Journal of Petrology

Journal of Petrology Advance Access originally published online on June 22, 2007
Journal of Petrology 2007 48(8):1543-1567; doi:10.1093/petrology/egm029

This Article Abstract FREE Full Text (PDF) Supplementary data All Versions of this Article: 48/8/1543    most recent egm029v1 Alert me when this article is cited Alert me if a correction is posted Services Email this article to a friend Similar articles in this journal Similar articles in ISI Web of Science Alert me to new issues of the journal Add to My Personal Archive Download to citation manager Request Permissions Google Scholar Articles by Suzuki, Y. Articles by Nakada, S. Search for Related Content GeoRef GeoRef Citation Social Bookmarking          What's this?

© The Author 2007. Published by Oxford University Press. All rights reserved. For Permissions, please e-mail: journals.permissions@oxfordjournals.org

Remobilization of Highly Crystalline Felsic Magma by Injection of Mafic Magma: Constraints from the Middle Sixth Century Eruption at Haruna Volcano, Honshu, Japan

Yuki Suzuki^1,* and Setsuya Nakada²

¹Institute of Mineralogy, Petrology and Economic Geology, Graduate School of Science, Tohoku University, Aoba-Ku, Sendai, 980-8578, Japan
²Earthquake Research Institute, University of Tokyo, 1-1-1, Yayoi, Bunkyo-Ku, Tokyo, 113-0032, Japan

RECEIVED JULY 21, 2006; ACCEPTED MAY 15, 2007

	ABSTRACT

TOP ABSTRACT INTRODUCTION BACKGROUND OF FUTATSUDAKE... SAMPLES ANALYTICAL METHODS ANALYTICAL RESULTS DISCUSSION CONCLUSION SUPPLEMENTARY DATA REFERENCES

 
The latest eruption of Haruna volcano at Futatsudake took place in the middle of the sixth century, starting with a Plinian fall, followed by pyroclastic flows, and ending with lava dome formation. Gray pumices found in the first Plinian phase (lower fall) and the dome lavas are the products of mixing between felsic (andesitic) magma having 50 vol. % phenocrysts and mafic magma. The mafic magma was aphyric in the initial phase, whereas it was relatively phyric during the final phase. The aphyric magma is chemically equivalent to the melt part of the phyric mafic magma and probably resulted from the separation of phenocrysts at their storage depth of 15 km. The major part of the felsic magma erupted as white pumice, without mixing and heating prior to the eruption, after the mixed magma (gray pumice) and heated felsic magma (white pumice) of the lower fall deposit. Although the mafic magma was injected into the felsic magma reservoir (at 7 km depth), part of the product (lower fall ejecta) preceded eruption of the felsic reservoir magma, as a consequence of upward dragging by the convecting reservoir of felsic magma. The mafic magma injection made the nearly rigid felsic magma erupt, letting low-viscosity mixed and heated magmas open the conduit and vent. Indeed the lower fall white pumices preserve a record of syneruptive slow ascent of magma to 2 km depth, probably associated with conduit formation.

KEY WORDS: high-crystallinity felsic magma; magma plumbing system; multistage magma mixing; upward dragging of injected magma; vent opening by low-viscosity magma

	INTRODUCTION

TOP ABSTRACT INTRODUCTION BACKGROUND OF FUTATSUDAKE... SAMPLES ANALYTICAL METHODS ANALYTICAL RESULTS DISCUSSION CONCLUSION SUPPLEMENTARY DATA REFERENCES

 
Knowledge of the triggering mechanisms of eruptions is important for the forecasting of future eruptions. The injection of new magma into a magma reservoir is one of the major mechanisms. Theoretical aspects of the triggering process include the increase in pressure as a result of (1) an increase in the magma reservoir resulting from the simple volumetric addition of injected magma (Blake, 1984), and (2) the intrinsic increase in volume owing to the exsolution of volatile phases from the melt. Exsolution can be generated either by cooling-induced crystallization of the injected high-temperature magma (Folch & Martí, 1998) or by heating of the low-temperature reservoir magma (Sparks et al., 1977). Furthermore, petrological studies on the ejecta have focused on magma mixing that results from the injection. To link the magma injection and eruption, some recent petrological studies have investigated the timescale from mixing to ejection (e.g. Nakamura, 1995; Venezky & Rutherford, 1999). Also, petrological studies have proposed a new triggering process for the case of highly crystalline felsic magma eruption: the injected low-viscosity mafic magma remobilizes the highly viscous felsic magma, reducing the magma viscosity as a result of mixing (e.g. Pallister et al., 1996; Venezky & Rutherford, 1997; Murphy et al., 2000; Takeuchi & Nakamura, 2001).

Although petrological study of ejecta is a powerful tool for studying the triggering mechanisms of eruptions, it has certain drawbacks. There are only a few studies that consider the migration of magma around the reservoir (e.g. Venezky & Rutherford, 1997; Cottrell et al., 1999). Although studies of groundmass microlites can provide information about magma ascent at shallow levels in the conduit (e.g. Hammer et al., 1999; Nakada & Motomura, 1999; Suzuki et al., 2007), they do not address migration around the reservoir. When an eruption is triggered by magma injection, the migration of magma around the reservoir is associated with chemical modification of the magmas, and is thus important for understanding the whole picture of the eruption triggering processes.

The most important factors that characterize and control the migration of the magma are the depth, shape and dimension of the magma storage system. However, it is sometimes difficult to resolve these storage systems, because of the still poor spatial resolution of geophysical imaging methods and incomplete knowledge of the end-member magmas that are mixed (especially the mafic magma, which comes from a deeper level than the felsic magma). The poor resolution of the magma storage systems makes it difficult to evaluate magma migration around the reservoir. To ascertain the locations of magma storage we need to determine the bulk and melt compositions of the end-member magmas, for example, using mass balance of phenocrysts for the mixing of phyric and aphyric magmas (Nakamura, 1995), and the compositions of melt inclusions (Cottrell et al., 1999). If magma storage conditions can be better ascertained, this would help in understanding the interaction of different magmas at the time of injection, through the estimation of their physical properties. At the same time, detailed study of phenocryst zoning could help us to reveal the progress of magma migration and mixing. If no modern observation record exists for a volcano, eruption triggering processes and the depths of magma storage reservoirs, which are obtained from petrological approaches, are particularly important in predicting the future eruptions based on geophysical observations.

With this motivation and background, we have investigated the mid sixth century Futatsudake eruption of the Haruna volcano in central Honshu (Fig. 1). Although no historical documentation is available, the eruption sequence is well known (Oshima, 1983; Soda, 1989, 1996). In this study we show that a highly crystalline (50 vol. %) felsic (andesitic) magma and two mafic magmas (aphyric and phyric) were involved in the eruption. Although the mafic magmas did not erupt without mixing, we reveal that the aphyric mafic magma is chemically equivalent to the melt fraction of the phyric mafic magma and that the two magmas are from a related storage system. We propose that injection of the mafic magma into the felsic reservoir triggered the eruption by forming low-viscosity magmas, such as heated felsic magma and mixed magmas. Based on separate periods of mafic magma ejection (in the first and last phases of the eruption), we conclude that upward dragging of the injected magma can come to an end (see Snyder & Tait, 1996). With regard to magma migration around the reservoir, it is shown that the first erupted magma was associated with conduit formation and ascended slowly. We propose further, using data from detailed compositional zoning in phenocrysts, a multistage mixing model for the formation of homogeneous mixed magma with a minor contribution from mafic magma.

View larger version (67K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 1. Distribution of ejecta from the Futatsudake eruptions and earlier ejecta of Haruna volcano (after Oshima (1986) and Soda (1989)). Earlier ejecta include those of all stages (Stage 1–5) in Haruna activity. Inset shows the location of Haruna volcano in the NE Japan arc. Samples of the middle sixth century eruption were collected at Owazawa and Ohinata. Only fall deposits are observed at Ohinata. Py-flow, pyroclastic flow; VF (inset), volcanic front.

 

	BACKGROUND OF FUTATSUDAKE VOLCANISM

TOP ABSTRACT INTRODUCTION BACKGROUND OF FUTATSUDAKE... SAMPLES ANALYTICAL METHODS ANALYTICAL RESULTS DISCUSSION CONCLUSION SUPPLEMENTARY DATA REFERENCES

 
Haruna volcano is located in the southern end of the northeast Japan arc (Fig. 1). The activity of this volcano (>300 ka) can be divided into five stages (Oshima, 1983, 1986). The ejecta is andesitic during Stages 1 and 2, andesitic–dacitic in Stage 3, and dacitic during Stages 4 and 5. Based on the SiO₂–FeO * /MgO diagram (Miyashiro, 1974), Watanabe & Takahashi (1995) indicated that rocks from Stage 1 activity can be classified as belonging to the tholeiitic series whereas those from Stages 2 to 5 belong to the calc-alkali series. During the first two stages, the formation and destruction of the stratovolcano were repeated around the present summit (Fig. 1), forming the base of the present volcanic edifice. During Stage 3 magma intruded into the flank of the stratovolcano. Stage 4 is characterized by two caldera-forming eruptions (40 500 ± 3500 years BP for the most recent; Oshima, 1986) at the present summit. During Stage 5, more than five lava domes (< 0·1 km³) were formed inside the caldera and on the eastern flank.

The latest activity occurred at the place where the present Futatsudake lava domes are located (Fig. 1). Three eruptive phases have been identified archaeologically, during the fifth century and in the early and middle sixth century (Machida et al., 1984; Sakaguchi, 1986, 1993). Soda (1989) proposed that the early sixth century event started with a phreatomagmatic eruption, followed by the eruption of pyroclastic flows (Fig. 1; 0·1 km³ in total). The middle sixth century event started with a Plinian eruption, followed by eruption of pyroclastic flows, and ended with lava dome formation. The pyroclastic fall and flow deposits (without any erosion hiatus) are directly covered by stratified fine ash layers from phreatomagmatic eruptions at the time of lava dome emplacement (Soda, 1993, 1996). This indicates that the time elapsed between the first two phases and lava dome emplacement was short. The volumes of the fall and flow deposits for the middle sixth century event are 1·3 km³ and 0·3 km³, respectively (Soda, 1996).

	SAMPLES

TOP ABSTRACT INTRODUCTION BACKGROUND OF FUTATSUDAKE... SAMPLES ANALYTICAL METHODS ANALYTICAL RESULTS DISCUSSION CONCLUSION SUPPLEMENTARY DATA REFERENCES

 
Localities
Pumice blocks of the Plinian and pyroclastic flow phases were sampled at Owazawa and Ohinata (Figs 1 and 2). The samples were classified as ‘white pumice’, ‘gray pumice’ and ‘banded pumice’ (a mixture of the magmas forming the white and gray pumices). In Owazawa, the pyroclastic fall deposit was divided into the lower, middle and upper fall deposits, each of which is distinctive in terms of the type and size of the pumice. The pyroclastic flow deposit has been divided into the lower flow, top 30–60 cm and top 0–30 cm. The lower fall unit includes gray and banded pumices along with white pumices. In Ohinata, the pyroclastic fall deposit has been divided into five units (Fig. 2). The gray and banded pumices are limited to the lowermost unit (unit 1; Fig. 2), similar to Owazawa. The samples from Ohinata were used only for bulk-rock analyses. The lava dome samples were collected from three peaks (Odake, Medake and Magodake).

View larger version (55K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 2. Columnar sections showing eruptive sequence and bulk-rock SiO2 contents of juvenile materials. Distance from Futatsudake is shown next to the site name. For the Owazawa site, SiO2 contents for white pumices are shown with different symbols, depending on the number of pumices used for a single analysis.

 
Petrography
Crystals in pumice and dome lava include phenocryst phases and microlites in the groundmass. The phenocryst phase can be classified as ‘phenocryst’ (>300 µm across) and ‘microphenocryst’ (100–300 µm across) (Fig. 3). The size difference between microphenocrysts and microlites is evident only in the white pumices and white parts of banded pumice (microlites of <20 µm across). Amphibole exists, but not as a microlite (Table 1).

View larger version (177K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 3. Photomicrographs of representative ejecta from the middle sixth century eruption at Futatsudake (plane-polarized light). All scale bars represent 1 mm. (a) Banded pumice from the lower fall deposit; (b) white pumice from the lower fall deposit; (c) white pumice from flow top 30–60 cm; (d) dome lava (Odake). In (a), the dark gray part (G) is adjacent to the white part (W), and the boundary between the two is sharp. Cpx, clinopyroxene; Am, amphibole; Opx, orthopyroxene; Pl, plagioclase; Ox, Fe–Ti oxide. (Note clinopyroxene with different sizes from microphenocrysts (MPh) in (b) and phenocrysts in (d).)

 

View this table: [in this window] [in a new window]   Table 1: Phenocryst and groundmass phases

 
The common phenocryst phases for all ejecta are plagioclase (< 2 mm), orthopyroxene (< 2 mm), amphibole (< 5 mm), magnetite (< 1 mm) and ilmenite (< 700 µm) (Table 1; Fig. 3). Sometimes they form multi-phase aggregates (Fig. 3c), and contain inclusions of other minerals and glass. Olivine occurs as a phenocryst (< 2 mm) and microphenocryst in the dome lava (Table 1; Fig. 4), and is found as a microphenocryst (100 µm across) in the gray pumice and the gray part of the banded pumice (Table 1; Fig. 4). Clinopyroxene is found as a phenocryst (<500 µm; Fig. 3d) and as a microphenocryst in the dome lava, and as a microphenocryst (< 300 µm; Fig. 3b) in the white pumice and the white part of the banded pumice in the lower fall deposit (Table 1). Olivine and clinopyroxene contain inclusions of Fe–Ti oxides.

View larger version (109K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 4. Backscattered electron image of typical olivine (Ol) in ejecta from the middle sixth century eruption at Futatsudake. Scale bars represent 100 µm. (a) Phenocryst in dome lava (Odake); (b) microphenocryst olivine in gray pumice from the lower fall. (Note the size difference between (a) and (b), and euhedral outline of the crystal in (a) and (b) irrespective of the orthopyroxene (Opx) reaction rim (R.R.).)

 
Although phenocrysts and microphenocrysts are commonly euhedral and free from reaction rims, this is not the case for the following phases (Table 1). Ilmenite shows resorption in gray pumice and the gray part of banded pumice. Olivine has an orthopyroxene reaction rim (Fig. 4). Some plagioclase phenocrysts and microphenocrysts in the dome lavas and the lower fall (white, gray and banded) pumices have dusty zones (100 µm wide). Some orthopyroxene grains in the dome lava, gray pumice and gray part of banded pumice have clinopyroxene reaction rims (20 µm). Some of the grains of amphibole are mantled by aggregations of pyroxene, plagioclase, Fe–Ti oxide and glass (up to 100 µm wide), as seen in pumices from the lower fall, upper fall and flow top 30–60 cm. All amphibole grains are mantled by similar rim intergrowths in the dome lava.

	ANALYTICAL METHODS

TOP ABSTRACT INTRODUCTION BACKGROUND OF FUTATSUDAKE... SAMPLES ANALYTICAL METHODS ANALYTICAL RESULTS DISCUSSION CONCLUSION SUPPLEMENTARY DATA REFERENCES

 
Major and trace elements in minerals and glass were analyzed with a JXA-8800R EPMA at the Earthquake Research Institute (ERI), University of Tokyo. The beam diameter was set focused for minerals and at 10 µm for glass. Major elements were analyzed using an accelerating voltage of 15 kV and a beam current of 12 nA. The measuring time was 10 s for the peak and the two backgrounds, respectively, for every element. In glass analyses, Na, K and Si were measured in the first analytical cycle and Al in the second cycle, to minimize the evaporation of alkalis as a result of continuous beam exposure (Devine et al., 1995). For NiO in olivine, the accelerating voltage and counting time were set at 20 kV and 120 s, respectively. Microlites with a large enough size were analyzed by EPMA. At least, two samples of each type from the individual eruptive units were analyzed.

Water contents in melt inclusions in phenocrysts were determined using the EPMA difference method (Devine et al., 1995; Morgan & London, 1996), using the JXM-8800 EPMA in Tohoku University. Hydrous glass standards were used as reference, together with the unknown. Following the optimal conditions for hydrous glass analyses (Morgan & London, 1996), the beam diameter was set at 15 µm, the beam current at 2 nA and the measuring time at 30 s (for both the peak and the two backgrounds). The rest of the analytical conditions were similar to those used for the mineral and glass analyses. Vesiculated inclusions were not used for water content estimation.

Whole-rock compositions were analyzed by X-ray fluorescence (XRF) (PW2400) at ERI, using glass beads with 10 parts flux to one part sample. When pumice grains are not large enough for X-ray analysis, multiple grains having similar characteristics and from the same horizon were used (Fig. 2). We could not separate each band of the banded pumices and hence obtained their bulk compositions (Table 2). Four to seven dome lava samples were analyzed for each peak of Futatsudake.

View this table: [in this window] [in a new window]   Table 2: Representative whole-rock analyses for pumice and dome lava samples. All oxide values normalized to 100% and total iron as FeO

 

	ANALYTICAL RESULTS

TOP ABSTRACT INTRODUCTION BACKGROUND OF FUTATSUDAKE... SAMPLES ANALYTICAL METHODS ANALYTICAL RESULTS DISCUSSION CONCLUSION SUPPLEMENTARY DATA REFERENCES

 
Whole-rock chemistry
The SiO₂ content of the white pumice is nearly constant, independent of the eruptive unit (60·0–61· 6; Fig. 2). Gray and banded pumices have lower SiO₂ content (57·5–60·4 wt %) than the white pumice. The SiO₂ content of the banded pumice (57·5–60·4 wt %) is similar to that of the gray pumice (57·9–58·5 wt %) or higher. The SiO₂ content of the dome lava (59·7–60·4 wt %) is lower than that of the white pumice, but partly overlaps with that of the banded pumice (Fig. 5). All samples show a linear trend in most Harker diagrams (Fig. 5). In other diagrams (e.g. MgO–SiO₂, K₂O–Cr), however, the trend for gray and banded pumices is different from that of the dome lava.

View larger version (36K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 5. Variation of SiO2, TiO2 and CaO vs MgO (wt %), and Zr and Cr (ppm) vs K2O (wt %) of whole-rock data.

 
Mineral chemistry
Compositions of the major phenocryst phases are reported in Electronic Appendices 1–6, available as Supplementary Data at http://www.petrology.oxfordjournals.org. In the following discussion, ‘Rim’ represents the outermost part of a crystal adjacent to the groundmass or reaction rim. The part inside the rim is referred to here as the ‘core’, but we introduce different terms for orthopyroxene and dusty plagioclase that shows complex zoning. Cores of olivine, clinopyroxene, magnetite and ilmenite are homogeneous except near the rims, so the core data are from crystal centers in the thin sections.

Olivine
Olivine crystals all show normal zoning in Mg-number [100 x Mg/(Mg + Fe)] (Fig. 6a). The olivine microphenocrysts in the gray pumice and gray parts of banded pumice of the lower fall deposit have a core composition lower in Mg-number (78–76) than the phenocrysts and microphenocrysts in the dome lava (Fig. 6a). In the Mg-number vs NiO diagram, the composition of olivine in dome lava overlaps with that of the olivine microphenocrysts from the lower fall deposit (Fig. 6b).

View larger version (16K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 6. Chemistry of olivine in gray pumice, gray part of banded pumice (BP), and dome lava. (a) Mg-number [Mg/(Mg + Fe)] of cores vs Mg-number of rims; (b) Mg-number vs NiO. Legend in (a) also applies to (b). Data for (b) were obtained through line-scan analyses for representative grains.

 
Clinopyroxene
Phenocrysts and microphenocrysts in the dome lava have cores of Mg-number 80–75 and Wo [Ca/(Mg + Fe + Mn + Ca)] 46–40, and rims of Mg-number 74–70 and Wo 45–38, showing normal zoning (Fig. 7). Clinopyroxene in the overgrowth rim on orthopyroxene phenocrysts in gray pumices and gray parts of banded pumices has Mg-number 71–74 and Wo 35–41.

View larger version (15K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 7. Variation of Mg-number in cores and rims of clinopyroxene phenocrysts and microphenocrysts in dome lavas.

 
Orthopyroxene
Two types of orthopyroxene can be identified based on the zoning profile (Fig. 8). Type 1 has an Mg-number peak in the rim, whereas Type 2 has the peak inside the rim. Regardless of the type, each phenocryst has a homogeneous part (‘core’) accounting for most of the area. The cores of both types are almost identical in composition. For Type 2, we define the part between the ‘core’ and the ‘rim’ as ‘inner rim’. Occurrence of Type 2 is limited to the lower fall pumice and the dome lava (Figs 8 and 9). The core composition range (Mg-number 62·4–65·7 and Wo 0·8–1· 8) is uniform throughout the eruption. Rims of Type 1 have Mg-number 65·3–67·5 and Mg-number 65·3–67·8 for the gray and white parts of pumices (including banded pumices), respectively, for lower fall, and Mg-number 66·0–73·2 for the dome lava. These values are higher than those for others (Mg-number 63·8–66·0). Inner rims of Type 2 have higher Mg-number than the rims of Type 1 in each eruption phase. Orthopyroxene microlites in the groundmass overlap chemically with the phenocryst rims (Fig. 9).

View larger version (21K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 8. Results of line-scan profiles across representative orthopyroxene phenocrysts. (For explanation of phenocryst types, see the text.) ‘Cpx’ indicates reaction rim (no data shown). Although the core centers are not shown in these profiles, the compositions are almost uniform throughout the core parts.

 

View larger version (37K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 9. Mg-number vs Wo [Ca/(Mg + Fe + Mn + Ca)] mol% plot for orthopyroxene phenocryst (Ph) and groundmass microlites. (For explanation of types and definition of parts in phenocryst, see the text.) BP, banded pumice.

 
Amphibole
Most amphibole phenocrysts are ‘hornblende’ (Fig. 10). They have oscillatory zoning in Na + K (in A site) and Al(IV), both of which change in a sub-parallel fashion. The rims are relatively low in Na + K (Fig. 10) and Al(IV). No systematic chemical changes were found between eruption units.

View larger version (25K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 10. Si vs Na + K (in A site) diagrams for amphiboles in representative samples. Dotted line indicates the boundary between tschermakite (Ts) and hornblende (sénsu strícto) (Hb), after Deer et al. (1992). Data for dome lavas do not include those of oxyhornblende. We analysed crystals without breakdown rims (Table 1) for samples excluding the dome lava. Rim composition is not available for the dome lava because of breakdown (Table 1). The compositional range in the middle fall deposit and lower flow is slightly wider than in the lower fall and is probably due to the larger size of the phenocrysts. There are no large phenocrysts in the lower fall deposit because of fragmentation.

 
Plagioclase
For a dusty phenocryst, we define parts inside the dusty zone as ‘core’, and the parts between ‘core’ and the ‘rim’ as ‘inner rim’ (Figs 11 and 12). Cores of all the plagioclase phenocrysts show oscillatory zoning, ranging from 55 to 90 in An mol % (100 x Ca/(Ca + Na)] (Figs 11 and 12). The FeO and MgO contents are relatively homogeneous with 0·2–0·4 wt % and 0·05 wt %, respectively, except near the rims of some clear phenocrysts in lavas and lower fall pumices. Analytical uncertainties (relative percent) in FeO and MgO are c. 20% at 0·5 wt % and c. 40% at 0·05 wt %, respectively. The rims of clear phenocrysts vary in their chemistry within the range of their cores, except for the lower fall pumice and dome lava. Cores of both clear and dusty plagioclase phenocrysts in the lower fall pumice and dome lava mostly have lower FeO and MgO contents than in the rim and inner rim (Figs 11a, b and 12). Plagioclase microlites in the groundmass overlap chemically with the rim and inner rim of plagioclase phenocrysts (Fig. 12).

View larger version (68K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 11. Line profiles for representative plagioclase phenocrysts. Each line profile is shown by an arrow in the adjacent backscattered electron image. Contents of An, FeO and MgO are shown on the left and right vertical axes, respectively. Shaded area in (a) indicates the inner rim and rim of dusty plagioclase. Although the core centers are not shown in these profiles, the compositions are almost uniform throughout the core parts. (For explanation of phenocryst types (clear, dusty) and definition of parts in phenocrysts, see the text.) Pl, plagioclase; DZ, dusty zone.

 

View larger version (32K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 12. An (mol %) vs FeO and MgO (wt %) diagrams for plagioclase phases. FA, fall; LFA, lower fall; GP, gray pumice; WP, white pumice; BP, banded pumice; G, gray part; W, white part; FL, flow; DL, dome lava; Ph, phenocryst. (For explanation of phenocryst types (clear, dusty) and definition of parts in phenocrysts, see the text.)

 
The thickness of the final stage of growth is commonly larger in microphenocrysts than in phenocrysts because of the different surface areas. The wider domain makes it possible to determine precisely, by microprobe analysis, the compositions of the parts equilibrated just before the eruption. The An contents of microphenocryst rims are compared in Fig. 13. Microprobe backscattered electron images show that the microphenocryst rims are different chemically from the groundmass microlites in these samples, even if microlites exist. Distribution of An content does not change significantly from sample to sample. FAL-2 and FAL-5 from the lower fall deposit have the lowest An contents for both maximum and minimum values of the whole distribution (An₅₀ and An₇₀, respectively).

View larger version (35K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 13. An (mol %) histograms for plagioclase microphenocryst rims. The data in each histogram are from a single sample of white pumice. Samples with asterisks include plagioclase microlites in the groundmass. The dimensions of the vertical axes are the same in all the histograms. FA, fall; LFA, lower fall; FL, flow. However, the microprobe backscattering images show that the phenocryst rims are different chemically from groundmass microlites.

 
Magnetite
The cores of magnetite phenocrysts are chemically homogeneous, with x-usp 0·17–0·20 (where x-usp is the fraction of the ulvöspinel molecule) and Mg/Mn 3–8, except for some grains in the lower fall gray parts of the banded pumices with high x-usp (Fig. 14). Phenocryst rims show no difference in x-usp and Mg/Mn from their cores, except for the lower fall pumices and lavas. In the white parts of pumices (including banded pumices) of the lower fall and dome lavas, some of the phenocryst rims have higher x-usp, but have similar Mg/Mn, in comparison with the cores. In the gray pumice and gray part of the banded pumice, the rims of some phenocrysts have higher x-usp and Mg/Mn. The cores of the microphenocrysts seem to have crystallized simultaneously with the outer zones of the phenocrysts. For example, zoned margins (rim and near rim) of phenocrysts chemically resemble the cores of the microphenocrysts.

View larger version (40K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 14. Histograms for x-usp and Mg/Mn contents of magnetite phenocrysts (Ph) and microphenocrysts (MPh). C, core; R, rim; FA, fall; LFA, lower fall; GP, gray pumice; WP, white pumice; BP, banded pumice; G, gray part; W, white part; FL, flow; N, number. (Note the different x-usp axis values for the dome lava.)

 
Ilmenite
Cores of ilmenite phenocrysts are chemically homogeneous, with x-ilm 0·55–0·60 (where x-ilm is the fraction of ilmentite in the rhombohedral phase) and Mg/Mn 6–9 (Fig. 15). The rims of some phenocrysts in the lower fall unit (white, gray and banded pumices) have higher x-ilm and Mg/Mn (11) than the cores. However, such an increase is not found in the other eruptive phases. The cores of the microphenocrysts are similar in composition to the phenocryst cores.

Fe–Ti oxide thermometry and fO₂ barometry
Because ilmenite is resorbed (Table 1), the ilmenite rims in the gray part of the pumice (including banded pumice) were not used. The equilibrium between the ilm-mt pairs used for calculation was checked in terms of Mg/Mn (Bacon & Hirshmann, 1988). The algorithm of Andersen & Lindsley (1988) yields temperatures of 820–850°C and log fO₂ of –10·3 to –10·9 for the core pairs from all samples and the rim pairs without compositional changes from the cores. The algorithm yields a temperature of 880°C and log fO₂ of –9·9 to –10·8 for the rim pairs from the white pumice fraction (including banded pumice) of the lower fall deposit.

View larger version (25K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 15. Range of x-ilm compositions for ilmenite phenocrysts (Ph) and microphenocrysts (MPh). Data are not available for samples including dome lavas because of their exsolution (Table 1) and scarce amount of crystals. FA, fall; FL, flow. White part of pumice of the lower fall includes white pumice and white part of banded pumice. Similarly, gray part of pumice includes gray pumice and gray part of banded pumice. Data for banded pumice are connected with dotted lines.

 
Groundmass composition
We employed two methods to obtain the composition of the groundmass. For crystal-free groundmass samples (white pumice; Table 1) with relatively large glass domains (10 µm across) the analysis of the glass part was treated as the groundmass composition. For crystal-bearing groundmass (Table 1), we averaged analytical data for up to 200 point analyses. The latter method was limited to lavas (three representatives), as the highly vesiculated groundmass of pumice prevents random analyses. The estimated SiO₂ content of the groundmass for the white pumice is in the range of 77·4–78·4 wt % (Table 3); no systematic change between eruption units was found. The groundmass of the dome lava is less fractionated (SiO₂ of 72–73 wt %) than that of the white pumice.

View this table: [in this window] [in a new window]   Table 3: Representative groundmass compositions

 
Phenocryst phase volume
For highly vesiculated samples it was necessary to section many samples to obtain statistically reasonable estimates of the phenocryst mode. Accordingly, we employed chemical mass balance only for the white pumice (with high vesicularity) in all samples. The middle fall white pumice we used as representative (Table 5) because it is free from groundmass microlites, and its groundmass glass is considered equivalent to the melt part of the reservoir magma. The total volume of the phenocrysts (including microphenocrysts) decreases from the white pumice (c. 50%), to the dome lava (37·0–48·8%) and to the gray pumice (30·5–35·4%) (Tables 4 and 5). Ratios of individual phenocrysts to total phenocrysts barely change throughout the samples (e.g. plagioclase is consistently the most abundant phase). Olivine and clinopyroxene account for 0·6 vol. % in total for the dome lava (Table 4).

View this table: [in this window] [in a new window]   Table 4: Phenocryst phase volume proportions in gray pumice and dome lavas

 

View this table: [in this window] [in a new window]   Table 5: Phenocryst content in white pumice

 

	DISCUSSION

TOP ABSTRACT INTRODUCTION BACKGROUND OF FUTATSUDAKE... SAMPLES ANALYTICAL METHODS ANALYTICAL RESULTS DISCUSSION CONCLUSION SUPPLEMENTARY DATA REFERENCES

 
Mixed and heated magmas at the beginning and end of the eruption
Constraints from mineralogy
The lower fall pumice and the dome lava have characteristics that cannot be explained by simple equilibrium crystallization from one to the other. It is important to note here that the gray and white parts of the banded pumice are identical to the white pumice and gray pumice of the lower fall deposit. We stress the following properties:

1. coexistence of normally and reversely zoned mafic phases (olivine, clinopyroxene, and orthopyroxene) in the gray parts of the pumice (including banded pumice) and the dome lava (Figs 6–9);
   
2. reversely zoned orthopyroxenes in the white parts of the pumice (including the banded pumice) (Figs 8 and 9);
   
3. chemical disequilibrium between olivine and orthopyroxene in the gray parts of the pumice (including the banded pumice) and the dome lava (Fig. 16a);
   
4. enrichment in FeO and MgO at and near the rims of plagioclase (Figs 11a, b and 12);
   
5. Mg/Mn increasing at and near the rims of both magnetite (gray parts of pumice including banded pumice; Fig. 14) and ilmenite (gray and white parts of pumice including banded pumice).
   

View larger version (28K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 16. Mg–Fe distribution between cores of olivine and pyroxene. (a) olivine Mg-number vs orthopyroxene Mg-number; (b) olivine Mg-number vs clinopyroxene Mg-number. Data are plotted for both phenocrysts and microphenocrysts. Each diagram shows the average with the compositional range. Ranges of equilibrium were calculated from KD [(Fe/Mg)crystal/(Fe/Mg)melt] of 0·27–0·33 for olivine (Roeder & Emslie, 1970; Ulmer, 1989), 0·22–0·31 for clinopyroxene (Sisson & Grove, 1993a, 1993b) and 0·245–0·323 for orthopyroxene (Beattie, 1993).

 
Compositional zoning, as in (1), (2), (4) and (5), was formed just before eruption, because it is found near the rims. The characteristics (2), (4) and (5) occur when crystals encounter less evolved melt. The most plausible process for explaining these characteristics would be magma mixing. The onset of amphibole and magnetite precipitation (Gill, 1981) or an increase in fO₂ may produce reverse zoning of mafic minerals. However, in such cases, the lines of evidence (4) and (5) favor magma mixing instead. FeO in plagioclase can increase as a result of an increase in fO₂ (Hattori & Sato, 1996). However, a simultaneous increase of MgO does not support a change in fO₂. Formation of a dusty zone in plagioclase requires an encounter with more calcic melt, an increase in temperature (e.g. Nakamura & Shimakita, 1998), or perhaps both, in accordance with the increases in FeO and MgO. The increase in Mg/Mn supports a temperature increase as a result of mixing with high-temperature mafic magma. Dissolution of ilmenite in the gray pumice and the gray part of banded pumice (Table 1) could be related to the mixing process. Normally zoned phenocrysts of olivine and clinopyroxene were derived from the high-temperature magma, whereas plagioclase and Fe–Ti oxides are from the low-temperature magma (Fig. 17). Microphenocryst olivines in the gray pumice and the gray part of the banded pumice are not crystals that nucleated upon and after mixing, because they have reaction rims of orthopyroxene. The fact that the microphenocryst olivine in the gray part of the pumice (100 µm) is much smaller in size than olivine in the dome lava (Fig. 4) may represent a shorter growth time prior to mixing; the mafic magma with olivine microphenocrysts (100 µm) having been aphyric until just before mixing. Amphibole is from the low-temperature magma, because of the association with aggregates of plagioclase and Fe–Ti oxides. However, their rims lack a record of heating, as indicated by increases in Na + K and Al(IV) (Blundy & Holland, 1990; Scaillet & Evans, 1999; Rutherford & Devine, 2003), in the lower fall white pumice (Fig. 10). This may be due to the amphibole becoming unstable and no longer growing after the mixing event (Table 1). The above phenocryst assemblages in the end-member magmas are also supported by chemical equilibrium or phase stability between the phenocrysts. That is, olivine and clinopyroxene in the dome lava show equilibrium Mg–Fe distributions (Fig. 16b). The white pumice and the white part of the banded pumice from the lower fall deposit contains clinopyroxene microphenocrysts (Table 1) that probably nucleated during or after mixing, because they have no chemical zoning or textures indicative of mixing. Even in the gray pumice and the gray part of the banded pumice, plagioclase was not present in the high-temperature end-member (Fig. 17). If so, some plagioclase microphenocryst cores should have higher MgO and FeO than the plagioclase rims crystallized from the mixed magma (absent in Fig. 12).

View larger version (24K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 17. Schematic diagram showing the range of magmas involved in the Futatsudake eruption, together with their phenocryst assemblages. ‘White’ indicates the white part of pumice including banded pumice; ‘Gray’ indicates the gray part of pumice including banded pumice; Ol, olivine; Cpx, clinopyroxene; Opx, orthopyroxene; Am, amphibole; Mt, magnetite; Il, ilmenite; Pl, plagioclase.

 
White pumice from the middle fall deposit to the flow top 0–30 cm probably maintains the original characteristics of the low-temperature end-member magma (Fig. 17), considering the phenocryst assemblage and the chemical compositions. Thus, the magma erupted in the middle stage represents the low-temperature end-member of the mixed magmas. In the pumice, a slight reverse zoning of orthopyroxene (Figs 8 and 9) is observed, whereas enrichments of FeO, MgO and Mg/Mn in the phenocryst rims of plagioclase and magnetite are not found (Figs 11, 12 and 14). The following two scenarios may explain this. The first is that the reverse zoning of orthopyroxene resulted from an increase in fO₂. The loss of H₂ gas resulting from devolatilization causes oxidation (e.g. Czamanske & Wones, 1973), which is likely to occur during syneruptive degassing processes. However, an increase of fO₂ should increase the FeO content of plagioclase (Hattori & Sato, 1996), which is not found in the present case (Figs 11 and 12). The other possibility may be that the reverse zoning in orthopyroxene records a previous heating event not shown by the magnetite because of rapid diffusion re-equilibration. In such a scenario, it is possible that reverse zoning of orthopyroxene was preserved, whereas plagioclase rims that once had been enriched in FeO and MgO disappeared because of diffusion. At the temperature of the felsic magma (850°C), the diffusion coefficient of Mg in plagioclase with An₇₀ [= 1·3 x 10^–18 m²/s, calculated using the data of Costa et al. (2003)] is found to be larger than the Fe²⁺–Mg interdiffusion coefficient in orthopyroxene (2·0 x 10^–20 m²/s; Tomiya & Takahashi, 2005). We shall defer discussion on the possible heating of the felsic magma erupted in the middle stages of the eruption until the section ‘Possible pre-heating of the felsic reservoir magma’.

Constraints from groundmass and bulk-rock compositions
The linear bulk-rock trends for the gray pumices and dome lavas (Fig. 5) support their origin as binary-mixing products. However, the trends for the lavas and gray pumices intersect on the high K₂O side of the K₂O–Cr diagram, indicating the same low-temperature (high K₂O) mixing end-member but different high-temperature end-members (phyric and aphyric mafic magmas; Fig. 17).

Although the white pumice is considered in terms of bulk chemistry to be the felsic end-member of mixing (Fig. 5), the white pumice of the lower fall deposit has some characteristics of a mixed magma; for example, the properties (2), (4) and (5) in the previous section. The white parts of the banded pumices have the same mineralogical characteristics as the lower fall white pumices. For these units, we suggest heating in a chemically closed system instead of mixing (Fig. 17). The heating can make the melt composition less evolved as a result of dissolution of earlier formed crystals, which leads to the formation of the crystal textures and compositional zoning observed in mixed magmas. Neighboring mixed magma (gray pumices and gray parts of banded pumices) or its mafic end-member should have heated the felsic magma. Irrespective of the heating that may result in the partial dissolution of phenocrysts, the white pumices from the lower fall deposits have similar groundmass compositions to the white pumices of the middle eruption (Table 3). This can be explained by the possibility that the lower fall felsic magma experienced slow syneruptive ascent and resultant crystallization at shallow level in the conduit (discussed below).

In major element variation diagrams the banded pumices plot on a mixing line between white and gray pumices (Fig. 5). This is consistent with their mineralogical characteristics; the gray and white parts of the banded pumices resemble the white pumice and gray pumice of the lower fall, respectively, and the resultant banded pumice is equivalent to a mixture of the white and gray pumice.

Heterogeneous reactions of phenocrysts upon heating and mixing in the lower fall deposits and dome lavas
Patterns of compositional zoning and the degree of compositional change are variable in each product, and are even discernible at the hand-specimen scale. Examples are Types 1 and 2 orthopyroxenes (Figs 8 and 9), clear and dusty plagioclases (Figs 11 and 12) and magnetites with variable x-usp and Mg/Mn values (Fig. 14). These indicate heterogeneous physicochemical changes and/or variable times from such changes to the magma quenching. Because of the extent of magma mingling or mixing, this heterogeneity should be anticipated and it might reflect the proximity of the contrasting magmas. The coexistence of variable phenocrysts in a hand specimen implies effective stirring in each mixed and heated magma.

Timescale from mixing and heating to final ejection for the lower fall deposit
We provide a rough estimate of mixing timescales by using the zoning pattern of magnetites in the white pumices and white parts of banded pumices from the lower fall deposit. The zoning profile in the magnetite grains (200 µm across) would be homogenized in 1·5–12 years at 800–900°C (Tomiya & Takahashi, 2005). Magnetite microphenocrysts in the white pumices and in the white parts of the banded pumices from the lower fall deposit have relatively homogeneous cores (x-usp 0·17–0·20; Fig. 14), surrounded by reversely zoned rims that are as thin as <5 µm. The core compositions are similar to those of the white pumices from the middle eruption, whose magma represents the low-temperature end-member magma in the mixing (Fig. 17). The cores are believed to record the condition before the heating of the lower fall magma. The preservation of the core composition represents a timescale from heating (mixing) to eruption that is shorter than 1·5–20 years. This indicates that the heating and mixing are caused by an event unrelated to an eruption that took place in early sixth century at Haruna, 50 years before the eruption studied here. Devine et al. (2003) showed experimentally that magnetites in the Soufrière Hills magma became zoned within 2–10 days after temperature changed from 835–850 to 860–880°C. This condition may provide a minimum timescale for the white pumice and the white part of banded pumice.

Characteristics of the mafic end-member magmas
Chemical and genetic relationships between the mafic magmas
The linear trend in Fo–NiO of olivines (Fig. 6b) suggests that the aphyric and phyric mafic magmas (Fig. 17) have a close genetic relationship. However, the switch from aphyric to phyric magma during an eruption does not simply result from the temporal evolution of aphyric mafic magma as a result of crystallization. If so, small olivine crystals in the gray pumices and gray parts of banded pumices should have the same core composition as the phenocrysts in the dome lavas (not observed in Fig. 6a) and mixed magmas, formed from the common felsic end-member magma (Fig. 17), should have the same bulk-rock compositional trend (not observed in Fig. 5). It is instead considered that the aphyric mafic magma is chemically equivalent to the melt part of the phyric mafic magma. The cores of the small olivine microphenocrysts in the gray pumice and the gray part of the banded pumice are chemically close to the rim of the olivine phenocrysts in the dome lava (Fig. 6a). This indicates that small olivines in the gray pumices and the gray parts of the banded pumices nucleated from a melt of similar composition to that which crystallized the rim of the phenocrysts.

Bulk composition of the aphyric mafic magma
The bulk chemistry of the aphyric mafic magma was estimated using phenocryst volumes and the major oxide compositions of the mixed magma (gray pumice) and felsic magma (white pumice). Assuming 0 vol. % phenocrysts in the aphyric magma, the estimated compositions are shown in Table 6 (A and B; c. 52 wt % SiO₂). The chemistry resembles that of basalt erupted during the early stage (Stage 1) of Haruna activity (Table 6), which has chemical characteristics common to this volcano. This estimation is supported by observed equilibrium between the melt (aphyric bulk) and the cores of olivine microphenocrysts in the gray pumice and the gray part of the banded pumice. K_D for the Fe–Mg distribution [(Fe/Mg)_Ol/(Fe/Mg)_melt] is 0·31–0·36 for the two compositions of aphyric mafic magma (A and B in Table 6). We used 0·1 for Fe³⁺/(Fe²⁺ + Fe³⁺) in the aphyric mafic magma at quartz–fayalite–magnetite (QFM) and nickel–nickel oxide (NNO) oxygen buffer conditions, based on estimations from Sack et al. (1980). For equilibrium olivine–melt pairs, the K_D can be 0·3 ± 0·03 for a wide range of melt compositions, temperatures and water contents (e.g. Roeder & Emslie, 1970; Ulmer, 1989).

View this table: [in this window] [in a new window]   Table 6: Estimated bulk-rock compositions for mafic magmas

 
Phenocryst content and bulk composition of the phyric mafic magma
Mass-balance calculations for the formation of the dome lava through the addition of aphyric magma (melt) and olivine and clinopyroxene phenocrysts (derived from the phyric mafic magma) to the white pumice magma yields their mixing ratios (Table 7). The ratios of felsic magma : aphyric mafic magma (melt) : olivine : clinopyroxene are about 85–86 : 12–13 : 0·3 : 0·3 (in wt %). The weight per cent ratios of olivine and clinopyroxene are roughly consistent with their phenocryst volumes in the dome lava (Table 4), when considering their relative densities. In the calculation, phenocrysts of felsic magma origin account for 42–44 vol. % of the mixture, consistent with the total phenocryst volumes in the dome lavas (37–48 vol. %; Table 4). The calculated proportion of phenocrysts in the phyric mafic magma is about 6 wt % (Table 7). These results allowed us to estimate the bulk composition of the phyric mafic magma (Table 6).

View this table: [in this window] [in a new window]   Table 7: Mass-balance calculation to form average dome lava composition

 
Figure 18 shows the relationships between mixed magmas (gray and banded pumices, and dome lavas) and the mafic end-members (aphyric and phyric), in which the phyric and aphyric mafic magmas lie at the ends of the mixing lines. The enrichment of MgO and Cr content in the phyric mafic magma compared with the aphyric mafic magma can be explained by the concentration of olivine and clinopyroxene in the former, in both of which Cr is highly compatible (Gill, 1981).

View larger version (21K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 18. Mixed magmas (dome lava, gray pumice) and the mafic end-members of mixing in MgO–SiO2 (wt %) and K2O(wt %) –Cr (ppm) variation diagrams. Each regression line determined for the mixed magmas (gray pumices or dome lavas) and the felsic end-member (white pumice, not shown in this figure) connects mafic and felsic end-member magmas. Phyric mafic magma is compositionally equivalent to a mixture of aphyric mafic magma and phenocrysts (Table 7). (a) A and B are given in Table 6. (b) Mafic magmas are plotted, based on estimated K2O contents (Table 6, average of A and B).

 
Storage conditions of the magmas
Storage condition of the mafic magmas
Because olivine and clinopyroxene were stable in the phyric mafic magma, the stability conditions were estimated using the MELTS algorithm (Ghiorso & Sack, 1995), using the composition of the melt fraction (aphyric mafic magma; Table 6) with different water contents. The estimated liquidus surface shown in Fig. 19 indicates near-liquidus coexistence of olivine and clinopyroxene at pressures of more than 4 kbar and with 2–6 wt % of water, varying with pressure. Under these conditions, the compositions of the crystallized olivine and clinopyroxene (Mg-number of 77–79 and 77–80, respectively) are in the composition range of the dome lavas (Figs 6a and 7). The calculated liquidus temperature (1130–1150°C) shows no correlation with pressure. The clinopyroxene barometer (Nimis & Ulmer, 1998) using phenocryts in the dome lavas yields 3–4 kbar pressure at 1130°C. This value is in agreement with the requirement from Fig. 19. Thus, we consider this value of 4 kbar as the storage pressure for the mafic magma (Fig. 20).

View larger version (40K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 19. Liquidus surface of aphyric magma (Table 6, A), estimated using MELTS (Ghiorso & Sack, 1995) as a function of pressure and H2O content. Redox conditions appropriate to QFM were assumed because the oxygen fugacity of typical island arc magma is in the range QFM–NNO + 2 (Carmichael & Ghiorso, 1990). Although a decrease in fO2 expands the field of olivine relative to orthopyroxene, the stabilities of clinopyroxene and plagioclase are independent of fO2 (e.g. Berndt et al., 2005). Therefore, we may have overestimated the olivine stability field if the fO2 is more oxidizing than assumed. Pl, plagioclase; Px, pyroxene; Cpx, clinopyroxene; Opx, orthopyroxene; Ol, olivine. The shaded zone shows olivine–clinopyroxene coexistence on the liquidus.

 

View larger version (37K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 20. Plausible magmatic processes before and during the Futatsudake eruption. LFA, lower fall; FA, fall; FL, flow. The erupted felsic magma originated from a part of the reservoir that was poorer in crystals and less viscous. The less viscous part might have been formed by pre-heating of the reservoir. After injection of mafic magma from beneath (step 1), mixed and heated magmas were forced to move upward by the convection driven by the injection of the hot mafic magma (step 2). The movement stopped because of cessation of injection and because of cooling (step 3). After establishment of an open conduit, crystal-rich felsic magma was erupted (step 4). Dome lavas sampled at the surface are the products of magma mixing. However, we cannot discount the existence of rocks as felsic as white pumice in the lower part of the lava dome (arrow in step 5). (For possible scenarios of dome lava formation (step 5), see text.)

 
The estimated densities of melt, olivine and clinopyroxene under the above conditions are 2600 kg/m³, 3400 kg/m³ and 3300 kg/m³. This means that there is a high possibility of density-segregation of mafic minerals from the melt in the magma reservoir. It is plausible that aphyric mafic magma was involved during the early eruption phase (Figs 17 and 20), provided that it was in the upper portion of the reservoir. Assuming a homogeneous upper crust having a density of 2700 kg/m³(e.g. granite–diorite, Best & Christiansen, 2001), a pressure of 4 kbar would correspond a depth of c. 15 km (Fig. 20).

Pressure in the felsic magma reservoir
As the white pumices in all eruptive phases are homogeneous in terms of their phenocryst cores, they are considered to be derived from the same reservoir. We have determined the water content in the reservoir magma by analyzing the water content in melt inclusions in phenocrysts in two representative samples (middle fall and lower flow). The average water content of vesiculation-free melt inclusions in orthopyroxene and plagioclase phenocrysts is 5·3 ± 0·4 wt % (n = 14). This corresponds to the solubility of water in the felsic melt at 1·9 ± 0·3 kbar according to Burnham's model (Holloway & Blank, 1994). It is therefore reasonable to assume that the felsic magma was water-saturated, based on evidence from dacitic ejecta from melt inclusion and phase relation studies (Wallace & Anderson, 2000). The estimated water pressure of 1·9 ± 0·3 kbar satisfies the requirement for amphibole stability; requiring a P_H2O of more than 1·3 kbar at c. 850°C (e.g. Rutherford et al., 1985). A water pressure of 1·9 kbar indicates a felsic magma reservoir depth of c. 7 km.

Slow decompression of the lower fall deposit felsic magma
The lower fall deposit felsic magma experienced exceptional heating just prior to eruption (Fig. 17). The outer zones of plagioclase phenocrysts certainly crystallized during and after the heating event, because FeO and MgO in those parts are enriched compared with the inner parts of the crystals (Figs 11b and 12). Given that plagioclase growth in the lower fall felsic magma ended at the felsic reservoir (1·9 ± 0·3 kbar), similar to other felsic magmas, the An content of the plagioclase rims in the lower fall felsic magma should be higher compared with other felsic magmas, because of the heating event. However, if we focus on plagioclase microphenocryst rims, which provide a wider domain for last-stage growth, we find that the plagioclase rims of the lower fall deposit magmas have similar or lower An contents compared with those from other eruption phases (Fig. 13). If the rims are relatively low in An content (e.g. LFA-5 in Fig. 13), that part (low-An overgrowth) is clearly distinguished from the inner part. The above observations suggest that the lower fall felsic magma underwent syneruptive phenocryst crystallization probably as a result of the slow decompression. The mechanism of increased crystallization in slowly decompressed magma is due to the time lag between the physicochemical change and the crystallization response (e.g. Lasaga, 1981). Application of the plagioclase hygrometer (Housh & Luhr, 1991) to the lowest An content rim (An₅₀) of the plagioclase phenocrysts in the lower fall white pumice (Fig. 13; Table 3) yields 3·5 wt % H₂O (Ab model basis) and 2·0 wt % H₂O (An model basis) at 880°C, at a reservoir pressure of 1·9 kbar or less. Given that the reservoir magma was saturated with water, the water content indicates the crystallization of the An₅₀ rims at c. 0·5 kbar. This represents the final pressure of slow decompression (Fig. 20 step 3), assuming equilibrium. The width of the low-An overgrowths (5 µm), which are observed on the rims of the lower fall deposit plagioclase, indicates that decompression occurred just before eruption; an overgrowth of 5 µm can be formed within 6 days after fast decompression to 0·5 kbar in a dacitic melt at 900°C (Suzuki et al. 2007).

Decompression can be achieved by either magma and/or volatile discharge from the reservoir (e.g. Druitt & Sparks, 1984) or by magma ascent. In the former case, the reservoir magma may experience general decompression, resulting in the formation of low-An plagioclase rims throughout the eruptive materials, which is not found in the present case. Thus, the slow decompression, favored here, corresponds to the slow ascent to the surface (Fig. 20, step 3). Only magma that erupted earliest (i.e. the lower fall deposit) can ascend slowly, compared with succeeding magma batches (Fig. 20, step 3), probably because it takes time to form or initially open the conduit or vent.

Clinopyroxene microphenocrysts found only in the white pumices and the white parts of the banded pumices of the lower fall deposit (Fig. 3b; Table 1) could have been stabilized upon decompression. The clinopyroxene liquidus temperature increases in the felsic magma with progressive decompression (Rutherford et al., 1985).

Eruption mechanisms
Physical properties of the magmas
In this section we focus on the physical properties of the magma in the felsic reservoir, and specifically where the felsic magma had an interface with mafic magma. By taking into account the heating of the felsic magma (lower fall deposit) by the mafic magma, we estimated the physical properties of the felsic magma before and after the heating (Table 8). To calculate melt viscosity, we used the formulation for a water-bearing melt of Shaw (1972). The effect of phenocryst presence can be introduced by multiplying the melt viscosity by (1 – a)ⁿ (Marsh, 1981), where a and n are constants [a = 1· 6, n = –2·5, is crystallinity (< 1)]. The viscosity of the felsic magma at 880°C is (1· 0–1· 5) x 10⁸ Pa/s (Table 8). The maximum viscosities of the aphyric and phyric mafic magmas are 7·2 x 10¹ Pa/s and (9·5–9·7) x 10¹ Pa/s, respectively. Thus, the viscosity of the felsic magma is about seven orders of magnitude larger than that of the mafic magma.

View this table: [in this window] [in a new window]   Table 8: Physical conditions of end-member magma

 
The density of the mafic magma decreased little as a result of vesiculation before its injection into the felsic magma reservoir, because the mafic reservoir magma was undersaturated with water (2 wt % H₂O at 4 kbar) and vesiculation would never occur unless the magma is decompressed to 0·5 kbar (Fig. 19). The densities of the aphyric and phyric mafic magmas are 2530–2540 kg/m³ and 2580–2590 kg/m³, respectively (Table 8). Moreover, the density of the felsic magma is similar to that of the mafic magmas, because of its high crystallinity (about 50 vol. %).

Necessary conditions for mixing
In the present case, mixed magmas (gray pumice, gray part of banded pumice and dome lava) appear to have been formed through complete hybridization (Sparks & Marshall, 1986). Sparks & Marshall (1986) showed that hybridization usually requires a small temperature difference between the end-member magmas and a high proportion of the high-temperature magma (>50 wt %). In the scenario we envisage at Futatsudake, the mixing ratio of the mafic magma is 29–40 wt % for the gray pumices, as discussed above, and 13–14 wt % for the dome lavas (Table 7), clearly smaller than this value.

We propose a multistage mixing model for the formation of the mixed magmas at Futatsudake, considering the coexistence of phenocrysts with different zoning patterns such as orthopyroxene in Fig. 8. The product of mixing is, subsequently, mixed further with the felsic magma. First, mafic magma in a higher proportion than that estimated from the final product was mixed with the felsic magma, producing a homogeneous magma with intermediate temperature between the mafic and felsic magmas. This mixing product is mixed again with the remaining felsic magma. In this scenario, phenocrysts with different histories (zoning profiles) can coexist in the final products. First, phenocrysts that originated from the mafic magma show normal zoning. Second, phenocrysts from the felsic magma that experienced two-stage mixing show strong reverse zoning in their inner parts (Type 2 opx in Fig. 8), whereas those experienced one-stage mixing are zoned slightly (Type 1). It is accepted that second-stage mixing involves mingling (mafic inclusion formation), if the disaggregation of mafic inclusions (Feeley & Dungan, 1996; Clynne, 1999) had proceeded completely. No trace of mafic inclusions has been found in the ejecta of the present study.

Assuming that the temperature of the mixed magma was broadly proportional to the mixing ratio of the mafic magma, the ratio for the initial mixing is found to be higher in the gray pumice than in the dome lava. Inner rims of Type 2 orthopyroxene in the gray part of the pumice and dome lava record temperatures just after the first mixing event. These orthopyroxene domains are inferred to have crystallized simultaneously with clinopyroxene that is overgrown on Type 2 orthopyroxene in the gray part of the pumice, and the rims of clinopyroxene phenocrysts in the dome lava. This is based on the observation that the compositions of these opx–cpx pairs fall on isothermal lines of similar temperature in the pyroxene quadrilateral (Lindsley, 1983), in each gray part of the pumice and dome lava (Fig. 21). The Wells (1977) two-pyroxene geothermometer yielded temperatures of 1100–1150°C and 1000–1050°C for clinopyroxene and orthopyroxene pairs from the gray part of the pumices and the dome lava, respectively.

View larger version (22K): [in this window] [in a new window] [Download PowerPoint slide]   Fig. 21. Pyroxene phenocryst rim compositions in the Di–En–Fs–Hd quadrilateral, to confirm their equilibrium just after the first mixing. Isothermal lines are based on Lindsley (1983). Clinopyroxene compositions of the dome lava are for the phenocryst rims, whereas those of the lower fall gray pumice and gray part of banded pumice are for overgrowth rims on orthopyroxene phenocrysts. Orthopyroxene compositions are for inner rims of Type 2 phenocrysts.

 
Possible pre-heating of the felsic reservoir magma
As discussed above, ejecta from the middle sixth century eruption records two types of heating. The lower fall white pumices record heating associated with mafic magma injection into the felsic magma reservoir and resultant mixed magma formation. White pumices from the middle eruption (middle fall to flow top 30–60 cm) probably record heating at a different stage (pre-heating), as observed only in reversely zoned orthopyroxene rims. This pre-heating is not related to mafic magma injection and mixed magma formation, because the white pumices during the middle eruption phase did not erupt together with mixed magma (Fig. 17). The pre-heating might have helped to remobilize the erupted felsic magma, as heating decreases the viscosity of the magma.

Eruption sequence
The middle sixth century eruption was triggered by the input of mafic magma into the felsic magma reservoir (Fig. 20, step 1). Heterogeneity in phenocryst abundance (aphyric and phyric parts) in the mafic magma might have been established before its input (Fig. 20, step 1). Cooling and/or decompression of the mafic magma progressed, resulting in crystallization of small olivines found in the gray pumices and the gray parts of the banded pumices (Fig. 4b). The lower fall ejecta (white, gray and banded pumices) indicate that each of the gray and white parts is homogeneous, independent of their dimensions. We envisage that this scenario may occur if the heated felsic magma (white part) and mixed magma (gray part) were formed and then stored separately, followed by their mixing with each other just before the eruption.

It should be noted that the eruption of the mixed magma (gray pumice and the gray part of the banded pumice) preceded that of the felsic magma of the middle fall to flow top 0–30 cm. Based on experimentation, Snyder & Tait (1996) proposed that convection of overlying magma, driven by heating from newly injected hot magma, can result in partial entrainment of the injected magma (Fig. 20, step 2). In the case of Futatsudake, the felsic magma seems too viscous to convect immediately after the mafic magma injection because the crystallinity of the felsic magma (50 vol. %; Table 5) is close to the critical crystallinity (55 vol. %) above which magma rheologically behaves as a solid (Marsh, 1981). Thus, it is considered that the mafic magma first intruded into the felsic reservoir, forming a dyke (Fig. 20, step 1). Following the intrusion, the felsic magma surrounding the mafic magma was reheated, becoming less viscous, and then was eroded by disaggregation into the mafic magma. The erosion was enhanced not only by heat transfer but also by dyke propagation, and accompanied by mixing between mafic and felsic magmas. Once a considerable volume of mobile felsic magma had accumulated and begun to convect, the upward entrainment of the mafic magma may have proceeded as suggested by Snyder & Tait (1996) (Fig. 20, step 2). The similar densities of the mafic and felsic magmas studied here (Table 8) could have enhanced the entrainment process. Upward entrainment could have been localized at the center of the reservoir (Fig. 20, step 2).

If the entrainment process had continued, mixed magma should have been erupted not only at the beginning of the eruption but also throughout. However, the occurrence of mixed magma is limited to the beginning and end of the eruption. Therefore, we infer that entrainment did not continue in the Futatsudake reservoir (Fig. 20, step 3). Snyder & Tait (1996) showed that a high injection rate was needed to maintain convection and entrainment.

Takeuchi & Nakamura (2001) showed that the viscous crystal-rich felsic magma (10⁷ Pa/s) of Komagatake volcano did not erupt at first, and the mixed magma (10⁴ Pa/s) could have been as a consequence of dyke propagation. Eruption of mixed and heated magmas as in the lower fall deposit may have caused vesiculation of the magma remaining in the reservoir–conduit system through decompression (Fig. 20, step 4) in the event of a conduit opening to the surface, which allowed eruption of the immobile felsic magma (middle fall–flow top 0–30 cm).

The dome lavas (Fig. 20, step 5) could have been formed from the phyric mafic magma that remained at the base of the felsic magma reservoir (Fig. 20, step 3). The eruption of the mixed (lava dome) magma in the final stages of the eruption can be explained by any of the following models: (1) simple turbulent mixing in the conduit during eruption of the remaining felsic magma and phyric mafic magma; (2) new injection of mafic magma and its entrainment, as in the first injection; (3) gravitational overturn of a stratified magma reservoir (through cooling-induced vesiculation of the lower-level phyric mafic magma), leading to mixing throughout the reservoir. As the transition from the felsic magma (pumice flow) to the mixed magma (dome lava) is not continuous, as manifested by the different events, different processes from previous eruption phases might have occurred in the reservoir just before the dome lava emplacement. However, the dome lava samples were collected only from the surface of the large dome complex. Therefore, we cannot preclude the possible existence of rocks as felsic as white pumice in the lower parts or interior of the lava dome (Fig. 20, step 5).

	CONCLUSION

TOP ABSTRACT INTRODUCTION BACKGROUND OF FUTATSUDAKE... SAMPLES ANALYTICAL METHODS ANALYTICAL RESULTS DISCUSSION CONCLUSION SUPPLEMENTARY DATA REFERENCES

 
1. We have investigated the latest Futatsudake eruption, which occurred in the middle sixth century in the sequence Plinian fall, pyroclastic flows and lava dome emplacement. We have revealed the magmatic processes leading to the eruption of the highly crystalline felsic magma, as well as the nature of the magma plumbing system beneath the Haruna volcanic complex (Fig. 20).

2. Gray pumices and the gray parts of banded pumices from the first Plinian phase (lower fall deposit) and dome lavas are the product of magma mixing. The mafic end-member switched from aphyric (lower fall) to phyric (lava dome). The felsic end-member corresponds to the white pumices found in all eruptive units of the Plinian and pyroclastic flow phases. The white pumices and white parts of banded pumices in the lower fall deposits record heating caused by the adjacent mixed magma and its mafic end-member.

3. The aphyric mafic magma chemically corresponds to the melt part of the phyric mafic magma and the two are closely related genetically. The two mafic magmas are derived from a storage system at around 15 km depth. These magmas were formed from an original phyric magma, through density-segregation of mafic phenocrysts in the mafic reservoir.

4. The aphyric part of the mafic magma was injected first into the felsic magma reservoir (c. 7 km depth) without affecting the main part that erupted after the lower fall deposit and before the lava dome emplacement. This is because thermal convection in the reservoir, driven by new magma injection, dragged the injected mafic magma upward. However, the phyric mafic magma was not dragged upward and remained at the base of the felsic reservoir, thereby resulting in the formation of the dome lavas.

5. Mafic magma injection made the high-viscosity felsic magma erupt, through the formation of less viscous mixed and heated magmas (lower fall ejecta) and by the opening of the conduit and vent by the low-viscosity magmas. Indeed, the lower fall white pumices have records of syneruptive slow ascent to a depth of 2 km depth, related to conduit formation.

6. Formation of homogeneous mixed magmas (gray pumices, gray parts of banded pumices, and dome lavas) with low mafic contributions requires multistage mixing (including disaggregation of mafic inclusions), which is supported by phenocryst zoning patterns.

	SUPPLEMENTARY DATA

TOP ABSTRACT INTRODUCTION BACKGROUND OF FUTATSUDAKE... SAMPLES ANALYTICAL METHODS ANALYTICAL RESULTS DISCUSSION CONCLUSION SUPPLEMENTARY DATA REFERENCES

 
Supplementary data for this paper are available at Journal of Petrology online.

	ACKNOWLEDGEMENTS

 
This research was mainly undertaken in the Earthquake Research Institute, University of Tokyo (ERI) as a part of the doctoral thesis of Y.S. We are deeply grateful to Associate Professor A. Yasuda (ERI) for help with the microprobe work. We would also like to thank Professor T. Fujii (ERI), Associate Professor T. Ishii, Professor H. Nagahara and Associate Professor H. Iwamori (University of Tokyo) for reading the first version in the thesis. Dr O. Oshima (now retired from University of Tokyo) is thanked for information on Haruna volcano. Also, Y.S. thanks Professor T. Yoshida and Associate Professor M. Nakamura (Tohoku University) for advice. Dr S. Okumura (Tohoku University) is thanked for lending Y.S. hydrous glass standards. Finally, the manuscript was greatly improved by the insightful comments from Dr Dougal Jerram, Dr Charles Bacon, an anonymous reviewer and Professor John Gamble. Professor John Gamble also corrected aspects of English language. This work was partly supported by Grant-in-Aid from MEXT to S.N. (no. 12304033). Additionally, Y.S. was supported by The 21st Century COE Program in Tohoku University.

---

*Corresponding author. Telphone: +81-22-795-5786. Fax: +81-22- 795-7763. E-mail: suzukiyk{at}ganko.tohoku.ac.jp

	REFERENCES

TOP ABSTRACT INTRODUCTION BACKGROUND OF FUTATSUDAKE... SAMPLES ANALYTICAL METHODS ANALYTICAL RESULTS DISCUSSION CONCLUSION SUPPLEMENTARY DATA REFERENCES

 
Andersen DJ, Lindsley DH. Internally consistent solution models for Fe–Mg–Mn–Ti oxides. American Mineralogist (1988) 73:714–726.[Abstract]

Bacon CR, Hirshmann MM. Mg/Mn partitioning as a test for equilibrium between coexisting Fe–Ti oxides. American Mineralogist (1988) 73:57–61.[Abstract]

Beattie P. Olivine–melt and orthopyroxene–melt equilibria. Contributions to Mineralogy and Petrology (1993) 115:103–111.[CrossRef][Web of Science]

Berndt J, Koepke J, Holtz F. An experimental investigation of the influence of water and oxygen fugacity on differentiation of MORB at 200 MPa. Journal of Petrology (2005) 46:135–167.[Abstract/Free Full Text]

Best MG, Christiansen EH. Igneous Petrology (2001) Oxford: Blackwell Science.

Blake S. Volatile oversaturation during the evolution of silicic magma chambers as an eruption trigger. Journal of Geophysical Research (1984) 89:8237–8244.

Blundy JD, Holland TJ. Calcic amphibole equilibria and a new amphibole–plagioclase geothermometer. Contributions to Mineralogy and Petrology (1990) 104:208–224.[CrossRef][Web of Science]

Burnham CW, Davis NF. The role of H₂O in silicate melts I. P–V–T relations in the system NaAlSi₃O₈–H₂O to 10 kilobars and 1000°C. American Journal of Science (1971) 270:54–79.[Abstract]

Carmichael ISE, Ghiorso MS. The effect of oxygen fugacity on the redox state of natural liquids and their crystallizing phases. Modern Methods of Igneous Petrology: Understanding Magmatic Processes. Mineralogical Society of America, Reviews in Mineralogy—Nicholls J, Russell JK, eds. (1990) 24:191–212.

Clynne MA. A complex magma mixing origin for rocks erupted in 1915, Lassen Peak, California. Journal of Petrology (1999) 40:105–132.[CrossRef][Web of Science]

Costa F, Chakraborty S, Dohmen R. Diffusion coupling between trace and major elements and a model for calculation of magma residence times using plagioclase. Geochimica et Cosmochimica Acta (2003) 67:2189–2200.[CrossRef][Web of Science]

Cottrell E, Gardner JE, Rutherford MJ. Petrologic and experimental evidence for the movement and heating of the pre-eruptive Minoan rhyodacite (Santorini, Greece). Contributions to Mineralogy and Petrology (1999) 135:315–331.[CrossRef][Web of Science]

Czamanske GK, Wones DR. Oxidation during magmatic differentiation, Finnmarka Complex, Oslo area, Norway: Part 2, The mafic silicates. Journal of Petrology (1973) 14:349–380.[Abstract/Free Full Text]

Deer WA, Howie RA, Zussman J. In: An Introduction to the Rock-forming Minerals (1992) 2nd. Harlow: Longman.

Devine JD, Gardner JE, Brack HP, Layne GD, Rutherford MJ. Comparison of microanalytical methods for estimating H₂O contents of silicic volcanic glasses. American Mineralogist (1995) 80:319–328.[Abstract]

Devine JD, Rutherford MJ, Norton GE, Young SR. Magma storage region processes inferred from geochemistry of Fe–Ti oxides in andesitic magma, Soufrière Hills Volcano, Montserrat, W.I. Journal of Petrology (2003) 44:1375–1400.[Abstract/Free Full Text]

Druitt T, Sparks RSJ. On the formation of calderas during ignimbrite eruptions. Nature (1984) 310:679–681.[CrossRef]

Feeley TC, Dungan MA. Compositional and dynamic controls on mafic–silicic magma interactions at continental arc volcanoes: evidence from Cordon El Guadal, Tatara–San Pablo Complex, Chile. Journal of Petrology (1996) 37:1547–1577.[Abstract/Free Full Text]

Folch A, Marti J. The generation of overpressure in felsic magma chambers by replenishment. Earth and Planetary Science Letters (1998) 163:301–314.[CrossRef][Web of Science]

Ghiorso MS, Sack RO. Chemical mass transfer in magmatic processes IV. A revised and internally consistent thermodynamic model for the interpolation and extrapolation of liquid–solid equilibria in magmatic systems at elevated temperatures and pressures. Contributions to Mineralogy and Petrology (1995) 119:197–212.[Web of Science]

Gill JB. Orogenic Andesites and Plate Tectonics (1981) Berlin: Springer.

Hall A. Igneous Petrology (1987) 2nd. Harlow: Longman.

Hammer JE, Cashman KV, Hoblitt RP, Newman S. Degassing and microlite crystallization during pre-climactic events of the 1991 eruption of Mt. Pinatubo, Philippines. Bulletin of Volcanology (1999) 60:355–380.[CrossRef][Web of Science]

Hattori K, Sato H. Magma evolution recorded in plagioclase zoning in 1991 Pinatubo eruption products. American Mineralogist (1996) 81:982–994.[Abstract]

Holloway JR, Blank JG. Application of experimental results to C–O–H species in natural melts. Volatiles in Magmas. Mineralogical Society of America, Reviews in Mineralogy—Carroll MR, Holloway JR, eds. (1994) 30:187–230.

Housh TB, Luhr JF. Plagioclase–melt equilibria in hydrous systems. American Mineralogist (1991) 76:477–492.[Abstract]

Lange RA. The effect of H₂O, CO₂ and F on the density and viscosity of silicate melts. Volatiles in Magmas. Mineralogical Society of America, Reviews in Mineralogy—Carroll MR, Holloway JR, eds. (1994) 30:331–369.

Lasaga AC. Implications of a concentration-dependent growth rate on the boundary layer crystal–melt model. Earth and Planetary Science Letters (1981) 56:429–434.[CrossRef]

Lindsley DH. Pyroxene thermometry. American Mineralogist (1983) 68:477–493.[Abstract]

Machida H, Arai F, Oda S, Endo K, Sugihara S. Tephra with special reference to Japanese archeology. In: Preservation and Cultural and Natural Science for Cultural Properties—Watanabe N, ed. (1984) Kyoto: Dohosha Shuppan. 865–928. (in Japanese).

Marsh BD. On the crystallinity, probability of occurrence, and rheology of lava and magma. Contributions to Mineralogy and Petrology (1981) 78:85–98.[CrossRef][Web of Science]

Miyashiro A. Volcanic rock series in island arcs and active continental margins. American Journal of Science (1974) 274:321–355.[Abstract]

Morgan G. B. IV, London D. Optimizing the electron microprobe analysis of hydrous alkali aluminosilicate glasses. American Mineralogist (1996) 81:1176–1185.[Abstract]

Murphy MD, Sparks RSJ, Barclay J, Carroll MR, Brewer TS. Remobilization of andesite magma by intrusion of mafic magma at Soufrière Hills Volcano, Montserrat, West Indies. Journal of Petrology (2000) 41:21–42.[Abstract/Free Full Text]

Nakada S, Motomura Y. Petrology of the 1991–1995 eruption at Unzen: effusion pulsation and groundmass crystallization. Journal of Volcanology and Geothermal Research (1999) 89:173–196.[CrossRef][Web of Science]

Nakamura M. Continuous mixing of crystal mush and replenished magma in the ongoing Unzen eruption. Geology (1995) 23:807–810.[Abstract/Free Full Text]

Nakamura M, Shimakita S. Dissolution origin and syn-entrapment compositional change of melt inclusion in plagioclase. Earth and Planetary Science Letters (1998) 161:119–133.[CrossRef][Web of Science]

Nimis P, Ulmer P. Clinopyroxene geobarometry of magmatic rocks Part 1: An expanded structural geobarometer for anhydrous and hydrous, basic and ultrabasic systems. Contributions to Mineralogy and Petrology (1998) 133:122–135.[CrossRef][Web of Science]

Oshima O. Geology, petrology, and mineralogy of Haruna volcano. In: Doctoral thesis (1983) University of Tokyo. 280.

Oshima O. Haruna volcano. (1986) In: Editorial Committee of Kanto, Part 3 of Regional Geology of Japan (ed.) Regional Geology of Japan, Part 3 Kanto Region. Tokyo: Kyoritsu Shuppan, pp. 222–224 (in Japanese).

Pallister JS, Hoblitt RP, Meeker GP, Knight RJ, Siems DF. Magma mixing at Pinatubo: petrographic and chemical evidence from the 1991 deposits. In: Fire and Mud: Eruptions and Lahars of Mt. Pinatubo—Newhall C, Punonhbayan R, eds. (1996) Seattle: University of Washington Press. 687–731.

Roeder PL, Emslie RF. Olivine–liquid equilibrium. Contributions to Mineralogy and Petrology (1970) 29:275–289.[CrossRef][Web of Science]

Rutherford MJ, Devine JD. Magmatic conditions and magma ascent as indicated by hornblende phase equilibria and reactions in the 1995–2002 Soufrière Hills magma. Journal of Petrology (2003) 44:1433–1453.[Abstract/Free Full Text]

Rutherford MJ, Sigurdsson H, Carey S. The May 1, 1980 eruption of Mount St. Helens: 1, Melt composition and experimental phase equilibria. Journal of Geophysical Research (1985) 90:2929–2947.

Sack RO, Carmichael ISE, Rivers M, Ghiorso MS. Ferric–ferrous equilibria in natural silicate liquids at 1 bar. Contributions to Mineralogy and Petrology (1980) 75:369–376.[CrossRef][Web of Science]

Sakaguchi H. Haji and Sue-ki detected below FA and FP erupted from Haruna-Futatsudake. (1986) In: Educational Board of Gunma Prefecture (ed.) Reports on the Arato-Kitahara Site, the Imai-Jinja Kofun Group and the Arato-Aoyagi Site. Maebashi: Gunma prefecture, pp. 103–119.

Sakaguchi H. Methods to infer the age and the season of volcanic disasters. In: Tephro-archeology—Arai F, ed. (1993) Tokyo: Kokon Shoin. 54–82. (in Japanese).

Scaillet B, Evans BW. The 15 June 1991 eruption of Mount Pinatubo. I. Phase equilibria and pre-eruption P–T–fO₂–fH₂O conditions of the dacite magma. Journal of Petrology (1999) 40:381–411.[CrossRef][Web of Science]

Shaw HR. Viscosities of magmatic silicate liquids: an empirical method of prediction. American Journal of Science (1972) 272:870–893.[Abstract]

Sisson TW, Grove TL. Experimental investigations of the role of H₂O in calc-alkaline differentiation and subduction zone magmatism. Contributions to Mineralogy and Petrology (1993a) 113:143–166.[CrossRef][Web of Science]

Sisson TW, Grove TL. Temperatures and H₂O contents of low-MgO high-alumina basalts. Contributions to Mineralogy and Petrology (1993b) 113:167–184.[CrossRef][Web of Science]

Snyder D, Tait S. Magma mixing by convective entrainment. Nature (1996) 379:529–531.[CrossRef]

Soda T. Two 6th century eruptions of Haruna volcano, central Japan. The Quaternary Research (JAPAN) (1989) 27:297–312. (in Japanese with English abstract).

Soda T. Haruna-Futatsudake eruptions in the Kofun period. In: Tephro-archeology—Arai F, ed. (1993) Tokyo: Kokon Shoin. 128–150. (in Japanese).

Soda T. Explosive activities of Haruna volcano and their impacts on human life in the sixth century A.D. Geographical Reports of Tokyo Metropolitan University (1996) 31:37–52.

Sparks RSJ, Marshall LA. Thermal and mechanical constraints on mixing between mafic and silicic magmas. Journal of Volcanology and Geothermal Research (1986) 29:99–124.[CrossRef][Web of Science]

Sparks RSJ, Sigurdsson H, Wilson L. Magma mixing: a mechanism for triggering acid explosive eruptions. Nature (1977) 267:315–318.[CrossRef][Web of Science]

Suzuki Y, Gardner JE, Larsen JF. Experimental constraints on syneruptive magma ascent related to the phreatomagmatic phase of the 2000 A.D. eruption of Usu volcano, Japan. Bulletin of Volcanology (2007) 69:423–444.[CrossRef][Web of Science]

Takeuchi S, Nakamura M. Role of precursory less-viscous mixed magma in the eruption of phenocryst-rich magma: evidence from the Hokkaido-Komagatake 1929 eruption. Bulletin of Volcanology (2001) 63:365–376.[CrossRef][Web of Science]

Tomiya A, Takahashi E. Evolution of the magma chamber beneath Usu volcano since 1663: a natural laboratory for observing changing phenocryst compositions and textures. Journal of Petrology (2005) 46:2395–2426.[Abstract/Free Full Text]

Ulmer P. The dependence of the Fe²⁺–Mg cation-partitioning between olivine and basaltic liquid on pressure, temperature and composition: an experimental study to 30 kbars. Contributions to Mineralogy and Petrology (1989) 101:261–273.[CrossRef][Web of Science]

Venezky DY, Rutherford MJ. Preeruption conditions and timing of dacite–andesite magma mixing in the 2·2 ka eruption at Mount Rainier. Journal of Geophysical Research (1997) 102:20069–20086.[CrossRef]

Venezky DY, Rutherford MJ. Petrology and Fe–Ti oxide reequilibration of the 1991 Mount Unzen mixed magma. Journal of Volcanology and Geothermal Research (1999) 89:213–230.[CrossRef][Web of Science]

Wallace P, Anderson A. T. Jr. Volatiles in magmas. In: Encyclopedia of Volcanoes—Sigurdsson H, ed. (2000) New York: Academic Press. 149–170.

Watanabe Y, Takahashi M. Temporal variations of major elemental chemistry of the Haruna volcano. (1995) In: Abstracts, 1995 Fall Meeting of the Volcanological Society of Japan, p. 62 (in Japanese).

Wells PRA. Pyroxene thermometry in simple and complex systems. Contributions to Mineralogy and Petrology (1977) 62:129–139.[CrossRef][Web of Science]

CiteULike    Connotea    Del.icio.us    What's this?

This Article Abstract FREE Full Text (PDF) Supplementary data All Versions of this Article: 48/8/1543    most recent egm029v1 Alert me when this article is cited Alert me if a correction is posted Services Email this article to a friend Similar articles in this journal Similar articles in ISI Web of Science Alert me to new issues of the journal Add to My Personal Archive Download to citation manager Request Permissions Google Scholar Articles by Suzuki, Y. Articles by Nakada, S. Search for Related Content GeoRef GeoRef Citation Social Bookmarking          What's this?

Online ISSN 1460-2415 - Print ISSN 0022-3530

Copyright © 2009 Oxford University Press

Oxford Journals Oxford University Press

• Site Map
• Privacy Policy
• Frequently Asked Questions

Other Oxford University Press sites: Oxford University Press Oxford Journals China Oxford Journals Japan American National Biography Booksellers' Information Service Children's Fiction and Poetry Children's Reference Corporate & Special Sales Dictionaries Dictionary of National Biography Digital Reference English Language Teaching Higher Education Textbooks Humanities International Education Unit Law Medicine Music Online Products Oxford English Dictionary Reference Rights and Permissions Science School Books Social Sciences Very Short Introductions World's Classics