annotated false color BSE image of a type 1 Plag, with An compositions from spot analyses. (c) Same for type 2 Plag, with An compositions from spot analyses. (d) Same for type 3 Plag, with An compositions from spot analyses. (e) Annotated BSE image of a main stage Plag with the analyzed traverse from rim to core shown by a red arrow. (f) Traverse analysis of An content from rim to core, denoted by the red arrow.</figcaption></figure>

Mag. Magnetite in main-stage samples shows a narrow range of X<sub>usp</sub> = 0.28–0.30 (Table 5), whereas magnetite in late-stage pumice shows a bimodal distribution at about 0.30 and at 0.31–0.34 (Fig. S4a), with no overlap. Ilmenite also displays a bimodal distribution, with a first mode at X<sub>ilm</sub> = 0.88–0.91 and a second one X<sub>ilm</sub> = 0.91–0.92 (Fig. S4b).

## Pre-eruptive magma storage conditions

### Pyroxene temperature constraints

The pyroxene-melt equilibrium temperatures from the main-eruptive stage samples show a wide range, from about 950 to 1100°C across textural groups. Opx-melt temperatures vary between about 980 to

1050°C, and the crystal core temperatures overlap or are somewhat higher than those obtained from the rims. The high temperature range from Opx-melt equilibria is close to or overlaps with the whole-rock liquidus temperatures (T<sub>liquidus</sub>) calculated with MELTS at NNO with 3–4 wt.% of water in the melt, and about 50°C lower the liquidus with 2 wt.% water in the melt (Fig. 9a). Moreover, temperatures recorded from individual core to rim crystal traverses record a temperature decrease of about <50°C, even for crystals that are reversely zoned in Mg-number (Fig. S5). Opx microlite-melt temperatures overlap with those of the phenocryst rims. Cpx-melt temperatures overlap or a somewhat higher than those of Opx, with a mode at ~1050°C. The temperatures of crystal cores and rims overlap, except for reversely zoned (in Mg-number) Cpx traverses,
<figure><img src="image_8.png" /><figcaption>**Fig. 7** (a) Histograms and representative traverses of Opx compositions, subdivided according to the textural groups: cores of phenocrysts, overgrowth Opx cores and microlites; rims of zoned phenocrysts; and glomerocrysts. Red arrows indicate the direction of traverse analysis. (b) BSE image of a reversely zoned main-stage Opx. (c) Traverse analyses from rim to core of Mg-number of a main-stage Opx. (d) BSE image of a normally zoned late-stage Opx. (e) Traverse analysis from rim to core of Mg-number of a late-stage Opx.</figcaption></figure>

where core temperatures are generally higher than the rims. Despite this apparent core to rim temperature decrease, the crystal cores in zoned Opx and Cpx are resorbed, and suggest a period of crystal dissolution between cores and rims induced by temperature and compositional effects. Finally, Opx-Cpx temperatures partly overlap or are slightly higher than those of Opx and Cpx-melt with one main mode at ~1000 (± 25°C) and smaller one at ~1080 (± 25°C). Temperature ranges from core-core pairs, rim-rim pairs, and microlites overlap.

Opx and Cpx core-melt temperatures for the late-stage deposits are similar and overlap with those of the main-stage (1040–1070°C), and are somewhat higher than those of the rims and microlites (970–1020°C) (Fig. 9a). The cores of zoned Opx and Cpx crystals tend also to be resorbed as in the main-stage samples. Opx-Cpx pair temperatures (990–1030°C) overlap between core and rims (Figure 9). Opx-Cpx temperatures for the microlites in late-stage

pumice overlap with those for crystal rims but are lower than those inferred for Opx-Cpx pairs in main-stage pumice.

**Magnetite-ilmenite oxybarometer**

Temperatures obtained from Mag-Ilm pairs from the late-stage pumice are 780–850°C, i.e. 100 to 150°C lower than those inferred from silicate minerals (Fig. 9a). Temperatures obtained using the Ghiorso and Evans (2008) model have a mode at 825°C, and an fO₂ of about 0.5–2 log units above the NNO buffer. The model of Sauerzapf et al. (2008) model yields somewhat lower temperatures (T = 780–800°C), but a similar range in fO₂. Such low temperatures indicate that the late-stage magmas experienced at some point in their history cooler storage conditions than the main-stage magma, which is consistent with their higher crystallinity, and the presence of low temperature minerals such as quartz.
<table><caption>**Table 4:** Opx and Cpx compositions and textures in pumices representative for the main and late eruption stages</caption><thead><tr><th></th><th>PL1-1-<br>px-z1a</th><th>PL1-1-<br>px-MX1</th><th>90B-<br>opxcpx17</th><th>PL1-<br>site1-<br>px4</th><th>PX6-<br>c-only</th><th>glom1-<br>c-only</th><th>PL2-<br>2px7</th><th>PL2-<br>site1-<br>px7</th><th>PL1-1-<br>px6</th><th>90-23-<br>px18</th><th>90B-<br>opxcorepxri<br>m-1</th><th>PL1-<br>site1-<br>px8</th><th>PX5-<br>core-<br>only</th><th>glom4-<br>c-only</th><th>PL2-px7</th><th>PL2-<br>site1-<br>px2</th></tr><tr><th></th><th>Opx</th><th>Opx</th><th>Opx</th><th>Opx</th><th>Opx</th><th>Opx</th><th>Opx</th><th>Opx</th><th>Opx</th><th>Opx</th><th>Opx</th><th>Opx</th><th>Opx</th><th>Opx</th><th>Opx</th></tr><tr><th>Pheno</th><th>Main</th><th>Main</th><th>Main</th><th>Main</th><th>Late</th><th>Late</th><th>Late</th><th>Late</th><th>Late</th><th>Late</th><th>Late</th><th>Main</th><th>Late</th><th>Late</th><th>Late</th></tr><tr><th>Pheno</th><th>Glom</th><th>Over-<br>growth</th><th>Microlite</th><th>Pheno</th><th>Glom</th><th>Over-<br>growth</th><th>Microlite</th><th>Pheno</th><th>Glom</th><th>Over-<br>growth</th><th>Microlite</th><th>Pheno</th><th>Glom</th><th>Over-<br>growth</th><th>Microlite</th></tr></thead><tbody><tr><th>wt.%</th><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><th>SiO<sub>2</sub></th><td>53.36</td><td>52.02</td><td>52.1</td><td>52.2</td><td>53.09</td><td>52.43</td><td>53.1</td><td>51.3</td><td>52.48</td><td>51.8</td><td>50.81</td><td>51.23</td><td>51.4</td><td>51.7</td><td>51.4</td></tr><tr><th>TiO<sub>2</sub></th><td>0.18</td><td>0.21</td><td>0.2</td><td>0.2</td><td>0.12</td><td>0.18</td><td>0.1</td><td>0.2</td><td>0.33</td><td>0.4</td><td>0.32</td><td>0.37</td><td>0.3</td><td>0.4</td><td>0.4</td></tr><tr><th>Al<sub>2</sub>O<sub>3</sub></th><td>1.07</td><td>1.75</td><td>1.4</td><td>1.8</td><td>1.01</td><td>1.61</td><td>0.8</td><td>1.6</td><td>1.71</td><td>2.0</td><td>1.82</td><td>1.59</td><td>2.0</td><td>2.2</td><td>1.7</td></tr><tr><th>FeO*</th><td>21.51</td><td>21.90</td><td>20.1</td><td>18.9</td><td>18.87</td><td>19.32</td><td>18.5</td><td>21.2</td><td>10.85</td><td>11.0</td><td>9.84</td><td>10.06</td><td>9.4</td><td>9.5</td><td>9.4</td></tr><tr><th>MnO</th><td>0.85</td><td>0.71</td><td>0.9</td><td>0.7</td><td>0.73</td><td>0.71</td><td>0.7</td><td>0.7</td><td>0.52</td><td>0.4</td><td>0.48</td><td>0.48</td><td>0.4</td><td>0.4</td><td>0.4</td></tr><tr><th>MgO</th><td>23.23</td><td>23.39</td><td>23.4</td><td>24.5</td><td>23.07</td><td>22.74</td><td>23.6</td><td>22.9</td><td>14.57</td><td>14.7</td><td>14.53</td><td>14.73</td><td>14.6</td><td>14.3</td><td>14.3</td></tr><tr><th>CaO</th><td>1.42</td><td>1.34</td><td>1.2</td><td>1.5</td><td>1.57</td><td>1.56</td><td>1.5</td><td>1.6</td><td>20.75</td><td>19.4</td><td>20.67</td><td>20.27</td><td>21.0</td><td>20.9</td><td>21.2</td></tr><tr><th>Na<sub>2</sub>O</th><td>0.01</td><td>0.01</td><td>0.0</td><td>0.1</td><td>0.02</td><td>0.02</td><td>0.0</td><td>0.2</td><td>0.3</td><td>0.3</td><td>0.29</td><td>1.38</td><td>0.3</td><td>0.3</td><td>0.3</td></tr><tr><th>Total</th><td>101.60</td><td>101.30</td><td>99.2</td><td>99.9</td><td>98.50</td><td>98.60</td><td>98.4</td><td>99.7</td><td>101.5</td><td>99.9</td><td>98.8</td><td>100.1</td><td>99.5</td><td>99.7</td><td>99.1</td></tr><tr><th>Oxide</th><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><th>Mg</th><td>65.81</td><td>65.6</td><td>67.5</td><td>69.8</td><td>68.5</td><td>67.7</td><td>69.4</td><td>65.8</td><td>70.5</td><td>70.4</td><td>72.5</td><td>72.3</td><td>73.4</td><td>72.9</td><td>73.0</td></tr><tr><th>Number</th><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><th>Wo</th><td>2.81</td><td>2.63</td><td>2.4</td><td>2.9</td><td>3.24</td><td>3.23</td><td>3.1</td><td>3.1</td><td>41.9</td><td>40.1</td><td>42.59</td><td>41.7</td><td>43.2</td><td>43.4</td><td>43.8</td></tr><tr><th>En</th><td>63.96</td><td>63.84</td><td>65.8</td><td>67.7</td><td>66.32</td><td>65.53</td><td>67.2</td><td>63.7</td><td>41.0</td><td>42.2</td><td>41.57</td><td>42.2</td><td>41.7</td><td>41.3</td><td>41.1</td></tr><tr><th>Fs</th><td>33.23</td><td>33.53</td><td>31.7</td><td>29.3</td><td>30.43</td><td>31.23</td><td>29.7</td><td>33.1</td><td>17.1</td><td>17.8</td><td>15.84</td><td>16.2</td><td>15.1</td><td>15.3</td><td>15.2</td></tr></tbody></table>

FeO* and structural end-members based on Deer et al. (1992). Mg Number = Mg/(Mg + Fetotal). Please see supplementary material for more analyses.
<figure><img src="image_9.png" /><figcaption>**Fig. 8.** (a) Histograms and representative traverses of Cpx compositions, subdivided into the same textural groups as Opx. Traverse directions shown by red arrows. (b) BSE image of a reversely zoned main-stage Cpx. (c) Traverse from rim to core of Mg-number of a main-stage Cpx. (d) BSE image of a normally-zoned late-stage Cpx. (e) Traverse analysis from rim to core of Mg-number of a late-stage Cpx.</figcaption></figure>

**Amphibole-melt geothermobarometer**

We obtained temperature and pressure estimates from the two Amph phenocrysts from the main-stage samples (Fig. 10). The Putirka (2016) and Li et al. (2021) combined formulation yields 960–980°C, and 400–550 MPa. Slightly lower temperatures (920–970°C) but higher pressures (480–590 MPa) are derived from the Ridolfi & Renzulli (2012) Amph-only formulation, whereas significantly higher temperatures (1000–1030°C) and pressures (570–830 MPa) result when applying the formulation of Ridolfi et al. (2010). These temperature estimates for amphibole either overlap with or are lower than pyroxene-based temperatures for crystal rims and microlites.

**Clinopyroxene-melt geobarometer**

We obtained pressure estimates from the Cpx-melt pairs from the main-stage samples using the formulations by Neave and Putirka (2017) and Putirka (2008) (Fig. S10). Both formulations yield

overlapping pressures of 100–350 MPa, with comparable peaks at ~150 MPa. These pressure estimates overlap with saturation pressure estimates, but are lower than estimates from Amph based geobarometers.

**Dissolved water content, saturation pressures, and exsolved fluids**

We assessed the concentration of dissolved water in the melt with the plagioclase-melt equilibrium model of Waters & Lange (2015), using the Plag rim and interstitial glass compositions, and the temperature obtained from Opx rim-melt (Table S3). Main-stage plagioclase rims of An60–70 yield ~2–4 wt.% water in the melt, consistent with the water content required to crystallize amphibole. In comparison, the late-stage plagioclase rims have lower anorthite content, with An40 that gave higher values of ~4–6 wt.% water in the melt. Cassidy et al.
<table><caption>**Table 5:** Representative Amph and Fe-Ti oxides compositions in main- and late-stage pumices</caption><thead><tr><th rowspan="3"></th><th>KELUD1990-PL1new-1-amph2</th><th>KELUD1990PL1-amph1_8</th><th>90-23-Mag1</th><th>Mag-Ilm-13</th><th>Mag-Ilm-26</th></tr><tr><th>Amph</th><th>Amph</th><th>Mag</th><th>Mag</th><th>Ilm</th></tr><tr><th>Main</th><th>Main</th><th>Main</th><th>Late</th><th>Late</th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>41.46</td><td>42.59</td><td>TiO<sub>2</sub></td><td>8.7</td><td>9.4</td><td>43.9</td></tr><tr><td>TiO<sub>2</sub></td><td>1.41</td><td>1.55</td><td>Al<sub>2</sub>O<sub>3</sub></td><td>2.9</td><td>1.2</td><td>2.4</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>14.92</td><td>14.01</td><td>Cr<sub>2</sub>O<sub>3</sub></td><td>0.0</td><td>0.1</td><td>0.0</td></tr><tr><td>Cr<sub>2</sub>O<sub>3</sub></td><td>0.00</td><td>0.02</td><td>FeO</td><td>39.6</td><td>41.2</td><td>34.2</td></tr><tr><td>Fe<sub>2</sub>O<sub>3</sub></td><td>5.10</td><td>8.61</td><td>Fe<sub>2</sub>O<sub>3</sub></td><td>44.7</td><td>44.3</td><td>14.6</td></tr><tr><td>FeO</td><td>7.04</td><td>3.45</td><td>MnO</td><td>0.5</td><td>0.6</td><td>1.0</td></tr><tr><td>MnO</td><td>0.15</td><td>0.17</td><td>MgO</td><td>2.1</td><td>1.0</td><td>2.3</td></tr><tr><td>MgO</td><td>13.28</td><td>14.35</td><td></td><td>0.11</td><td>0.1</td><td>0.1</td></tr><tr><td>CaO</td><td>12.50</td><td>11.68</td><td></td><td></td><td></td><td></td></tr><tr><td>Na<sub>2</sub>O</td><td>2.29</td><td>2.25</td><td>Total</td><td>99.0</td><td>97.9</td><td>98.68</td></tr><tr><td>K<sub>2</sub>O</td><td>0.31</td><td>0.29</td><td></td><td></td><td></td><td></td></tr><tr><td>F</td><td></td><td>0.165</td><td>Structural cations***</td><td></td><td></td><td></td></tr><tr><td>Cl</td><td></td><td></td><td>Ti</td><td>1.70</td><td>2.0</td><td>6.9</td></tr><tr><td>H<sub>2</sub>O<sub>calc</sub></td><td>1.93</td><td>1.87</td><td>Al</td><td>0.89</td><td>0.4</td><td>0.3</td></tr><tr><td>Total</td><td>100.39</td><td>101.01</td><td>Cr</td><td>0.01</td><td>0.0</td><td>0.0</td></tr><tr><td></td><td></td><td></td><td>Fe<sup>2+</sup></td><td>12.44</td><td>13.1</td><td>7.4</td></tr><tr><td></td><td></td><td></td><td>Fe<sup>3+</sup></td><td>3.19</td><td>3.0</td><td>0.9</td></tr><tr><td>T site</td><td></td><td></td><td>Mn</td><td>0.11</td><td>0.1</td><td>0.2</td></tr><tr><td>Si</td><td>6.00</td><td>6.07</td><td>Mg</td><td>0.81</td><td>0.4</td><td>0.7</td></tr><tr><td>Al<sup>IV</sup></td><td>2.00</td><td>1.93</td><td>Si</td><td>0.03</td><td>0.0</td><td>0.0</td></tr><tr><td>Ti</td><td>0.0</td><td>0.00</td><td></td><td></td><td></td><td></td></tr><tr><td>Al<sup>VI</sup></td><td>0.55</td><td>0.43</td><td></td><td></td><td></td><td></td></tr><tr><td>Ti</td><td>0.15</td><td>0.17</td><td></td><td></td><td></td><td></td></tr><tr><td>Cr</td><td>0.00</td><td>0.00</td><td></td><td></td><td></td><td></td></tr><tr><td>Fe<sup>3+</sup></td><td>0.56</td><td>0.92</td><td></td><td></td><td></td><td></td></tr><tr><td>Mg</td><td>2.87</td><td>3.05</td><td></td><td></td><td></td><td></td></tr><tr><td>Fe<sup>2+</sup></td><td>0.85</td><td>0.41</td><td></td><td></td><td></td><td></td></tr><tr><td>Mn</td><td>0.02</td><td>0.02</td><td></td><td></td><td></td><td></td></tr><tr><td>B site</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Fe<sup>2+</sup></td><td>0.00</td><td>0.00</td><td></td><td></td><td></td><td></td></tr><tr><td>Ca</td><td>1.94</td><td>1.78</td><td></td><td></td><td></td><td></td></tr><tr><td>Na</td><td>0.06</td><td>0.22</td><td></td><td></td><td></td><td></td></tr><tr><td>A site</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Na</td><td>0.58</td><td>0.41</td><td></td><td></td><td></td><td></td></tr><tr><td>K</td><td>0.06</td><td>0.05</td><td></td><td></td><td></td><td></td></tr><tr><td>Mg/(Mg + Fe<sup>2+</sup>)</td><td>0.84</td><td>0.88</td><td></td><td></td><td></td><td></td></tr><tr><td>Mg Number</td><td>78.02</td><td>76.75</td><td></td><td></td><td></td><td></td></tr><tr><td>P (MPa)*</td><td>508</td><td>484</td><td></td><td></td><td></td><td></td></tr><tr><td>T (°C)*</td><td>912</td><td>939</td><td></td><td></td><td></td><td></td></tr><tr><td>End-member**</td><td colspan="2">Magnesio-Hastingsite</td><td colspan="2">Tschermakite-Pargasite</td><td></td><td></td></tr></tbody></table>

*Pressure and temperature calculated using formulation of Ridolfi & Renzulli (2012)

**Structural formulae according to Leake et al. (1997)

***Structural cations calculated using Droop (1987)

(2016) studied the 2014 eruption deposits and found that plagioclase-hosted melt inclusions contained 0.4–2.3 wt.% water and up to 336 mg/kg CO₂. Assuming that this CO₂ content is also representative of the 1990 melts, we calculated volatile saturation pressures

for the main-stage pumices of 110 to 170 MPa (~4–7 km depth), whereas for the late-stage pumices we obtained 150 to 230 MPa (~6–9 km depth) (Fig. 11; Table S3). The whole-rock, pre-eruptive temperature, mineral assemblage, and phenocryst content of the 1990
<figure><img src="image_10.png" /><figcaption>**Fig. 9.** (a) Histograms of calculated temperatures for the main- and late-stage pumice samples constrained by different geothermometers and for various textures (cores, rims, microlites, Mag-Ilm pairs). Geothermometers include: two-pyroxene QUIIF geothermometer (Andersen et al., 1993); Opx- and Cpx-melt geothermometers (eq. 28a and eq. 33, respectively) (Putirka 2008), using core-whole rock pairs and rim-interstitial glass pairs; and Mag-Ilm geothermometer for late-stage pumice (GE: Ghiorso & Evans [2008]; S: Sauerzapf et al. [2008]). Whole rock liquidus temperatures ($T_{liquidus}$) at various water contents are also provided for comparison and were calculated using rhyolite-MELTS (Gualda et al., 2012; Ghiorso & Gualda, 2015) and the main-stage pumice KELUT90B used as input material. (b): $D^{Mg-Fe2+}K_d$ vs. $DH_2O$, as recorded by main-stage Opx cores and rims calculated using an empirical relationship from experimental results by Waters & Lange (2017).</figcaption></figure>

main pumices and the 2014 pumice are similar and thus suggest that they had similar pre-eruptive content (Utami *et al.*, in preparation).

We also assessed the existence of an exsolved pre-eruptive volatile phase by comparing the SO₂ mass that was degassed during the eruption as measured by satellite data (e.g. Carn 2019), and that which can be produced by degassing of the melt alone (e.g. by the petrologic method, see methods section above). We used the total SO₂ gas detected by satellite over the course of the eruption from February to March 1990 (~0.17 Mt of SO₂; Carn et al. 2015). After correcting for post-entrapment crystallization, pyroxene-hosted melt inclusions have 440–1110 mg/kg SO₂ (Table 8) whereas the interstitial melt has about 60–90 SO₂ mg/kg. Considering a crystal content of 30–47 vol.%, and a DRE volume of 0.13 km³ we find that about 0.05–0.18 Mt SO₂ could have been outgassed from the melt during the eruption (Table 8). These values partly overlap with but is lower than the amount detected from space-borne sensing. This implies that about 0.01–0.12 Mt SO₂ could be stored in an exsolved fluid

phase at pre-eruptive conditions. We acknowledge that these values are likely underestimates of the true amount of exsolved fluid phase from the eruption. Uncertainties include maximum melt inclusion SO₂ content being unknown, errors related to the SO₂ detected from satellite, instrumental errors during analyses of glasses on the EPMA, and uncertainties related to erupted volume estimation from deposits used to scale the degassed SO₂ content.

## Timescales of magmatic processes

We find that Mg/Fe concentration profiles of reversely zoned Opx crystals in the main- and late-stage pumices yield overlapping timescales (Fig. 12). Timescales of Opx crystals range from about a couple of weeks to up to three months using the Mg-Fe inter-diffusion coefficient from Dohmen *et al.* (2016), which are about a factor of 3 shorter when using the diffusion coefficient from Ganguly & Tazzoli (1994). On the other hand, timescales from Cpx profiles vary from
<figure><img src="image_11.png" /><figcaption>**Fig. 10.** P–T storage condition of the 1990 magma estimated from amphibole composition in pumices from the main eruptive stage using the different geothermometers. Storage conditions were computed by using the Amph- melt geothermometer of Putirka (2016) (± 30°C; ± 100 MPa), as well as the Amph chemometric formulations by Ridolfi et al. (2010) (± 22°C; ± 210 MPa) and Ridolfi & Renzulli (2012) (± 24°C; ± 91 MPa).</figcaption></figure>

one to three months. Bearing this in mind, the Opx timescales better overlap better with those of Cpx when using the diffusion coefficient of Dohmen et al. (2016) rather than Ganguly & Tazzoli (1994). Hereafter we refer to the computed timescales from Dohmen et al. (2016), which appear more consistent with the recorded length of the pre-eruptive signals.

DISCUSSION

A range of pre-eruptive storage conditions in Kelud's plumbing system

Kelud volcano has consistently erupted basaltic andesite magma for over a hundred years (Bourdier et al., 1997; Humaida, 2013; Jeffery et al., 2013). The main-stage and late-stage pumice samples from the 1990 eruption have both this same composition and display miner- alogical and textural features broadly similar to previously erupted products. However, the late-stage pumices are more crystal-rich, have a more evolved glass composition, and record lower temperatures than the main stage pumices, and thus the magma was stored at cooler pre-eruptive conditions.

Our geothermobarometry results indicate that the highest temper-
atures are about 1050°C and recorded by Opx and Cpx cores in both
main-stage and late-stage pumices (Fig. 8), and match the liquidus
temperature of the whole-rock with 3 wt.% water in the melt. These
temperatures are also consistent with the estimates of Cassidy et al.
(2019, 2016) for the 2014 eruption using the Plag-melt equilibria
and experiments. Pyroxene rim compositions give lower temperatures
of around 980°C. Isobaric MELTS modelling at 100 MPa and with
3 wt.% water in the melt indicates that at the above temperature
range plagioclase is An₈₄-₈₉, which is the composition of the mantle

zones of the Plag phenocrysts. Amphibole temperatures calculated
following Putirka (2016) are similar to that of pyroxene rims (about
970°C), but much lower temperatures of about 800°C were obtained
with the Mag-Ilm equilibria in the late-stage samples. Colder storage
conditions for the late-stage magma are also indicated by the presence
of quartz and the lower An (= 40) of Plag rims compared to those
from the main-stage Plag rims (An₅₈-₇₀; Fig. 7a). Oxygen fugacity
constrained from the Fe-Ti oxide equilibria in the late-stage pumice
indicate relatively oxidizing conditions of about 1 log unit above the
NNO buffer.

Mineral-melt equilibria data indicates that the melt could have
contained up to 3–4 wt.% dissolved water, which is high enough
to stabilize amphibole in a range of typical arc compositions (Costa
et al., 2013; Sisson & Grove, 1993). Although amphibole phenocrysts
were found in the rocks, amphibole pseudomorphs are more com-
mon, and thus magmas were stored close to the stability field of
amphibole. Our pre-eruptive pressure estimates from Amph and Plag
hygrometry-H₂O/CO₂ saturation suggest that Kelud's plumbing sys-
tem also has a significant vertical extent. Volatile saturation pressures
(110–170 MPa) correspond to 4–9 km depth (using an average local
crustal density of 2640 kg m⁻³; Smyth et al., 2007). A caveat to
these depths is that given the minimal value represented CO₂ content
used in these calculations from Cassidy et al. (2016), these pressures
and depth estimates are likely minima values. Recalculation of these
saturation and pressures with the scaled fluxed CO₂ content from
VolatileCalc and zoned Opx (437 ppm) increases the saturation
pressure to 160–190 MPa, which corresponds to 6–8 km depth
using the same local crustal density (Table 6). On the other hand,
pressure estimates from amphibole phenocrysts correspond to about
15–19 km depth (Fig. 10). Previous studies have suggested that the
magma plumbing system of at Kelud may extend to 30 km depth
beneath the volcano (Jeffery et al., 2013) so it is possible that part
of the crystal cargo from the 1990 eruption could have formed
deeper than we have estimated. Thus, our data and those from
the literature suggest that the magma plumbing system of Kelud
is extended vertically and includes thermally zoned magma storage
levels of at least 15 km (see Fig. 11). In addition, mass balance
calculations of the sulphur budget (see section 4.3) indicate that the
magma coexisted with an exsolved volatile phase at pre-eruptive
conditions.

In summary, we propose that before the 1990 eruption, the Kelud plumbing system contained melts and crystals from about 4 km but up to 19 km depth below the crater, with temperatures ranging from 980-1030°C, and water contents in the melt of 2-4 wt.% from which the main-stage magma originated (Fig. 11). The reservoir also contained an exsolved fluid phase before magma ascent towards the surface. The geometry of the plumbing system was such that allowed for the coexistence of magmas with a range of crystallinities and interstitial melt composition. The hotter and less crystal-rich of the central parts erupted as the main phase magma, whereas colder (down to ~800°C), and more crystal-rich parts of the margins produced the late-stage magmas. Hence, it is likely that the magma reservoir beneath Kelud is thermally zoned, with storage temperatures varying from ~750°C (lowest estimate from Mag-Ilm) at its margins to ~980°C in more inner zones (average of pyroxene-melt rim temperatures), and to ~1030°C in its central part (average of pyroxene-melt core temperatures). Such range of temperatures can also explain the abundant presence of a variety of glomerocrysts that could be parts of the higher crystallinity and colder regions of the plumbing system (e.g. Nakada et al., 1994).
<figure><img src="image_12.png" /><figcaption>Fig. 11. Schematic cartoon of Kelud’s plumbing system according to petrologic observations and geothermobarometric estimates and of the pre-eruptive processes that most likely drove the 1990 eruption. The magma reservoir is likely to be zoned both in temperature (°C) and water content (wt.%). Inset white box summarizes the corresponding mineral textures and compositions. Distinct magma types from the main and late stages along with their crystal cargoes are shown in the drawing. The fluxing of CO₂ dominated fluid from depth as well as the creation of an exsolved fluid phase in the shallow parts of the system are also highlighted and considered key for driving the short-lived explosive eruptions of Kelud. The brick-like pattern represents the volcaniclastic crustal sediments and carbonates that are found underneath Kelud (Jeffery et al., 2013 and references therein).</figcaption></figure>

## Magmatic processes and timescales related to the 1990 eruption

### Relating mineral textures and chemistry to pre-eruptive processes

Many studies have proposed that intrusion of primitive magma and mixing are the main processes commonly driving volcanic eruptions, as the influx of more mafic hotter and volatile-rich magma can build enough overpressure (Costa & Chakraborty, 2004; Folch & Martí, 1998; Huppert et al., 1982). Moreover, mafic magma intrusion has also been proposed based on textural evidence of the crystal cargo, including multimodal core and rim mineral compositions, reverse zoning of Mg-number in mafic minerals such as pyroxenes, and of An in plagioclase (Streck, 2008). We found that Opx and Cpx show limited range in compositional zoning in Mg-number or are reversely zoned, with rims that have higher Mg-number than the cores. This is especially true for Opx in the main-stage magma (23% of the Opx population surveyed). Although the reverse zoning in Opx could be interpreted as evidence of input of more primitive melt before the eruption, we would also expect to find reverse zoning in plagioclase, but this is not the case. We also did not find the typical bimodal crystal populations that appears in many systems where intrusion of primitive magma has been recognized. Moreover, the calculated temperature of the Opx rims and the cores overlaps, or the rims yield somewhat lower temperatures despite having high Mg-number. An alternative possibility in relating these observations with their magmatic processes is that the reverse zoning in Opx reflects an

increase in oxygen fugacity and/or a decrease of the water content in the melt. Waters & Lange (2017) found a systematic increase of the Mg-number of the Opx as the water content in the melt decreases, and they experimentally calibrated a relationship between the two variables. We have exploited this finding by applying the experimental $Fe^{2+}-MgK_d$ relationship to find the change in water composition corresponding to the change observed between rim and core compositions i.e.:

$$ \Delta Fe^{2+}-MgK_d = \left( \left[ \left( X_{MgO}^{liquid} X_{FeO}^{opx} \right) / \left[ X_{FeO}^{liquid} X_{MgO}^{opx} \right] \right)_{core} - \left( \left[ \left( X_{MgO}^{liquid} X_{FeO}^{opx} \right) / \left[ X_{FeO}^{liquid} X_{MgO}^{opx} \right] \right)_{rim} \right) \quad (2) $$

From this, it is apparent that a decrease of water in the melt of up to ~2 wt.% from the Opx core to the rim (Fig. 9b) can explain the increase of Mg-number of the Opx cores (67–70) to the rims (70–72). Such a water decrease would also lead to crystallization and normal zoning in plagioclase, and thus would explain several of the textural and compositional features of Opx and Plag. The observed water decrease in the melt could be related to magma degassing and ascent towards the surface, and/or fluxing of the reservoir by CO₂ dominated fluids from depth as we discuss below.
<figure><img src="image_13.png" /><figcaption>**Fig. 12.** (a) Histogram of calculated Opx and Cpx diffusion timescales (recalculated dates before the 1990 eruption day) compared to the pre-eruptive temporal evolution of (b) crater lake temperature (Vandemeulebrouck *et al.*, 2000); (c) seismicity (black histogram = number of VT events per hour, red line = cumulative seismic energy) (Lesage and Surono, 1995); and (d) hydroacoustics (audiometric and ultrasonic bands) signals (Vandemeulebrouck *et al.*, 2000). Vertical bands highlight the timing of seismic swarms and ae used a visual guidance and comparison to changes in other data. The missing data interval in lake temperature and hydroacoustics was due to power supply failure.</figcaption></figure>

### Fluid fluxing from depth and creation of an exsolved volatile phase

In recent years it has become apparent that in addition to replenishment by a more primitive melt, magma plumbing systems may be fluxed by hot volatiles arising from the deeper part of the magma system (Mastin and Witter, 2000; Spilliaert et al., 2006; Giuffrida et al., 2017; Flaherty et al., 2018; Hartung et al., 2019). Petrological analysis of melt inclusions and modeling support this possibility, and such fluid addition from depth can also account for the permanent volcanic plumes of many mafic volcanoes. Given the dependency of different volatile elements with pressure, the deeper parts of the system are typically CO₂-rich, because CO₂ is much less soluble than water and SO₂, and can thus degas earlier and flux through the upper parts of the system, with mechanisms that may include channeling (Caricchi et al., 2018; Collombet et al., 2021; Giuffrida et al., 2017; Spilliaert et al., 2006). CO₂ fluxing can have widespread effects on a shallow magmatic system that vary depending on the pressure, but include the crystallization of high-Mg, high-Ca pyroxene rims (Giuffrida et al., 2017). Thus, we propose that addition of CO₂-rich fluids from depth may have decreased the water fugacity and thus water content of the melt in the shallower parts of the Kelud magmatic system and may explain the reverse zoning of Opx, normally zoned Plag, and the Cpx overgrowths and reverse zoning of some Cpx in our samples (Fig. 4d).

### Mass balance constraints on fluid fluxing

Quantifying the CO₂ that may have fluxed through the system is challenging because it is a dynamic process. There are no measurements of pre-eruptive/syn-eruptive CO₂/SO₂ and H₂O/CO₂ values of the gases for the 1990 eruption, although some fumaroles were measured in 1978 (H₂O/CO₂ = 17) and after the eruption in August 1990 (H₂O/CO₂ = 7.7) (Global Volcanism Program, 1990). Moreover, although melt inclusion compositions can be in principle used to identify the CO₂ fluxing (Wallace, 2005) it is becoming increasingly apparent that a significant part of the CO₂ can be trapped in bubbles inside the inclusion, and thus the data is also difficult to interpret in terms of fluid fluxing. To test the CO₂ fluxing hypothesis we propose, we have performed a series of calculations based on some observations and mass balance as we describe below.

The change in Mg-Fe in reversely zoned Opx can be explained by a water loss of the melt from about 4 wt.% to about 2 wt.% (section 5.2.1, Fig. S6), which we propose it is due to addition of CO₂ of mafic magma from depth or limestone assimilation. Assuming that the melt is fluid saturated, we calculated the concentration of CO₂ that would need to be added to the melt to decrease the water content in the melt from about 4 to 2 wt.% (Table 6). For this we used two end-member cases of the initial (pre- fluxing) CO₂ concentration in the melt, one close to zero and another of 336 ppm as measured
<table><caption>**Table 6:** mass balance calculations for CO₂ dominated fluid flushing and comparison of accumulation times using flux rates Javanese volcanoes</caption><thead><tr><th colspan="2">1. Melt volatile content</th></tr></thead><tbody><tr><td><strong>1a. Initial melt volatile content</strong></td><td>Volatile content</td></tr><tr><td>H₂O (wt.%)</td><td>3.8–5</td></tr><tr><td>CO₂ (ppm)</td><td>336–0</td></tr><tr><td><strong>1b. Final melt volatile content</strong></td><td></td></tr><tr><td>H₂O (wt.%)</td><td>2.0</td></tr><tr><td>CO₂ (ppm)</td><td>759</td></tr><tr><td><strong>1c. Change in melt volatile content</strong></td><td></td></tr><tr><td>H₂O (wt.%)</td><td>1.8–3.0</td></tr><tr><td>CO₂ (ppm)</td><td>423–759</td></tr></tbody><thead><tr><th colspan="2">2. Mass balance of amount of CO₂ fluxed</th></tr></thead><tbody><tr><td>Erupted volume (DRE) (km³)</td><td>0.13</td></tr><tr><td>Rock density (kg m⁻³)</td><td>2402</td></tr><tr><td>Mass of magma (Mt)</td><td>312</td></tr><tr><td>Melt proportion</td><td>0.58</td></tr><tr><td>Mass of melt (Mt)</td><td>180</td></tr><tr><td></td><td>Range</td></tr><tr><td>Fluxed CO₂ (ppm)</td><td>244–437</td></tr><tr><td>Fluxed CO₂ (Mt)</td><td>0.08–0.14</td></tr></tbody><thead><tr><th colspan="2">3. Amount of limestone vs. mafic magma degassing needed for CO₂ flushing</th></tr></thead><tbody><tr><td><strong>3a. Effect of limestone assimilation</strong></td><td></td></tr><tr><td>CO₂ in pure limestone (wt.%)</td><td>44</td></tr><tr><td>Mass of limestone (Mt)</td><td>0.13–0.23</td></tr><tr><td>Volume of limestone& (m³)</td><td>84 840</td></tr><tr><td><strong>3b. Effect of degassing basaltic magma<sup>#</sup></strong></td><td></td></tr><tr><td>CO₂ content in wt.%</td><td>Mass of basaltic melt (Mt)</td></tr><tr><td>0.5</td><td>15–31</td></tr><tr><td>1</td><td>27–55</td></tr><tr><td></td><td>Volume of basaltic melt* (m³)</td></tr><tr><td>0.5</td><td>5034–10 403</td></tr><tr><td>1</td><td>9060–18 456</td></tr></tbody><thead><tr><th colspan="2">4. Time needed to produce total 0.08–0.14 CO₂ using degassing data of Javanese volcanoes<sup>‡</sup></th></tr></thead><tbody><tr><td>Volcano</td><td>CO₂ degassing (tonnes/day)</td></tr><tr><td>Merapi</td><td>104</td></tr><tr><td>Bromo-Semeru</td><td>2184</td></tr><tr><td>Raung-Ijen</td><td>1111</td></tr><tr><td></td><td>Time (days)</td></tr><tr><td>Merapi</td><td>1346</td></tr><tr><td>Bromo-Semeru</td><td>64</td></tr><tr><td>Raung-Ijen</td><td>126</td></tr></tbody></table>

in melt inclusions from the 2014 magma (Cassidy et al., 2016). We found that for these two end members, the melt would have gained about 420 ppm to 760 ppm CO₂. Taking into account the crystal proportions and volume of the eruption (DRE) (which corresponds to 180 Mt of erupted magma) gives that between about 0.08 to 0.14 Mt of CO₂ should have been added to the melt.

To further test whether these are realistic amounts of CO₂ we also calculated: (1) the mass of pure limestone and basaltic magma that would be need to degas to generate such a mass of CO₂, and (2) the time that would be required to degas such amount of CO₂ using the degassing fluxes from other open vent volcanoes in Java. We describe these below in two sections., and where our calculations are summarized in Table 6. Such sources of CO₂-dominated fluids could come from the degassing of deep-seated basaltic magma, for example at Cerro Negro (Lowenstern, 2001) or Galunggung (Sisson & Bronto, 1998), or from assimilation of limestone or volcaniclastic crustal sediment underlying Kelud that commonly underlie active volcanoes along the Sunda Arc. Calc-silicate skarn xenoliths have been found in rocks from the more recent 2007 dome (Jeffery et al., 2013), and crustal limestones also outcrops in some localities around East Java (e.g. Smyth et al., 2007). Jeffery et al. (2013) suggested that crustal assimilation likely came from volcaniclastic crustal sediment making up most of the country rock (~25% contamination), with limestone and skarns possibly making up smaller portions, as well. This contrasts with other volcanoes such as Merapi, where limestone is the dominant source of crustal contaminant (Handley et al., 2014). The process of crustal CO₂ assimilation reaction at volcanoes such as Merapi (Central Java, Indonesia) is shown to enhance the intensity of explosive eruptions by a factor of 3–5 (Troll et al., 2012), and could make a significant contribution to the overall volatile budget of the magma (Deegan et al., 2010). Thus, we calculated the amount of mass contribution of either limestone assimilation or basalt degassing from depth to produce the change in CO₂ in the melt that we have estimated (Table 6). Here we assume that the entirety of the CO₂ exsolved originates either from limestone assimilation or degassing of basaltic magma. We used the stoichiometric concentration of CO₂ in pure limestone, and basaltic magma containing 0.5–1 wt.% CO₂. We find that it would require up to 0.13–0.23 Mt of limestone or CO₂ from 15–55 Mt basaltic magma, although this is likely an underestimation. These masses are much smaller than the total amount of magma erupted, and thus they seem realistic.

Finally, it is also worth evaluating whether accumulating such an amount of CO₂ fluid before eruption is realistic. To do this, we consider CO₂ fluxes from passively degassing volcanic plumes of volcanoes that are similar to Kelud in terms of magma composition, tectonic setting, and geographical location, as we do not have CO₂ fluxes from Kelud itself. Aiuppa et al. (2019) reported 104 tonnes/day of CO₂ for Merapi, 1.1 kt/day for Raung-Ijen, and 2.1 kt/day for Bromo-Semeru. Using a range of these values (Table 6) we find that to flux 0.14 Mt would require CO₂ fluxes between that of Bromo-Semeru (2.2 kt/day or 64 days) and Raung-Ijen (1.1 kt/day or 126 days). These times are short compared to the inter-eruptive repose at Kelud of several years, and thus our estimates of accumulation time also seem realistic. In summary, we believe that it is possible that the estimated amount of fluids can flux through Kelud's magmatic system and accumulate months before eruption.

Regardless of its source, volatile fluxing can thus affect the textures and compositions of minerals without necessarily affecting the bulk rock composition, if no crystal cargo were brought up with the volatiles. Most importantly, CO₂ fluxing favors the exsolution of water and thus a dehydration of the melt (Spilliaert et al., 2006), thereby promoting enhanced magma crystallization as well as the

<footnote><sup>#</sup>Maximum CO₂ content taken from Cassidy et al. (2016)</footnote>

<footnote><sup>‡</sup>Range of CO₂ content taken from arc basalts (min CO₂ from modal CO₂ content from Galunggung [Sisson and Bronto, 1988]; max CO₂ content from Cerro Negro [Lowenstern, 2001])</footnote>

<footnote><sup>‡</sup>CO₂ fluxes from Javanese volcanoes Aiuppa et al. (2019)</footnote>

<footnote><sup>‡</sup>Density of pure calcite = 2711 kg m⁻³ (https://www.mindat.org/min-859.html)</footnote>

<footnote><sup>*</sup>Density of wet basalt = 2980 kg m⁻³ (Sharma, 1997, p. 17)</footnote>
<table><caption>**Table 7:** Petrologic calculations of outgassed SO₂ with satellite (TOMS) detection of SO₂ emissions, following the methodology of Jay *et al.* (2014)</caption><thead><tr><th></th><th>KELUD1990-PL1</th><th>90-23</th><th>Units</th></tr></thead><tbody><tr><td colspan="4"><strong>1. Kelud's excess SO₂ gas</strong></td></tr><tr><td>Erupted volume (DRE)</td><td>0.13</td><td>0.13</td><td>km³</td></tr><tr><td>Rock density</td><td>2352</td><td>2402</td><td>kg/m³</td></tr><tr><td>Mass of magma</td><td>306</td><td>312</td><td>Mt</td></tr><tr><td>Concentration of SO₂ melt Inclusion</td><td>1110</td><td>440</td><td>ppm</td></tr><tr><td>Concentration of SO₂ interstitial glass</td><td>60</td><td>90</td><td>ppm</td></tr><tr><td>Glass proportion</td><td>0.69</td><td>0.53</td><td></td></tr><tr><td colspan="4"><strong>2. Petrologic calculations</strong></td></tr><tr><td>SO₂ degassed from the melt</td><td>0.18</td><td>0.05</td><td>Mt</td></tr><tr><td colspan="4"><strong>3. Volcanic gas satellite measurements</strong></td></tr><tr><td>Atmospheric SO₂</td><td>0.17</td><td>0.17</td><td>Mt</td></tr><tr><td>Excess SO₂ exsolved from melt</td><td>0.01</td><td>0.12</td><td>Mt</td></tr></tbody></table>

accumulation of an exsolved fluid phase. Fluid addition dominated by CO₂ from depth or through crustal contamination plus crystallization at the shallow parts of the plumbing system can lead to melt supersaturation in volatiles, and the creation of an exsolved volatile phase, for which we have found also evidence using the SO₂ budget of the eruption (Table 7). Given that Kelud is a semi-plugged stratocone (Whelley *et al.*, 2015), fluid accumulation in the reservoir is possible and may lead to the overpressure necessary to promote explosive eruption (Stock *et al.*, 2016). Various studies of the 2014 eruption have also proposed that this eruption of Kelud was probably driven by volatile accumulation rather than magma replenishment, in a similar manner to that of the 2014 eruption and the Calbuco 2015 eruption in Chile (Arzilli *et al.*, 2019; Cassidy *et al.*, 2019; Namur *et al.*, 2020; Pardini *et al.*, 2018).

## Time scales of processes and relation to monitoring data

The monitoring records that we have compiled may also be related to fluid movement and accumulation prior to the 1990 eruption and which were captured as unrest signals. Vandemeulebrouck *et al.* (2000) showed that the baseline low frequency hydroacoustic noise started to increase as far back as April 1989. Although the baseline increase was lower than the unrest observed between November and February, this could indicate that the volatile pressure had increased enough to allow gases to leak to the surface and reach the bottom of the lake. In the November-February unrest period there is a stepwise increase in lake temperature, with a temporary dip in January due to a short rainfall episode (Fig 12b). This trend in lake temperature is accompanied by a significant increase in the hydroacoustic noise level in the ultrasonic band after the November swarm (Fig 12d). However, ultrasonic noise levels declined after the January swarm until just before the eruption, in contrast with the gradual increase in audiometric noise level. Vandemeulebrouck *et al.* (2000) interpreted these variations as the hydrothermal system opening up to prime for an eruption, with the drop in January reflecting a large outgassing episode into the lake.

In addition to the hydroacoustic record, the unrest include two major VT seismic swarms in November and January, accompanied by sharp pulses in seismic energy release (Fig 12c) (Lesage & Surono, 1995). Between these swarms, there is incremental increase in both

number of VT earthquakes and a minor spike in seismic energy release. Lesage & Surono (1995) interpreted these patterns in seismicity as originating from the brittle plug fracturing and opening of the system. Thus, it seems possible that the plug was fractured from a combination of pulses of fluid fluxing (seismic swarms) and accumulation of fluids exerting overpressure on the plug (incremental increase in seismicity). Our interpretation of the changes in Mg-number of pyroxene (reversely zoned crystals) is that they record a decrease in the water content in the melt due to CO₂-dominated fluids from depth (basaltic magma or limestone degassing) rather than more primitive magma intrusion. Our models of equilibration of the Fe/Mg zoning of these same pyroxenes gave times of 2-3 months before the eruption, which is within the duration of unrest signals in both the seismicity (Lesage & Surono, 1995), crater lake temperature, and hydroacoustics (Vandemeulebrouck *et al.*, 2000) (Fig. 12a-d). However, we do not observe an exact match between timing of the crystals and the changes of the monitoring parameters. Crystal zoning timescales precede the recorded seismic swarms or spikes in seismic energy by up to a week, and steep increases of the lake temperature and/or hydroacoustic signals by up to two weeks. This is different from other studies where the match between the crystal zoning and seismicity has been reported (Costa *et al.*, 2020; Kahl *et al.*, 2011; Saunders *et al.*, 2012). Although we do not have enough data in crystal zoning to make it a statistically robust observation, such a time difference could be related to the lag of time between the phenomena of fluid percolation from depth and the processes that are responsible for the changes in seismic data and water lake changes. Finally, we emphasize that our diffusion timescales are considerably longer than what would be expected for a simple process of magma decompression and ascent from depth and thus the zoning recorded in the pyroxenes is more likely due to the fluid fluxing than simple magma ascent. As noted in the previous section, it seems possible to add the necessary mass of CO₂ to the system in times of a few months and thus it is compatible with duration of the unrest and also with the time scales of the crystals.

## A general model for Kelud eruptions and their relation to other volcanoes

We believe that a simple conceptual model to explain the explosive (e.g. VEI 4) but short-lived eruptions, that have limited unrest until
very close to the eruption is the accumulation of a pre-existing volatile phase over time, and especially the addition of fluids from depths that may overpressure the system to fracture the plug. The resulting eruption is highly explosive but very brief, because it is mainly driven by accumulated fluids in the upper parts of the system. The monitoring data show that for volcanoes with a water lake, it might be possible to identify when the fluid pressure has become large enough for gases to start leaking towards the surface and affect the hydroacoustics and lake temperature. Although changes in seismicity occur much closer to eruption, especially the exponential increases that are typically associated with run-away processes and failure of the rocks (e.g. Nakamichi et al., 2017).

Fluid accumulation may also explain why these eruptions also show limited unrest in terms of signals and duration, whereas we would expect renewed intrusions with more primitive magma and the significant addition of mass to have much stronger signals, especially seismicity as seen at other volcanoes such as Soufrière Hills, Montserrat (Murphy et al., 1998; Plail et al., 2018), Tungurahua (Myers et al., 2014; Samaniego et al., 2011), and Ruapehu (Conway et al., 2018; Kilgour et al., 2013). From a petrologic and monitoring perspective, we can consider the 2014 eruption of Kelud as an intra-volcanic analogue to the 1990, as it was quite similar in many volcanological and petrological attributes to the 1990 eruption. In turn, the 1990 and 2014 eruptions are similar to the VEI 4 explosive eruption of Calbuco volcano in 2015. Prior to the 2015 eruption, Calbuco also had a protracted repose time with the last recorded activity in 1996, yet experienced a short and subdued unrest of only 2–3 hours before erupting explosively (VEI 4) (Global Volcanism Program, 2015). The similarities between these different eruptions and volcanoes may indicate that they operate by similar processes, and that fluid fluxing from depth accompanied by accumulation of a pre-existing volatile phase can be a major scenario to be considered for future events (e.g. Petrelli et al., 2018).

## CONCLUSION

The combined petrological and monitoring data we report for the Kelud 1990 eruption suggest that the explosive but short-lived event was probably driven by accumulation of exsolved fluids at the top of the plumbing system. Fluxing of CO₂-dominated fluids from depth was likely a main contributor to the over pressures that eventually broke a plugged conduit necessary for the eruption to occur. The fluid fluxing promoted partial dehydration of the magma, leading to crystallization and second boiling and additional fluid exsolution. Diffusion profiles in zoned pyroxene crystals allow us to infer fluid fluxing and overpressure build up in the shallow magmatic system began 2–3 months before the eruption. This time frame which matches with the onset and duration of both seismic and crater lake unrest signals. The conceptual model we propose is likely applicable to other short-lived but explosive eruptions with limited unrest such as the 2014 Kelud eruption or the 2015 Calbuco in Chile.

## SUPPLEMENTARY DATA

Supplementary data are available at *Journal of Petrology* online.

## ACKNOWLEDGEMENTS

We thank C.Widiwijayanti, S.Jenkins, J.Herrin, J. Oalmann, L. Cheng, G. Fabbro, R.I. Made, D. Schonwalder, L. Siang Joo, and E. Y. Yin for assistance

with sample preparation, analyses and discussions on the Kelud volcano system. We also thank student assistants W. Wijaya and J. Fu for digitizing historical unrest records. Thanks also go to J. Vandemeulebrouck for providing the original dataset for the hydroacoustic measurements. We thank Fabio Arzilli, Frances Deegan, and an anonymous reviewer for their insightful comments that improved our manuscript considerably. This research was supported by the Earth Observatory of Singapore via its funding from the National Research Foundation Singapore and the Singapore Ministry of Education under the Research Centres of Excellence initiative. F. Costa was funded by NRFi Award (NRF-NRFI2017-06). The data underlying this article are available in the article and in its online supplementary material.

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