RESEARCH ARTICLE

Experimental constraints on syneruptive magma ascentrelated to the phreatomagmatic phase of the 2000ADeruption of Usu volcano, Japan

Yuki Suzuki & James E. Gardner & Jessica F. Larsen

Received: 30 March 2005 /Accepted: 4 June 2006 / Published online: 30 August 2006# Springer-Verlag 2006

Abstract We experimentally studied the dacitic magmaejected during the first event in the Usu 2000 eruption toinvestigate the conditions of syneruptive magmatic ascent.Geophysical data revealed that the magma reached underWest Nishiyama, the location of the event’s craters, afterrising beneath the summit. Prior study of bubble-sizedistributions of ejecta shows two stages (stage 1 and stage2) with different magma ascent rates, as the magmaaccelerated beneath West Nishiyama with the start of thesecond stage. To simulate ascent of stage 1 from the mainreservoir, which was located at a depth of 4–6 km(125 MPa) to 2 km (50 MPa) beneath West Nishiyama,decompression experiments were conducted isothermally at900°C following two paths. Single step decompression(SSD) samples were decompressed rapidly (0.67 MPa/s) totheir final pressure and held for 12 to 144 hours. Multiplestep decompression (MSD) samples were decompressedstepwise to their final pressure and quenched instantly. InMSD, the average decompression rates and total experi-

mental durations varied between 0.01389 to 0.00015 MPa/sand 1.5 to 144 hours, respectively. Syneruptive crystalliza-tion was confined to stage 1, and the conditions of ascentwere determined by documenting similarities in decom-pression-induced crystallization between ejecta and experi-ments. Core compositions, number densities, and shapes ofexperimental microlites indicate that ascent to 2 km depthoccurred in less than 1.5 h. Volumes and number densitiesof experimental microlites from the SSD experiments thatbest replicate the decompression rate to 2 km indicate thatthe magma remained at 2 km for approximately 24 h beforethe eruption. Stagnation at a depth of 2 km correspondswith horizontal transport through a dike from beneath thesummit to West Nishiyama, according to geodetic results.The total magma transport timescale including stage 2 istens of hours and is shorter than the timescale of precursoryseismicity (3.5 days), indicating that the erupted magma didnot move out of the reservoir for the first 2 days. This isconsistent with the temporal change in numbers of earth-quakes, which reached a peak after 2 days.

Keywords Usu 2000 eruption . Syneruptive magma ascent .

Decompression-induced crystallization .

Groundmass microlite . Textual analyses .

Decompression experiments . Eruption trigger

Introduction

The rate and manner of syneruptive magma ascent from areservoir are crucial parameters that influence eruptivebehavior. For example, eruptive style can switch betweenexplosive and effusive, depending on magma ascent rate(Eichelberger et al. 1986; Jaupart and Allègre 1991; Woodsand Koyaguchi 1994; Jaupart 1998). In ascending magma,

Bull Volcanol (2007) 69:423–444DOI 10.1007/s00445-006-0084-3

Editorial responsibility: DB Dingwell

Y. Suzuki (*)Institute of Mineralogy, Petrology and Economic Geology,Graduate School of Science, Tohoku University,Aoba-ku, Sendai 980-8578, Japane-mail: suzukiyk@ganko.tohoku.ac.jp

J. E. GardnerDepartment of Geological Sciences, Jackson Schoolof Geosciences, The University of Texas at Austin,1 University Station C1100,Austin, TX 78712-0254, USA

J. F. LarsenGeophysical Institute, University of Alaska Fairbanks,903 Koyukuk Drive,Fairbanks, AK 99775-7320, USA

a gas phase evolves in response to the decreasing volatilesolubility and gas expansion. How easily this gas phaseescapes from magma before reaching the surface partlydetermines the eruption style. Along with the permeabilityof magma and the conduit walls, magma ascent ratecontrols the ease of degassing. Understanding the styleand rate of magmatic ascent is useful for eruptionforecasting especially for the initial phase of an eruption,because such knowledge can reveal how an eruption mayevolve. Thus, quantification of the rate and manner ofascent is important for a better understanding of how escapeof gas influences eruption style, and it helps reveal theeruption mechanism.

One approach to investigating syneruptive magma ascentis to quantify the crystal and bubble textures of ejectaformed in response to water exsolution from melt (e.g.,Rutherford and Gardner 2000). Groundmass microlitescommonly crystallize during magmatic ascent, if ascent isslow enough. Cashman (1992) studied microlite texturesand showed that the amount of undercooling experiencedby magma during ascent is determined by water exsolutionrate, which is almost proportional to magma ascent rate.Experimental (e.g., Swanson 1977; Lofgren 1980) andtheoretical studies (e.g., Kirkpatrick 1981) found that thedegree of undercooling controls crystal nucleation rate andform. Since then, microlite textures in felsic ejecta havebeen related to eruptive parameters, such as repose intervaland effusion rate (e.g., Gardner et al. 1998; Nakada andMotomura 1999; Hammer et al. 1999, 2000).

Recently, experimental studies have quantified the kinet-ics of decompression-induced crystallization in rhyoliticmelts (e.g., Geschwind and Rutherford 1995; Hammer andRutherford 2002; Martel and Schmidt 2003; Couch et al.2003), but only a few of them examined magma ascentduring a particular eruption through experimental replicationof microlite textures (e.g., Geschwind and Rutherford 1995;Couch et al. 2003). The style and rate of magmatic ascentobtained through such work could improve our interpretationof eruption mechanisms by enabling us to combineobservations from the eruptive products and time-resolvedgeophysical observations, or in turn by providing us with anadditional method of interpreting geophysical results.

We focus here on the phreatomagmatic phase of the 2000AD eruption of the Usu volcano in Japan, because previousstudies have outlined the conditions of syneruptive ascent ofthe erupted dacitic magma (referred to as Usu 2000 magmahereafter). This will allow us to design experiments that willbest replicate the natural conditions of magma ascent. First,geophysical observations revealed that Usu 2000 magmaarrived beneath the western foot of Nishiyama (WestNishiyama; Fig. 1 hereafter) after rising beneath the summit.In West Nishiyama, craters from the phreatomagmatic phaseexist (Nishiyama craters in Fig. 1). Referring to the

geophysical results, a previous petrological and texturalinvestigation on ejecta discussed the evolution of magmaascent, including the period after it passed below WestNishiyama, and the resulting groundmass textures (Suzukiand Nakada 2001, 2002). Suzuki and Nakada (2002) showedthat the syneruptive magma ascent can be divided into twostages: stage 1, as the magma ascended from the reservoir tobelow West Nishiyama, and stage 2, with whose start themagma accelerated below West Nishiyama.

Some questions about the ascent of Usu 2000 magmaremain. For example, it was difficult to infer the mannerand timescale of magma ascent for stage 1 from thereservoir to below West Nishiyama, because precursoryseismicity was concentrated over a similar region before theeruption (Oshima et al. 2000; Oshima and Ui 2003). Also,storage conditions of the Usu 2000 magma were not knownwell. Although Tomiya and Miyagi (2002) proposed that asimilar magma plumbing system as in historical eruptionswas active in 2000, a phase equilibrium study would betterconstrain the locations of the magma storage before theeruption and reveal the magmatic ascent conditions.

The purpose of this study is to experimentally replicatemagma ascent before the phreatomagmatic phase of theUsu 2000 eruption focusing on stage 1. We first review theessential points of Suzuki and Nakada (2001, 2002), whichform the basis for this study. Then, newly obtained crystal-size distribution (CSD) for microlites in the ejecta areexamined to better infer that the decompression-inducedcrystallization is limited to stage 1 and the textures replicatedexperimentally correspond to those found in the naturalsamples. We propose that magma ascended from 4–6 to2 km beneath the summit in less than 1.5 h, then stayed at thatdepth for about 24 h. The repose at 2 km corresponds tohorizontal transport from below the summit to West

Fig. 1 Location of Usu volcano and distribution of fall deposits (withdate) after Urai et al. (2001)

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Nishiyama, as suggested by recent geodetic evidence(Okazaki et al. 2002). In addition, we also show the 2-daylag period between the start of precursory seismicity and theexperimentally constrained ascent start for erupted magma.Although this study focuses on the Usu 2000 eruption, weemphasize that the combined textural and compositionalconstraints developed through this work are useful to examinesyneruptive magma ascent in other similar magmas.

Usu 2000 eruption and its background

Sequence of the 2000 eruption

The 2000 eruption is the eighth historical eruption since1663 AD after a dormancy of several thousands of years. Uiet al. (2000, 2002) and Nakada (2001) have outlined thesequence of the 2000 eruption. After ∼3.5 days ofprecursory earthquakes and ground deformation (Table 1),the eruption started at 1:07 p.m. (JST) on 31 March fromthree newly formed craters at the western foot of Nishiyama(West Nishiyama) (Fig. 1). The 31 March event wasphreatomagmatic and the most voluminous and intenseevent during the 2000 eruption. Ohno et al. (2002) and

Takarada et al. (2002) report that during this event anintense gray eruption column and less intense white columnappeared alternately, accompanying ejection of ash andcinders and the formation of the craters. The gray eruptioncolumn reached as high as 3,500 m above the craters. It isnot known when this event stopped, as weather conditionsworsened after 5:25 p.m. The total mass of ejecta in the 31March event is about 2.2–2.4×108 kg. Tomiya et al. (2000)proposed that pumice and micropumice from this event arejuvenile, and Nakagawa et al. (2002) showed that they aredacitic (Table 2). The volume of micropumice in the ashranges from 20–30% (Nakagawa et al. 2002) to 50% (Tomiya etal. 2001). After the first event, intermittent phreatic eruptionsbegan in West Nishiyama and Konpirayama (Fig. 1) andcontinued until September 2001 (Takarada 2003). Craters inKonpirayama were newly formed from 1:40 p.m., 1 April2000.

Geophysical observations before and shortly after 31 March2000

Precursory seismicity before 31 March can be divided intothree phases based on frequency and distribution (Oshimaet al. 2000; Oshima and Ui 2003; Table 1). Earthquakes

Table 1 Timetable of geophysical observations and experimental constraints for Usu 2000 magma

Deep(10km) Shallow (4-6km) Deep (10km)

3/27

3/28 Phase-1over 6 to 4km deep

beneath summit

No data Slight inflation Stayed in reservoir Mostly stayed in reservoir

3/29 Phase-2 (ref. geodesy to reveal

same depth, deep phenomena)

also beneath

outside of summit

3/30 Phase-3 Ascent up to 2km deep in less than 1.5hourPeak of seismicity, Magma transport ← beneath summit (ref. seismicity) Ascent to shallower depth

almost same depth, through dike Deflation of large triggered by rise of shallow magma

original region plus (Summit to magnitude(~4/3) Stayed at 2km deep for ab. 24hour (ref. experimental result

3/31 extension toward W Nishiyama) transport through dike and geodesy for deep phenomena)

southwest and (ref. geodesy)

Eruption southeast a After Oshima et al. (2000) and Oshima and Ui (2003)b After Takahashi et al. (2000) and Okazaki et al. (2002) for shallow magma and Murakami et al. (2001) for deep magmac Note that the erupted magma (Usu 2000 magma) was supplied from shallow reservoir. Italic indicates geophysical result

reffered in interpretation of experimental results and other geophysical observations

Geodesy bSeismicitya Experimental constraints on magma transferc

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began at 6 km beneath the western end of the summit in theevening of 27 March. Soon, the distribution extended to adepth of 4 km, while the lower end remained at 6 km.Hypocenters were distributed only below the western endof the summit (phase 1, ∼the evening of 28 March). Then,in addition to the main activity beneath the western end ofthe summit, seismicity near the northern end of the summitbegan (phase 2, ∼the early evening on 29 March), but thedepths of the hypocenters were the same as in phase 1. Inphase 3, seismic activity reached a peak, but most earthquakesstill occurred at 4–6 km depth with a slight extension up to2 km. Vertical projections indicate that hypocenters duringphase 3 expanded to the southeast and southwest from belowthe western end of the summit, but the distribution to the northwas continuously fixed around the summit.

In addition, geodetic studies revealed possible magmaticmovement at different depths, depending on the density ofGPS observatory stations around the volcano. Murakami etal. (2001) measured broadscale crustal deformation, whichrecorded inflation and deflation of a source located at10 km deep beneath the summit (Table 1). To predictshallow magma activity, additional GPS stations wereinstalled 2 days after the start of precursory seismicity(Table 1). As a result, Takahashi et al. (2000) preliminarilyproposed that magma ascended in a dike that reachedbeneath West Nishimaya by 31 March. Additionally,Matsushima et al. (2000) recorded tilt changes indicating

that the magma remained 0.5 km below West Nishiyamajust before the event on 31 March.

Previous results on magma storage at Usu volcano

Tomiya (1995) proposed that dacitic and rhyolitic reservoirs,located at 4–6 and 10 km depth, respectively, have suppliedall magma ejected at Usu since 1663 AD. On the basis ofgeophysical data, Tomiya and Miyagi (2002) proposed thatthe deformation source at 10 km depth (Table 1; Murakamiet al. 2001) represents a deeper rhyolitic reservoir. They alsopropose that a shallower dacitic reservoir is confirmed by (1)a low P-wave velocity region (6 km deep; Onizawa et al.2002), (2) an aseismic region (4–6 km deep; Oshima et al.2000; Oshima and Ui 2003), and (3) long-period volcanictremor (5 km deep; Yamamoto et al. 2002). They alsoproposed that Usu 2000 magma was derived solely from theshallower reservoir, based on the unimodal compositionaldistribution of magnetite phenocrysts. The lack of magmamixing was supported by additional analyses of otherphenocryst phases (Suzuki and Nakada 2001).

Tomiya (1995) also showed that magma temperatureincreased constantly during historic times. Although magmaserupted after 1853 lack magnetite–ilmenite pairs, correlationsbetween temperature and Mg/Mn ratios in magnetite frompre-1853 magma makes it possible to infer post-1853 magmatemperatures based on the calibration by Bacon andHirschmann (1988). Mg/Mn ratios in magnetite of the 2000ejecta are higher than those of the dacite magma (850–870°C)erupted in 1853 (Suzuki and Nakada 2001), indicatingthat the temperature of 2000 magma was 880°C or more.

Petrography of juvenile material erupted on 31 March2000

Suzuki and Nakada (2001) described juvenile micropumice(≦ several millimeters in length) and larger pumice (1.5 cmin maximum length), which were deposited at Field Site A(Fig. 1). Juvenile material can be divided into phenocrysts(≦4 mm, e.g., Fig. 2a) and groundmass (microlites plusglass, e.g., Fig. 2b, c). The abundance and composition ofphenocrysts and groundmass are the same between micro-pumice and pumice. This indicates that the erupted Usu2000 magma was homogeneous. Microlites are present ingreater number densities than phenocrysts, suggesting thatmicrolites nucleated during the ascent from the reservoir.The groundmass is interpreted to be the melt part of themagma just before ascent from the reservoir.

Phenocrysts (≦4 mm, plagioclase) and microphenocrysts(≦400 μm, plagioclase + orthopyroxene + magnetite + apatite)are euhedral, indicating equilibrium with melt before ascent.

Table 2 Compositions of juvenile material in Usu 2000 eruption fromwhich starting materials for phase equilibrium and decompressionexperiments were made

Wt%a Whole rockb GMSc GMS glassd

SiO2 69.43 74.73 79.80 (0.42)TiO2 0.48 0.43 0.43 (0.04)Al2O3 15.15 13.57 11.50 (0.20)FeO* 4.01 3.12 2.13 (0.13)MnO 0.17 0.14 0.10 (0.02)MgO 1.02 0.85 0.25 (0.03)CaO 4.05 2.92 1.57 (0.11)Na2O 4.61 3.19 2.88 (0.27)K2O 0.92 1.05 1.31 (0.05)P2O5 0.16 n.a. n.a.Ca/Na n.c. 0.51 n.c.N – – 59

aAll oxide values normalized to 100% and total iron as FeObAverage of XRF analyses (personal communication with Prof. M.Nakagawa)cGroundmass (GMS) composition was obtained using average volumefraction and composition of groundmass phases in micropumiceand pumice (after Suzuki and Nakada 2001). Composition of eachgroundmass phase was obtained with microprobe.dAverage composition and standard deviation were determined withmicroprobe analysis on several fragments of micropumiceand pumice (after Suzuki and Nakada 2001).n.a. Not analyzed, n.c. not calculated, N number of analyses

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Only plagioclase phenocrysts can be divided into two parts:sieved cores and clear mantles. In addition, for all phenocrystsand microphenocrysts, the outermost parts are defined as rims.The compositional ranges of plagioclase phenocryst mantlesare similar to those of plagioclase microphenocrysts. Theyshow compositional oscillation ranging from An47.0 to An76.0(with most below An65, Fig. 3), and the normal zoning toAn44.4–56.4 in the rims (up to 5 μm in width, Fig. 3).

Microlites (Fig. 2b, c) are, on average, 15.5 vol%plagioclase, 2.6 vol% pigeonite, and 0.3 vol% magnetite.The microlites are mostly euhedral, but plagioclase micro-lites are skeletal (hopper and swallowtail; see Fig. 2b, c).Because of their small size, microlite compositions areavailable only for core regions. Core compositions matchthose of phenocryst and microphenocryst rims. This similar-ity implies that the growth of phenocryst rims was partlysynchronous with nucleation and growth of microlites. Thus,for plagioclase, a zone just inside the rim (average An59.5 fornine analyses, 1σ=1.5) is considered to coexist with melt justbefore the ascent from the reservoir (Fig. 3).

Mass balance of the average compositions of groundmassmicrolites and glass (Table 2) yield a bulk groundmasscomposition of 74.7 wt% in SiO2 (Table 2). Compared toglass inclusions in microphenocrysts, the bulk groundmasshas higher MgO and lower SiO2 contents (Suzuki andNakada 2001). When the proposed increase in magmatictemperature since 1663 (Tomiya 1995) is considered, glassinclusions may not represent the melt composition justbefore syneruptive ascent related to Usu 2000 eruption.

Previous results for Usu 2000 magma ascent

Work by Suzuki and Nakada (2001) demonstrated that theUsu 2000 eruption produced pumice with a higher

Fig. 2 a–c Plagioclase and groundmass (GM) textures of ejecta inbackscattered electron images. a Microphenocryst (MP) from amicropumice (18.0 vol% vesicularity). b, c GM from a micropumice(47.0 vol% vesicularity) and a pumice, respectively. Microlites aremagnetite (Mt), pigeonite (Pig), and plagioclase in order of decreasingbrightness. Skeletal plagioclase microlite appears to show a swallowtail form (S.T.) or to have a hollow filled with glass (H.L.) dependingon the section of each crystal. Bubble coalescence can be detected bythe presence of waists and glass films connecting the waists

Fig. 3 Representative line analysis on plagioclase phenocryst (a) andmicrophenocryst (b). Modified after Suzuki and Nakada (2001). Solidand dotted lines represent data from pumice and micropumice,respectively. The arrows (not for rims) indicate parts that were inequilibrium with melt just before syneruptive ascent, so rims grew insyneruptive ascent (for detail, see text)

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vesicularity than micropumice (80 vs 15–65 vol%). Inaddition, the vesicularity of the micropumice variesinversely with the water content dissolved in the ground-mass glass (increases from 15 to 60 vol%, with decreasefrom 1.7 to 0.4 wt% H2O; Suzuki and Nakada 2001). Webelieve these differences resulted from variable quenchingand fragmentation of the magma during its interaction withshallow-level aquifers (Fig. 4a).

Vesicle coalescence is prominent in pumice and micro-pumice with ≧29 vol% vesicularities (Figs. 2b, c and 5a).Micropumices that are less vesicular lack coalescence(Fig. 5b), and they have bubble-size distributions consistingof two linear parts (Fig. 5a). Such distributions can beinterpreted to indicate that bubble nucleation was continu-ous but that the rate increased (Marsh 1988, 1998). Webelieve that the increased nucleation rate occurred when the

Fig. 4 Magma plumbing system beneath Usu volcano and ascentprocess (a) and textural evolution (b) of Usu 2000 magma suppliedfrom shallow reservoir. Modified from Suzuki and Nakada (2002). Fordeeper reservoir, see Tomiya (1995) and Tomiya and Miyagi (2002).There exist aquifers over less than 200 m deep below West Nishiyama(Yahata 2002). Similar crystallinity of micropumices solidified atvarious depths around the aquifer indicates no crystallization over thedepth of aquifers (Suzuki and Nakada 2001) (b). The mechanism forno crystallization is the time lag between the generation of theundercooling and crystallization response (Lasaga 1981; Lofgren

1980). For a depth below West Nishiyama, Suzuki and Nakada(2002) preferred 2 km (Oshima et al. 2000; Oshima and Ui 2003)rather than 0.5 km (Matsushima et al. 2000). Below West Nishiyama,water content in melt was equal to solubility at this depth because ofequilibrium degassing in stage 1 (a). In view of this, the 0.5-kmcontradicts with water content in groundmass glass of least-vesiculat-ed micropumice (1.7 wt%). Square brackets indicate our experimentalresults. In addition, this study clarified that syneruptive crystallizationis limited to stage 1

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magma accelerated just after passing below West Nishiyama(Fig. 4a). Ascent before that time was probably slow enough toallow equilibrium H2O exsolution from the melt, as suggestedby the precursory seismicity period (3.5 days). After passingbeneath West Nishiyama, however, ascent was too fast to allowequilibrium H2O exsolution from the melt, as shown by somemicropumice groundmasses containing more water than wouldbe expected if they had degassed in equilibrium up to the timethey were quenched by the aquifer. Hence, stage 1 and 2ascents differ by the acceleration of magma.

Materials and methods

Phase equilibria experiments

Isobaric and isothermal phase equilibria experiments wereconducted using René-style cold-seal pressure vessels at theExperimental Petrology Laboratory at the University ofAlaska Fairbanks. Starting materials for these experimentswere prepared by crushing 10 to 20 pumice blocks to a finepowder (crushed pumice; C.P. in Table 3). The reason forusing crushed, whole natural powders as starting materials,rather than fused glasses, is that inclusion of brokenphenocrysts best approximates the state of the magma.The magma reservoir conditions are better approximated if

phenocryst cores remain in the starting materials (J.Hammer, personal communication, 2005).

All experiments were run at water-saturated conditionswith fO2 of 0.5 to 1 log unit above NNO, fixed by areaction between the Ni-alloy vessel, Ni filler rod, and thewater pressurizing medium (Geschwind and Rutherford1992; Gardner et al. 1995). Experiments were quenched bya stream of compressed air followed by immersion in coldwater. Estimated quenching rates to reach the glasstransition using this method are tens of seconds. A set ofexperiments was run to verify the magmatic temperatureestimation of 880°C or more derived from previous work(Suzuki and Nakada 2001). These experiments determinedthe stability fields of major phenocryst phases over atemperature range of 850 to 900°C and a pressure range of100 to 200 MPa. In a second set of experiments, reversalexperiments were run to further verify the positions of thestability curves using melt-rich and crystal-rich experiments(Usu00-13 and Usu00-4 and 9, respectively; Table 3) toform melting and crystallization experimental pairs.

Ag70Pd30 tubing was used. The diameter of the tubingwas 2 mm, except for Usu00-13 (5 mm) to make startingmaterial for crystallization experiments together. Afterwelding one end of a capsule, starting material and enoughwater to ensure water saturation were added (∼2.0 or2.5 mg H2O added to ∼30 mg starting material for 100 and

Fig. 5 a Bubble-size distribu-tion (BSD) and b binary imageof bubble texture. Modifiedfrom Suzuki and Nakada (2002).a Numbers in legend showvesicularity. BSDs for less ve-siculated micropumices(∼26.0 vol%) are composed oftwo linear parts. Smaller bubblesdecrease and larger bubbles in-crease with increasing vesicu-larity of more than 29.1 vol%.Combined with texture(Fig. 2b), this can be interpretedto reflect coalescence. b Numberindicates vesicularity. E showsthe edge of micropumice

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200 MPa runs, respectively). The capsule was then weldedshut. To confirm sealing, the capsule was weighed beforeand after welding and then after heating on a hotplate to120°C for at least 30 min. Experimental durations were 92to 164.5 h (Table 3), which is sufficient to achieve localchemical equilibrium.

Decompression experiments

Decompression experiments employed the same finepowder of pumice (Table 2), capsule preparation tech-niques, fO2, and quenching procedure as the phaseequilibria runs described above. Small phenocryst frag-ments did not dissolve completely during 3–5 day runs at900°C and 125 MPa, our determined preeruptive startingconditions. This introduces larger crystal melt boundariesthan in the natural magma before ascent, which willenhance crystal growth at the expense of nucleation (Fokinet al. 1999). Thus, preparation of starting material fordecompression consisted of superheating the powder at900°C and 200 MPa for 2–3 days, ensuring that most of thephenocryst fragments are dissolved. After superheating,pressure was dropped, and the sample held for 4–5 days at900°C and 125 MPa, and then quenched. We started fromnatural pumice so that the reequilibration at the most

plausible storage condition (900°C and 125 MPa) canreproduce similar composition and amount of crystal andmelt as in natural reservoir magma. The preparation processof the starting material is similar to the run condition ofUsu00-19 (Table 3), so for detail, its result can be referred.In each capsule, enough material was equilibrated toprovide starting material for several decompression runs.After quenching, the experiments were gently broken apartinto several separate chunks that provided the startingmaterial for several decompression runs.

For each decompression run, ∼30 mg of the starting materialplus water was loaded into a 2-mm diameter Ag70Pd30 tubing,which was then welded shut. After thermally reequilibrating atstorage conditions for several hours, the experiments weredecompressed isothermally to 50 MPa. This pressure waschosen to replicate ascent to the 2 km depth (Fig. 4a). Thedecompression runs followed two pathways, after Hammerand Rutherford (2002), to replicate different possible styles ofmagmatic ascent. Each decompression path was initiated bydropping pressure using a hand-operated intensifier. The firstseries consisted of single-step decompressions (SSD), wheresamples were decompressed rapidly and continuously to50 MPa, and then held at that pressure for 12 to 144 h(Table 4). The decompression rate to 50 MPa was about0.67 MPa/s. Thus, the time for decompression (about 2 min)

Table 3 Conditions of isobaric experiments for Usu 2000 magma

Run Starting materiala Temperature (°C) Pressure (MPa) Duration (h) Products

Usu00-2 C.P. 850 200 141.5 G, Plb, Opx, Ox, HbUsu00-3 C.P. 900 100 159.75 G, Pl, Opx, OxUsu00-4 C.P. 875 100 159.75 G, Pl, Opx, OxUsu00-6 C.P. 875 200 143.5 G, Opx, OxUsu00-7 C.P. 900 150 119.5 G, Plb, Opx, OxUsu00-9c C.P. 850 150 116 G, Pld, Opx, OxUsu00-11 C.P. 860 200 113 G, Plb, Opx, OxUsu00-13 C.P. 900 200 92 GUsu00-14 Usu00-13 890 125 143.5 G, Pl, Opx, OxUsu00-15 Usu00-13 865 175 141 G, Pl, Opx, OxUsu00-16 Usu00-13 875 125 159.75 G, Pl, Opx, OxUsu00-17 Usu00-13 850 175 158.5 G, Ple,Opx, Ox, HbUsu00-18-1c C.P. 865 150 163.5 G, Pld, Opx, OxUsu00-18-2 Usu00-13 865 150 163.5 G, Ple, Opx, OxUsu00-19 Usu00-13 900 125 164.5 G, Pl, Opx, OxUsu00-20-1 Usu00-13 872 175 149 G, Ple, Opx, OxUsu00-20-2 Usu00-9 872 175 149 G, Plb, Opx, OxUsu00-21-1 Usu00-13 890 150 150 G, Pl, Opx, OxUsu00-21-2 Usu00-4 890 150 150 G, Plb, Opx, Ox

G Glass, Pl plagioclase, Opx orthopyroxene, Ox Fe–Ti oxides, Hb hornblendeaStarting material was either crushed pumice (C.P.) or run product.bRecognition of grown part was difficult.cPlagioclase can be stable (see text for detail).dNot euhedraleToo small to analyze (less than 4 μm in width)

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was much shorter than that at 50 MPa. The second seriesemployed multiple step decompressions (MSD), wheresamples were decompressed to 50 MPa in 10 MPa incrementsafter an initial 5 MPa decrease from 125 to 120 MPa. Sampleswere held at each lower pressure step for varying times toapproximate total decompression durations of 1.5 to 144 h(Table 4). Decompression at each step was conducted at about0.67 MPa/s. The average decompression rate during an MSDexperiment, derived by dividing the 75-MPa total pressuredrop by the time spent below 125 MPa, ranges from 0.00015to 0.01389 MPa/s (Table 4). We did not hold samples at50 MPa after slow decompression as conducted in MSD. Thisis because the conducted decompression experiments clarifiedthat decompression to 50 MPa should be as fast as in SSD.

Analytical and textural methods

Glass and mineral compositions were analyzed using aCameca SX-50 electron microprobe at University of AlaskaFairbanks. Major elements were determined using anaccelerating voltage of 15 kV and a beam current of10 nA. The beam diameter was set to 1 and 10 μm formineral and hydrous glass analyses, respectively. Countingtimes were 10 s for both peak and background. In hydrousglass analyses, a self-calibrating routine to correct for Namigration and increasing Al and Si counts was employed,which involved counting Na first and correcting theelement counts on each detector to the initial contents(Devine et al. 1995).

Backscattered electron (BSE) images of both naturalejecta and experimental products were used to determinevolume percent, number density, and size distribution of

minerals. BSE images for ejecta were taken using a JEOLJXA-8800R microprobe at the Earthquake Research Insti-tute, University of Tokyo, using conditions of 15 kV and12 nA. BSE images for experimental products were takenusing a Cameca SX-50 electron microprobe at University ofAlaska Fairbanks, under operating conditions of 20 kV and10 nA. To ensure the identification of the small microlites,BSE images were taken at 1,000× (at the University ofTokyo) or 800× (at the University of Alaska). Imageresolution is ∼0.5 μm per pixel. The BSE images wereanalyzed using NIH Image, which makes it possible toestimate microlite areas, lengths, widths, and numbers ineach binary image. Producing a binary image using NIHimage can be difficult when plagioclase has similargrayscale pixel values as the surrounding glass. In thiscase, the original BSE was manually edited using AdobePhotoshop so that the plagioclase crystals were clearlyseparated from the surrounding glass, before converting theimage to binary.

In all cases, BSE images were analyzed over a bubble-free reference area of 30,000–800,000 μm2, which issufficient to characterize each ejecta and experimentalproduct. The number of crystals analyzed for size distribu-tions and number densities ranged between 239 and 839,for plagioclase, and between 12 and 850 for Fe–Ti oxides.For the volume percent estimates, two-dimensional mea-surements were used because microlites were not oriented.In converting two-dimensional size distributions into threedimensions, we used the method of Sahagian and Proussevich(1998) for the nearly equidimensional Fe–Ti oxides andCSD correction 1.3.2 (Higgins 2000, 2002) for orthogonalplagioclase. To obtain two-dimensional size distributions

Table 4 Conditions of decompression experiments from 125 to 50 MPa at 900°C

Run Time held at 50 MPa (hour)a Time below 125 MPa (hour)b Time held at each step (hour) Decompression rate (MPa/s)c

Single step decompression(SSD)SSD-144 144 – – 0.67SSD-96 96 – – 0.67SSD-24 24 – – 0.67SSD-12 12 – – 0.67Multiple step decompression(MSD)MSD-144 – 144 20.6 0.00015 (0.67)MSD-96 – 96 13.7 0.00022 (0.67)MSD-60 – 24 8.6 0.00035 (0.67)MSD-12 – 12 1.7 0.00174 (0.67)MSD-6 – 6 0.9 0.00347 (0.67)MSD-1.5 – 1.5 0.2 0.01389 (0.67)

aTime for decompression to 50 MPa (about 2 min for all runs) is excluded.bTime for each decompression step (about 2 min through a run) is excluded.cFor SSD, decompression rate to 50 MPa is shown. For MSD, average rate through an MSD experiment and a rate used at each step (inparenthesis) are shown.

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for plagioclase, we measured the crystal widths. Applica-tion of this correction to plagioclase requires findingthe ratio of the short, intermediate, and long axes of thecrystals, following Higgins’ (1994) method, using thecrystal widths. In running CSD correction 1.3.2, the modelshape was set to be block, and the fabric was set to bemassive.

Results

Groundmass microlite CSD from juvenile fragments

Crystal form data are listed in Table 5. Microlite CSDsfrom different juvenile fragments resemble one another forboth plagioclase and Fe–Ti oxide (Fig. 6), suggesting that

Table 5 Microlite textural data for ejecta and experimental products

Sample no. Plagioclase Pyroxene Fe–Ti oxides

Vol% Number per unitarea (mm−2)

Number density(mm−3)

Form(S:I:L)a

Vol% Vol% Number per unitarea (mm−2)

Number density(mm−3)

Juvenile fragmentsMicropumice-14.6%b

16.0(5.0)c

4,476 (1439) 526,030 1:4:4 2.1 (0.5) 0.3(0.2)

347 (85.8) 64,807

Micropumice-29.1%

14.1(3.8)

5,499 (695) 518,883 1:5:5 3.2 (0.9) 0.3(0.3)

303 (43.0) 53,814

Micropumice-47.0%

16.3(2.2)

7,610 (1,378) 832,872 1:5:5 2.9 (0.6) 0.5(0.4)

342(117.1) 91,019

Pumice-82.0%

15.0(6.1)

5,106 (1,870) 1,295,170 1:3:3 2.3 (0.8) 0.2(0.5)

287 (99.7) 55,978

Average 15.5(0.9)d

5,673 (1,358) 793,239(365,229)

n.d. 2.6 (0.4) 0.3(0.1)

320 (29.0) 66,405 (17085)

Decompression productsSSD-144 18.6

(3.8)7,947 (3,922) 932,100 1:1:3 2.0 (1.4) 1.0

(0.4)925(454) 368,265

SSD-96 15.8(5.6)

8,850 (3,811) 1,647,680 1:3:3 2.2 (0.8) 1.2(0.3)

2,279 (663) 1,082,567

SSD-24 15.4(3.6)

5,764 (795) 497,940 1:1:3 1.8 (1.1) 0.6(0.3)

931 (719) 482,596

SSD-12 15.4(4.1)

7,546 (1,170) 750,149 1:1:3 1.1 (0.7) 0.1(0.1)

54 (80) 22,824

MSD-144 12.0(n.c.)e

3,419 (n.c.) (n.d.)f 1:2:2 0.8 (n.c.) 1.1(n.c.)

442 (n.c.) 136,714

MSD-96 19.6(n.c.)

5,381 (n.c.) (n.d.) 1:1:3 1.1 (n.c.) 1.4(n.c.)

2,414 (n.c.) 835,023

MSD-60 11.5(5.0)

7,221 (1,872) 2,962,090 1:2:2 1.1 (0.9) 1.4(0.3)

2,053 (254) 916,028

MSD-12 11.1(1.8)

15,504 (6,061) 7,132,250 1:2:2 3.2 (1.3) 1.8(0.6)

3,211 (597) 1,562,853

MSD-6 15.4(3.7)

34,667 (19,859) 13,212,880 1:2:2 3.9 (1.3) 1.5(0.4)

4,553 (599) 2,923,059

MSD-1.5 3.3(2.9)

8,583 (3,822) 1,324,950 1:1:3 0.4 (0.3) 0.7(0.3)

1,952 (891) 1,139,210

Volume and number are values per groundmass (for ejecta) and matrix (for run product). Volumes for juvenile fragments are from Suzuki andNakada (2001). For MSD-144 and MSD-96, the number of plagioclase microlite was calculated in a different way from other samples, asmicrolite cannot be distinguished. Firstly, the number in the whole analyzed area (preexisting crystal + matrix) was obtained by subtracting thenumber of preexisting crystal (300 mm−2 inferred from starting material) from number of all plagioclase. Then, this data was corrected withmatrix ratio in run product (91.8 vol%; inferred from starting material). For the same samples, plagioclase microlite volumes were calculated in asimilar way, using plagioclase volume in starting material (7.1 vol%). Standard deviation for number density is not shown, because each numberdensity was calculated from a CSD (Fig. 6).

aRatio of shortest, intermediate, and longest axes in three dimensionsbVesicularity for juvenile fragmentscStandard deviation for a set of data of the same number as the photographs analyzeddStandard deviation for four data from same number of juvenile fragmentseNot calculated, because one BSE image was too small to conduct the above calculationfNo data because CSD itself cannot be obtained

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plagioclase and Fe–Ti oxide microlite volumes in ground-mass do not change significantly among juvenile fragments,as previously suggested by Suzuki and Nakada (2001).CSD for plagioclase microlites are mostly linear, except forthe smallest sizes (Fig. 6). On the other hand, CSD forFe–Ti oxide microlites is negatively sloped at large sizesand positively sloped at small sizes (Fig. 6). Microlitenumber densities range from 5.2×105 to 1.3×106 mm–3 and5.4 to 9.1×104 mm−3 for plagioclase and Fe–Ti oxide,respectively (Table 5).

Phase equilibria experiments

Fe–Ti oxides and orthopyroxene were crystallized in allruns, except at superliquidus conditions of 900°C and

200 MPa (Table 3; Fig. 7). At slightly lower temperaturesand pressures, plagioclase is stabilized. In reversal exper-iment using melt-rich and crystal-rich experiments (Usu00-13 and Usu00-4 and 9, respectively; Table 3), stable phaseswere mostly in accordance with those using crushedpumice, except for two experiments. In Usu00-9 and 18-1,which used crushed pumice, plagioclase in the run product isnot euhedral (Table 3), although the conditions are far belowthe plagioclase stability curve, as deduced from crystalliza-tion experiments. This may be because plagioclase fragmentsof ca. max. An65 existed in the starting material of crushedpumice (Fig. 3). Anorthite-rich plagioclase may not be ableto react at relatively low temperatures and pressures,preventing crystallization of stable plagioclase. Equilibriumbetween crystal surface and melt should have been complet-

Fig. 6 Microlite CSDs forejecta (a) and decompressionproducts (b, c). For ejecta,numbers in legend show vesic-ularity. For plagioclase, thereexist additional classes at ahigher end only for ejecta. Thehorizontal axis represents thelongest axis of crystal in threedimensions. No CSD for MSD-144 (see text for detail). Grayareas represent ranges for ejec-ta. For Fe–Ti oxide, the samelines and symbols are used as inplagioclase

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ed in other experiments using crushed pumice. Hornblende,which is not found in ejecta, grew at relatively lowtemperatures and high water pressures (Fig. 7).

Chemical analysis was conducted focusing on plausiblemagmatic temperature (880°C or more). Stable plagioclaserims were analyzed when large enough, primarily in thecrystallization experiments, except when recognition of growthwas difficult in experiments using crushed pumice or in themelting experiments (Table 3). Compositions are usuallyhomogeneous to within <3 mol% An. In general, the averageAn content increases with water pressure and temperature(Figs. 7 and 8a). In addition, SiO2 content in experimentalglass decreases with increasing pressure at a fixed temperature(Fig. 8b, Table 6). Both An content and glass SiO2 content inthe experiment match those in the ejecta at similar conditions(between 125 and 150 MPa) at 890°C and 900°C (Fig. 8).Except for runs where high An fragment in crushed pumicehindered growth of equilibrium plagioclase (Usu00-9 and18-1; Table 3), we do not find a contradiction betweenresults from experiments using crushed pumice and crystal-

lization experiments employing Usu 00-13 (Table 3). Forexample, stable plagioclase composition changes systemat-ically with pressure and temperature (Fig. 7).

Results of decompression experiments

For run products, we define those crystals that nucleatedduring decompression as microlites, crystals that existedbefore decompression (equivalent to phenocryst in ejecta)as preexisting crystals, and residual glass + microlites(equivalent to groundmass in ejecta) as matrix. Alldecompression experiments contain plagioclase, clinopy-roxene, and Fe–Ti oxides as microlites (Fig. 9). Runproducts are homogeneous and microlite textures do notchange with respect to position within a given sample. Inmost decompression runs, microlites can be distinguishedfrom preexisting crystals by abrupt changes in size andnumber per unit area. We could not, however, clearlydistinguish plagioclase microlites in MSD-96 and MSD-144, because of the continuity in size and number density.

Fig. 7 Phase equilibria for Usu 2000 magma under H2O-saturatedcondition. Circles represent the use of crushed pumice. Left- and right-pointing triangles indicate crystallization and melting experiments,respectively. Each phase is stable over the low pressure side of thephase boundary, except hornblende. Filled symbols indicate thathornblende was stable. Fe–Ti oxides (Ox), orthopyroxene (Opx),plagioclase (Pl), and hornblende (Hb). Numbers below experimentalpoint indicate average An mol% of plagioclase rims (1σ of ∼1.5 for∼14 crystals). Plagioclase composition for 890°C and 150 MPa isfrom the crystallization experiment (Table 3). Based on probabletemperature and a fact that higher pH2O is required for a given Anmol% with decrease in temperature, the hatched region can be apossible storage condition

Fig. 8 Compositional variations of plagioclase rims (a) and glass (b)in experimental products as a function of pressure at 890–900°C. Eachplot indicates the average with standard deviation (±1σ). Plots a and bused the same data set as shown in Fig. 7 and Table 6, respectively.The shaded region in a represents composition (±1σ, N=9) ofplagioclase, which were in equilibrium with melt in 2000 magma justbefore syneruptive ascent. The broken line in b indicates groundmasscomposition of ejecta (Table 2)

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For these runs, we subtracted the number of plagioclase inthe starting material from those in the runs to infer thenumber of plagioclase nucleated during decompression.

Microlite shapes change systematically with changingdecompression conditions. Plagioclase microlites in SSDproducts are skeletal, whereas skeletal plagioclase is notfound in the MSD products (Fig. 9). Excluding MSD-1.5, in which pyroxene microlites are too small to observetheir shapes, pyroxene microlites are generally prismatic,with hopper pyroxene limited to SSD runs and MSD-6 andMSD-12 (Fig. 9). Fe–Ti oxide microlites are mostly four-sided or trigonal in form, excluding short-duration SSDand MSD runs (anhedral; Fig. 9, SSD-12). In productswhere microlites are skeletal or hopper, skeletal or hoppergrowth can also be seen on the termination of preexistingcrystals.

Volumes of microlites in matrix

In the SSD experiments, plagioclase volumes (15.4–18.6 vol%; Table 5) do not change significantly with time(Fig. 10a). On the other hand, volumes of pyroxene andFe–Ti oxide (1.1–2.2 and 0.1–1.2 vol% respectively;Table 5) increase with increasing experimental time(Fig. 10a). In the MSD experiments, volume changesinconsistently with increasing time (Fig. 10b). Pyroxeneand Fe–Ti oxides (0.4–3.9 and 0.7–1.8 vol% respectively;Table 5) initially increase, but then decrease in volume withtime. An initial increase over a short time can also be seenfor plagioclase, as volume increases to ∼15 vol% at 6 h.

Number density and CSD of experimental plagioclasemicrolites

We determined crystal-size distributions (CSD) for all runsexcept MSD-144 and MSD-96. CSD for plagioclasemicrolites are mostly linear regardless of the decompressionpath (Fig. 6). In SSD products, population densities for thesmallest sizes (<3 μm) fall off a linear distribution, whichdoes not appear to occur in the MSD products (Fig. 6).

Plagioclase microlite CSD for four SSD productsresemble one another (Fig. 6), similar to their microlitevolumes (Fig. 10a). Number densities for SSD products(5.0×105–1.6×106 mm−3; Table 5) also do not change withexperimental time (Fig. 11a). In contrast, CSDs for MSDplagioclase microlites vary with experimental time (Fig. 6).The inclinations of the CSDs change systematically withexperimental time, increasing from MSD-60 to MSD-6 andthen decreasing to MSD-1.5. Number densities of plagio-clase microlites (1.3×106–1.3×107 mm−3; Table 5) alsochange systematically with time. They initially increasewith increasing experimental time between 1.5 and 6 h andthen decrease between 6 and 60 h (Fig. 11b).

Number density and CSD of experimental Fe–Ti oxidemicrolite phases

Regardless of decompression style, CSDs for the Fe–Tioxide microlites are mostly linear (Fig. 6). In general, theSSD experiment Fe–Ti oxide CSDs resemble one another,except for SSD-12, which has a less steep CSD compared

Table 6 Matrix glass compositions in products of isobaric experiments and starting material for decompression experiments

Run no. Usu00-7 Usu00-14 Usu00-19a Usu00-21-1b

Temperature (°C) 900 890 900 890Pressure (MPa) 150 125 125 150

Wt%c

SiO2 73.88 (0.44) 75.63 (0.52) 74.67 (0.48) 73.55 (0.26)TiO2 0.41 (0.13) 0.34 (0.08) 0.35 (0.12) 0.41 (0.17)Al2O3 14.69 (0.12) 14.54 (0.17) 14.74 (0.20) 15.21 (0.06)FeO* 1.49 (0.08) 1.65 (0.11) 1.45 (0.05) 1.64 (0.10)MnO 0.13 (0.12) 0.16 (0.08) 0.17 (0.06) 0.17 (0.07)MgO 0.69 (0.06) 0.57 (0.07) 0.55 (0.04) 0.81 (0.09)CaO 3.31 (0.11) 2.41 (0.13) 3.19 (0.06) 2.81 (0.10)Na2O 4.25 (0.28) 3.49 (0.20) 3.76 (0.42) 4.34 (0.19)K2O 1.14 (0.08) 1.20 (0.10) 1.14 (0.11) 1.05 (0.09)Ca/Na n.c. n.c. 0.47 n.c.Totald 94.53 (0.90) 95.90 (0.68) 96.21 (1.15) 94.49 (0.83)N 7 5 7 5

Standard deviations are in parentheses.n.c. Not calculated, N number of analysesaReferred to as starting material in decompression experimentbProduct of crystallization experiment was selected for 890°C and 150 MPa, because of larger glass areas than the product of melting experiment(Usu00-21-2; Table 3)cAll oxide values were normalized to 100% and total iron as FeO.dTotal before normalization

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to the other runs. In the SSD experiments, Fe–Ti oxidenumber density increases from 2.3×104 to 1.1×106 mm−3

with experimental time (Table 5; Fig. 11a).In the MSD experiments, the Fe–Ti oxide CSDs

resemble one another, excluding those from MSD-144(Fig. 6). The size range is generally smaller in runs ofshorter duration. Fe–Ti oxide number densities for MSDexperiments range between 1.4×105 and 2.9×106 mm−3

(Table 5), and thus do not change greatly with time, exceptthat products from longer runs have fewer crystals(Fig. 11b).

Matrix glass compositions

Among SSD products, the matrix glass in SSD-12 is leastfractionated (Fig. 12). Excluding that run, glass composi-

Fig. 9 Backscattered electronimages of decompression prod-ucts. The label for each imageindicates the style of decom-pression and experimental time(hour). For details, see the textand Table 4. Phases are Fe–TiOxides (Ox), pyroxene (ortho-pyroxene, Opx, and clinopyrox-ene, Cpx), plagioclase (Pl), glassand bubble in order of decreas-ing brightness. Most crystals aremicrolite. Preexisting crystalsare labeled with pe (abbreviationof preexisting). Plagioclase pe inMSD-1.5 is fragmented. Notethe increase of Fe–Ti oxidemicrolite in SSD products withincrease of experimental time.Plagioclase microlites in SSDproducts are judged to be skele-tal as they show swallow tailform (S.T., e.g., in SSD-12) orhollow (H.L., e.g., in SSD-24)depending on the section of eachcrystal

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tions of the SSD experiments do not change systematicallywith experimental duration (Fig. 12). These observationsare consistent with results on microlite volume, where SSD-12 has the lowest microlite crystallinity (Fig. 10a).

Among MSD products, matrix glass could not beanalyzed in MSD-6 because glass areas are smaller than∼10 μm. MSD-1.5 is the only product whose matrix glassis compositionally similar to that of the starting material(Fig. 12), which is consistent with its microlite volumebeing the smallest (4.4 vol%; Fig. 10b). Glass compositionsfor other MSD products vary with microlite volume (Table 5and Fig. 12).

Plagioclase compositions

For plagioclase microlites, crystal cores were analyzed as inthe ejecta. Data for MSD-1.5, MSD-6, and MSD-12 are notavailable because microlites are too small to analyze with afocused electron microprobe beam. For MSD-96 and MSD-144, where distinction between microlite and preexistingcrystal is not clear, we analyzed relatively small crystals.For SSD products, no difference in An content is found(Fig. 13; Table 7). This is reasonable, because microlitecores grow at an early stage of the run, and thesesamples experienced the same physicochemical changes,

Fig. 10 a, b Volumes ofmicrolite phases and duration ofexperiments for SSD (a) andMSD (b). Pl Plagioclase, Pypyroxene, Ox Fe–Ti oxide. Barfor each data indicates ±1 SDfor data of same number asanalyzed photographs. Thisstandard deviation is not shownfor MSD-96 and MSD-144, be-cause there exists only one data(Table 5). The line and shadedregion represent average and±1σ (N=4) in ejecta, respective-ly. The standard deviation ismore variable in pyroxene thanin Fe–Ti oxides even in the casewhen the volume proportionsare comparable, because numberdensities are smaller in pyroxenethan in Fe–Ti oxides (Fig. 9)

Fig. 11 Number density (permillimeter3) of microlite andduration of experiments for SSD(a) and MSD (b). Pl plagioclase,Ox Fe–Ti oxide. The line andshaded region represent the av-erage and standard deviation(±1σ) for several fragments ofejecta. Standard deviation is notshown for experimental databecause each number densitywas calculated from a CSD

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except that one was held longer at 50 MPa. In MSDproducts, microlites in MSD-60 and small crystals inMSD-96 and MSD-144 mostly have the same An content(Fig. 13; Table 7), except that there are more anorthiticcores in MSD-96 and MSD-144 (Fig. 13). We suspect thatthese high-An crystals are grains that existed beforedecompression.

Analyses of rims around preexisting crystals provide anadditional estimate on plagioclase compositions that grewduring decompression. Rims surrounding the preexistingcrystals are lower in An content than the crystal rims in thestarting material for both SSD products (45–50 in Anmol%; Fig. 14c) and MSD products (50–55 in An mol%;Fig. 14a, b) excluding MSD-1.5 (Fig. 14b). These low-Anovergrowths indicate growth during decompression. Whenmicrolite composition is available (Fig. 14a, c), the low-An

overgrowths compositionally resemble coexisting microlitecores, indicating that the overgrowth occurred together withcrystal nucleation. In MSD-12, low-An overgrowths (An51;Fig. 14b) compositionally resemble microlites in otherMSD products. In MSD-1.5, rims of preexisting crystalsshow no compositional difference from plagioclase rims inthe starting material (Fig. 14b). Because the degree ofcrystallization in MSD-1.5 is the lowest, as inferred frommatrix glass composition (Fig. 12) and microlite volumes(Fig. 10b), we believe that no growth occurred. Insummary, it is clear that the An content of plagioclaseformed during decompression is lower in SSD runs than inMSD runs if we exclude the shortest MSD (MSD-1.5) run(Figs. 13 and 14; Table 7).

Discussion

Interpretation of groundmass microlites in juvenilefragments

We find for juvenile fragments that there are relatively fewof the smallest sizes of microlites CSDs for both plagio-clase and Fe–Ti oxides (<5 μm), and hence CSD trendsdeviate from linear relative to larger crystals (Fig. 6). Adrop in numbers may result from a decreasing nucleationrate and/or Ostwald ripening. In juvenile fragments, clearer

Fig. 12 SiO2 variation diagrams. Except for groundmass, each plot isfrom up to 10 data and its ±1σ is shown. There exists only onegroundmass data (see “Petrography” in text). For starting material,data from Usu00-19 (Table 6) is used. SSD-12 and MSD-1.5 datapoints can be identified by comparison with the other diagrams thatindicate the same SiO2 contents. GMS Groundmass, gl. glass

Fig. 13 Core compositionaldistribution of plagioclasemicrolite and small plagioclase(MSD-144 and MSD-96). Datafor MSD-1.5, MSD-6, andMSD-12 are not available be-cause microlites are too small.Averages for MSD-144 andMSD-96 are not calculated be-cause data is from both preex-isting crystal and microlite.Compositional range of plagio-clase rims before decompressionis also shown (white arrow)referring Usu00-19. Blackarrows indicate corrected com-positional ranges

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deviation for Fe–Ti oxide microlites (Fig. 6) would reflectmore extensive nucleation suppression and/or Ostwaldripening. The linear size distribution of the larger sizeclasses (>5 μm), for both plagioclase and Fe–Ti oxidegroundmass microlites, suggests that they nucleated andgrew at constant rates. Even if nucleation were to stop, theCSD will remain linear as long as the growth rate does notchange with crystal size.

We suggest that microlite crystallization was limited tostage 1 (Fig. 4). If crystallization had occurred in stage 2before reaching the aquifer, we would expect to find kinkedsize distributions, with steeper negative slopes over smallerclasses, as found for bubbles (Fig. 5a). Although CSDs arenot available for pyroxene microlites, we assume thatpyroxene crystallization was also limited to stage 1. We

also assume that the rims on phenocrysts grew in stage 1.Accordingly, the conditions of Stage 1 ascent can bedetermined by documenting similarities in microlitesbetween ejecta and the experiments.

Storage conditions of Usu 2000 magma

If we assume from natural magnetite compositions that theUsu 2000 magma was initially at 880–900°C, we find thatpH2O of 125–150 MPa is the most plausible reservoirpressure (Fig. 4a), based upon the agreement between theexperimental and natural plagioclase and glass composi-tions (Figs. 7 and 8; Tables 2 and 6). At these conditions, nodifference exists between glass compositions for all elementsexcept between FeO contents, which is lower in theexperimental glass relative to that of natural groundmass(Tables 2 and 6).

Several important points must be considered here. Thelower FeO content of the experimental glass can beexplained by the shift from NNO in the natural magmasto NNO + 0.5 to 1 log units in the experiments. The oxygenfugacity for the Usu 2000 magma cannot be estimateddirectly because it lacks ilmenite. If it were similar to thoseof the historical magmas after 1663, then fO2 would havebeen equal to NNO (Tomiya and Takahashi 1995).According to phase equilibria results over a range of fO2

(e.g., Eggler and Burnham 1973; Rutherford et al. 1985;Martel et al. 1999), Fe–Ti oxide stability is very sensitive tochanges in fO2. Fe–Ti oxide stability increases withincreasing fO2, and the magnetite crystallization decreasesthe FeO in the melt. Weight fraction of FeO in the melt isnot large (e.g., in comparison with Al2O3 and SiO2;Table 6); thus, the decrease in FeO did not cause asignificant difference for other elements. However, pH2Oestimation is unaffected by this increased Fe–Ti oxidestability because the anorthite content of plagioclase doesnot depend on fO2 (Scaillet and Evans 1999; Martel et al.1999).

Table 7 Representative core compositions of plagioclase microlites in decompression products

Run no. SSD-24 SSD-144 MSD-60 MSD-96 MSD-144

Wt%SiO2 58.30 58.75 56.71 57.00 55.53Al2O3 26.81 26.70 27.24 27.91 27.64FeO*a 0.70 0.56 1.05 0.60 0.42MgO 0.05 0.07 0.09 0.10 0.06CaO 9.50 8.98 10.26 10.52 10.70Na2O 5.87 6.15 5.33 5.32 5.29K2O 0.11 0.11 0.12 0.06 0.10Total 101.32 101.33 100.80 101.50 99.73An mol% 47.2 44.7 51.6 52.2 52.8

aTotal iron as FeO

Fig. 14 a–c Line-scan analysis for preexisting plagioclase indecompression products. The hatched section indicates the crystalrim range before decompression, referring Usu00-19 (Fig. 7)

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Geophysical evidence, such as a low P-wave velocityzone, aseismic region, and long-period volcanic tremor,indicates the Usu 2000 magma storage at 4–6 km. If weassume a crustal density of 2,500–3,000 kg·m−3 (Wada etal. 1988) and a lithostatic pressure gradient, our experi-mentally constrained water pressure of 125–150 MPacorresponds to a depth of 4–6 km. From this agreement,we deduce that the magma was nearly water-saturated. Wecannot confirm this, however, because of the lack ofreliable glass inclusions. Regardless of that, our experi-ments support the conclusion of Tomiya and Miyagi (2002)that Usu 2000 magma came from a reservoir located atapproximately 4–6 km.

Constraints on the rate and path of ascent of Usu 2000magma

Before we can compare our decompression results tonatural ejecta, we must first adjust our measured plagio-clase microlite data because there exists a Ca/Na ratiodifference between ejecta groundmass and glass of thestarting material for decompressions (Tables 2 and 6). Themelt composition is one parameter that changes the Ancontent of liquidus plagioclase. The Ca/Na ratio in the meltexerts a strong control on An content (e.g., Sisson andGrove 1993; Panjasawatwong et al. 1995). Thus, theplagioclase that crystallized in our decompression experi-ments has lower An contents than that expected from thehigher Ca/Na ratio of the natural groundmass (Table 2).

To compare the experimental products with ejecta, wecorrected the experimental plagioclase microlite composi-tion for each run product. We first calculated a partitioncoefficient for Ca/Na between experimental plagioclasemicrolite cores and melt [hereafter (Ca/Na)pl/(Ca/Na)melt;Table 8]. In this calculation, we considered that microlitecores coexisted with the melt of the same composition asthe melt before decompression. In this process, wecalculated a range of partition coefficients, using maximumand minimum observed Ca/Na ratios of plagioclase micro-lite cores and one Ca/Na ratio for the melt (Table 8). Using

those partition coefficients and Ca/Na data of the naturalgroundmass (Table 8), we calculated a compositional rangeof plagioclase crystallizing from a melt with the naturalgroundmass composition (Table 8). We did not correctresults of MSD-96 and MSD-144. This is because our dataare partly from crystal fragments that existed beforedecompression (Fig. 13); thus, we cannot accuratelydetermine compositional ranges of nucleated plagioclase.Corrected An ranges for microlites (Table 8) are shown inFig. 13.

The rate of magmatic ascent to 2-km depth can beestimated using the run product’s texture and compositionnot affected by residence time at 50 MPa. The corecompositions of plagioclase microlites in the rapidlydecompressed SSD experiments appear to reproduce theAn content of plagioclase microlites in the ejecta (Fig. 13).The relatively slow MSD decompressions produce moreanorthitic microlites than those seen in the ejecta (Fig. 13).Because An contents of plagioclase microlite cores are notavailable for MSD-1.5, MSD-6, and MSD-12 (Fig. 13), wecannot be sure that those experiments did not replicate thenatural microlite compositions.

The fastest SSD decompression experiments replicatenumber densities of plagioclase and Fe–Ti oxide in the ejecta,although Fe–Ti oxide number density was replicated only inshort run (Fig. 11a). This agrees with the overlappingplagioclase microlite compositions. To best match theobserved decrease in An content with increasing decompres-sion rate (Fig. 13), the shorter decompression duration ismore reasonable. Here, the experimental plagioclase numberdensities are similar to that in ejecta at 1.5 h in the MSDseries, but become ten times greater over 6 to 12 h(Fig. 11b). In addition, Fe–Ti oxide microlite number inMSD experiments also suggest a decompression timescale ofless than 1.5 h, because the number density at 1.5 h exceedsthat in the ejecta (Fig. 11b). These results indicate thatmagma ascended to shallow levels (50 MPa) in ∼1.5 h.

To determine how long magma resided at 50 MPa, weuse SSD products whose decompression rate to 50 MPafulfill ∼1.5 h ascent constrained by plagioclase microlite

Table 8 Ca/Na distribution between plagioclase microlite cores and melt before decompression

Run no. Experimental Ca/Na Experimental(Ca/Na)pl/(Ca/Na)melt

Ca/Na of groundmassin ejectac

Correction for Pld

Pla Meltb Ca/Na An mol%

MSD-60 1.00–1.29 0.47 2.13–2.76 0.51 1.08–1.39 51.9–58.2SSD-144 0.80–0.90 0.47 1.70–1.91 0.51 0.86–0.97 46.2–49.1SSD-24 0.78–0.89 0.47 1.66–1.91 0.51 0.84–0.97 45.6–49.1

aPlagioclase. The same data sets as shown in Fig. 13 were used.bAlso see Table 6.cAlso see Table 2.dCalculated for a case where melt before decompression has the same Ca/Na as groundmass in ejecta and Ca/Na distribution between plagioclaseand melt has the same value as in experiment.

440 Bull Volcanol (2007) 69:423–444

compositions and microlite number densities. Numberdensities of Fe–Ti oxide microlites in SSD runs held at50 MPa for 12 to 24 h best replicate those in the ejecta(Fig. 11a). An experimental timescale of around 24 h bestreplicates the microlite volumes in ejecta for all phases(Fig. 10a), which agrees with the overlap in Fe–Ti oxidenumber density data. Direct comparisons between the ejectaand SSD-24, however, reveal a difference. The smallestsizes (∼4 μm) of the Fe–Ti oxide microlites deviate fromthose found in the ejecta CSDs (Fig. 6). That deviation inexperiment may have resulted from suppression of nucle-ation and/or Ostwald ripening. The difference betweenejecta and experimental products is systematic (Fig. 6),suggesting different conditions between magma and exper-imental melt. However, despite those differences, webelieve our ascent rate estimates and duration of repose at50 MPa are reasonable, because they are based uponagreement between the plagioclase microlite compositionsand microlite number densities (for ascent rate) and Fe–Tioxide microlite number density and microlite volumes (forduration at 50 MPa). In fact, even if nucleation of Fe–Tioxides had ceased in the magma, their original numberdensity would be preserved. In summary, we propose thatmagma ascended in less than 1.5 h up to a depth of ∼2 kmbelow West Nishimaya (Fig. 4a) and stayed there foraround 24 h; stage 1 includes stagnation of magma ascent(Fig. 4).

We note that the longer duration MSD run (MSD-144)also reproduces the volume, number density, and CSDs ofmicrolites seen in the natural ejecta (Figs. 6, 10, and 11).However, the pyroxene and Fe–Ti oxide volumes in thatrun are significantly different from those in ejecta(Fig. 10b). Furthermore, the SSD products replicateplagioclase microlite shapes in ejecta (skeletal; Fig. 2),while MSD–144 reproduces euhedral and tabular plagio-clase (Fig. 9). These data, in combination with the goodagreement between microlite core compositions and de-compression rate in the shorter duration SSD runs, help torule out the longer duration MSD runs as relevant toconstraining magmatic ascent at Usu in 2000.

Okazaki et al. (2002) proposed a new, improved modelof the dike propagation that includes the time evolutionfrom 6:00 p.m. on 29 March to the start of the eruption(Table 1). During the early stage from 6:00 p.m. to 12:00p.m. on 29 March, the dike extended between the westernend of the summit and Nishiyama. The dike tip propagatedwestward and reached West Nishiyama just before theeruption, between 12:00 a.m. and 6:00 a.m. on 31 March.The vertical extent (2 km) and shallow depth (0.5–0.25 km)of the dike are consistent with the 2-km depth constrainedby our experiments. It is likely that their observation startedat the start of the lateral transport, because the Usu 2000magma was supplied from a reservoir beneath the summit;

and the eastern end of the dike was originally beneath thesummit (Fig. 4a; Table 1).

The apparent 24-h period of stagnation of the magma at50 MPa that we proposed from our experimental resultsprobably corresponds to the lateral magma transportthrough the dike (Table 1; Fig. 4a). Geophysical resultssuggest that the upper end of the dike remained at constantdepth (0.5–0.25 km; Okazaki et al. 2002), which matchesour conclusion of isobaric stagnation. Our model requiresthat all erupted magma was held isobarically in the dike for∼24 h. The total volume of erupted magma is estimated at∼1 to 2×10−4 km3, which is a maximum, as it includesaccessory and accidental lithic material. Okazaki et al.(2002) suggest that the dike was ∼2 × 2,000 × 2,500 m indimensions, which is ∼1×10−2 km3. Hence, the erupted partcorresponds to only ∼1–2% of the total; therefore, it isreasonable that it all could have come from the tip of thedike (Fig. 4a).

Thus, Usu 2000 magma moved from its reservoir (4–6 km depth) to beneath West Nishiyama during stage 1 intens of hours (Fig. 4a). Experiments, starting from 50 MPaand replicating vesiculation textures seen in Fig. 5b(Suzuki, unpublished data), indicate that stage 2 (Fig. 4a)took less than 10 s. Thus, because stage 1 accounts foralmost all of syneruptive ascent, we conclude that theascent from the reservoir started about 2 days after theinitiation of precursory seismicity (Table 1). That timecorresponds to a peak in precursory seismicity (phase 3;Table 1).

We believe that the peak in seismicity records theinitiation of magma ascent, which is supported by a lackof migration of hypocenters during the first two phases ofthe precursory seismicity (Oshima and Ui 2003). Murakamiet al. (2001) showed that the deeper reservoir (10 km deep)inflated slightly simultaneously with the start of precursoryseismicity (evening of 27 March) and lasted until the earlyevening of 29 March (Table 1). Although its volumechange probably reflects a balance between mass inputand output, we believe that the inflation resulted from newmagma being injected from below. From the early eveningof 29 March to 3 April, a larger deflation of the deepreservoir occurred simultaneously with the start of magmaascent from the shallow reservoir (Table 1). We proposethat the removal of the magma from the shallow reservoir(4–6 km) triggered magma ascent from the deeper one(Table 1).

Conclusions

To constrain the syneruptive dynamics of the daciticmagma expelled during the first phase of 2000 AD eruption

Bull Volcanol (2007) 69:423–444 441

of Usu volcano, Japan, we used decompression experimentsto replicate the ascent of that magma. This study constrainsboth the shallow level storage conditions of the Usu 2000magma and the timescale of magmatic transport during thefirst stage of ascent (Fig. 4a and Table 1). Our mainconclusions are the following:

1. Assuming that the temperature of the magma was 890to 900°C and that the magma was H2O-saturated,storage pressure is estimated to have been 125–150 MPa. That pressure implies that a magma reservoirwas located at 4–6 km deep, which overlaps with thestorage region inferred from previous petrological andgeophysical studies (Tomiya 1995; Onizawa et al.2002; Oshima et al. 2000; Oshima and Ui 2003;Yamamoto et al. 2002).

2. The linear pattern of the microlite size distributions inthe ejecta implies that the decompression inducedcrystallization during the first magma ascent (stage 1)but no crystallization in the final ascent (stage 2), whichfollowed the lateral migration of magma from thecenter of Usu volcano to beneath the eruption site(West Nishiyama). This idea was supported by decom-pression experiments.

3. Plagioclase microlite compositions, shapes, and micro-lite number densities in decompression products indi-cate that the ascent from 4–6 km to 2 km took less than1.5 h, but that the magma then stalled for ∼24 h at2 km.

4. Stage 2 represents the ascent from beneath WestNishiyama to the surface (Fig. 4a), and it is muchshorter than stage 1. When our transport timescale forstage 1 was compared with that of the precursoryseismicity (3.5 days), the expelled magma was found tobegin its ascent 2 days after precursory seismicitystarted, and that the beginning of the ascent coincideswith when seismicity reached its peak.

Acknowledgements This study was mainly carried out in Universityof Alaska Fairbanks (UAF). We express our thanks to Dr. Bill Witteand Dr. Ken Severin at UAF for assistance with computer problemsand microprobe. Yuki Suzuki (Y. S.) is deeply grateful to Prof. SetsuyaNakada at Earthquake Research Institute, University of Tokyo, fororiginally giving her the chance to work on the Usu 2000 eruption,and for encouragement and advice through this work. Y.S. is alsoindebted to Prof. John Eichelberger at UAF for support andencouragement during her stay at UAF. Prof. Takeyoshi Yoshida,Associate Prof. Michihiko Nakamura (Tohoku University) and Prof.Hiroaki Sato (Kobe University) are thanked for comments. Also, Prof.Mitsuhiro Nakagawa at Hokkaido University is thanked for valuableinformation on bulk rock composition of juvenile material in Usu2000 eruption. Finally, the manuscript was greatly improved by theinsightful comments of two anonymous reviewers. This work wassupported by grants to Jessica Larsen (NSF EAR-0106658) and to JimGardner (NSF EAR-0400745). In addition, Y.S. got support from The21st Century COE Program, Tohoku University in summarizing thispaper.

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