Long-period seismicity at Redoubt Volcano, Alaska, 1989–1990 related to magma degassing

JotuTlalofvokanol0~ andgeothemdreseacch

ELSEVIER

Journal of Volcanology

and Geothermal

Research75 (1997) 321-335

Long-period seismicity at Redoubt Volcano, Alaska, 1989- 1990 related to magma degassing Meghan M. Morrissey U.S. Geological Survey,

*

345 Middlefield Road, MS 977, Menlo Park, CA 94025, USA

Received 26 January

1996; accepted 8 May 1996

Abstract The mass of exsolved magmatic H,O is estimated and compared to the mass of superheated steam (25-50 Mtons) released through the resonating crack producing the December 13- 14, 1989 swarm of long-period seismic events at Redoubt Volcano. Results indicate degassing of a H,O-CO,-SO,-saturated magma upon ascending from at least 12 km to 3-4 km beneath the crater as the source of the superheated steam. The mass of exsolved H,O (3.2-2.50 Mtons) is estimated from solubility diagrams of H,O-CO,-saturated silicate melts for the ascent history of the Redoubt magmas. Crystal size distribution, seismological, petrological, and geochemical data are used to constrain the ascent history of the two andesitic magmas prior to the eruption. Two stages of crystallization are inferred from crystal size distributions of plagioclase crystals in andesites erupted in December 1989. The first stage occurred 30-150 years before the eruption in both magmas and the second stage occurred at least 8 years and 15 years before the eruption in the dacitic andesite and rhyolitic andesite, respectively. The depths of crystallization are constrained from the spatial and temporal variations of volcano-tectonic earthquakes locations (Lahr et al., 1994) and from the P-wave and S-wave velocity structures (Benz et al., 1996). These data suggest that the rhyolitic andesite magma ascended to a depth of 7-8 km within at least 15 years of the eruption. Within at least 8 years of the eruption, the dacitic andesite magma migrated to a depth just below the other magma body where it resided until hours to days of the eruption. At this time, the dacitic andesite magma mixed with the rhyolitic andesite magma and established the reservoir for the eruption. Near the top of the reservoir, some of the mixed magma was displaced into fractures which extended 4-5 km toward the surface. This displaced magma created the eruption conduit and released the fluids related to the resonating crack. This scenario is consistent with the trends in major- and trace-element chemistry, and the stability of hornblende in the pre-eruption Redoubt magmas. It also provides a source for the SO, and CO, emissions measured during the eruption. Keywords: long-period

seismicity;

magma degassing;

Redoubt Volcano;

1. Introduction Seismic events observed at a volcano and gas emissions measured above a volcano are data com-

* Fax: + 1.415.329-5163; e-mail: [email protected] 0377.0273/97/$1700 Copyright PZZSO377-0273(96)00033-9

crystal size distribution

monly used to remotely assess the level of magmatic activity and eruption potential. Two types of seismic events, volcano-tectonic earthquakes and long-period events, provide information on the movement of fluids within the edifice. Volcano-tectonic earthquakes which are high-frequency, broadband events, are thought to reflect stress relaxation in regions of

0 1997 Elsevier Science B.V. All rights reserved

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the crust where magma has been removed (Scandone and Malone, 1985). Long-period events which are low-frequency, narrow banded events with a monochromatic waveform, are thought to be seismic energy radiating from a resonating fluid-driven crack triggered by pressure disturbances in the fluid (Chouet, 1986, 1988, 1992). Gas emissions such as SO, are measured with either a correlation spectrometer (COSPEC) or a total ozone mapping spectrometer (TOMS). Fluctuations in emissions data are thought to reflect the degassing behavior of magma caused by changes in properties of the host rock or movement of the magma at depth (Fischer et al., 1993). Hence, coupling these data may provide a means of monitoring the flow of fluids in and out of the edifice. There have been several case studies in which both seismic observations and measured gas emissions have been used to infer magmatic activity beneath an active volcano including Redoubt Volcano in Alaska, 1989 and Galeras Volcano in Colombia, 1993. During the six months of eruptive activity at Redoubt which began on December 14, eleven of the eighteen major eruptive events were preceded by detectable seismicity (Power et al., 1994). In March 1990, successive daily measurements were made with COSPEC which happened to span over a time when an eruption occurred with precursory seismicity (Casadevall et al., 1994). SO, emissions began to decrease three days before the eruption on March 23. Immediately after the eruption, an increase in SO, emissions was observed. Seismicity recorded during this time was characterized by a series of long-period events in which individual events occurred more frequently before the eruption. The observed decrease in gas emissions suggested that a plug of magma was solidifying in the vent (Casadevall et al., 19941, which could have sealed the upper part of the edifice causing pressurization and the observed seismicity. Similar trends in SO, emissions and seismicity were observed prior to eruptions on January 14 and 23 at Galeras Volcano (Fischer et al., 1993). Fischer et al. (1993) attributed the decrease in gas emissions to a reduction of permeability in the edifice which was caused by hydrothermal alteration or rheological changes of the magma. A reduction in permeability would have steepened the hydrostatic pressure gradi-

ent causing unsteady flow through fractures which thereby increased the level of long-period seismicity. In both case studies, it is inferred that fluctuations in gas emissions and the occurrence of long-period seismicity were related to the movement and pressurization of gases within the edifice. The primary assumption for these inferences was that the gases that were either emitted from the volcano or contained within the resonating crack beneath the volcano were magmatic in origin. Gerlach et al. (1994) demonstrated that the only viable source of measured SO, emissions at Redoubt was excess vapor in the pre-eruption magma. Other sources of sulfur in the edifice, such as the decomposition of sulfides and anhydride by hydrothermal water, were ruled out because the observed volumes of these mineral phases in the erupted materials were insufficient. Their approach has been used at other volcanoes e.g., Pinatubo (Gerlach et al., 19941, to constrain the source of measured gas emissions. To date, there is no method available to quantitatively constrain the source of the gases inside the resonating crack. A magmatic source is commonly inferred when the hypocenters of long-period events are within 1 km of a proposed magma body (Fischer et al., 1993; Casadevall et al., 1994). In this paper, crystal size population, tomographical, petrological, and geochemical data are used, in addition to observed seismicity and measured gas emissions, to test the hypothesis that the source of fluid in the resonating crack at Redoubt is in fact magmatic. The approach taken involves estimating the mass of exsolved H,O in the pre-eruption magma and comparing it to the mass of H,O estimated from a numerical model for the triggering mechanism for each of the long-period events in the December 13-14 precursory swarm (Morrissey, 1994; Morrissey and Chouet, in press). The mass of exsolved H,O from Redoubt magmas is calculated from the solubility of H,O-CO, as a function of pressure, fluid composition, and volume of H,O-CO, saturated magma assuming closed system degassing (Holloway and Bank, 1994). The calculation requires the ascent history of the magma through the crust. Three models for the magma ascent history have been proposed using different combinations of petrological, geochemical, and seismological data. One model is based on constraints from whole-rock

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chemistry, and earthquake locations (Nye et al., 1994), the second is based on constraints from mineral chemistry and textures, and earthquake locations (Swanson et al., 1994), and the third is based on constraints from tomography data (Benz et al., 1996). The ascent model used here is a hybrid of these three models and is consistent with all compiled data. The spatial distribution of magma within the upper 12 km of the crust prior to the eruption is constrained by integrating results from size distributions of plagioclase crystals, mineral and glass textures (Swanson et al., 1994) trends in major- and trace-element data (Nye et al., 1994) temporal variation of earthquake locations (Lahr et al., 1994), and tomographic data (Benz et al., 1996). What follows is a brief summary of data for the 1989-1990 eruptions at Redoubt Volcano including petrologic, geochemical and seismic observations. Next, results from a crystal size distribution analysis are presented along with an inferred magma ascent model. The third section presents the calculations for the mass of degassed steam and for the total mass of steam emitted through the resonating crack (Morrissey, 1994; Morrissey and Chouet, in press). A discussion follows on how shallow level magma degassing can explain the observed long-period seismicity, the eruptive onset, and total SO, and CO, emissions during the eruption.

2. Background 2.1. Summary of 1989-1990 Redoubt eruptions The 1989-1990 eruption at Redoubt Volcano erupted 0.05 km3 to 0.13 km3 of magma over a six month period (Gardner et al., 1994). Roughly half of this volume was erupted within the first two weeks when activity in the crater was mostly explosive (Miller, 1994). The initial eruption on December 14 produced tephra fall deposits that covered a > 200 km radius mainly to the NE (Scott and McGimsey, 1994). On the following day, four major explosive events occurred ejecting a much larger volume of material which produced pyroclastic and ice-supported debris flow deposits down the northern valleys. In the ensuing months, eruptive activity was typified by repetitive dome-building which included

the growth and destruction of 13 domes. Juvenile materials erupted were mainly medium-K calc-alkaline andesitic magma with more mafic andesitic magma erupted during the December eruptions (Nye et al., 1994). The December 14 and 15 events erupted banded pumices and frothy dome blocks indicating a heterogeneous magma body as reported by Gardner et al. (1994). The pumices were comprised of two types of porphyritic andesites distinguished primarily by the matrix glass compositions (Nye et al., 1994). The two andesites were found to have similar whole rock compositions and to contain the same phenocrysts and groundmass mineral assemblages. The matrix glass composition of one of the andesites was rhyolitic (77 wt.% SiO,) (Swanson et al., 1994) which will be referred to hereafter as rhyolitic andesites. The other andesite has dacitic glass (68-69 wt.% SiO,) (Swanson et al., 1994) and will be referred to hereafter as dacitic andesite. 2.2. Petrological

and geochemical

constraints

The whole-rock major oxides and trace-element chemistry of the andesites from the 1989-1990 Redoubt eruptions were analyzed by Nye et al. (1994). The andesites from the December eruptions were observed to have the widest range of SiO, values (58.2 to 62.5 wt.%), whereas, in andesites from post-December eruptions, the SiO, contents converge on 60 wt.% (fig. 5 of Nye et al., 1994). Trends in variation diagrams of the major oxides versus SiO, were all tight, linear arrays. These trends were interpreted by Nye et al. (1994) as the mixing arrays of two andestic magmas. This was further supported by trace-element data. Relative enrichment of highly incompatible elements varied substantially along the arrays which they argue could not have occurred by fractional crystallization of the magma. In addition, the ratio of Zr to Hf was observed to increase along the array which provided support for the magma mixing. Mineral and matrix glass chemistry and textures in the andesites from the 1989- 1990 Redoubt eruption were analyzed by Swanson et al. (1994). They found that the matrix glasses in the dacitic andesites had more Al, Na, Fe, Mg and Ca and less K than those in rhyolitic andesites. Intermediate compositions were detected in matrix glasses from banded

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pumices. As mentioned above, dacitic compositions were not detected in matrix glasses from andesites erupted after December 1989 (Swanson et al., 1994). In andesites from post-December eruptive events, matrix glass compositions were homogeneous and contained 78-79 wt.% SiO,. In all samples, Swanson et al. (1994) found plagioclase to be the most abundant mineral phase with hornblende and augite being the dominant mafic phases. Homblendes were observed to be subordinate to pyroxenes and Fe-Ti oxides only in the dacitic andesites. Mineral phases observed as phenocrysts were also found as microlites in most samples. Plagioclase crystals in both types of andesites had similar textural features, many were zoned and several had sieve-like textures with calcic overgrowths. The composition of the homblendes comprise two distinct populations, pargasitic hornblende and magnesiohomblende, which were found in both types of andesites. Reaction coronas appeared along the rims of hornblende crystals in andesites erupted after December 1989. In andesites erupted in December, very thin reactions rims were observed only on some homblendes in rhyolitic andesites. Similarities in phenocryst assemblages, textures, and bulk composition of the two andesites suggest that the magmas were derived from a parental magma that formed from the homogenization of an andesite magma and a more mafic magma. The two distinct matrix glasses were interpreted as reflecting the subsequent ascent and storage of the two andesitic magmas at different locations in the crust. The banded pumices and linear arrays of the major elements indicated that the two magmas mixed a second time at a shallower level just before the eruption. From the rarity of reaction coronas on hornblende crystals in December andesites, they argue that the shallow mixing event could not have occurred much later than hours to days before the eruption. Their argument assumes a mixing depth of < 7-8 km (2 kbar) which is the level at which hornblende would begin to break down in pre-eruption Redoubt magma (Rutherford et al., 1985). However, Nye et al. (1994) argue that matrix glass compositions indicate that a mafic magma intruded into a more silicic magma and mixed for a period of time sufficiently long (> days) to allow the composition of the melt to evolve towards rhyolite.

2.3. Seismological

constraints

The December 13-14 swarm of long-period seismicity was the unequivocal evidence that Redoubt Volcano was reactivating after 23 years of dormancy. The swarm was comprised by over 4,000 events (Stephens et al., 1994) which were recorded within 23 h of the eruptive onset. The fact that all the waveforms originated from 1.4 + 0.5 km below the crater and that the peak frequencies did not change with time suggested a single seismic source (Chouet et al., 1994). The peak frequencies characteristic of the long-period waveforms were modelled by Chouet et al. (1994) as resonance modes in a fluid-driven crack; the location and dimensions are shown in Fig. 1. In addition to the December 13-14 swarm of long-period events, several tremor events and a number of episodes of volcano-tectonic earthquakes were recorded in December 1989 (Power et al., 1994). Shallow tremor events lasting less than an hour preceded many of the explosive events that occurred on December 1.5. Episodes of volcano-tectonic earthquakes occurred after each of the eruptions. These earthquakes originated from depths immediately below the cluster of long-period events ranging from 1.5 km to 3 km below the crater as shown in Fig. 2.

Fig. 1. Sketch illustrating the spatial relationship between the crack acting as the source mechanism of the long-period seismicity, the magma reservoir, and the hydrothermal reservoir at Redoubt Volcano prior to the 1989-1990 eruption (from Chouet, 1996). Seismicity is caused by crack resonance trigger by the recurrence of a pressure transient in a flow of high-speed steam inside the crack (Chouet et al., 1994; Morrissey, 1994; Morrissey and Chouet, in press).

M.M. Morrissey / Journal of Volcanology and Geothermal Research 75 (1997) 321-335 N

S

111 0

5

10

15

20

DISTANCE (KM)

Fig. 2. North-south vertical profile showing earthquake locations and velocity structures in the upper 11 km beneath Redoubt Volcano taken from Benz et al. (1996). High velocity region is interpreted as a remnant conduit from earlier eruptions. The low velocity region is interpreted as the 1989-1990 eruption conduit. Hybrid seismic&y is defined by seismic events that have waveforms with features characteristic of both VT and LP events.

After December 15, the locations of the volcanotectonic earthquakes migrated down to depths exceeding 5-6 km. The distribution of earthquake locations was observed to form a narrow, vertical concentration of activity bending slightly to the northeast at a depth of 7 km (fig. 10 of Lahr et al., 1994). The overall pattern of concentrated seismic activity was interpreted by Lahr et al. (1994) to define a conduit in which magma migrated to the surface from a deep reservoir. A narrow linear feature is observed in the P-wave and S-wave velocity structures in the upper 11 km beneath the volcano imaged over a lo-km-wide area from 1,ooO volcano-tectonic earthquakes; 919 of which were recorded between December 14, 1989 and January 13, 1990 (Benz et al., 1996). Similar trends were observed in velocity variations in vertical cross sections of both P-wave and S-wave velocities. Within the upper 3 km, velocity variations were found to correlate well with the local geology exposed at the surface. At greater depths, velocity variations became more complex to the north of the summit and less so towards the south. Benz et al. (1996) noted that most of the deeper seismic&y occurred along the northern flank where velocity heterogenities are highest. The most apparent feature observed in this region was a cylindrical structure

325

approximately 1 km in diameter extending from 4 km to 9 km beneath the summit. In the S-wave sections, this structure appeared as a low-velocity (3.4 km/s) body bounded by two high-velocity (3.6-3.8 km/s) bodies at depths between 3 and 9 km. This structure appears in the P-wave section as a weak low velocity (5.8 km/s) body extending between 5 and 7 km. Another anomalous feature appears at a depth near 2.5 km beneath the crater, just to the north of the cluster of long-period events. This feature occupies a 2-3 km3 volume and has a velocity ratio of 1.85-1.87 reflecting reduced S-wave velocities. The combination of P-wave and S-wave velocity structures and the spatial and temporal variations of seismic events indicates that the cylindrical velocity structure defines the pathway of pre-eruption magma migration during its ascent from greater than 9 km to 3-4 km beneath the crater (Benz et al., 1996). The 2-3 km3 low S-wave velocity anomaly at depths of 2.5 km represents the uppermost extension of the conduit system. Earthquakes originating from depths of O-3 km align with the lower extension of the cylindrical structure and are interpreted as stresses related to the withdrawal of magma from the upper conduit after the onset of the eruption as observed in Fig. 2. The lack of precursory volcano-tectonic earthquakes supports the presence of the top of the magma column at a 3-4 km depth prior to the eruption.

3. Crystal size distributions An analysis is conducted on the size distribution of plagioclase crystals in andesites from the December 1989 events to study the crystallization history of pre-erupted magma during its ascent. When magma rises through the crust it experiences changes in the degree of undercooling, the difference between the ambient temperature and liquidus temperature. In general, the degree of undercooling governs the rates of crystal nucleation and growth (Kirkpatrick, 1981). For example, fewer crystals tend to nucleate when the degree of undercooling is small but they grow to a much larger size than the many small crystals that nucleate and grow when the degree of undercooling

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is large. Hence, if a magma resides at more than one depth during ascent through the crust, it will experience multiple stages of crystallization which will be recorded in the crystal size distribution (Marsh, 1988).

Crystal size distribution (CSD) theory defines, quantitatively, the nucleation and growth rates of crystals from the size population distribution (Marsh, 1988). The theory states that any change in the number of crystals per size per volume with time

(a)

Fig. 3. (a) Binary image of a dacitic andesite showing the area percent of plagioclase (blue). Matrix glass, voids and less silicic phases are highlighted in dark blue and oxides are highlighted in white. (b) Example of geometric parameters measured from a binary image for the CSD analysis of plagioclase crystals. Colors denote crystal size ranges, largest crystals are shown in blue and smallest crystals are shown in yellow.

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reflects changes in the flux of crystals in and out of the volume, and in the rate of crystals growing into and out of the size population. For steady state conditions in the crystallizing system, which is commonly assumed, the crystal size population density becomes an exponential function of the crystal size and growth rate (see Eq. (5) of Marsh, 1988). The population density, being a function of size, can then be estimated from the slope of the cumulative frequency distribution. By plotting the log of the population density versus the crystal size, the number of crystal nuclei and crystal growth rates can be calculated from the intercept and the slope of the line, respectively, if the residence time is known. Crystal sizes are obtained by measuring dimensions of the crystal phase of interest in rock samples.

18

I

1

6

-

181

I

P

5

l4

2

10

3 6

3.1. Samples and methods The sizes of plagioclase crystals were measured in 6 samples of andesites from the December 15, 1989 eruptive event. The samples were supplied by the Alaska Volcano Observatory staff. As described in Table 1, the samples are mostly banded frothy blocks of rhyolitic andesites and dacitic andesites. A total of eleven thin sections were made from these samples; five were made from bands of rhyolitic glass, and six were made from bands of dacitic glass. Lengths and widths of > 2000 plagioclase crystals were measured in each thin section using a backscatter electron image analyzer; microlites were not considered in this study so crystal < 55 pm in length were not counted. The ISI-OS130 Dual Stage scanning electron microscope at Los Alamos National Laboratory,

Table 1 Samples from December

15, 1989 eruptive event

Sample a

Location

Description

Thin sections b

NR90-30A RV90-49 RV90-50 RV90-53 RV90-57 RV90-72

RDN ’ CG d CG CG CG CG

Lt grey pumice Green-grey banded blk Lt grey rhyolitic blk Dk grey dacitic blk Lt-dk grey banded blk Lt grey banded blk

1R 1R, 2D 1R 1D 2R, 1D 2D

a In accordance with the AVO b The number of rhyolitic (R) ’ RDN seismic station located dPy rot lt’ as tc avalanche deposit

sample log. and dacitic (D) sections. on N flank. at Crescent Glacier.

6

61”“““‘*“““” 0 0.02

-..* 0.04

0.06

0.06

LENGTH (CM) Fig. 4. Crystal size population density versus crystal length plots for plagioclase crystals measured in rhyolitic andesites: (a) NR9030A; (b) RWO-49(Rl); (c) RV90-50; (d) RWO-57(Rl); (e) RWO-57(R2). Dashed and solid lines are the linear fits of the microphenocrysts (crosses) and phenocrysts (open circles), respectively.

operating at an acceleration voltage of 20 kV, was used to obtain backscatter images. Before measuring crystal dimensions, a qualitative electron dispersive energy (EDS) analysis was made of the area of interest to verify the composition of the matrix glass. Energy intensity ratios of Fe:Si, K:Si, Ca:Si, and Al:Si were used to distinguish dacitic glasses from rhyolitic glasses. The intensity for each element is the amount of energy emitted in the wavelength characteristic of that particular element. Dacitic glass have ratios that were 3-4 times higher than those

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10 14 10 6 18

6

0

0.02

0.04

0.06

0.08

0

0.02

0.04

0.06

0

0.08

0.02

0.04

0.06

0.08

LENGTH (CM) Fig. 5. Crystal size population density versus crystal length plots for plagioclase crystals measured in dacitic andesites: (a) NR90-49(Dl); (b) RV90-49(D2); (c) RV90-53; (d) RV90-57(Dl); (e) RV90-57(Dl); (f) RV90-72(D2). Dashed and solid lines are. the linear fits of the microphenocrysts (crosses) and phenocrysts (open circles), respectively.

measured in rhyolitic glass except for K:Si which was less in the dacitic glasses. Similar trends were observed in the major oxides measured in samples from December 15, 1989 by Swanson et al. (1994). Dimensions were measured in a binary image of plagioclase crystals with a Tracer Northern TN-8000 image analyzer as demonstrated in Fig. 3. The binary image was created from a backscatter image highlighting plagioclase crystals with all other phases set to background levels. Each image covered a 6.25 mm2 area of a thin section and contained > 200 crystals; 5-7 images were made of each thin section.

coefficients between 0.50 and 0.88. In dacitic andesites, the phenocryst population fit linear trends with correlation coefficients between 0.37 and 0.97. The scatter in phenocryst measurements can be attributed to the limited number found in each sample (Cashman, 1988). The slopes and intercepts of the two size populations in the CSD plots in Figs. 4 and 5 are listed in Tables 2 and 3. The slopes and intercepts of the microphenocrysts in the five rhyolitic andesites show little variation except for section RV90-57(Rl) (Table 3; Fig. 4d). The mean slope is - 281 cm-’

3.2. Results Plots of the natural log of the population density versus crystal size (CSD plots) are made from the cumulative frequency of the lengths of plagioclase crystals measured in each thin section. Two size populations of plagioclase crystals are identified from linear curve fitting of the population distribution curves in all plots (Figs. 4 and 5). In rhyolitic andesites, populations of microphenocrysts are defined by linear fits with correlation coefficients of at least 0.90 for crystal lengths between 60 pm and loo-160 pm (Table 2). The microphenocryst populations in dacitic andesites are defined by linear curve fits with correlation coefficients of at least 0.91 for crystal lengths between 60 pm and 160-210 pm (Table 3). The phenocryst populations in the rhyolitic andesites have linear fits with correlation

Table 2 Slooe and interceot values for plagioclase Sample

Size ( pm)

l/G7 (cm-‘)

phenocrvsts In no (cm-“)

R2

Rhyolitic glass andesites NR90-30A > 150 > 120 RV90-49(Rl) > 130 RV90-50 .lOO RV90-57(Rl) RV90-57(R2) > 160

- 83.6 -94.1 - 93.8 - 82.7 - 53.4

14.17 14.69 14.14 14.23 13.52

0.76 0.88 0.85 0.85 0.50

Dacitic glass andesites RV90-49(Dl) > 160 >200 RV90-49(D2) > 210 RV90-53 >200 RV90-57(D 1) > 210 RV90-72(Dl) RV90-72(D2) >200

- 134.4 - 66.3 -87.1 - 89.3 - 60.2 - 111.7

15.40 13.29 14.34 14.50 13.90 13.90

0.72 0.97 0.93 0.37 0.83 0.97

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(GT = 0.00355 cm) and the mean intercept is 16.69 cmP4 (no = 18.0 X lo7 cm-4). The steeper slope and smaller population size of the microphenocrysts in section RV90-57(Rl) can be partially attributed to the limited number of crystals analyzed in this section. The phenocrysts in these sections have shallower slopes with an average slope of - 81.5 cm-’ (GT= 0.0123 cm) and a mean intercept of 14.15 cmm4 (no = 1.4 X lo6 cmW4>(Table 2). As for the microphenocryst slopes and intercepts in the dacitic andesite, these too do not vary significantly and the slopes are slightly steeper than those in the rhyolitic sections; the opposite is true for the phenocryst populations. The population size of the microphenocrysts in the dacitic sections is on average (< 197 pm) larger than that in the rhyolitic sections ( < 132 pm). The microphenocrysts have a mean slope of - 254 cm- ’ (Gr = 0.00393 cm) and a mean intercept of 17.25 cme4 (no = 31.0 X lo7 cmW4). The phenocryst populations have an average slope of -91.5 (Gr= 0.0109 cm) and a mean intercept of 14.23 cmW4 (n” = 2.0 X lo6 cmm4>. Crystal nucleation and growth rates are calculated for the two size populations of plagioclase crystals in the two andesites from the average slopes and intercepts from the CSD plots. The slopes are the inverse products of the growth rate and residence time as defined in Eq. (23) in Marsh (1988). Therefore, growth rates can be calculated as functions of residence times which are assumed to represent the duration of crystal growth at a certain depth in the crust. Given the two size populations in samples of both andesites, the phenocrysts are interpreted to represent an early stage of crystallization, whereas the microphenocrysts represent a later stage of crystallization. This stage of crystallization did not occur during the final ascent that accompanied the December eruptions. The presence of microlite plagioclase ( < 60 pm in length) in the groundmass represents syn-eruptive crystallization or crystallization that accompanied magma mixing (Swanson et al., 1994). The residence time related to each stage of pre-eruption crystallization is constrained by comparing the calculated growth rates to values obtained by either direct measurements (e.g., Kirkpatrick, 1981) or estimated values from CSD data for plagioclase crystals that have been radiometrically dated (e.g., Cashman, 1988).

For the phenocrysts, the related residence time is constrained by comparing growth and nucleation rates calculated for time range between 1 day and 250 years, to values obtained from CSD data for radiometrically dated plagioclase phenocrysts in Mount St. Helens dacites (Cashman, 1988). The upper time limit of 250 years coincides with the last major historic eruption at Redoubt (Nye et al., 1994). For a residence time of 1 day, the growth rates for phenocrysts in the rhyolitic andesites and dacitic andesites are 1.42 X 1O-7 ems-, and 1.26 X 10e7 cm s-, respectively. For a residence time of 250 years, the growth rates for plagioclase in the rhyolitic andesites and dacitic andesites are 1.5 X lo- l2 cm sand 1.38 X lo-l2 ems-, respectively. The related nucleation rates (.I) for phenocrysts in the two andesites are 1.99 X 10-l cmW3s- and 2.52 X 10-l cmm3 s-, respectively for a residence time of 1 day, and 2.18 X 10m6 cme3 s- and 3 X lop6 cme3 s-, respectively, for a residence time of 250 years. Growth and nucleation rates of plagioclase crystals measured by Cashman (1988) in the samples of Mount St. Helens dacite from CSD data for similar size population are 3.0-10.46 X lo-‘* ems- and 4 .7-21 . 1 X lO-‘j cmm3 s-, respectively. These rates were determined from the mean slope and intercept (no) of CSD plots of these phenocryst populations were - 103.56 cm-’ (GQ-= 9.66 X 10m3 cmm3> and 1.19 X lo6 cmm4, respectively, and a residence time of 30-150 years constrained from uranium and thoTable 3 Slope and intercept values for plagioclase microcrysts Sample

Size

l/G7 (cm-‘)

In no km-4)

R2

Rhyolitic glass andesites NR90-30A < 150 RWO-49tRl) < 120 RV90-50 < 130 RV90-57(Rl)
- 200.9 - 282.0 - 276.4 -451.4 - 196.1

16.30 17.15 15.96 17.62 16.42

0.95 0.90 0.95 0.93 0.93

Dacitic glass andesites RV90-49(Dl) < 160 RV90-49(D2) <200 RWO-53 <210 RWO-57(Dl) <200 RV90-72(Dl) < 210 RV90-72(D2) <200

- 287.0 - 260.8 - 213.5 - 294.5 -215.0 - 255.4

17.65 17.24 16.64 17.83 16.77 17.36

0.92 0.95 0.97 0.91 0.97 0.95

(

pm)

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1000 Dacitic Andesite

3

5

100

Rhyolitic Andesite

T

Id

0

7

0 c

10 r

r d $ B

cl

1Y ‘:

i

0.0

/

/ I

I

0.1 e 0.01 1

/

lOY

1ooy

I

I

I

0.2

I-

I

0.4

I

I

I

0.6

I 0.8

I 1.0

CRYSTAL FRACTION Fig. 6. Crystal growth rates as a function of crystal volume fraction. Data points are measured values of plagioclase crystals growing in the Makaopuhi basaltic lake in Hawaii (Kirkpatrick, 1981). Vertical lines show the growth rates for specified times at the two measured volume fractions from the slopes of the CSD plots for the Redoubt andesites. The dashed segment of the basalt line is extrapolated from the trend between crystal fractions of 0.35 and 0.40.

rium age dating method. The consistencies in CSD data for phenocrysts in samples from these two volcanoes suggest that the Redoubt magmas may have started crystallizing at least 30-150 years before the eruption. The residence time corresponding to the second stage of crystallization is constrained by comparing growth rates calculated from the CSD results as a function of time to measured values for plagioclase crystals crystallizing in basalt (Kirkpatrick, 1981). Calculated growth rates for the microphenocryst populations in the rhyolitic andesites and dacitic andesites range from 5.0 X lo-* emsand 4.5 X lo-* ems-, respectively, for a residence time of 1 day, to 5.4 X lo-l3 ems- and 5.0 X lo-l3 ems-, respectively, for a residence time of 250 years. The range of growth rates for each andesite is plotted in Fig. 6 along with those measured as a function of crystal fraction for plagioclase crystals growing in Hawaiian lava lakes (Kirkpatrick, 1981). For the two Redoubt andesitic magmas, the crystal fractions used to construct Fig. 6 are roughly one half of the observed crystal fractions in the erupted andesites, based on the fraction of phenocrysts that comprise the total crystal population. The minimum time of

crystallization is estimated from the intersection of the measured growth rates in basalt with the ranges of growth rates for the two andesites. For the dacitic andesite, the intersection is at 15 years assuming a crystal fraction of 0.26. For the rhyolitic andesite, the intersection occurs at 8 years assuming a crystal fraction of 0.35. The crystallization rates in more silicic magmas can be at least an order of magnitude lower than in more mafic magmas. Crystals tend to grow faster in magmas with lower viscosities and with a smaller degree of undercooling, whereas the number of crystals nucleated is higher in a magma with a higher viscosity and with a larger degree of undercooling (Kirkpatrick, 198 1). Therefore, the

s

(b) 0

5

10

15

20

DISTANCE (KM) Fig. 7. North-south vertical profiles showing the locations of two Redoubt magmas along with earthquake locations and velocity structures observed by Benz et al. (1996) in the upper 11 km of the edifice. (a) shows the inferred locations of the two magmas at least 8-15 years before the eruption during the second stage of crystallization. (b) shows the dacitic andesite magma mixing with the rhyolitic andesite within hours to days of the eruption and the conduit extending from the reservoir containing mixed magma.

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M.M. Morrissey/ Journal of Volcanology and Geothermal Research 75 (1997) 321-335

maximum residence time for the microphenocrysts is constrained at 150 years in the dacitic andesite and 80 years in the rhyolitic andesite. This upper limit is probably too high given that the composition of the melt fraction was less evolved at the onset of this crystallization stage. Each stage of crystallization is assumed to represent a period of time during which the magma resided at a certain depth in the crust. Depths are constrained from petrological data as well as seismic data. The lack of reaction rims on hornblende phenocrysts and microphenocrysts indicate that this phase was stable prior to the eruption. From experimentally determined phase relations (Rutherford et al., 1985>, the stability of hornblende in the Redoubt andesites limits the level of ascent to 7-8 km (2 kbar). Using this depth as limiting storage depth, the storage depths related to the first stage of crystallization for both magma bodies are constrained at depths below 10 km. As shown in Fig. 7a, the second stage of crystallization in the rhyolitic andesite magma occurs within at least 15 years of the eruption at a depth of 7-8 km which coincides with the magma plexus defined in Fig. 2. The second stage of crystallization in the dacitic andesite magma begins within at least 8 years of the eruption upon ascending to a depth located just below the other magma body (8-9 km). The dacitic andesite eventually moves up into the overlying magma body and mixes with the more evolved magma within hours to days of the eruption as shown in Fig. 7b. This mixing scenario is suggested by the banded pumice reported by Gardner et al. (19941, the presence of dacitic andesite melt inclusions in phenocrysts contained in rhyolitic andesite reported by Gerlach et al. (19941, and the trends in major and trace element chemistry reported by Nye et al. (1994). The invasion of the dacitic andesite magma into the more evolved magma displaces mixed magma near the top of the reservoir into shallower regions of the crust. The observed downward migration of volcano-tectonic earthquakes from O-3 km to 6-10 km within the first two days of the eruption suggests that the first magma erupted was situated at depths between at least 3 km and 6 km. The pathway that the displaced magma took followed the one related to the 1966 erupted magma which would explain the lack of appreciable volcano-tectonic seismicity prior to the eruption. This

pathway then became the main conduit for the eruption.

4. Degassed H,O calculations To date, no data are available for the water content in the Redoubt magmas. However, the mass of exsolved H,O contained in the pre-eruption magma can be estimated from its solubility in magma given the magma ascent path and the volume of displaced magma. The presence of stable hornblende in the mineral assemblage in andesites erupted in December indicates that the magmas contained dissolved water and a coexisting H,O-rich vapor phase (Rutherford et al., 1985). From phase equilibrium experiments on magma compositions similar to the Redoubt magmas (Rutherford et al., 1985), the amount of dissolved H,O had to be at least 4.7 wt.% and the vapor phase has to contain at least 0.67 mole fraction of H,O. The vapor phase of the Redoubt magmas has also been shown by Gerlach et al. (1994) to have contained SO, and CO,, based on melt inclusion chemistry and COSPEC emission measurements. The appropriate solubility diagram for these calculations is that for a H,O-SO,-CO,saturated magma. For the purpose of the following calculations, it is assumed there is no significant

800

IT 8

600 400

0

1

2

3

4

5

6

Wt % Hz0

Fig. 8. H,O-CO, solubility diagrams in vapor saturated rhyolite as a function of pressure and mole fraction of water (X,zo) in the vapor phase (from Holloway and Bank, 1994). The ascent paths of the rhyolitic andesite (R) and dacitic andesite (D) magmas assuming closed system degassing are defined by the numbered points starting from a depth of 12 km (3 kbar) as described in the text.

332

M.M. Morrissey / Journal of Volcanology and Geothermal Research 75 (1997) 321-335

amount of pre-existing gas phase present at the time of the ascent induced degassing. To calculate the amount of exsolved H,O associated with the various stages of magma ascent, closed system degassing is assumed. Fig. 8 shows the solubility of a H,O-CO,-saturated rhyolitic magma as a function of pressure and fluid (vapor) composition. Solubility diagrams for intermediate magma compositions are not available; the use of the solubility diagram for an evolved magma composition will result in conservative estimates. In these calculations, the effect of SO, is accounted for by roughly doubling the concentration of CO,; this is indicated by the available CO, and S solubility changes with pressure (Carrol and Rutherford, 1985). Plotted in the figure are the ascent paths for each magma which start with 0.67 mole fraction of H,O in the coexisting fluid (Rl and Dl). Upon ascending from a depth of at least 12 km (3 kbar) to 7-8 km (2 kbar) where it experiences its second stage of crystallization, the rhyolitic andesite magma exsolves 0.2 wt.% of H,O and 350 ppm of CO,. This segment of the ascent path is defined by points Rl and R2 (Fig. 8). The dacitic andesite exsolves an estimate of 0.07 wt.% H,O and 180 ppm of CO, from its ascent from at least 12 km (3 kbar) to 8-9 km (2.5 kbar) as defined by points Dl and D2. An additional 0.1 wt.% of H,O and 160 ppm of CO, are exsolved from this magma when it moves up into and mixes with the overlying magma. D2 and D3 define this ascent path.

The last ascent of magma in the crust before eruption is the fraction that is displaced when the two magmas mix (Fig. 7b). Assuming an average storage depth of 4 km (1 kbar) for the displaced magma, approximately 0.6 wt.% Hz0 and 200 ppm of CO, are exsolved; R2 to R,,3 in Fig. 8. The total mass of exsolved H,O related to each step along the ascent path depends on the volume of magma. This relation is expressed by the following equation which is used to tabulate the masses of exsolved fluids listed in Table 4: Mass exsolved = 10m9A,p, +V

(1)

where A, is the concentration of exsolved H,O or CO,, p is the density of the andesite which is assumed to be 2.30 kg/m3, and 4 is the melt fraction in the magma. For simplicity, a value of 0.55 and 0.75 are chosen for 4 for the rhyolitic andesite and dacitic andesite, respectively, which are the fractions of matrix glass measured by Swanson et al. (1994) in the December andesites. The volume of displaced magma V is constrained from the volume of erupted magma, which sets the lower limit at 0.05-0.13 km3. An upper limit is estimated from the volume of the conduit extending from the 7-8-kmdeep magma reservoir defined by the seismicity (Benz et al., 1996; Lahr et al., 1994). Assuming a cylindrical geometry with a radius of 0.5 km and length of 5 km, approximately 3.9 km3 of magma would occupy this conduit. Equal volumes are assumed for the two andesites, therefore V is equal to

one-half the upper and lower volume limits. Table 4 Volatile exsolution

along the ascent paths of Fig. 8 0.065 km3

1.95 km3

Exsolved Hz 0 (Mtons) Rl to R2 0.63 R2toR,,3 1.90 Dl to D2 0.30 D2 to D3 0.39

1.64 4.93 0.78 1.oo

49 148 23 30

Total

8.30

250

Ascent path

0.025 km3

3.20

Exsolved CO, lMtonsl Rl to R2 0.110 R2toR,3 0.063 Dl to D2 0.078 D2 to D3 0.069

0.288 0.164 0.202 0.179

8.63 4.93 6.05 5.38

Total

0.833

24.99

0.320

5. Discussion Decompression degassing of a H,O-CO,-SO, saturated magma from lo-12 km to 3-4 km is found to yield 3.2 to 250 Mtons of H,O (Table 4) which would adequately supply the resonating crack with the quantity of superheated steam required by the fluid dynamic model for triggering resonance (Chouet et al., 1994; Morrissey, 1994; Morrissey and Chouet, in press). Using results from their numerical model of unsteady choked flow through a crack, Morrissey and Chouet (in press) demonstrate how an oscillating shock front can produce all the characteristics of the pressure transient constrained by the fluid-filled crack

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of Volcanology and Geothermal Research 75 (1997) 321-335

model (Chouet et al., 1994). The fluid in the flow model is assumed to be superheated steam having an average sound speed of 850 m/s which is within the range constrained by the Chouet et al. (1994) seismic model. This sound speed is also within the range of sound speeds corresponding to the temperatures and pressures of pre-eruption Redoubt magmas. The temperatures of the Redoubt magmas have been estimated at 840°C to 950°C from Fe-Ti oxides (Swanson et al., 1994) which according to fig. 2 of Kieffer (1977), correspond to sound speeds of 850-900 m/s at pressures I 220 bar. The flow of superheated steam in the model is assumed to be continuous over the entire 23 h period before the eruption. Over this time period, the total mass of steam transported through the crack is estimated at 25-50 Mtons, based on a mass flux of 300-600 tons/s at the crack inlet. From Table 4, decompressing a combined total of magma of at least 0.2 km3 from a depth of 12 km would yield the necessary amount of H,O. This volume corresponds to a conduit with a diameter of 180 m which is twice the average diameter of all 14 domes that grew during the eruption as reported by Miller (1994) and more realistic than that constrained by seismicity. For comparison, the diameter of the feeder conduit for the Mount St. Helens domes was estimated at 25 m (Swanson and Holcomb, 1990). An alternative source of fluid for the unsteady choked flow model would be hydrothermal water. Seismic observations described in an earlier section, indicate that the top of the magma conduit was at 3-4 km beneath the crater prior to the eruption. Hydrothermal waters at depths of l-3 km have temperatures of lOO-300°C in the presence of a heat source such as the shallow magmabody (S.E. Ingebritsen, U.S. Geological Survey, written communication). At this temperature range and depth, supercritical water has a sound speed of 900 to 1500 m/s, which is within the range required for the fluid-driven crack model of Chouet et al. (1994). Although hydrothermal waters have the essential sound speed, the circulation rate of a thermally-driven flow is up to ten orders of magnitude lower than the inflow velocity required for the flow model. The flow rate driven by a thermal anomaly can be estimated by a form of Darcy’s Law for a one-dimensional, steady flow. Assuming the pressure gradient

333

is near the hydrostatic gradient, Darcy’s Lawis expressed as (Turcotte and Schubert, 1982): u:= -[‘V,g+-T,)]/p

(2)

where K is the permeability, p,, is a reference density, g is gravity, and LYand p are the fluid thermal expansivity and dynamic viscosity, respectively. Taking values of the difference between the magmatic temperature and surrounding temperature (T T’) as 300°C and permeabilities of 10e6 to lo- l6 m2, give flow rates between lop9 and 10 m/s. Appropriate circulations rate are achieved for permeabilities > 10m6 m* which are characteristic of fracture flow. What makes hydrothermal circulation an undesirable source for the fluid is that the unsteady choked flow model requires a mass flux of 300-600 tons/s to be sustained for over 23 h (Morrissey, 1994; Morrissey and Chouet, in press). The highest mass fluxes in a hydrothermal system are observed at geysers which have been estimated to be < 0.5 tons/s sustained for less than 10 min (White, 1967). The mass flux and duration of continuous flow required to supply the fluid for the unsteady choked flow model are not characteristic for a hydrothermal system. These values are characteristic of discharge rates of high pressure fluids such as gases released from a confined body of magma. Degassing of magma in the conduit extending from the 6-lo-km-deep reservoir to at least 3 km within hours to days of the eruption not only explains the source of H,O flowing through the resonating crack but it also explains the source of total SO, and CO, emissions measured during the eruption. A total of 0.93-1.03 Mtons of SO, emissions were measured using COSPEC and TOMS (Casadevall et al., 1994). From the SO, measurements, a total of 1.30-1.44 Mtons of CO, was estimated by Gerlach et al. (1994) using a weight ratio of 1.4 CO,/SO,. Gerlach et al. (1994) have demonstrated that syn-eruptive degassing of the volume of erupted magma would yield a total mass of SO, and CO, that would be two order of magnitude lower than that measured by TOMS and COSPEC (Casadevall et al., 1994). They propose that the source of emitted gases is an excess vapor phase that accumulates in the magma at depth. Combining this explanation with the ascent model present here, the volume of excess CO,-SO, vapor phase could form

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in the mixed magma as it ascends into the conduit from the 6-10 km reservoir within hours to days of the eruption as depicted in Fig. 7b. Based on the solubility of CO, and H,O saturated magma (Table 4), the total volume of displaced magma 0.2 km3 would release approximately 2.5 Mtons of CO, which assumes a concentration ratio of CO, to SO, of 2: 1, and breaks down to 1.66 Mtons of CO, and 0.83 Mtons of SO,.

Acknowledgements I thank all at the Alaska Volcano Observatory whose samples I used in this study, especially Tina Neal who selected and shipped all samples. Much of the SEM work was conducted at Los Alamos National Laboratory under the support of Ken Wohletz. The manuscript benefitted from the careful reviews by Herbert Shaw, Terry Gerlach, Malcolm Rutherford and Jake Lowenstem.

6. Conclusions Shallow level degassing of a H,O-CO,-SO, saturated mixed magma has been demonstrated to be the probable source of superheated steam related to the swarm of long-period events that occurred prior to the 1989-1990 eruptions at Redoubt Volcano. The volume of mixed magma 0.05-3.9 km3 situated in the conduit beneath the crater is found to yield 3.2-250 Mtons of H,O via ascent of magma from at least 12 km based on solubility data of H,O and CO, saturated magmas. This mass is comparable to the 25 to 50 Mtons of superheated steam required for the unsteady choked flow model (Morrissey, 1994; Morrissey and Chouet, in press) suggesting that the swarm of long-period seismicity was related to the flow of magmatic gases through a fracture 1.4 rfr 0.5 km below the crater. The ascent history of the pre-eruption Redoubt magmas inferred from coupling petrological, geochemical, and seismological data provides insight towards understanding the seismicity at Redoubt Volcano prior to December 14, 1989. The lack of appreciable volcano-tectonic earthquakes and the rapid onset of long-period events within days of the eruption is interpreted as indicating that magma was actively degassing at depths as shallow as 3 km. The continuous flow of magmatic fluids from the conduit and through the resonating crack interacted with the hydrothermal system and subsequently pressurized the shallow dome structure which led to its ultimate failure. Based on the unsteady choked flow model (Morrissey, 1994; Morrissey and Chouet, in press), this pressurization explains the trends in magnitude and time interval between each long-period event. The eruptive scenario described above is in agreement with that proposed earlier by Chouet et al. (1994).

References Benz, H.M., Chouet, B.A., Dawson, P.B., Lahr, J.C., Page, R.A. and Hole, J.A., 1996. Three-dimensional P and S-wave velocity structure of Redoubt Volcano, Alaska. J. Geophys. Res., 101: 8111-8128. Carrol, M.R. and Rutherford, M.J., 1985. Sulfide and sulfate saturation in hydrous silicate melts. J. Geophys. Res., 90: 601-612. Casadevall, T.J., Doukas, M.P., Neal, C.A., McGimsey, R.G. and Gardner, C.A., 1994. Emission rates of sulfur dioxide and carbon dioxide from Redoubt Volcano, Alaska during the 1989-1990 eruptions. In: T.P. Miller and B.A. Chouet (Editors), The 1989-1990 Eruptions of Redoubt Volcano, Alaska. J. Volcanol. Geothetm. Res., 62: 519-530. Cashman, K.V., 1988. Crystallization of Mount St. Helens 198086 dacite: A quantitative textural approach. Bull. Volcanol, 50: 194-209. Chouet, B.A., 1986. Dynamicsof a fluid-driven crack in three dimensions by the finite difference method. J. Geophys. Res., 91: 13,967-13,992. Chouet, B.A., 1988. Resonance of a fluid-driven crack: Radiation properties and implications for the source of long-period events and harmonic temor. J. Geophys. Res., 93: 4375-4400. Chouet, B.A., 1992. A seismicmodel for the source of long-period events and harmonic tremor. In: P. Gasparini, R. Scarpa and K. Aki (Editors), Volcanic Seismology, IAVCEI Proc. in Volcanology. Springer-Verlag, Berlin, pp. 3,133-3,156. Chouet, B.A., 1996. Long-period volcano seismicity: Its source and use in eruption forecasting. Nature, 380: 309-3 16. Chouet, B.A, Page, R.A., Stephens, C.D., Lahr, J.C. and Power, J.A., 1994. Precursory swarms of long-period events at Redoubt Volcano (1989-1990), Alaska: Their origin and use as a forecasting tool, In: T.P. Miller and B.A. Chouet (Editors), The 1989-1990 Eruptions of Redoubt Volcano, Alaska. J. Volcanol. Geotherm. Res., 62: 95-136. Fischer, T.P., Monissey, M.M., Calvache, M.L.V., Gomez, D.M., Torres, R.C., Stix, J. and Williams, S.N., 1993. Correlation between SO, flux and long-period seismicity at Galeras volcano. Nature, 368: 135-137. Gardner, C.A., Neal, C.A., Waitt R.B., and Janda, R.J., 1994. Proximal pyroclastic deposits from the 1989- 1990 eruption of

M.M. Morrissey/ Journal of Volcanology and Geothermal Research 75 (1997) 321-335 Redoubt Volcano, Alaska - stratigraphy, distribution, and physical characteristics. In: T.P. Miller and B.A. Chouet (Editors), The 1989-1990 Eruptions of Redoubt Volcano, Alaska. J. Volcanol. Geotherm. Res., 62: 213-250. Gerlach, T.M., Westrich, H.R., Casadevall, T.J. and Finnegan, D.L., 1994. Vapor saturation and accumulation in magmas of the 1989-1990 eruption of Redoubt Volcano, Alaska. In: T.P. Miller and B.A. Chouet (Editors), The 1989-1990 Eruptions of Redoubt Volcano, Alaska. J. Volcanol. Geotherm. Res., 62: 317-339. Holloway, J.R. and Bank, J.G., 1994. Applications of experimental results to C-O-H species in natural melts. In: M.R. Carrel and J.R. Holloway (Editors), Reviews of Mineralogy. Mineral Sot. Am., 30: 281-323. Kieffer, S.W., 1977. Sound speed in liquid-gas mixtures: Waterair and water-steam. J. Geophy. Res., 82: 2895-2905. Kirkpatrick, R.J., 1981. Kinetics of crystallization in igneous rocks In: A.C. Lasaga and R.J. Kirkpatrick (Editors), Reviews of Mineralogy, Mineral Sot. Am., 8: 321-398. Lahr, J.C., Chouet, B.A., Stephens, C.D., Power, J.A. and Page, R.A., 1994. Earthquake location and error analysis procedures for a volcanic sequence: Application to the 1989-1990 eruptions at Redoubt Volcano, Alaska. In: T.P. Miller and B.A. Chouet (Editors), The 1989-1990 Eruptions of Redoubt Volcano, Alaska. J. Volcanol. Geotherm. Res., 62: 137-152. Marsh, B.D., 1988. Crystal sizedistribution (CSD) in rock and the kinetics and dynamics of crystallization I: Theory. Contrib. Mineral. Petrol., 99: 277-291. Miller, T.P., 1994. Dome growth and destruction during the 1989-1990 eruption of Redoubt Volcan. In: T.P. Miller and B.A. Chouet (Editors), The 1989-1990 Eruptions of Redoubt Volcano, Alaska. J. Volcanol. Geotherm. Res., 62: 197-212. Morrissey, M.M., 1994. Magmatic fluids and long-period seismicity: A geological and fluid dynamic perspective. Ph.D. Thesis, Arizona State University, 124 pp. Morrissey, M.M. and Chouet, B.A., in press. A numerical investigation of choked flow dynamics and its application to the triggering mechanism of long-period events at Redoubt Volcano, Alaska. J. Geophys. Res. Nye, C.J., Swanson, S.E., Avery, V.F. and Miller, T.P., 1994. Geochemistry of the 1989-1990 eruption of Redoubt Volcano: Part II. Whole-rock major and trace chemistry. In: T.P. Miller

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and B.A. Chouet (Editors), The 1989-1990 Eruptions of Redoubt Volcano, Alaska. J. Volcanol. Geotherm. Res., 62: 429-452. Power, J.A., Lahr, J.C., Page, R.A., Chouet, B.A., Stephens, CD., Harlow, D.H. and Murray, T.L., 1994. Seismic evolution of the 1989-1990 eruption at Redoubt Volcano, Alaska. In: T.P. Miller and B.A. Chouet (Editors), The 1989-1990 Eruptions of Redoubt Volcano, Alaska. J. Volcanol. Geotherm. Res., 62: 69-94. Rutherford, M.J., Sigurdsson, H., Carey, S. and Davis, A., 1985. The May 18, 1980, eruption of Mount St. Helens melt composition and experimental phase equilibria. J. Geophy. Res., 90: 2929-2947. Scandone, R. and Malone, S.D. 1985. Magma supply, magma discharge and readjustment of the feeding system of Mount St. Helens during 1980. J. Volcanol. Geotherm. Res., 23: 239-262. Scott, W.E. and McGimsey, R.G., 1994. Character, mass, distribution, and origin of tephra-fall deposits of the 1989-1990 eruption of Redoubt Volcano, south-central Alaska. In: T.P. Miller and B.A. Chouet (Editors), Tbe 1989-1990 Eruptions of Redoubt Volcano, Alaska. J. Volcanol. Geothenn. Res., 62: 251-272. Stephens, C.D., Chouet, B.A., Page, R.A., Lahr, J.C. and Power, J.A., 1994. Seismological aspects of the 1989-1990 eruptions at Redoubt Volcano, Alaska: the SSAM perspective. In: T.P. Miller and B.A. Chouet (Editors), The 1989-1990 Eruptions of Redoubt Volcano, Alaska. J. Volcanol. Geotherm. Res., 62: 153-182. Swanson, D.A. and R.T. Holcomb, 1990. Regularities in growth of the Mount St. Helens dacite dome 1980-1986. In: J.H. Fink (Editor), Lavas and Domes. Springer-Verlag, New York, pp. 3-24. Swanson, SE., Nye, C.J., Miller, T.P. and Avery, V.F., 1994. Magma mixing in the 1989-1990 eruption of Redoubt Volcano: Part II. Evidence from mineral and glass chemistry. In: T.P. Miller and B.A. Chouet (Editors), The 1989-1990 Eruptions of Redoubt Volcano, Alaska. J. Volcanol. Geotherm. Res., 62: 453-468. Turcotte, D.L. and Schubert, G., 1982. Geodynamics. John Wiley and Sons, New York. White, D.E., 1967. Some principles of geyser activity, mainly from Steamboat Springs, Nevada. Am. J. Sci., 265: 644684.