Excessive degassing of Izu-Oshima volcano: magma convection in a conduit (1994)(en)(10s)

Bull Volcanol (1994) 56:207-216

Volc ology 9 Springer-Verlag 1994

Excessive degassing of Izu-Oshima volcano: magma convection in a conduit Kohei Kazahaya, Hiroshi Shinohara, Genji Saito Geological Survey of Japan, 1-1-3 Higashi, Tsukuba, Ibaraki 305, Japan, Fax: Japan-298-56-8725 or Japan-298-54-3533 Received: July 29, 1993/Accepted: December 18, 1993

Abstract. Excess degassing of magmatic H20 and SO2 was observed at Izu-Oshima volcano during its latest degassing activity from January 1988 to March 1990. The minimum production rate for degassed magma was calculated to be about I x 104 kg/s using emission rates of magmatic H20 and SO2, and H20 and S contents of the magma. The minimum total volume of magma degassed during the 27-month period is estimated to be 2.6x 108 m 3, This volume is 20 times larger than that of the magma ejected during the 1986 summit eruption. Convective transport of magma through a conduit is proposed as the mechanism that causes degassing from a magma reservoir at several kilometers depth. The magma transport rate is quantitatively evaluated based on two fluid-dynamic models: Poiseuille flow in a concentric double-walled pipe, and ascent of non-degassed magma spheres through a conduit filled with degassed magma. This process is further tested for an andesitic volcano and is concluded to be a common process for volcanoes that discharge excess volatiles. Key words: Izu-Oshima volcano - degassing - magma convection - conduit - volatile flux - reservoir

Introduction Compositions and fluxes of volcanic gases provide important information about magmas beneath volcanoes. COSPEC, which is now a well-known tool for the monitoring of degassing volcanoes, enables the quantitative measurement of SO2 flux to be undertaken (Stoiber et al. 1983). The SO2 flux data are used not only for monitoring of active volcanoes but also for the estimation of the magma volume that degassed the SO2 (Rose et al. 1982; Andres et al. 1991), if all the SO2 comes directly from the magma. The volume of magma has been calculated by dividing the released amount of Correspondence to: K. Kazahaya

SO2 by the sulfur content of the magma, which is estimated by melt inclusion analyses (Anderson 1974; Devine et al. 1984). The estimated volume of degassed magma often exceeds the volume of magma that erupted during the same period (Rose et al. 1982; Casadevall et al. 1983). A recent survey of SO2 fluxes from Chilean volcanoes suggested that the volume of degassed magma is up to a few hundred times greater than that of the erupted magma (Andres et al. 1991). This excess SO2 has been attributed to either degassing from a large magma reservoir (Rose et al. 1982; Casadevall et al. 1983), degassing of a magma with a high S content (Andres et al. 1991) or a large amount of sulfur-rich vapor phase in a reservoir (Luhr 1990; Westrich and Gerlach 1992). However, mechanisms for excess SO2 degassing have not yet been discussed quantitatively. Excessive degassing has also been observed at IzuOshima volcano during passive degassing in recent active periods: 1950-1974 and 1986-1990. This study documents the recent passive degassing activities of IzuOshima volcano. The water and sulfur content of melt inclusions in plagioclase were analyzed as the representatives of the volatile content of the magma. By combining the volatile content and flux of magmatic gases, we can estimate the corresponding volume of degassed magma. Mechanisms of excess degassing of magma are discussed on the basis of several models, and the significance of the excess degassing for volcanic activity and the growth process of the volcano are evaluated.

Recent activity at Izu-Oshima volcano Izu-Oshima is an active basaltic volcano in Japan. The activity of the volcano is characterized by a repetitive cycle consisting of (1) a dormant period, (2) a lava fountaining period, and (3) an active degassing period (Shinohara and Kazahaya 1990). Detailed observations were conducted during the periods of the last two volcanic activities: 1950-1974 and 1986-1990 (e.g. Issiki

208 1984; Ando 1992). The characteristics of the two periods are quite similar. The volcanic activity during 1986-1990 is summarized as follows. After the summit eruption (lava fountaining) in November 1986, the summit crater was filled with a lava, which crusted over; little degassing activity was observed for about one year. After the formation of a new pit crater by a rapid drain-back of the pooled magma in November 1987, degassing activity increased. The active degassing phase started in January 1988, following small eruptions releasing ashes and gases (Kazahaya et al. 1993). During the active degassing period, volcanic gas was continuously released with a large flux from the summit crater without any eruptions. The degassing rate decreased in March 1990, and weak degassing activity continued through to April 1993. Volcanic tremor was observed during the active degassing, and ceased in March 1990, consistent with the cessation of the active degassing phase. Red-hot magma was sometimes exposed at the bottom of the summit crater during the active degassing stage, and its reflection in the sky (a reddish glow) was called 'Gojinka' (God's fire) by the local population. The existence of a traditional expression for this phenomenon suggests that the exposure of red-hot magma at the summit is typical for this volcano. The 'Gojinka' was frequently observed during 1970-1974 (Ando 1992). In 1972, a lava lake was observed at the bottom of the central pit crater, about 300 m below the summit (Kimura 1988). Waves were often seen traveling across the lake surface. The amplitude of these waves increased during and after intermittent degassing episodes from the lava surface (I Takagi personal communication, 1989), indicating that the volcanic gas was transported in large bubbles through the magma near the surface. The 'Gojinka', however, has not been seen during the latest activity of the volcano because a collapse of the pit-crater wall has covered the molten magma head.

Degassing of magma

Steam flux estimation Heat and steam fluxes through volcanic plumes have been estimated from the plume height with an assumption of a point source (e.g. Morton et al. 1956; Kagiyama 1978). At Izu-Oshima, however, the plume is discharging from a pit crater with a diameter of 350 m and depth of 120 m after intensive mixing with air by convection in the pit. Therefore, the conventional method is not applicable for this volcano, and the steam flux was estimated from a volume flux of the white plume (measured by a video camera) and the water content of the plume by the following procedure. The picture of the white plume was taken perpendicularly to the flow direction of the plume at 2 km from the pit crater. The volumetric plume flux (VpI) was estimated by tracing the plume movement, assuming that the cross section of the plume is a circle. The steam flux (QH2o_p]) through the plume is calculated as:

=

Vp,

(1)

where Wpl is the water content of the plume. Since the plume is a mixture of the volcanic gas and air, the flux of the volcanic water (Qa2o-~o,o) is obtained as the difference between QH2o-pl and the airderived steam flux through the plume (QH2o-~ir): QHzo-volc= QHzo-pl- QEzo-air = (Wp, - Wair) Vpl = Wvo,cVp,

(2)

where Wvo~cis the volcanic water content and Wai~ is the water content of air at the plume altitude. Wpl is estimated from a saturation vapor pressure (P~o-~at) at a plume temperature (Tp3 with an assumptioff that the amount of liquid water droplets is negligible: % , = M.~o PH:o-~t (rp,)

(3)

mpl where R is the gas constant and MH2O is the molecular weight of water. Similarly, W ~ is written as:

Wair = M,~o PH20-air (/"air) Rrair

(4)

where Pmo-ai,(T~i~) is the water vapor pressure of the air at a temperature (T, i0 at the plume altitude. P~2o-ai,(Tair) was calculated with an assumption that the mole fraction of water vapor in the air at the plume altitude (X~o-aG Eq. 5) is the same as that of air at an observational site (XH~o-w~;Eq. 6):

XHzO_air PH20-air (Tair) =

P,:o -air

Xn~o-ws - PH~O-w~(Two)

(5) (6)

etot-ws where Ptot-air and Ptot-~ are atmospheric pressure at the plume altitude and at the Oshima weather station (180 m asl), respectively. T~ir was calculated with Tw~ on the basis of the standard vertical temperature structure of 0.006 ~C/re. With an assumption that Tp,--T~i,, the flux of volcanic water (Qr~o-vo~c) can be obtained. The results for the steam flux observed fromDecember 1988 to March 1990 are listed in Table 1. Since the amount of water existing as liquid droplets in the white plume is neglected, the calculated flux (QH~o-~ol~) gives minimum estimates. Temperature in the plume may be higher than the ambient temperature (Tpt> T~i,), because the volcanic water is supplied as a high temperature vapor. Therefore, the assumption that Tp,=Tair may also causes the underestimation of Qi~o-~oac.

Flux of magmatic gases The minimum flux of steam released from the central pit crater was estimated to be from 180 to 550 kg/s (av. 370 kg/s; Table 1). Hydrogen and oxygen isotopic compositions of water were measured for summit fumarolic

209 Table 1. Estimated H20 flux by volume flux measurements of the white plume Date

Steam flux (kg/s)

Temp.* (~

Humid.* (%)

Plume altitude** (m)

Atm. Press.* (rob)

Plume flux** (m3/s)

21Dec. 1988 22Dec. 1988 22Dec. 1988 24Dec. 1988 16Dec. 1989 17Dec. 1989 18Dec. 1989 19Dec. 1989 9Mar. 1990 average

180 520 550 220 550 520 210 200 350 370

6.2 4.2 3.8 6.6 7.5 7.6 8.8 1.6 6.0

72 59 60 64 77 69 58 69 60

970 950 1010 850 960 1020 780 830 950

901 911 904 922 908 905 933 930 915

0.80 x 10s 1.9 x105 2.0 x105 0.74 x 105 2.7 x105 1.9 xl0 s 0.53 x 105 1.1 x105 1.1 x105

* Calculated with assumptions (see text) at a height where volume flux of the white plume was estimated ** Estimated using a video camera

gases collected during passive degassing, indicating a mixture of local meteoric water, seawater and islandarc type magmatic water (Kusakabe and Matsubaya 1986) with a varying mixing ratio with time (Kazahaya et al. 1993). Since subsidence of the magma head estimated with gravity measurements coincides with the change in the isotopic composition of water in fumarolic gases, the meteoric water and seawater are inferred to be entrained over the magma head. The proportion of magmatic water was calculated to range from 0.35 to 0.50 for summit fumarolic gases (Kazahaya et al. !993). This magmatic contribution to the summit fumarolic gases is assumed to be the same as that of the volcanic gases released from the bottom of the pit crater. Thus, the flux of magmatic water is estimated to range from 63 to 280 kg/s (av. 160 kg/s). The flux of SO2 during the latest degassing period was measured by COSPEC (Ohta et al. 1988; Kyushu University 1990). The averaged flux from May 1988 to July 1989 was 5 kg/s (range 2-10 kg/s), which decreased to 0.7 kg/s in March 1990 (Kyushu University 1990). An SO2 flux of 4 kg/s from the lava lake was measured in 1971 during the earlier active degassing period (Okita and Shimozuru 1974). The sulfur isotopic ratio of fumarolic SO2 was + 1%o in 1988 indicating that SO2 is magmatic, and the proportion of S from seawater (~34S=20%o) to the discharged SO2 is inferred to be little (Kazahaya et al. 1993).

Volatile contents of magma The volatile content of a magma can be estimated by analyses of melt inclusions in phenocrysts (Anderson 1974; Devine et al. 1984). The volatile content of IzuOshima magma was estimated by analyzing melt inclusions in plagioclase phenocrysts in scoria of the 1986 eruption.

Water. Water contents were analyzed for doubly polished melt inclusions by Fourier transform infrared spectroscopy (Nicolet 20SXC spectrometer). The wa-

ter contents were determined on the basis of the BeerLambert law from absorbance intensities at 3600 cm-1 (Stolper 1982). The molar absorption coefficient of 63+5 (1 mo1-1 cm -1) was used for the 3600 cm -~ band (Dixon et al. 1988). Analyses of Mariana trough basaltic glass (1826-1), whose water content is 1.35 +0.04 wt% (G Saito, unpublished data), yielded 1.40+0.01 wt%, and Juan de Fuca basaltic glass (TT152-21), whose water content is 0.36+0.01 wt% (Dixon et al. 1988), yielded 0.37 + 0.01 wt%. The water analyses were reproducible to within 3%, and the accuracy of the water analyses are thought to be about 10% based on comparison of these water analyses with the accepted values. We analyzed five melt inclusions. These glassy inclusions contained water from 0.89 to 1.1 wt% (av. 1.0+0.1 wt%). Therefore, we assume the water content of the magma prior to degassing to be 1.0 wt%.

Sulfur. Sulfur contents were also measured for melt inclusions in plagioclase phenocrysts using an electron microprobe (JEOL JXA-5A). The operating conditions were as follows; 8 txm beam diameter, 15 kV accelerating voltage, i x 10-7 amp beam current. Celestite (grSO4) was used as the sulfur standard and data were corrected by the method of Bence and Albee (1968). Ten electron microprobe analyses of a sodalime glass (NBS-SRM-621) gave 410+20 (lo-) ppm, whereas a bulk analysis using the strong phosphoric acid decomposition technique yielded 460 ppm (M Kusakabe personal communication, 1989). The detection limit for sulfur by electron mircoprobe was estimated to be 70 ppm (3o-). The sulfur contents of the melt inclusions range from 240 to 340 ppm (av. 270 ppm), whereas that of matrix glass is below the detection limit. This indicates that the sulfur species effectively degassed from the melt during the lava fountaining eruption. Decomposition of the anhydrite has been considered as one of the sources of huge SO2 fluxes from some volcanoes during eruptions (Luh r et al. 1984; Fournelle 1990; Bernard et al. 1991). Excess degassing

210 of SO: during or prior to the explosive eruptions of El Chichtn and Pinatubo were attribued to a sulfur-rich gas phase which existed in a magma reservoir prior to the eruption of volatile-rich and highly oxidized magmas (Luhr 1990; Westrich and Gerlach 1992). However, Izu-Oshima magma is water-unsaturated and relatively reduced (Iwasaki et al. 1960), and neither sulfate minerals nor a sulfur-rich gas phase would have existed in a magma reservoir of Izu-Oshima. Therefore, all the SO2 emitted from Izu-Oshima volcano during passive degassing is assumed to be derived from the magma that contained 270 ppm sulfur. 1

Mass rate o f magma degassing

Because there was no extrusion of magma during the passive degassing phase, the huge amounts of gases released are inferred to have come from a magma beneath the volcano (Rose et al. 1982; Casadevall et al. 1983; Andres et al. 1991). Mass rates of the magma degassing (Qdegas) are estimated independently from water and sulfur data as follows: Q dH~g~ =

102QH~o

(7)

and: QSOz 10 6Ms Qso~ degas -Mso2 ACs

(8)

where QH~0 and Qso~ denote emission rates of magmatic water and SO2 (kg/s), AC~I2O and ACs represent the decrease in water content (wt%) and sulfur content (ppm) of magma by the degassing, and Ms and Ms% represent molecular weights of S and SO2, respectively. The ACH~o and ACs values may vary with degassing pressures. Higher degassing pressures result in these values being smaller. Complete degassing of the magma is assumed and the water content (1.0 wt%) and the sulfur content (270 ppm) of the magma are used for ACH~o and ACs respectively. The assumption of complete degassing results in minimum Qdegas values. The mass rate of the magma degassing during 19881989 was calculated to be (6.3-28)x103 kg/s (av. 16 x 103 kg/s) with HzO data and (3.7-18)x 103 kg/s (av. 9.3 x 103 kg/s) with SO2 data. These two independently estimated values agree with each other. A value of 1 x 104 kg/s or 4 m3/s is used as a minimum estimate for the mass rate of the magma degassing (Table 2). The minimum total mass of magma degassed during the two years is 7 x 1011 kg or 2.6 x 10 g m 3 (Table 2). The mass of magma erupted from the summit crater in 1 s content of 270 ppm is much lower than the estimates made for other subduction-related basalts (e.g. Devine et al. 1984), suggesting that the plagioclase grew from already degassed melt, although the mechanism for the S depletion is not yet dear. The estimation of the mass rate of the magma degassingwill not be changed, however, as long as the melt inclusion represents the melt in a resevoir during passive degassing.

November 1986, during the main lava fountaining activity, is 3.5 x 101~ kg (Nagaoka 1988). The total mass of degassed magma for the period 1986-1990 was 20 times greater than that of erupted magma (Table 2). The mass rate of the magma degassing and total amount of degassed magma were also calculated from Eq. (8) for the previous activity, from 1950 to 1974 (Table 2). Although the mass of magma erupted from the summit crater during the lava fountaining of this earlier period (1950-1951) was about twice that of the 1986 activity, the mass ratio of degassed magma and erupted magma was 14, similar to the ratio for the later activity. Such excess degassing of SOz has been reported for many other volcanoes (summarized by Andres et al. 1991), and seems to be a common feature of active volcanoes.

Degassing mechanism of magma Degassing of the magma requires oversaturation of volatiles in melt, therefore, effective degassing occurs at lower pressure conditions. However, the depth of the magma reservoir is estimated to be 4 km on the basis of Mogi's ground deformation model using deformation data during a rapid drain-back on 18 November 1987 (Ida et al. 1988). Therefore, the huge amount of volatiles needs to be transported from a magma reservoir at 4 km depth to the surface. Gas flow through a fracture from the magma reservoir to the surface will be rapid because of the low density and viscosity of the gas phase. The conduit of Izu-Oshima volcano, however, is suggested to be filled with molten magma by the observation of 'Gojinka', and gas transfer in the conduit needs to occur as bubble ascent through the molten magma or as an ascent of volatile-bearing magma. In this chapter, we discuss mechanisms that permit the estimated mass rate of the magma degassing from water and sulfur. Two possible mechanisms - bubble ascent and magma convection in a conduit - are compared.

Bubble ascent

The saturation pressure for basaltic magma containing 1.0 wt% water is only 19 MPa (Burnham 1979), which corresponds to 800 m depth under lithostatic conditions. Although Izu-Oshima melt is undersaturated with H20 at the magma reservoir (4 km depth), CO2rich bubbles may exist in the magma reservoir even at higher pressure conditions,~because the solubility of CO2 in the melt is much lower than that of water (Stolper and Holloway 1988). CO2-rich bubbles in the magma reservoir ascend because of their low density. However, bubble ascent through the molten magma can transport only part of the volatiles - the supersaturated portion - in the magma. In the case of a basaltic magma with an water content of 1.0 wt%, only 0.007-0.034 wt% of water can exsolve from melt to CO2-rich bubbles at 1 kb when the initial COz content is 0.1-0.4

211 Table 2. List of parameters for eruptive and degassing activity of Izu-Oshima volcano during 1950-1974 and 1986-1990

Magmatic activity parameters

1950-1974

1986-1990

Eruptive product (kg)

6.6 x 101~* (Jul. 1950-Apr. 1951) 1400"** (Aug. 1970-Jun. 1974) --4.0 # # 270s$ > 7 x 103

3.5 x 101~ (Nov. 1986) 800 # (Jan. 1988-Mar. 1990) 160 1.0 5.05 270 > 1 x 10 4

> 9 x 10~

> 7 x 10~

> 14

> 20

Duration of degassing (days)

H20 flux (kg/s) H20 content of magma (wt%)" S02 flux (kg/s) S content of magma (ppm) Mass rate of the magma degassing (kg/s) Total mass of degassed magma (kg) Unerupted/erupted ratio

* Morimoto (1958), ** Nagaoka (1988), *** Ando (1992), # Kazahaya et al. (1993), # # Okita and Shimozuru (1974), $ Kyushu University (1990), and *$value assumed to be equal to that of 1986-1990

w t % (Shinohara and K a z a h a y a 1994). T h e r e f o r e water cannot be transported effectively by the bubble ascent f r o m the m a g m a reservoir. Although there are several studies on the sulfur solubility in silicate melts coexisting with pyrrhotite or anhydrite (Luhr 1990; Carrol and R u t h e r f o r d 1985, 1987), the solubility data as a function of partial pressure of sulfur species are not yet available. F o r m a t i o n of the sulfur-rich gas phase was suggested to occur at low pressures; sulfur content in submarine basalt decreases with depths shallower than 200 m, which corresponds to 2 M P a (Moore and Schilling 1973). Volcanic gas studies in Hawaii suggest that SO2 degasses only u n d e r lower pressure conditions (Gerlach 1986). Since the effective degassing of the sulfur-species will occur only at low pressures, the bubble ascent f r o m the deep reservoir m a y not transport the sulfur effectively. Consequently b o t h water and sulfur need to be carried dissolved in the melt.

Convection in conduit Effective degassing of the deep m a g m a reservoir can occur by m e a n s of m a g m a convection in the conduit that transports non-degassed m a g m a to a shallow part of the conduit ( K a z a h a y a et al. 1989). Because m a g m a at the surface of m a g m a h e a d r e m a i n e d in a molten state without crusting over f r o m 1970 to 1974, heat supply f r o m the depths is also necessary to c o m p e n s a t e for the heat loss f r o m the glowing surface. Convection of m a g m a in the conduit can explain not only the large mass rate of the m a g m a degassing but also the continuous glowing of m a g m a at the b o t t o m of a pit crater.

Driving force of convection T h e driving force for convection is p r o b a b l y the density increase of the m a g m a by the degassing at shallow levels, because m a g m a density is an inverse function of

0

Bubbleseparation

~"

/

Iltg

..t%/ i 1 O0

ft. /

200

/i

0

I ~

~

E g, I E

-o

= ,3

2

9

................ ~, ................... g, o

....l:i.. 1

4

9

g

~

~

I ~

t

|

o

~

j

2400

2600 Density

6

8 2800

(kg/rm)

Fig. 1. Density changes of ascending non-degassed magma and descending degassed magma as functions of depth. Arrows schematically show a path for the convective transport of magma caused by this density change. Numerals show water content in wt%

water content ( B u r n h a m and Davis 1972) 2. In Fig. 1, melt density with various water contents was calculated at 1150~ (Fujii et al. 1988) as a function of pressure using partial m o l a r volume of oxides in melt (Lange and Carmichael 1987), compressibility of melts (Kress and Carmichael 1991), and partial m o l a r volume of water in melt ( B u r n h a m and Davis 1971). Density of the degassed m a g m a would also increase by bubble sepa-

2 The effect of a small concentration of water on melt density is not well established yet. For example, Silver et al. (1990) estimated a zero or even a slightly negative molar volume of molecular water in a rhyolitic melt. However, this does not readily imply a density increase by water dissolution, because a larger molar volume of hydroxyl could decrease density. In fact, they found a decrease in density of a quenched rhyolitic glass with increasing total water content.

212

~-

ration and crystallization induced by degassing. Because degassed magma is denser than non-degassed magma, the latter can rise buoyantly through the degassed magma in the conduit.

Radius of magma conduit

y

The radius of the magma conduit was estimated by fluid-dynamic analysis of a sudden drain-back of magma on 18 November 1987. The magma was suggested to be connected to the magma reservoir at 4 km depth because the tilt change indicated the inflation of the reservoir (Ida et al. 1988). The volume of the lava drained back on 18 November is 7 x 106 m 3 (Yamaoka et al. 1989). Duration time of the drain-back event was estimated to be 78 min from the duration of tilt change (Nation Res Inst Earth Sci Disas Prev 1988). Thus, the average drain-back rate (Qdrain-back) is estimated to be 1500 m3/s. Using parameters obtained for the drain-back event, we can estimate the conduit radius at the start of the active degassing phase. Magma flux (Qa~in-baek) is given by the following equation for Poiseuille flow: Qarai~-baok =

"B'r4cAPdrain_back 8/xL

(9)

where APdrai~_baok, r~, L and /x denote overload pressure at the magma head during the event, radius of the conduit, length of the conduit and viscosity of the magma respectively. The overload pressure is estimated by the following equation:

APdrain-back -----pmgh

rc

(10)

where, Pm is the density of magma pooled in the pit crater, g is the acceleration due to gravity, and h is the difference in depth between the pooled'magma and the magma head after the drain-back event (120 m; Yamaoka et al. 1989). Thus, a value for APdrain_back of 3.1 MPa is obtained at the beginning of the drain-back. The lava pooled in the crater for about a year is assumed to be at 1150ec and completely degassed. L is assumed to be 4000 m. Viscosity of the dry basaltic melt at 1150~C is calculated by the method of Persikov (1990) to be 100 Pa.s. Since the pooled lava must have been cooled, the viscosity estimate at 1150~ gives a minimum, resulting in an underestimate of the re. Furthermore a constant APd~ai~-U,ck is assumed for Eq. 9, even though APar~i._b,ek decreases with subsidence, and this also leads to an underestimate of the r~. Thus, the radius of the conduit (re) is calculated to be wider than 4.7m.

Model calculation of convection Magma flux by convection in a conduit can be calculated with fluid-dynamic models as functions of conduit radius and differences in the density and viscosity

Descending degassed magma

Ascending no'n-degassedmagma

a

,,b

I

y,

Descending Ascending degassedmagma non-degassedmagma

5

Fig. 2a, b. Schematic diagrams showing magma transport models: a Poiseui]]e flow in a concentric double-walled pipe and b ascent of non-degassed magma spheres through the degassed magma column

between the ascending and descending magmas. Two simple models are evaluated: Poiseuille flow through a concentric double-walled pipe, and ascent of non-degassed magma spheres through dense degassed magma (Fig. 2).

Poiseuille flow model The magma is transported by Poiseuille flow in a concentric double-walled pipe, where the non-degassed magma ascends in the inner pipe and the degassed magma descends along the outer part (Fig. 2a). Single-cell convection without mixing between the ascending and descending magmas is assumed for simplicity. The ascent rate of non-degassed magma through the conduit is taken to be equal to the mass rate of the magma degassing. The volume flux of ascending magma (Qa~oena) in the inner pipe is expressed as Poiseuille flow driven by buoyancy due to the density difference between the descending and the ascending magma (2Xpa_a): Qascena- frApd-agr4a

(11)

8/*a where ra is the radius of the inner ascending flow, and /-~a is the viscosity of the ascending flow. The volume flux of the descending magma (Qdescend) is expressed as follows (Lamb 1959; p 587):

"n"Apd.ag Ir 4 r 4 Odeseend --

8/Z'--~ [ c -- a

(/'2 -- ra2)21 In (rJr~) J

(12)

where /za is the viscosity of the descending magma. Since the Q ..... a is equal to the Qa..... d under a steady-state condition, Eqs. 11 and 12 are combined to give: /z_~a= f 4

IXa where:

1

(f2--1) a

In f

(13)

213 f = r~ ra

(14)

Viscosities of the Izu-Oshima tholeiitic basalt at 1150~ with water contents of 1.0 and 0 wt% (/Xa and /xa) were calculated by the method of Persikov (1990) to be 63 and 100 Pa.s, respectively. The ra was calculated from Eqs. 13 and 14 to be 2.4 m. The Apd_avalue is assumed to be 70 kg/m 3, which is the density difference between hydrous magma with 1.0 wt% water and anhydrous magma (Fig. 1). The assumed density difference of 70 kg/m 3 is far in excess of that needed to explain the estimated Q .... ,a value (Fig. 3). Even a small density difference of 2 kg/ m 3 can drive convection with the inferred Qas~nd of 4 m3/s (1 x 104 kg/s) through a conduit with re of 4.7 m (Fig. 3). This excess may be due to the following reasons: (1) degassing is not complete; (2) the real conduit is not straight and does not have a constant radius; (3) the ascending and descending magmas shear past and mix with each other; and (4) the real convection is not as simple as that modeled with a single cell, but multiple convection cells exist in the conduit. Stokes model. Ascent of non-degassed magma spheres through degassed magma is considered as an alternative model to the Poiseuille flow model (Fig. 2b). This model corresponds to a former model in which the inner ascending flow column was isolated as a number of spheres. The transition of ascending flow style from the Poiseuille flow to the ascent of isolated magma spheres was indicated during the model experiments of the magma mixing in a rising magma batch (Koyaguchi 1985). Therefore, the radius of the sphere (rs) is assumed to be equal to that of the inner ascending flow column obtained in the previous Poiseuille flow model: the rs is 2.4 m and the r~ is 4.7 m. The magma flux is calculated with the Stokes's equation assuming that the

effect of the conduit wall is negligible for simplicity. This neglect leads to the calculated higher ascending rate. Buoyancy of the non-degassed magma sphere (F) in the degassed magma column is given by: F = x4 ~rgrs3 Apd_a

The resistance force which acts on the sphere is described with the Stokes's equations as (Lamb 1959; p. 601): F=2crtxdr~v 2/za+ 3/Xa {1 + ~ R e }

~ 1988~1990 . . . . . . . . . . . .

j . . . . . . . . . . . .

10-1

/ /

.......

0.1

10

100

(16)

]'~d "[- ],~a

where v is the ascent velocity of the sphere relative to surrounding magma, and Re is the Reynolds number written by: 2rsVpd Re - - /xa

(17)

where pd is the density of the descending magma. Equation 16 contains the first and second spreading terms of Oseen's approximation to take the convection in the liquid sphere into account (Fig. 2b). This equation is applicable for viscous flow with a large Reynolds number ( R e > l ) , whereas the Stokes's approximation can only be used at Re ~ 1. The density and viscosity values for the ascending and descending flows in the previous model are used for those of the sphere and the surrounding magma. Combining Eqs. 15-17, we obtain v and Re as 0.66 m/s and 82 respectively. Because the descent velocity of the surrounding magma is much slower than the ascent velocity of the sphere, the absolute ascent velocity is close to 0.66 m/s. Volume flux of the ascending magma is calculated as: Qascend --

10 2

(15)

4 ~vr 3 3~-

(18)

where y denotes a length of the sphere interval (Fig. 2b). With a volume flux for the ascending magma of 4 roB/s, the sphere interval (y) is calculated as 10 m. Consequently, magma spheres with radius 2.4 m must rise from the reservoir at an average interval of 10 m or every 15 s to cause the observed flux of the magmatic gases. Both of the above simple models, which are based on quite different styles of magma transport, can explain the large flux of the non-degassed magma. Consequently, we infer that magma convection in the conduit allows the rapid upward transport of magma from the deep reservoir, and causes the excess degassing from Izu-Oshima volcano.

ro (m)

Fig. 3. Relationship between magma flux (Q..... a) and conduit radius (re) with varying density differences between ascending and descendingflows.The relationship is calculatedby Eq. 11 for the model of Poiseuille flow in a concentric double-walledpipe. Numerals indicate the density differencebetween ascending and descending magma in kg/m3

Degassing of the magma reservoir The convection process is likely to continue until the whole magma reservoir is degassed or convection is blocked by the collapse of the conduit, because magma

214 convection is driven by even a small density difference between the degassed and the non-degassed magma. The total volume of degassed magma was 2.6 • 108 m 3 during the active degassing from January 1988 to March 1990. This value may give an estimate of the volume of the Izu-Oshima magma reservoir, which fed the latest eruptive activity. Shinohara and Kazahaya (1990) and Kazahaya and Shinohara (1990) proposed a hydrostatic model of eruption processes in which the driving force for magma movement is controlled by the hydrostatic pressure balance between the magma and the crust. The model assumes that buoyancy of magma in the crust is a necessary condition for the start of an eruption, and that buoyancy is strongly controlled by the volatile (including bubbles) content of the magma. Resumption of eruptions from the same reservoir after the active degassing period is unlikely, since the magma remaining in the reservoir is degassed. In fact, lava fountaining activity has been observed only before the active degassing periods at Izu-Oshima. Consequently, convective transport of magma in the conduit may control not only the volatile flux from the volcano, but also the potential for resumed volcanic eruptions.

Destiny of degassed magma Izu-Oshima volcano has repeated similar cycles of activities for the last 15000 years (Tazawa 1980). Therefore, a large volume of degassed magma is likely to have been produced beneath the volcano. Since the degassed magma had a larger density and could not buoyantly rise and erupt, the degassed magma is likely to have accumulated as gabbroic intrusions under the volcano. The total volume of degassed magma was calculated to be 14-20 times larger than that of the erupted magma for the last two periods of activity (Table 2). By assuming the same volume ratio of nonerupted degassed magma to erupted magma for the last 15000 years, we can estimate the total volume of gabbroic intrusions. Each major eruption produced about 0.1 km 3 magma (Nakamura 1964), and eruptions have occurred about 100 times during these 15000 years (Tazawa 1980). Therefore, the total volume of erupted magma is about 10 km 3, and the volume of corresponding degassed gabbroic intrusives is estimated to be 140-200 k m 3. Hasegawa et al. (1987) estimated a density structure model for Izu-Oshima based on explosion seismic measurements. The model includes a high density layer and equivalent gabbroic rocks beneath the center of the volcano at a depth greater than 2 km. Because the crustal structure of Izu-Oshima has not been modeled at depths greater than 3 km, the volume of the gabbroic intrusives cannot be estimated. However, the existence of the high density layer is consistent with our estimates of the large volume gabbroic intrusions. Degassing of a deep reservoir causes the large volume of gabbroic intrusions under the volcano, because passive degassing decreases the potential for eruption

from the reservoir (Shinohara and Kazahaya 1990). Therefore, this degassing mechanism appears to control the volume ratio of extrusion/intrusion and the growth history of the volcano.

Application to silicic volcanoes Excess S O 2 emission has been noted not only for basaltic volcanoes but also for many andesitic to dacitic volcanoes (Rose et al. 1982; Andres et al. 1991). The proposed mechanism of magmatic convection and degassing is also applicable to these more silicic volcanoes. Although the larger viscosity of more silicic magma lowers the rate of convection, a larger conduit radius, expected for silicic volcanoes, can easily compensate for the difference in viscosity. According to Eq. 11, magma flux is proportional to the 4th power of ra (radius of ascending magma column) and is inversely proportional to the viscosity (/x~). Therefore a four orders of magnitude differences in viscosity is compensated by a ten times difference in the conduit radius. An example of convection at an andesitic volcano is given below for Sakurajima volcano, Japan. Sakurajima volcano is an andesitic stratovolcano with a height of 1170 m in southern Kyushu island, Japan. The volcano has had active degassing and frequent Strombolian to Vulcanian eruptions since 1955. The conduit diameter of Sakurajima volcano is estimated to be a few hundred meters, based on the distribution of volcanic earthquakes' hypocenters (Ishihara 1988). Viscosities of these magmas are calculated to be ]-~d= 103"4 and ].La : 10 3.2 Pa.s at 1000~C, and a density difference (Apa-a) is calculated to be 70 kg/m 3 assuming that the ascending magma contains 1 wt% of water, and the descending magma is completely degassed. The convection rates in the Sakurajima conduit are calculated to be 7.4 • 1 0 4 m3/s (2 • 10 s kg/s) when ro is 50 m and 1.6 • 103 m3/s (4 • 106 kg/s) when rc is 20 m with the Poiseuille flow model. The SO2 flux of Sakurajima volcano during active degassing has been measured at about 20 kg/s on average by COSPEC (Ohta et al. 1988). Thus, degassing of only 0.1-2.5 ppmS in the magma is sufficient to cause the observed SOz flux for the conduit radius from 20 to 50 m. This is much smaller than the sulfur contents of many andesitic melts (100-500 ppmS) estimated by melt inclusion analyses (Devine et al. 1984; Palais and Sigurdsson 1989). Consequently, convective transport of magma in a conduit is effective enough to cause the excess SO2 degassing observed not only at basaltic volcanoes but also at more silicic volcanoes.

Summary The total volume of magma degassed during active degassing periods at Izu-Oshima volcano is estimated to be 14-20 times larger than the erupted volume of magma, based on emission rates of magmatic water and SO2, and water and sulfur contents of the magma.

215 M e c h a n i s m s f o r excess d e g a s s i n g o f m a g m a d u r i n g t h e p a s s i v e d e g a s s i n g p h a s e a r e d i s c u s s e d o n t h e basis of s e v e r a l m o d e l s a n d c o n v e c t i v e t r a n s p o r t of n o n - d e g a s s e d m a g m a f r o m t h e d e e p e r r e s e r v o i r to s h a l l o w d e p t h is c o n c l u d e d to b e t h e m o s t likely. N u m e r i c a l a n a l y s e s o f t w o m o d e l s ( P o i s e u i l l e flow in a c o n c e n t r i c d o u b l e wall, a n d a s c e n t o f m a g m a s p h e r e s t h r o u g h d e gassed magma) reveal that the density increase of magm a t h a t r e s u l t s f r o m d e g a s s i n g c a n l e a d to m a g m a c o n vection. Convection causes degassing of the deep magm a r e s e r v o i r a n d r e s u l t s in t h e f o r m a t i o n o f g a b b r o i c intrusives. T h e c o n v e c t i o n m e c h a n i s m is a p p l i c a b l e n o t o n l y to b a s a l t i c s y s t e m s b u t also to m o r e silicic m a g m a systems. T h e style o f t h e m o d e l l e d m a g m a flows in a c o n d u i t is n o t c o n s t r a i n e d e i t h e r b y e x p e r i m e n t a l o r t h e o r e t i c a l studies. I n p a r t i c u l a r , t h e n e g l e c t of t h e m i x i n g o f asc e n d i n g a n d d e s c e n d i n g m a g m a s is l i k e l y to b e a signifi c a n t s o u r c e o f e r r o r s in t h e c a l c u l a t e d r a t e s o f c o n v e c tion. T h e r e f o r e , f u r t h e r w o r k o n t h e f l o w r e g i m e d u r ing d e g a s s i n g a n d c o n v e c t i v e t r a n s p o r t o f m a g m a a r e r e q u i r e d to q u a n t i f y this c o n v e c t i o n m o d e l .

Acknowledgements. We thank RJ Andres, JB Lowenstein and M Takahashi for their critical comments and constructive suggestions on the manuscript. The manuscript was improved with the reviews and comments by AT Anderson, RJS Sparks and an anonymous reviewer, whose assistance is acknowledged.

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