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Factors controlling the temporal evolution of explosive eruptions
Joumalof volcanology and geothennal research
ELSEVIER
Journal of Volcanology and Geothermal Research 72 (1996) 71-83
Factors controlling the temporal evolution of explosive eruptions Roberto Scandone
I
Dipartimento di Geofisica e Vulcanologia, Universita di Napoli" Federico II", Largo San Marcellino 10, 80138 Naples, Italy
Received 23 September 1994; accepted 3 December 1995
Abstract A simple model, based on observations of recent explosive eruptions, is proposed for the patterns of variable discharge rate during explosive eruptions. Such eruptions are characterized by a slowly increasing waxing phase and a phase of more rapid waning activity. This behaviour is associated with volatile saturation of the magma and a rigid response of the magma chamber walls until brittle fractures occur. The opening of the reservoir and the escape of the vapour phase cause an initial decompression of the magma. The emptying of the reservoir and the rigid behaviour of the wall rocks permit a further decompression of the magma during the course of the eruption. The waxing phase is related to a progressive increase in the rate of vesiculation of saturated melt that re-equilibrates the decompression of the magma. The following increase of discharge rate permits an early Plinian phase which culminates in collapse of the eruption column and emplacement of pyroclastic flows. The emission of abundant lithic fragments, making up the roof of the chamber, signals the beginning of the waning phase. The eruption lasts until the vesiculation, caused by the initial decompression and emptying of the chamber, can counterbalance the lithostatic load. Collapse of the chamber occurs as soon as the pressure becomes lower than the lithostatic load by an amount similar to the strength of rocks. The collapse tends to re-establish the original pressure conditions preventing further vesiculation and thereby resealing the magma chamber.
1. Introduction
A schematic model of an erupting volcano is that of a subterranean reservoir connected with the surface through a narrow conduit. The magma flow is driven by the pressure gradient between the reservoir and the surface whereas the physical properties of the magma and the geometry of the conduit control the peculiarities of the flow. Several authors have analyzed the effect of the
I Present address: III Universita di Roma, Dipartimento di Fisica "E. Amaldi", Via Vasca Navale 84, 00146 Rome, Italy; e-mail: [email protected]
0377-0273/96/$15.00 Published by Elsevier Science B.V. SSDI 0377-0273(95)00086-0
geometry of the conduit on the eruptive character of the eruption (see, for example, Wilson et aI., 1980; Papale and Dobran, 1994). These models hypothesize a constant pressure in the reservoir so that the mass flux is dictated by the difference between the pressure in the reservoir and the pressure at the surface. If the geometry of the pathway does not flare out somewhere by a required amount, then the flow may be choked, the exit pressure is greater than atmospheric and the discharge rate is less than it would be in the non-choked case (Wilson et aI., 1980; Wilson and Head, 1981; Giberti and Wilson, 1990). A flaring of the conduit, caused, for example, by vent erosion at the surface, may allow a transition to supersonic flow with a reduction of the exit
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R. Scandone / Journal of Volcanology and Geothermal Research 72 (1996) 71-83
pressure and therefore an increase of the overall pressure gradient and the discharge rate (Wilson et al., 1980; Wilson and Head, 1981). In this paper I take the different approach of analyzing the variations of the pressure gradient caused by factors independent of the geometry of the conduit, but related to the "structural" model of a volcano under observation. An overpressure above the lithostathic value in a magma reservoir is often invoked as necessary to start an eruption (Blake, 1981; Parfitt et al., 1993). The overpressure may originate by different mechanisms: (1) ex solution of a volatile phase or intrusion of new magma into a reservoir (Blake, 1981, 1984; Tait et al., 1989) with development of an overpressure above the strength of the rocks surrounding the magma chamber which causes the opening of fractures (Rubin, 1993); (2a) instability of fluid-filled cracks driven by the Peach-Koehler pseudo-buoyancy force (Weertman, 1971) due to ex solution of a gas phase or crack propagation by stress corrosion (Anderson and Grew, 1977); (2b) positive density contrast of liquids lighter than surrounding rocks (Ryan, 1987) deriving, for example, from the differentiation in a magma reservoir with the accumulation of a light differentiate at the top of chamber and the development of stress above the strength of the rocks. I will try to show that the overpressure, available to drive the magma out of the chamber, is controlled by the mechanical behaviour of the chamber walls. I will discuss two extreme cases: (I) The "elastic reservoir", already proposed by Machado (1974) and Wadge (1981), is characterized by an elastic readjustment of the reservoir walls as magma is removed. (II) The "rigid reservoir" has rigid walls at least for some time after the beginning of the eruption. I will use the observations made during the eruptions of May 18, 1980 of Mt. St. Helens and of June 15, 1991 of Mt. Pinatubo to illustrate the evidence for and consequences of this assumption. Obviously, in nature there is no such clear-cut difference in the behaviour of actual volcanoes, but these examples may help to illustrate the different physical processes that govern the development of
the eruption if one mechanical behaviour prevails over the other.
2. The elastic model Here I will recall only the main characteristics of the model as Machado (1974) and Wadge (1981) give a more detailed discussion. A perfectly elastic medium, by definition, transmits instantaneously the elastic strain stored by the rocks surrounding the magma chamber and produces a continuous squeezing of the chamber until the overpressure is exhausted. The elastic energy is, by definition, made immediately available upon the opening of the reservoir, so the maximum pressure within the magma chamber is observed right at the beginning of the eruption. Some basaltic eruptions show an unsteady behaviour with a rapid waxing phase and a slower waning in the rate of magma discharge (Fig. la) [several cases are shown in Scandone (1979) and an extensive bibliography is provided by Wadge (1981), although contrasting examples of effusion rate variations have been exhibited during the Pu'u '0'0 eruption (Wolfe et al., 1988)]. The rapid initial waxing phase of these effusive
Magma discharge rate max 100 -10 3 m3 /sec
time (days-months) Effusive
type
Magma discharge
1O~::: I 3
m /sec
)\~
~
time (hours-days) Explosive
type
Fig. 1. Examples of trends of variation of magma discharge during effusive and explosive eruptions.
R. Scandone / Journal of Volcanology and Geothermal Research 72 (J996) 71-83
eruptions may be interpreted as due to the opening of a magma-filled crack as soon as it reaches the surface. The high initial discharge is a result of the high overpressure immediately provided through elastic readjustment (Wadge, 1981). The overpressure is rapidly accommodated by the drainage of magma. The exponential waning of magma discharge is interpreted as due either to the elastic decompression of the magma (Machado, 1974), or of the rocks surrounding it (Wadge, 1981). The deformations of the flanks of volcanoes observed prior to several eruptions are well explained in terms of an elastic deformation of the rocks surrounding a pressure source (Mogi, 1958) so supporting the model proposed by Wadge. A purely elastic model is used by Stasiuk et aI. (1993) to account for second-order variations of flow rate in basaltic eruptions. They explain departures from the trend of exponential decay of flow rate as due to the relative changes in viscosity, conduit dimensions and thickness of lava over the vent. They estimate that the overpressure necessary to drive the eruptions of several volcanoes is of the order of 10-20 MPa. During the Pu'u '0'0 eruption of Kilauea there was a typical behaviour that contrasted with the one shown previously. A gradual buildup in the eruption rate occurred after the start of the eruption and then there was an approximately exponential decrease. This peculiar feature has been explained by Parfitt and Wilson (1994) as due to a non-newtonian behaviour of the magma. One possible explanation of the relevance of the rheological properties of magma during this eruption is that the volcano is in an open conduit condition since the beginning of the eruption in 1983. In this case, the magma flow, driven by minor variations of stress, is strongly controlled by the magma yield strength and geometry of the conduit (Parfitt and Wilson, 1994). The summit inflation of Kilauea and the subsequent deflation related to the eruptive episodes of, the Pu'u '0'0 vent are of the order of 5-30 microradians (Parfitt and Wilson, 1994) while those observed after longer periods of inactivity are of the order of hundreds of microradians (Banks et aI., 1989). The larger stresses implied during these other episodes likely produce high initial eruption rates not influenced by the yield strength of the magma.
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3. The rigid model 3.1. Introduction In this model I hypothesize that the magma chamber walls do not deform elastically in response to a pressure change within the reservoir, but react with a brittle failure when a sufficiently high stress is applied. This assumption is probably justified for silicic volcanoes. A long-lived magma chamber is often associated with silicic volcanoes (see, for example, Smith, 1979: Hildreth, 1981). The cooling of the liquid and crystallization at the walls (McBirney, 1980) may provide a carapace which behaves as a rigid envelope. The occurrence of a high-velocity layer at the top of the Mt. St. Helens magma chamber (Lees, 1992) is indicative of this occurrence. A similar occurrence is also reported at Vesuvius and at other volcanoes (P. Gasparini and S. Malone, pers. commun.). The important point to be noted is that a rigid behaviour of the walls of a magma chamber, after the start of the eruption, makes it impossible to squeeze the magma out of the chamber. A close analogy to this case is the opening of a champagne bottle: a sudden opening of cracks (the removal of the "cork") and escape of volatiles through the newly formed conduit cause the sudden decompression of the magma and it attempts to equilibrate to the new conditions. The overpressure driving the magma to the surface is made available, when the reservoir ruptures and is connected with the atmosphere, in the form of stored internal energy of the magma (see, for example, Burnham, 1972, 1985). This last mechanism occurs through expansion of already ex solved H 2 0, if present, and exsolution of additional H 2 0 upon decompression of the magma body. A geological analogy of this process is, for example, the flashing of a geothermal reservoir. 3.2. Chamber wall behaviour during the eruption of May 18, 1980 of Mt. St. Helens Scandone and Malone (1985) analyzed the seismicity of Mt. St. Helens occurring mostly after 17:49 on May 18, 1980 and identified an earthquake-free volume, whose top was at about 7 km depth, interpreted as the magma chamber of the volcano. These
R. Scandone / Journal o/Volcanology and Geothermal Research 72 (1996) 71-83
74
authors have shown that, during the major explosive eruptions of Mt. St. Helens of May 18 and 25, 1980 there was a readjustment of the chamber walls in response to the drainage of magma. The readjustment occurred as a brittle fracture of rocks producing earthquakes, characterized by a high-frequency signal around the reservoir. A similar phenomenon, occurred also after the explosive eruptions of June 12, July 22, August 7 and October 16, 1980. However, no occurrence of deep seismicity was observed during or after the effusive dome-building eruptions of the following years. Barker and Malone (1991) gave further support to the idea of chamber wall readjustment by analyzing the focal mechanism solutions of the post-eruption earthquakes. They found that the focal mechanisms were well explained by a stress concentration caused by a decrease in pressure within a cylindrical magma chamber at a depth of 7 -11 km, in the presence of the regional stress field. Shemeta and Weaver (1986) made a detailed analysis of the temporal readjustment of the chamber wall following the onset of the main eruption by analyzing the rate of earthquake occurrence on May 18 and 19. They subdivided this time interval into five periods as shown in Table 1. During the first two periods (Fig. 2), the earthquake release rate was low (low cumulative seismic moment) and the events were located in a volume below the volcano at
one
t. C!l
3
~6 ;5
9
~
12
C!lC!l
C!l
15 18
C!l
E
(!)
g>
(!)
(')
E
09
11
13
15
17 19 21 23 Hours (PDT)
01
03 05
07
Fig. 2. Depth-time distribution of earthquakes in the first day of the May 18, 1980 eruption of Mt. St. Helens. The vertical bar indicates the location of the magma chamber (slightly modified after Shemeta and Weaver, 1986).
depths ranging between 3 and 9 km, most earthquakes occurring between 3 and 6 km, with only a few events as deep as 12 km. During period 3, there was the maximum release of seismic energy (about 3/4 of the total), and a scatter in epicentral distribution as wide as 3 km, with most events occurring at depths between 2 and 8 km, and only a few between 8 and 12 km. During periods 4 and 5 there was a decrease in the release of seismic energy, and a deepening of events (4-14 km) with a lack in the shallow portions of the system. After 19:00 on May 18 nearly all events were located at depths between 5 and 12 km. I propose that the low rate of occurrence of
Table 1 The rate of earthquake occurrence on May 18 and 19, 1980 at Mt. St. Helens Time (PDT) May 18, 1980
Character of seismic activity
Character of volcanic activity
08.49-11.30
Minor seismic activity with most earthquakes between 3 and 6 km
Plinian column
2
11.30-14.15
Minor seismic activity with most earthquakes between 3 and 6 km; beginning of strong ground shaking at 11:40
Plinian column and beginning of pyroclastic flow emplacement
3
14.15-16.30
Maximum release of seismic energy with most earthquakes between 2 and 8 km
Pyroclastic flow emplacement; maximum output between 15:00 and 16:30
4
16.30-17.30
Decrease of seismic energy with deeper events (4-14 km)
Return to plinian column and decrease of magma discharge
5
17.30-08.30 of May 19
Decrease of seismic energy, after 19:00, all earthquakes are between 5 and 12 km
Final decrease of activity
Period
C!l
Q.
R. Scandone / Journal of Volcanology and Geothermal Research 72 (1996) 71-83
earthquakes during the first two periods is evidence of the absence of collapse. Extensive collapse of the magma chamber is signalled by the high rate of occurrence of deep earthquakes whose occurrence has been commonly observed after explosive eruptions at silicic volcanoes.
3.3. Expansion of magma within the chamber Johnson et al. (1994) suggest that volatile saturation occurs in magma chambers prior to silicic explosive eruptions. Strong evidence for an excess gas phase is found, for example, in the 1991 eruption of Mt. Pinatubo (Rutherford and Devine, 1991). Gerlach et ai. (in press) suggest that the dacite erupted on 15 June 1991 at Mt. Pinatubo was vapour saturated at depth prior to eruption to account for the immediate source of excess sulfur detected by remote sensing during the eruption. The occurrence of a vapour phase in the reservoir is invoked also for the magma that fed the 1989-1990 eruption of Redoubt (Gerlach et aI., 1994). Another strong experimental evidence of volatile saturation in magma chambers prior to explosive eruptions derives from the occurrence of primary gas bubbles as fluid inclusions in crystals (Roedder, 1979). Anderson et ai. (1989) found such inclusions in quartz crystals in the products of the Bishop Tuff eruption of Long Valley. Belkin and De Vivo (1993) found primary CO 2 -H 2 0 inclusions in the clinopyroxene of nodules erupted during plinian eruptions of Vesuvius. Other peculiar inclusions that can give informations on the pressure history of a magmatic system are hourglass inclusions (Anderson, 1991). These are glass inclusions in a quartz crystal characterized by a narrow neck connecting the body of the inclusion with the rim of the crystal (Anderson, 1991). They are probably formed "in gas-saturated decompressing magma as growing phenochrysts surround re-entrants of melt that contain a tiny bubble of gas" (Anderson, 1991). Different pressure stages are necessary to account for hourglasses: Stage (1) an hourglass forms in crystallizing magma at some pressure Po in a gas-saturated environment; Stage (2) the magma decompresses instantly to a new pressure PI' and remains there for some time t;
75
during this time some melt is lost through the neck because of decompression; Stage (3) the magma is instantly erupted and quenched. The study of the Bishop Tuff deposit suggests rapid decompression for the hourglasses found in the plinian pumice with a short ascent time; the hourglasses found in the late erupted ash, on the contrary, indicate that they formed at approximately 240 MPa; they were decompressed at approximately 110 MPa for at least a week while hourglasses evolved and bubbles of gas attained a 50 fLm diameter; then there was a rapid ascent to a pressure of 70 MPa with a thermal quenching as the magmatic foam disrupted (Anderson, 1991). Witham and Sparks (1986) hypothesize vesiculation in the magma chamber to explain the bi-modal size distribution of vesicles in the pumices of many explosive eruptions (including the May 18, 1980 eruption of Mt. St. Helens). They attributed the mode relative to larger bubbles to growth in the chamber; the second mode was due to growth in the conduit. The overpressures caused by volatile oversaturation or buoyancy of the magma may fracture the rocks overlying the magma chamber and eventually connect it to the surface. The excess gas can immediately expand to the surface if it forms a separate phase at the top of the chamber; steam explosions may results as a consequence of the interaction between gas and the water table. The connection with the surface and the initial emptying of the reservoir cause a decompression of the magma, its expansion and outpouring out of the chamber. The rigid behaviour of the wall rocks and the emptying of the reservoir permit the decompression of the magma. The decompression causes exsolution of volatiles, and the magma responds by growing bubbles, increasing its volume and attempting to re-establish the original pressure. The evolution of the eruption is different if the stored internal energy is made available immediately as an explosion or after a measurable delay time since the beginning of the eruption. Explosive nucleation of bubbles occurs when a large supersaturation pressure (10-70 Mpa) is sufficient for large nucleation rates (Hurvitz and Navon, 1994). However, slow growth of bubbles is favoured by a high value
R. Scandone / Journal of Volcanology and Geothermal Research 72 (1996) 71-83
76
of the ambient pressure (Proussevitch et al., 1993). The vesiculation rate is further controlled by the nature of the magma (Sparks, 1978), and the amount of oversaturation (Proussevitch et al., 1993). Proussevitch et al. (1993) made a parametric study describing diffusion-induced growth of closely spaced bubbles in magmatic systems. Numerical modelling, corresponding to an instantaneous depressurization of a saturated magma, shows that the bubbles start to grow with a measurable rate after a variable delay time (Proussevitch et al., 1993, p. 22,297). Some indicative quantitative conclusions can be drawn from this model (Proussevitch et al., 1993), although the lack of independent data with which to compare the results is an intrinsic difficulty of this as well as other numerical models (Sparks, 1994). Other things being equal, liquids with basaltic composition have a faster bubble growth rate than liquids of rhyolitic composition. The ambient pressure is the factor which predominantly controls the growth rate. The time for complete growth of bubbles in a silicic melt varies from a few seconds at atmospheric pressure to days or even a few weeks at pressures of 200-400 MPa. I report in Fig. 3 the results of Proussevitch et al. (1993) relative to the influence of the ambient pressure. It is possible to see that the bubbles have a growth time of the order of hours to days at a pressure corresponding to a depth of about 7 km.
3.4. Harmonic tremor The increase of the vapour phase within the magmatic reservoir is suggested also by other geophysi1000
1 MPa 10 MPa '"
::,
100
~
10
50 1O?z00
300 350 9
:::. ~ a:
0.1
10
10 3 Time (5)
lOS
Fig. 3. Diffusive growth of bubbles in rhyolitic melt for different values of ambient pressure from 0.1 to 400 MPa (saturation pressure for 8 wt.% of water). Redrawn after Proussevitch et al. (J 993).
cal peculiarities. During the Mt. St. Helens eruption, strong harmonic tremor was measured throughout the eruption (Malone et al., 1981). After 11 :40 (period 2), the tremor rapidly increased and saturated the entire seismic network of Washington State. Harmonic tremor is the most elusive seismic signal recorded in a volcanic environment. It is a continuous ground shaking with frequency content similar to that of low-frequency earthquakes (1-5 Hz), and several models have been proposed to explain its origin. Aki et al. (1977) proposed that tremor may originate by the repeated opening of cracks filled with magma. This model was developed to explain episodes of deep tremor at the basaltic volcano Kilauea on Hawaii. Crosson and Bame (1985) have proposed a simplified model based on a magma chamber filled with a gas phase. This model is able to produce a resonant radial oscillation that is dependent on the existence of a froth region inside the magma chamber and which explains many of the characteristics of harmonic tremor. The frequency is relatively insensitive to the size of the magma chamber. The synchronous oscillation of the magma chamber and surrounding rocks can exist at low frequency only when the froth region is sufficiently large. As magma is withdrawn, the amount of magma versus gas in the chamber decreases, with a consequent change in the seismic wave generation because of the large acoustic impedance between the magma-gas mixture and the country rock. Although I am not aware of systematic changes in the character of the tremor during the eruption of Mt. St. Helens, a systematic change of the frequency of tremor was observed, for example, during the 1944 eruption of Vesuvius at the transition between the low-rate, effusive-phase and the higher-rate explosive one (Imbo, 1952). According to the present model, the appearance of strong harmonic tremor during the May 18 eruption is indicative of extensive bubble nucleation within the magma chamber; drastic reduction of tremor occurs after 16:15 when the lithostatic pressure is restored by the collapse of the walls.
3.5. Variation of magma discharge during explosive eruptions The acceleration of the rate of bubble growth (Fig. 3) may result in an increase of magma dis-
R. Scandone / Journal of Volcanology and Geothennal Research 72 (J996) 71-83
charge. The importance of this process can be understood by the temporal evolution of the magma discharge during explosive eruptions. A slow increase of magma discharge and subsequent rapid waning is observed in the course of many explosive events (Fig. Ib). This pattern has been observed during the eruptions of Mt. St. Helens in 1980, El Chichon in 1982 and Pinatubo in 1991, having a Volcanic Explosivity Index (Newhall and Self, 1982) in the range of 4-6, but also several other smaller eruptions such as Pavlov and Redoubt in 1989-1990 display a similar behaviour (McNutt, 1987; Power et aI., 1994). The deposits of major eruptions have a typical succession of facies (Sparks et aI., 1973; Walker, 1985). The beginning of the eruption is commonly characterized by vent opening and a plinian phase with a reversely graded pumice-fall deposit. This phase is followed by a wavering plinian stage with partial column collapse and intraplinian ignimbrites or surge deposits. The climactic stage is reached with the eruption of an ignimbrite characterized, in proximal areas, by a lag-breccia deposit. The deposition of this typical sequence has been explained as due to conditions of magma discharge that allow a sustained eruption plume which later collapses giving rise to pyroclastic flows or surges (Sparks and Wilson, 1976; Sparks et aI., 1978). Reverse grading of the pumice deposit, as well as the generation of ignimbrites, suggest an increase in the discharge rate of magma until the emission of the pyroclastic flow, followed by a rapid waning in discharge rate. Fluiddynamical models, based on the analysis of quasisteady phases of eruptions, have explained that an increase in magma discharge can actually lead to the collapse of the eruption column (Wilson et aI., 1980; Wood, 1988). As mentioned before, the variation of conduit geometry by erosion can justify an increase of magma discharge by flaring of the vent as proposed by Wilson et aI. (1980). However, there is some questioning (Varekamp, 1993) about the field evidence for conduit erosion. Conduit erosion and widening of the vent is a requisite for eruptions with supersonic exit velocities for the gas-pyroclast mixture. Erosion probably occurs in the upper conduit, whereas the deeper parts of it, below the fragmentation level and the exsolution level, are likely to remain of the same
77
size. No relevant increase of discharge is allowed if a large part of the conduit remains of the same section, unless there is a substantial change of the pressure gradient driving the flux because of other causes. The pattern of volcanic activity during the May 18, 1980 eruption of Mt. St. Helens is analyzed by Criswell (1987) and Carey et aI. (1990). With reference to the already cited periods there are: Period 1 and most of period 2. The initial plinian phase occurred between 09:00 and 12:15; the progressive increase of magma discharge led to the generation of pyroclastic flows at 12: 15. At 11 :40 the seismographs around Mt. St. Helens began to show a gradual increase in strong ground shaking and by 13:30 it was impossible to discriminate individual earthquakes (Foxworthy and Hill, 1982; Scandone and Malone, 1985). Period 3. The maximum rate of magma output occurred between 15:00 and 16:30 with the widespread deposition of pyroclastic flows and the maximum release of seismic energy. Period 4. At 16:35 there was a decrease in magma discharge rate and a return to a plinian phase. The general pattern of the eruption described by these different observations is shown in Fig. 4. It is similar to fig. 3 of Criswell (1987), but the magma discharge is modified to take into account the observations of Carey et aI. (1990). Based on these and the previous observations I hypothesize that the eruption of silicic magma from shallow reservoirs is strongly controlled by the expansion of the magma + volatile phase because the rigid behaviour of the wall rocks does not permit the transmission of the elastic strain of the surrounding rocks. 3.6. Collapse of the chamber
The growth of bubbles tends to re-establish the original magmatic pressure (litho static or higher); it causes a steady increase in the discharge of magma until, eventually, conditions are attained to produce pyroclastic flows by non-convecting eruption columns. The process can last as long as the chamber behaves as a rigid envelope and as long as vesiculation of the water-rich melt occurs. The magma chamber becomes progressively filled with a higher proportion of gas and the density of the magma may
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R. Scandone / Journal of Volcanology and Geothermal Research 72 (J996) 71-83
Strong ground shaking
.~
f
1000
800
1800
1200
hour
Cumulative seismic moment
l~ 4*10
2*10
23 I'
I
800
1000 Magma Discharge
1400
1200 (kg/sec)
1800
1600
Pyroclastic Flows Generation
Late Plinian Phase
1" ....---«::00 - - - - - . ......
I
Early Plinian Phase
•
------~ ~~~-"" 800
1000
1200
1400
1600-
1800
Fig. 4. Variation of the main physical parameters of the eruption of May 18, 1980 of Mt. St. Helens. Tremor amplitude (top) (Scandone and Malone, 1985). Seismic strain release (middle) (Shemeta and Weaver, 1986). Magma discharge rate (bottom) (Carey et aI., 1990).
become sufficiently low that it becomes mechanically unstable (Tait et aI., 1989). The most voluminous and intense eruptions cause the formation of calderas. Scandone (1990) suggested that the mechanism of caldera formation related with the collapse of a magma chamber is due both to the size and aspect ratio of the reservoir as well as the nature of the rocks making up the roof of the chamber. The same collapse of the chamber is indicative of a brittle behaviour of materials, in fact an elastic readjustment of the rocks surrounding the magma chamber would cause a down-sagging of the surface (Walker, 1984), as observed for example at Campi Flegrei after the end of periods of unusual uplift (Lirer et aI., 1987). A rigid behaviour of the magma chamber walls would instead cause a collapse of the chamber when the pressure difference between the magma and the lithostatic load becomes sufficiently larger than the strength of the rocks (Druitt and Sparks, 1984). We can estimate, through mass conservation, the volume ~ V of drained magma producing a pressure decrease equal to the strength of the material +
eventual overpressure in a magma chamber containing a volume VI of melt: PIVI = P2VI
+ PI LlV
Ll V = VI PI - P2 PI
where PI and P2 are the density of the original magma and of the magma plus bubbles, respectively. The density of the magma plus bubbles can be estimated (Wilson and Head (1981) according to: 1
n
1- n
P
Pg
PI
-=-+-where n is the weight fraction of ex solved gas, and Pg and PI are the density of gas and liquid, respectively. The gas density can be estimated through the equation of state of the gas, assuming a good thermal contact between gas and magma:
RT
p=pg
m
R. Scandone / Journal of Volcanology and Geothermal Research 72 (J996) 7I -83
R = 8.3143 JK- 1 mol- 1 is the universal gas constant and m is the molecular weight for water (= 0.Q18 kg). The content of water for the mafic dacite of Mt. St. Helens has been estimated by Rutherford et al. (1985) at 4.6(± 1.0) wt.%. This estimate is between the value reported by Melson and Hopson (1982) of about 7%, and that of Eichelberger and Hays (1982) of 0.7-1.7%. Rutherford et ai. (1985) suggest that the magma was undersaturated in water; however, it is possible that a mixed volatile gas (C0 2 + H 20 + S02) was present (Johnson et aI., 1994). I will hypothesize, for modeling purposes, that the magma of Mt. St. Helens was gas saturated, and use the ex solution law for water provided by Burnham (1975) for rhyolite. The weight fraction of water dissolved in the melt is given by: n
=
4.11 X 10- 6 po.s (p in pascals)
A rhyolitic melt at a pressure of 185 MPa (depth of - 7 km) has a saturation content of 5.6 wt.% of water (the upper limit of Rutherford et aI., 1985). I use a total pressure decrease of 40 MPa (hypothesizing an overpressured magma chamber, for example 20 MPa, and a decrease of another 20 MPa below lithostatic pressure, corresponding to the strength of rocks). This decompression produces an ex solution of about 0.6 wt% of water, causing a density decrease from 2400 to 2282 kg/m3. The amount of magma necessary to be drained to produce such a pressure decrease is approximately 4.9% of the original volume. This value gives an estimate of 5.1-10.1 km 3 for the volume of the Mt. St. Helens magma chamber, using the volume of products erupted during the May 18 eruption (0.25-0.5 km 3 ) (Lipman et aI., 1981; Criswell, 1987). The value agrees well with the estimate of 10-20 km 3 made by Scandone and Malone (1985) for the earthquake-free zone considered to be the magma reservoir. The same method, applied to the eruption of Mt. Pinatubo in 1991, gives a result of a magma chamber of approximately 96 km 3 for an eruption of 5 km 3 of dacite (6.4% weight of water at 220 MPa and decompression of 40 MPa). The results is comparable with the 40-90 km 3 estimated from seismic evidence (Mori et aI., 1993). These rough estimates support the idea that highpressure vesiculation in the chamber and the amount
79
of saturated melt are important factors affecting the development of an eruption. 3.7. Deposit associated with the collapse
The collapse of the magma chamber tends to re-establish the lithostatic load. Collapse may not occur in a single phase as arching of the roof eventually determines a confining pressure less than the original one, but, as a rule, collapse is followed by a rapid waning in the rate of discharge. The violent earthquakes occurring around and immediately above the magma chamber possibly cause extensive fracturing and produce lithic breccia (Walker, 1985; Scandone, 1990; Varekamp, 1993). During the May 18 eruption of Mt. St. Helens, the emission of lithic breccia (lithic breccia zone of Criswell, 1987) between 15:00 and 16:30 occurs contemporaneously with the highest seismic energy release and peak magma discharge (Carey et aI., 1990). After this, there is a decrease in seismic energy release and magma discharge.
4. Discussion and conclusions Several other recent explosive eruptions closely monitored by seismic networks give support to the ideas proposed in this paper. The eruptions of Pinatubo in 1991, of Redoubt in 1989-1990 and of Mt. Spurr in 1992 provide further evidences of a close cause-effect relationship between the explosive emission of magma and the following swarms of deep earthquakes (Pinatubo Volcano Observatory Team, 1991; Wolfe, 1992; Power et aI., 1994; Alaska Volcano Observatory, 1993). The comparison of post-eruption seismicity observed at Mt. Pinatubo and Mt. St. Helens is shown in Fig. 5. The spacetime evolution of seismicity of the 1992 eruption of Mt. Spurr shows a close similarity with that of Mt. St. Helens shown in Fig. 2 (Alaska Volcano Observatory, 1993). Pre-eruption seismicity probably reflects either ascent of viscous magma because of its buoyancy (Mt. St. Helens and Pinatubo) and/or opening of fractures favoured by expansion of a vapour phase (as suggested for example by the occurrence of
80
R. Scan done / Journal of Volcanology and Geothermal Research 72 (1996) 71-83
low-frequency seismic events at Redoubt - Power et aI., 1994). Deep seismic activity around the magma chamber, which starts after the beginning of the eruption, reflects instead collapse of the chamber because of decompression of the magma and its readjustment to equilibrium conditions. This phenomenon can occur only if the walls of the chamber have an initial rigid behaviour. The basic ideas expressed in this paper suggest that the mechanical properties of the rocks surrounding the magma chamber control the mechanism of
the eruption. In some eruptions, especially basaltic ones, the driving pressure is furnished by the elastic strain of the medium surrounding the magma chamber. A purely elastic response produces a high pressure gradient, generally at the beginning of the eruption, and will result in a high initial discharge rate (in the case of constant conduit shape and water content of the magma). The gradual decrease of the overpressure does not permit extensive vesiculation within the magma chamber. The elastic mechanism is not permitted in the case
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R. Scandone / Journal of Volcanology and Geothermal Research 72 (J996) 71-83
of a rigid behaviour of the chamber walls (commonly during silicic eruptions). In this case, the pressure driving the eruption is furnished at the expense of the magma internal energy, mostly by an expansion of the volatile phase. The rate of bubble growth and expansion of the magma govern the rate at which magma is expelled by the chamber. Maximum eruption rate possibly occurs when the bubble growth rate attains high values. The high discharge favours the tapping of deeper levels in zoned magma chambers (Spera, 1984; Blake and Ivey, 1986) with possible eruption of magma with a lower gas content. These two factors greatly favour column collapse (Wilson et aI., 1980) with consequent pyroclastic flow activity, as observed at Mt. St. Helens (Carey et aI., 1990). The length of the eruption is controlled by the amount of volatiles in solution in the magma and by the mechanical stability of the reservoir. These two factors combine in controlling the amount of material that must be erupted before a decompression sufficient to cause a collapse of the wall rocks occurs. The work of Proussevitch et aI. (1993) suggests that the ambient pressure (depth) strongly controls the growth rate of bubbles, with higher pressures slowing the process. Within this frame, I suggest that deep chambers possibly favour a slow development of the eruption with a relatively long time span between the start of eruptions and their paroxysmal phases when there are maximum discharge rates and possible pyroclastic flow emplacement events. Shallower reservoirs would instead cause fast-developing eruptions because of higher bubble growth rates at lower confining pressures. Although this model makes a strong separation between the two modes of behaviour of surrounding rocks, it is probable that a mixed behaviour is the norm. However, the predominance of an elastic readjustment of the rocks seems more likely during basaltic eruptions; the rigid behaviour is instead favoured during the eruption of silicic magma.
Acknowledgements
My sincere appreciation goes to Christopher Kilburn, Lisetta Giacomelli and Grazia Giberti for critical reading and suggestions to an earlier version of
81
this manuscript. Steve McNutt, Elisabeth Parfitt, Steve Sparks and Lionel Wilson made thoughtful reviews of this paper; their comments and suggestions resulted in a substantial improvement of the manuscript. I am particularly grateful to Steve Sparks for drawing my attention to recent models of diffusive bubble growth. This work was made when the author was at Dipartimento di Geologia e Geofisica, Universita di Napoli. The financial support of grant MURST 60%, Universita di Napoli, is aknowledged.
References Aki, K., Fehler, M. and Das, S., 1977. Source mechanism of volcanic tremor: fluid-driven crack models and their application to the 1963 Kilauea eruption. J. Volcano!. Geotherm. Res., 2: 259-287. Alaska Volcano Observatory, 1993. Mt Spurr's 1992 eruptions. EOS, Trans. Am. Geophys. Union, 74: 217, 221-222. Anderson Jr., AT, 1991. Hourglass inclusions: theory and application to the Bishop Rhyolitic Tuff. Am. Minera!., 76: 530547. Anderson Jr., A.T., Newmann, S., Williams, S.N., Druitt, T.H., Skirius, C. and Stolper, E., 1989. H 2 0, CO, CI, and gas in plinian and ash flow Bishop rhyolite. Geology, 17: 221-225. Anderson, O.L. and Grew, P.C., 1977. Stress corrosion theory of crack propagation with applications to geophysics. Rev. Geophys. Space Phys., 15: 77-104. Banks, N.G., Tilling, R.I., Harlow, D.H. and Ewert, J.W., 1989. Volcano monitoring and short term forecasts. In: R. Tilling (Editor), Volcanic Hazards. (Short Course in Volcanology') Am. Geophys. Union, Washington, DC, pp. 51-80. Barker, S.E. and Malone, S.D., 1991. Magmatic system geometry at Mount St. Helens modeled from the stress field associated with posteruptive earthquakes. J. Geophys. Res., 96: 11,88311,894. Belkin, H.E. and De Vivo, B., 1993. Fluid inclusions study of ejected nodules from plinian eruptions of Mt SommaVesuvius. J. Volcanol. Geotherm. Res., 58: 89-100. Blake, S., 1981. Volcanism and the dynamics of open magma chamber. Nature, 289: 783-785. Blake, S., 1984. Volatile oversaturation during the evolution of silicic magma chambers as an eruption trigger. J. Geophys. Res., 89: 8237-8244. Blake, S. and Ivey, G.N., 1986. Density and viscosity gradients in zoned magma chambers, and their influence on withdrawal dynamics. J. Volcano!' Geotherm. Res., 30: 201-230. Burnham, C.W., 1972. The energy of explosive eruption. Earth Mineral. Sci. Pa. State Univ., 41: 69-70. Burnham, C.W., 1975. Water and magmas; a mixing mode!. Geochim. Cosmochim. Acta, 39: 1077-1084. Burnham, C.W., 1985. Energy release in subvolcanic Environments: Implications for breccia formation. Econ. Geol., 80: 1515-1522.
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Carey, S., Sigurdsson, H., Gardner, J.E. and Criswell, C.W., 1990. Variation in column height and magma discharge during the May 18, 1980 eruption of Mt St. Helens. J. Volcano!. Geotherm. Res., 43: 99-112. Criswell, C.W., 1987. Chronology and pyroclastic stratigraphy of the May 18, 1980 eruption of Mount St. Helens, Washington. J. Geophys. Res., 92: 10,237-10,266. Crosson, RS. and Bame, D.A., 1985. A spherical source model for low frequency volcanic earthquakes. J. Geophys. Res., 90: 10,237-10,247. Druitt, T.H. and Sparks, R.S.J., 1984. On the formation of calderas during ignimbrite eruptions. Nature, 310: 679-681. Eichelberger, J.C. and Hays, P.B., 1982. Magmatic model for the Mount St Helens blast of May 18, 1980. J. Geophys. Res., 87: 7727-7736. Foxworthy, B.L. and Hill, M., 1982. Volcanic eruptions of 1980 at Mount St Helens: the first 100 days. U.S. Geo!. Surv. Prof. Pap. 1249, 125 pp. Gerlach, T.M., Westrich, H.L., Casedewall, TJ. and Finnegan, D.L., 1994. Vapor saturation and accumulation in magmas of the 1989-90 eruption of Redoubt volcano, Alaska. J. Volcano!. Geotherm. Res., 62: 317-337. Gerlach, T.M., Westrich, H.L. and Symonds, RB., in press. Pre-eruption vapor in magma of the climactic Mount Pinatubo eruption: source of the giant stratospheric sulfur dioxide cloud. In: C.G. Newhall and RS. Punongbayan (Editors), A Volcano Awakens: Eruptions and Lahars of Mount Pinatubo Philippines. U.S. Geo!. Surv. Prof. Pap. 1600. Giberti, G. and Wilson, L., 1990. The influence of geometry on the ascent of magma in open fissure. Bul!. Volcano!., 52: 515-521. Hildreth, W., 1981. Gradients in silici magma chambers: implications for lithospheric magmatism. J. Geophys. Res., 86: 10,153-10,192. Hurvitz, S. and Navon, 0., 1994. Bubble nucleation in rhyolitic melts: experiments at high pressure, temperature and water content. Earth Planet. Sci. Lett., 122: 267-280. Imbo, G., 1952. Sismicita' del parossismo vesuviano del marzo 1944. Ann. Oss. Vesuviano, VI Ser., I: 59-268. Johnson, M.e., Anderson Jr., A.T. and Rutherford, M.J., 1994. Pre-eruptive volatile contents of magmas. In: M.R Carroll and lR Holloway (Editors), Volatiles in Magmas. Rev. Minera!., 30: 283-330. Lees, J.M., 1992. The magma system of Mt St. Helens: non-linear high resolution P-wave tomography. J. Volcano!. Geotherm. Res., 54: 103-116. Lipman, P.W., Nortin, D.R., Taggart, J.R., Brandt, E.L. and Engleman, E.E., 1981. Compositional variation in 1980 magmatic deposits. USGS Prof. Pap., 1250: 147-156. Lirer, L., Luongo, G. and Scandone, R., 1987. On the volcanological evolution of Campi Flegrei. EOS, Trans. Am. Geophys. Union, 68: 226-234. Machado, F., 1974. The search for magmatic Reservoir. In: L. Civetta, P. Gasparini, G. Luongo and A Rapolla (Editors), Physical Volcanology. Elsevier, Amsterdam, pp. 255-273.
Malone, S.D., Endo, E.T., Weaver, CJ., Ramsey, J.W., 1981. Seismic monitoring for eruption prediction. U.S. Geo!. Surv., Prof. Pap., 1250: 803-814. McBimey, AR, 1980. Mixing and unmixing of magmas. J. Volcano!. Geotherm. Res., 7: 357-371. McNutt, S.R, 1987. Eruption characteristics and cycles at Pavlov Volcano, Alaska and their relation to regional earthquake activity. J. Volcano!. Geotherm. Res., 31: 239-268. Melson, W.G. and Hopson, C.A, 1982. Petrologic model of the 1981-82 Mount St Helens magma chamber. EOS, Trans. Am. Geophys. Union, 63: 1144. Mogi, K., 1958. Relations between the eruptions of various volcanoes and the deformations of the ground surfaces around them. Bull. Earth Res. Inst., 36: 99-134. Mori, J., Eberhart-Phillips, D. and Harlow, D., 1993. Three-dimensional velocity structure at Mount Pinatubo Philippines: resolving magma bodies and earthquake hypocenters. EOS, Trans. Am. Geophys. Union, 74: 667. Newhall, C.F. and Self, S., 1982. The volcanic explosivity index (VEr): an estimate of explosive Magnitude for historical eruptions. J. Geophys. Res., 87: 1231-1238. Papale, P. and Dobran, F., 1994. Magma flow along volcanic conduit during the plinian and pyroclastic flow phases of the May 18, 1980, Mount St Helens Eruption. J. Geophys. Res., 99: 4355-4373. Parfitt, E.A and Wilson, L., 1994. The 1983-86 Pu'u '0'0 eruption of Kilauea volcano: a study of dyke geometry and eruption mechanisms for a long-lived eruption. J. Volcano!. Geotherm. Res., 59: 179-206. Parfitt, E.A., Wilson, L. and Head III, J.W., 1993. Basaltic magma reservoirs: factors controlling their rupture characteristics and evolution. J. Volcano!. Geotherm. Res., 55: 1-14. Pinatubo Volcano Observatory Team, 1991. Lessons from a major eruption: Mt Pinatubo, Philippines. EOS, Trans. Am. Geophys. Union, 72: 545-555. Power, J.A., Lahr, J.e., Page, B.A, Chouet, B.A., Stephens, C.D., Harlow, D.H., Murray, T.L. and Davies, IN., 1994. Seismic evolution of the 1989-1990 eruption sequence of Redoubt Volcano, Alaska. 1. Volcano!. Geotherm. Res., 62: 95-136. Proussevitch, A.A., Sahagian, D.L. and Anderson Jr., A.T., 1993, Dynamics of diffusive bubble growth in magmas: isothermal case. J. Geophys. Res., 98: 22,283-22,307. Roedder, E., 1979. Origin and significance of magmatic inclusions. Bull. Minera!., 102: 487-510. Rubin, A.M., 1993. Tensile fracture of rock at high confining pressure: implication for dyke propagation. J. Geophys. Res., 98: 15,919-15,935. Rutherford, MJ. and Devine, J., 1991. Pre-eruption conditions and volatiles in the 1991 Pinatubo magma. EOS, Trans. Am. Geophys. Union, 72: 62. Rutherford, MJ., Sigurdsson, H., Carey, S. and Davis, A., 1985. The May 18, 1980, eruption of Mount St. Helens 1. Melt composition and experimental phase equilibria. J. Geophys. Res., 90: 2929-2947. Ryan, M.P., 1987. Neutral buoyancy and the mechanical evolution
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of magmatic system. In: B.O. Mysen (Editor), Magmatic Processes: Physico-chemical Principles. Geochem. Soc. Spec. Pub!., 1: 259-287. Scandone, R, 1979. Effusion rate and energy balance of Paricutin eruption (1943-1952), Michoacan, Mexico. J. Volcano!. Geotherm. Res., 6: 49-59. Scandone, R, 1990. Chaotic collapse of calderas. J. Volcano!. Geotherm. Res., 42: 285-302. Scandone, R and Malone, S.D., 1985, Magma supply, magma discharge and readjustment of the feeding system of Mount St. Helens during 1980. J. Volcano!. Geotherm. Res., 23: 239-262. Shemeta, J.E. and Weaver, e.S., 1986, Seismicity accompanying the May 18, 1980 eruption of Mount St. Helens, Washington. In: S.AC. Keller (Editor), Mt St. Helens: Five Years Later. Eastern Washington Univ. Press, Cheney, WA, pp. 44-58. Smith, RL., 1979. Ash-flow magmatism. Geo!. Soc. Am. Spec. Pap., 180: 5-27. Sparks, R.SJ., 1978. The dynamics of bubble formation and growth in magmas: A review and analysis. J. Volcano!. Geotherm. Res., 3: 1-37. Sparks, RS.J., 1994. Comments on "Dynamics of diffusive bubble growth in magmas: isothermal case" by A.A Proussevitch, D.L. Sahagian, AT. Anderson. J. Geophys. Res., 99: 17,827-17,828. Sparks, R.SJ. and Wilson, L., 1976. A model for the formation of Ignimbrites by gravitational column collapse. J. Geo!. Soc. London, 132: 441-452. Sparks, R.SJ., Self, S. and Walker, G.P.L., 1973. Products of ignimbrite eruptions. Geology, 1: 115-118. Sparks, RSJ., Wilson, L. and Hulme, G., 1978. Theoretical modelling of the generation, movement and emplacement of pyroclastic flows by column collapse. J. Geophys. Res., 83: 1727-1739. Spera, FJ., 1984. Some numerical experiments on the withdrawal of magma from crustal reservoirs. J. Geophys. Res., 89: 8222-8236.
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Stasiuk, M.V., Jaupart, e. and Sparks, R.SJ., 1993. On the variation of the flow rate in non-explosive lava eruptions. Earth Planet. Sci. Lett., 114: 505-516. Tait, S., Jaupart, e. and Vergniolle, S., 1989. Pressure, gas content and eruption periodicity of a shallow crystallising magma chamber. Earth Planet. Sci. Lett., 92: 107-123. Varekamp, J.e., 1993. Some remarks on volcanic vent evolution during plinian eruptions. J. Volcano!. Geotherm. Res., 54: 309-318. Wadge, G., 1981. The variation of magma discharge during basaltic eruptions. J. Volcano!' Geotherm. Res., 11: 139-168. Walker, G.P.L., 1984. Downsag calderas, ring faults, caldera sizes, and incremental caldera growth. J. Geophys. Res., 89: 8407-8416. Walker, G.P.L., 1985. Origin of coarse lithic breccias near ignimbrite source vents. J. Volcano!. Geotherm. Res., 25: 157171. Weertman, J., 1971. Theory of water filled crevasses in glaciers applied to magma transport beneath oceanic ridges. J. Geophys. Res., 76: 1171-1183. Wilson, L. and Head, J.W., 1981. Ascent and eruption of basaltic magma on the Earth and Moon. J. Geophys. Res., 86: 29713001. Wilson, L., Sparks, R.SJ. and Walker, G.P.L., 1980. Explosive volcanic eruptions - IV. The control of magma properties and conduit geometry on eruption column behaviour. Geophys. J. R Astron. Soc., 63: 117-148. Witham, AG. and Sparks, RSJ., 1986. Pumice. Bull. Volcano!., 48: 209-223. Wolfe, E., 1992. The 1991 eruptions of Mount Pinatubo, Philippines. Earthquakes Volcanoes, 23: 5-37. Wolfe, E.W., Neal, e.A, Banks, N.G. and Duggan, TJ., 1988. Geologic observations and chronology of eruptive events. U.S. Geo!. Surv. Prof. Pap., 1463: 1-98. Wood, A.W., 1988. The fluid-dynamics and the thermodynamics of eruption columns. Bull. Volcano!., 50: 169-193.
ELSEVIER
Journal of Volcanology and Geothermal Research 72 (1996) 71-83
Factors controlling the temporal evolution of explosive eruptions Roberto Scandone
I
Dipartimento di Geofisica e Vulcanologia, Universita di Napoli" Federico II", Largo San Marcellino 10, 80138 Naples, Italy
Received 23 September 1994; accepted 3 December 1995
Abstract A simple model, based on observations of recent explosive eruptions, is proposed for the patterns of variable discharge rate during explosive eruptions. Such eruptions are characterized by a slowly increasing waxing phase and a phase of more rapid waning activity. This behaviour is associated with volatile saturation of the magma and a rigid response of the magma chamber walls until brittle fractures occur. The opening of the reservoir and the escape of the vapour phase cause an initial decompression of the magma. The emptying of the reservoir and the rigid behaviour of the wall rocks permit a further decompression of the magma during the course of the eruption. The waxing phase is related to a progressive increase in the rate of vesiculation of saturated melt that re-equilibrates the decompression of the magma. The following increase of discharge rate permits an early Plinian phase which culminates in collapse of the eruption column and emplacement of pyroclastic flows. The emission of abundant lithic fragments, making up the roof of the chamber, signals the beginning of the waning phase. The eruption lasts until the vesiculation, caused by the initial decompression and emptying of the chamber, can counterbalance the lithostatic load. Collapse of the chamber occurs as soon as the pressure becomes lower than the lithostatic load by an amount similar to the strength of rocks. The collapse tends to re-establish the original pressure conditions preventing further vesiculation and thereby resealing the magma chamber.
1. Introduction
A schematic model of an erupting volcano is that of a subterranean reservoir connected with the surface through a narrow conduit. The magma flow is driven by the pressure gradient between the reservoir and the surface whereas the physical properties of the magma and the geometry of the conduit control the peculiarities of the flow. Several authors have analyzed the effect of the
I Present address: III Universita di Roma, Dipartimento di Fisica "E. Amaldi", Via Vasca Navale 84, 00146 Rome, Italy; e-mail: [email protected]
0377-0273/96/$15.00 Published by Elsevier Science B.V. SSDI 0377-0273(95)00086-0
geometry of the conduit on the eruptive character of the eruption (see, for example, Wilson et aI., 1980; Papale and Dobran, 1994). These models hypothesize a constant pressure in the reservoir so that the mass flux is dictated by the difference between the pressure in the reservoir and the pressure at the surface. If the geometry of the pathway does not flare out somewhere by a required amount, then the flow may be choked, the exit pressure is greater than atmospheric and the discharge rate is less than it would be in the non-choked case (Wilson et aI., 1980; Wilson and Head, 1981; Giberti and Wilson, 1990). A flaring of the conduit, caused, for example, by vent erosion at the surface, may allow a transition to supersonic flow with a reduction of the exit
72
R. Scandone / Journal of Volcanology and Geothermal Research 72 (1996) 71-83
pressure and therefore an increase of the overall pressure gradient and the discharge rate (Wilson et al., 1980; Wilson and Head, 1981). In this paper I take the different approach of analyzing the variations of the pressure gradient caused by factors independent of the geometry of the conduit, but related to the "structural" model of a volcano under observation. An overpressure above the lithostathic value in a magma reservoir is often invoked as necessary to start an eruption (Blake, 1981; Parfitt et al., 1993). The overpressure may originate by different mechanisms: (1) ex solution of a volatile phase or intrusion of new magma into a reservoir (Blake, 1981, 1984; Tait et al., 1989) with development of an overpressure above the strength of the rocks surrounding the magma chamber which causes the opening of fractures (Rubin, 1993); (2a) instability of fluid-filled cracks driven by the Peach-Koehler pseudo-buoyancy force (Weertman, 1971) due to ex solution of a gas phase or crack propagation by stress corrosion (Anderson and Grew, 1977); (2b) positive density contrast of liquids lighter than surrounding rocks (Ryan, 1987) deriving, for example, from the differentiation in a magma reservoir with the accumulation of a light differentiate at the top of chamber and the development of stress above the strength of the rocks. I will try to show that the overpressure, available to drive the magma out of the chamber, is controlled by the mechanical behaviour of the chamber walls. I will discuss two extreme cases: (I) The "elastic reservoir", already proposed by Machado (1974) and Wadge (1981), is characterized by an elastic readjustment of the reservoir walls as magma is removed. (II) The "rigid reservoir" has rigid walls at least for some time after the beginning of the eruption. I will use the observations made during the eruptions of May 18, 1980 of Mt. St. Helens and of June 15, 1991 of Mt. Pinatubo to illustrate the evidence for and consequences of this assumption. Obviously, in nature there is no such clear-cut difference in the behaviour of actual volcanoes, but these examples may help to illustrate the different physical processes that govern the development of
the eruption if one mechanical behaviour prevails over the other.
2. The elastic model Here I will recall only the main characteristics of the model as Machado (1974) and Wadge (1981) give a more detailed discussion. A perfectly elastic medium, by definition, transmits instantaneously the elastic strain stored by the rocks surrounding the magma chamber and produces a continuous squeezing of the chamber until the overpressure is exhausted. The elastic energy is, by definition, made immediately available upon the opening of the reservoir, so the maximum pressure within the magma chamber is observed right at the beginning of the eruption. Some basaltic eruptions show an unsteady behaviour with a rapid waxing phase and a slower waning in the rate of magma discharge (Fig. la) [several cases are shown in Scandone (1979) and an extensive bibliography is provided by Wadge (1981), although contrasting examples of effusion rate variations have been exhibited during the Pu'u '0'0 eruption (Wolfe et al., 1988)]. The rapid initial waxing phase of these effusive
Magma discharge rate max 100 -10 3 m3 /sec
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type
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~
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Fig. 1. Examples of trends of variation of magma discharge during effusive and explosive eruptions.
R. Scandone / Journal of Volcanology and Geothermal Research 72 (J996) 71-83
eruptions may be interpreted as due to the opening of a magma-filled crack as soon as it reaches the surface. The high initial discharge is a result of the high overpressure immediately provided through elastic readjustment (Wadge, 1981). The overpressure is rapidly accommodated by the drainage of magma. The exponential waning of magma discharge is interpreted as due either to the elastic decompression of the magma (Machado, 1974), or of the rocks surrounding it (Wadge, 1981). The deformations of the flanks of volcanoes observed prior to several eruptions are well explained in terms of an elastic deformation of the rocks surrounding a pressure source (Mogi, 1958) so supporting the model proposed by Wadge. A purely elastic model is used by Stasiuk et aI. (1993) to account for second-order variations of flow rate in basaltic eruptions. They explain departures from the trend of exponential decay of flow rate as due to the relative changes in viscosity, conduit dimensions and thickness of lava over the vent. They estimate that the overpressure necessary to drive the eruptions of several volcanoes is of the order of 10-20 MPa. During the Pu'u '0'0 eruption of Kilauea there was a typical behaviour that contrasted with the one shown previously. A gradual buildup in the eruption rate occurred after the start of the eruption and then there was an approximately exponential decrease. This peculiar feature has been explained by Parfitt and Wilson (1994) as due to a non-newtonian behaviour of the magma. One possible explanation of the relevance of the rheological properties of magma during this eruption is that the volcano is in an open conduit condition since the beginning of the eruption in 1983. In this case, the magma flow, driven by minor variations of stress, is strongly controlled by the magma yield strength and geometry of the conduit (Parfitt and Wilson, 1994). The summit inflation of Kilauea and the subsequent deflation related to the eruptive episodes of, the Pu'u '0'0 vent are of the order of 5-30 microradians (Parfitt and Wilson, 1994) while those observed after longer periods of inactivity are of the order of hundreds of microradians (Banks et aI., 1989). The larger stresses implied during these other episodes likely produce high initial eruption rates not influenced by the yield strength of the magma.
73
3. The rigid model 3.1. Introduction In this model I hypothesize that the magma chamber walls do not deform elastically in response to a pressure change within the reservoir, but react with a brittle failure when a sufficiently high stress is applied. This assumption is probably justified for silicic volcanoes. A long-lived magma chamber is often associated with silicic volcanoes (see, for example, Smith, 1979: Hildreth, 1981). The cooling of the liquid and crystallization at the walls (McBirney, 1980) may provide a carapace which behaves as a rigid envelope. The occurrence of a high-velocity layer at the top of the Mt. St. Helens magma chamber (Lees, 1992) is indicative of this occurrence. A similar occurrence is also reported at Vesuvius and at other volcanoes (P. Gasparini and S. Malone, pers. commun.). The important point to be noted is that a rigid behaviour of the walls of a magma chamber, after the start of the eruption, makes it impossible to squeeze the magma out of the chamber. A close analogy to this case is the opening of a champagne bottle: a sudden opening of cracks (the removal of the "cork") and escape of volatiles through the newly formed conduit cause the sudden decompression of the magma and it attempts to equilibrate to the new conditions. The overpressure driving the magma to the surface is made available, when the reservoir ruptures and is connected with the atmosphere, in the form of stored internal energy of the magma (see, for example, Burnham, 1972, 1985). This last mechanism occurs through expansion of already ex solved H 2 0, if present, and exsolution of additional H 2 0 upon decompression of the magma body. A geological analogy of this process is, for example, the flashing of a geothermal reservoir. 3.2. Chamber wall behaviour during the eruption of May 18, 1980 of Mt. St. Helens Scandone and Malone (1985) analyzed the seismicity of Mt. St. Helens occurring mostly after 17:49 on May 18, 1980 and identified an earthquake-free volume, whose top was at about 7 km depth, interpreted as the magma chamber of the volcano. These
R. Scandone / Journal o/Volcanology and Geothermal Research 72 (1996) 71-83
74
authors have shown that, during the major explosive eruptions of Mt. St. Helens of May 18 and 25, 1980 there was a readjustment of the chamber walls in response to the drainage of magma. The readjustment occurred as a brittle fracture of rocks producing earthquakes, characterized by a high-frequency signal around the reservoir. A similar phenomenon, occurred also after the explosive eruptions of June 12, July 22, August 7 and October 16, 1980. However, no occurrence of deep seismicity was observed during or after the effusive dome-building eruptions of the following years. Barker and Malone (1991) gave further support to the idea of chamber wall readjustment by analyzing the focal mechanism solutions of the post-eruption earthquakes. They found that the focal mechanisms were well explained by a stress concentration caused by a decrease in pressure within a cylindrical magma chamber at a depth of 7 -11 km, in the presence of the regional stress field. Shemeta and Weaver (1986) made a detailed analysis of the temporal readjustment of the chamber wall following the onset of the main eruption by analyzing the rate of earthquake occurrence on May 18 and 19. They subdivided this time interval into five periods as shown in Table 1. During the first two periods (Fig. 2), the earthquake release rate was low (low cumulative seismic moment) and the events were located in a volume below the volcano at
one
t. C!l
3
~6 ;5
9
~
12
C!lC!l
C!l
15 18
C!l
E
(!)
g>
(!)
(')
E
09
11
13
15
17 19 21 23 Hours (PDT)
01
03 05
07
Fig. 2. Depth-time distribution of earthquakes in the first day of the May 18, 1980 eruption of Mt. St. Helens. The vertical bar indicates the location of the magma chamber (slightly modified after Shemeta and Weaver, 1986).
depths ranging between 3 and 9 km, most earthquakes occurring between 3 and 6 km, with only a few events as deep as 12 km. During period 3, there was the maximum release of seismic energy (about 3/4 of the total), and a scatter in epicentral distribution as wide as 3 km, with most events occurring at depths between 2 and 8 km, and only a few between 8 and 12 km. During periods 4 and 5 there was a decrease in the release of seismic energy, and a deepening of events (4-14 km) with a lack in the shallow portions of the system. After 19:00 on May 18 nearly all events were located at depths between 5 and 12 km. I propose that the low rate of occurrence of
Table 1 The rate of earthquake occurrence on May 18 and 19, 1980 at Mt. St. Helens Time (PDT) May 18, 1980
Character of seismic activity
Character of volcanic activity
08.49-11.30
Minor seismic activity with most earthquakes between 3 and 6 km
Plinian column
2
11.30-14.15
Minor seismic activity with most earthquakes between 3 and 6 km; beginning of strong ground shaking at 11:40
Plinian column and beginning of pyroclastic flow emplacement
3
14.15-16.30
Maximum release of seismic energy with most earthquakes between 2 and 8 km
Pyroclastic flow emplacement; maximum output between 15:00 and 16:30
4
16.30-17.30
Decrease of seismic energy with deeper events (4-14 km)
Return to plinian column and decrease of magma discharge
5
17.30-08.30 of May 19
Decrease of seismic energy, after 19:00, all earthquakes are between 5 and 12 km
Final decrease of activity
Period
C!l
Q.
R. Scandone / Journal of Volcanology and Geothermal Research 72 (1996) 71-83
earthquakes during the first two periods is evidence of the absence of collapse. Extensive collapse of the magma chamber is signalled by the high rate of occurrence of deep earthquakes whose occurrence has been commonly observed after explosive eruptions at silicic volcanoes.
3.3. Expansion of magma within the chamber Johnson et al. (1994) suggest that volatile saturation occurs in magma chambers prior to silicic explosive eruptions. Strong evidence for an excess gas phase is found, for example, in the 1991 eruption of Mt. Pinatubo (Rutherford and Devine, 1991). Gerlach et ai. (in press) suggest that the dacite erupted on 15 June 1991 at Mt. Pinatubo was vapour saturated at depth prior to eruption to account for the immediate source of excess sulfur detected by remote sensing during the eruption. The occurrence of a vapour phase in the reservoir is invoked also for the magma that fed the 1989-1990 eruption of Redoubt (Gerlach et aI., 1994). Another strong experimental evidence of volatile saturation in magma chambers prior to explosive eruptions derives from the occurrence of primary gas bubbles as fluid inclusions in crystals (Roedder, 1979). Anderson et ai. (1989) found such inclusions in quartz crystals in the products of the Bishop Tuff eruption of Long Valley. Belkin and De Vivo (1993) found primary CO 2 -H 2 0 inclusions in the clinopyroxene of nodules erupted during plinian eruptions of Vesuvius. Other peculiar inclusions that can give informations on the pressure history of a magmatic system are hourglass inclusions (Anderson, 1991). These are glass inclusions in a quartz crystal characterized by a narrow neck connecting the body of the inclusion with the rim of the crystal (Anderson, 1991). They are probably formed "in gas-saturated decompressing magma as growing phenochrysts surround re-entrants of melt that contain a tiny bubble of gas" (Anderson, 1991). Different pressure stages are necessary to account for hourglasses: Stage (1) an hourglass forms in crystallizing magma at some pressure Po in a gas-saturated environment; Stage (2) the magma decompresses instantly to a new pressure PI' and remains there for some time t;
75
during this time some melt is lost through the neck because of decompression; Stage (3) the magma is instantly erupted and quenched. The study of the Bishop Tuff deposit suggests rapid decompression for the hourglasses found in the plinian pumice with a short ascent time; the hourglasses found in the late erupted ash, on the contrary, indicate that they formed at approximately 240 MPa; they were decompressed at approximately 110 MPa for at least a week while hourglasses evolved and bubbles of gas attained a 50 fLm diameter; then there was a rapid ascent to a pressure of 70 MPa with a thermal quenching as the magmatic foam disrupted (Anderson, 1991). Witham and Sparks (1986) hypothesize vesiculation in the magma chamber to explain the bi-modal size distribution of vesicles in the pumices of many explosive eruptions (including the May 18, 1980 eruption of Mt. St. Helens). They attributed the mode relative to larger bubbles to growth in the chamber; the second mode was due to growth in the conduit. The overpressures caused by volatile oversaturation or buoyancy of the magma may fracture the rocks overlying the magma chamber and eventually connect it to the surface. The excess gas can immediately expand to the surface if it forms a separate phase at the top of the chamber; steam explosions may results as a consequence of the interaction between gas and the water table. The connection with the surface and the initial emptying of the reservoir cause a decompression of the magma, its expansion and outpouring out of the chamber. The rigid behaviour of the wall rocks and the emptying of the reservoir permit the decompression of the magma. The decompression causes exsolution of volatiles, and the magma responds by growing bubbles, increasing its volume and attempting to re-establish the original pressure. The evolution of the eruption is different if the stored internal energy is made available immediately as an explosion or after a measurable delay time since the beginning of the eruption. Explosive nucleation of bubbles occurs when a large supersaturation pressure (10-70 Mpa) is sufficient for large nucleation rates (Hurvitz and Navon, 1994). However, slow growth of bubbles is favoured by a high value
R. Scandone / Journal of Volcanology and Geothermal Research 72 (1996) 71-83
76
of the ambient pressure (Proussevitch et al., 1993). The vesiculation rate is further controlled by the nature of the magma (Sparks, 1978), and the amount of oversaturation (Proussevitch et al., 1993). Proussevitch et al. (1993) made a parametric study describing diffusion-induced growth of closely spaced bubbles in magmatic systems. Numerical modelling, corresponding to an instantaneous depressurization of a saturated magma, shows that the bubbles start to grow with a measurable rate after a variable delay time (Proussevitch et al., 1993, p. 22,297). Some indicative quantitative conclusions can be drawn from this model (Proussevitch et al., 1993), although the lack of independent data with which to compare the results is an intrinsic difficulty of this as well as other numerical models (Sparks, 1994). Other things being equal, liquids with basaltic composition have a faster bubble growth rate than liquids of rhyolitic composition. The ambient pressure is the factor which predominantly controls the growth rate. The time for complete growth of bubbles in a silicic melt varies from a few seconds at atmospheric pressure to days or even a few weeks at pressures of 200-400 MPa. I report in Fig. 3 the results of Proussevitch et al. (1993) relative to the influence of the ambient pressure. It is possible to see that the bubbles have a growth time of the order of hours to days at a pressure corresponding to a depth of about 7 km.
3.4. Harmonic tremor The increase of the vapour phase within the magmatic reservoir is suggested also by other geophysi1000
1 MPa 10 MPa '"
::,
100
~
10
50 1O?z00
300 350 9
:::. ~ a:
0.1
10
10 3 Time (5)
lOS
Fig. 3. Diffusive growth of bubbles in rhyolitic melt for different values of ambient pressure from 0.1 to 400 MPa (saturation pressure for 8 wt.% of water). Redrawn after Proussevitch et al. (J 993).
cal peculiarities. During the Mt. St. Helens eruption, strong harmonic tremor was measured throughout the eruption (Malone et al., 1981). After 11 :40 (period 2), the tremor rapidly increased and saturated the entire seismic network of Washington State. Harmonic tremor is the most elusive seismic signal recorded in a volcanic environment. It is a continuous ground shaking with frequency content similar to that of low-frequency earthquakes (1-5 Hz), and several models have been proposed to explain its origin. Aki et al. (1977) proposed that tremor may originate by the repeated opening of cracks filled with magma. This model was developed to explain episodes of deep tremor at the basaltic volcano Kilauea on Hawaii. Crosson and Bame (1985) have proposed a simplified model based on a magma chamber filled with a gas phase. This model is able to produce a resonant radial oscillation that is dependent on the existence of a froth region inside the magma chamber and which explains many of the characteristics of harmonic tremor. The frequency is relatively insensitive to the size of the magma chamber. The synchronous oscillation of the magma chamber and surrounding rocks can exist at low frequency only when the froth region is sufficiently large. As magma is withdrawn, the amount of magma versus gas in the chamber decreases, with a consequent change in the seismic wave generation because of the large acoustic impedance between the magma-gas mixture and the country rock. Although I am not aware of systematic changes in the character of the tremor during the eruption of Mt. St. Helens, a systematic change of the frequency of tremor was observed, for example, during the 1944 eruption of Vesuvius at the transition between the low-rate, effusive-phase and the higher-rate explosive one (Imbo, 1952). According to the present model, the appearance of strong harmonic tremor during the May 18 eruption is indicative of extensive bubble nucleation within the magma chamber; drastic reduction of tremor occurs after 16:15 when the lithostatic pressure is restored by the collapse of the walls.
3.5. Variation of magma discharge during explosive eruptions The acceleration of the rate of bubble growth (Fig. 3) may result in an increase of magma dis-
R. Scandone / Journal of Volcanology and Geothennal Research 72 (J996) 71-83
charge. The importance of this process can be understood by the temporal evolution of the magma discharge during explosive eruptions. A slow increase of magma discharge and subsequent rapid waning is observed in the course of many explosive events (Fig. Ib). This pattern has been observed during the eruptions of Mt. St. Helens in 1980, El Chichon in 1982 and Pinatubo in 1991, having a Volcanic Explosivity Index (Newhall and Self, 1982) in the range of 4-6, but also several other smaller eruptions such as Pavlov and Redoubt in 1989-1990 display a similar behaviour (McNutt, 1987; Power et aI., 1994). The deposits of major eruptions have a typical succession of facies (Sparks et aI., 1973; Walker, 1985). The beginning of the eruption is commonly characterized by vent opening and a plinian phase with a reversely graded pumice-fall deposit. This phase is followed by a wavering plinian stage with partial column collapse and intraplinian ignimbrites or surge deposits. The climactic stage is reached with the eruption of an ignimbrite characterized, in proximal areas, by a lag-breccia deposit. The deposition of this typical sequence has been explained as due to conditions of magma discharge that allow a sustained eruption plume which later collapses giving rise to pyroclastic flows or surges (Sparks and Wilson, 1976; Sparks et aI., 1978). Reverse grading of the pumice deposit, as well as the generation of ignimbrites, suggest an increase in the discharge rate of magma until the emission of the pyroclastic flow, followed by a rapid waning in discharge rate. Fluiddynamical models, based on the analysis of quasisteady phases of eruptions, have explained that an increase in magma discharge can actually lead to the collapse of the eruption column (Wilson et aI., 1980; Wood, 1988). As mentioned before, the variation of conduit geometry by erosion can justify an increase of magma discharge by flaring of the vent as proposed by Wilson et aI. (1980). However, there is some questioning (Varekamp, 1993) about the field evidence for conduit erosion. Conduit erosion and widening of the vent is a requisite for eruptions with supersonic exit velocities for the gas-pyroclast mixture. Erosion probably occurs in the upper conduit, whereas the deeper parts of it, below the fragmentation level and the exsolution level, are likely to remain of the same
77
size. No relevant increase of discharge is allowed if a large part of the conduit remains of the same section, unless there is a substantial change of the pressure gradient driving the flux because of other causes. The pattern of volcanic activity during the May 18, 1980 eruption of Mt. St. Helens is analyzed by Criswell (1987) and Carey et aI. (1990). With reference to the already cited periods there are: Period 1 and most of period 2. The initial plinian phase occurred between 09:00 and 12:15; the progressive increase of magma discharge led to the generation of pyroclastic flows at 12: 15. At 11 :40 the seismographs around Mt. St. Helens began to show a gradual increase in strong ground shaking and by 13:30 it was impossible to discriminate individual earthquakes (Foxworthy and Hill, 1982; Scandone and Malone, 1985). Period 3. The maximum rate of magma output occurred between 15:00 and 16:30 with the widespread deposition of pyroclastic flows and the maximum release of seismic energy. Period 4. At 16:35 there was a decrease in magma discharge rate and a return to a plinian phase. The general pattern of the eruption described by these different observations is shown in Fig. 4. It is similar to fig. 3 of Criswell (1987), but the magma discharge is modified to take into account the observations of Carey et aI. (1990). Based on these and the previous observations I hypothesize that the eruption of silicic magma from shallow reservoirs is strongly controlled by the expansion of the magma + volatile phase because the rigid behaviour of the wall rocks does not permit the transmission of the elastic strain of the surrounding rocks. 3.6. Collapse of the chamber
The growth of bubbles tends to re-establish the original magmatic pressure (litho static or higher); it causes a steady increase in the discharge of magma until, eventually, conditions are attained to produce pyroclastic flows by non-convecting eruption columns. The process can last as long as the chamber behaves as a rigid envelope and as long as vesiculation of the water-rich melt occurs. The magma chamber becomes progressively filled with a higher proportion of gas and the density of the magma may
78
R. Scandone / Journal of Volcanology and Geothermal Research 72 (J996) 71-83
Strong ground shaking
.~
f
1000
800
1800
1200
hour
Cumulative seismic moment
l~ 4*10
2*10
23 I'
I
800
1000 Magma Discharge
1400
1200 (kg/sec)
1800
1600
Pyroclastic Flows Generation
Late Plinian Phase
1" ....---«::00 - - - - - . ......
I
Early Plinian Phase
•
------~ ~~~-"" 800
1000
1200
1400
1600-
1800
Fig. 4. Variation of the main physical parameters of the eruption of May 18, 1980 of Mt. St. Helens. Tremor amplitude (top) (Scandone and Malone, 1985). Seismic strain release (middle) (Shemeta and Weaver, 1986). Magma discharge rate (bottom) (Carey et aI., 1990).
become sufficiently low that it becomes mechanically unstable (Tait et aI., 1989). The most voluminous and intense eruptions cause the formation of calderas. Scandone (1990) suggested that the mechanism of caldera formation related with the collapse of a magma chamber is due both to the size and aspect ratio of the reservoir as well as the nature of the rocks making up the roof of the chamber. The same collapse of the chamber is indicative of a brittle behaviour of materials, in fact an elastic readjustment of the rocks surrounding the magma chamber would cause a down-sagging of the surface (Walker, 1984), as observed for example at Campi Flegrei after the end of periods of unusual uplift (Lirer et aI., 1987). A rigid behaviour of the magma chamber walls would instead cause a collapse of the chamber when the pressure difference between the magma and the lithostatic load becomes sufficiently larger than the strength of the rocks (Druitt and Sparks, 1984). We can estimate, through mass conservation, the volume ~ V of drained magma producing a pressure decrease equal to the strength of the material +
eventual overpressure in a magma chamber containing a volume VI of melt: PIVI = P2VI
+ PI LlV
Ll V = VI PI - P2 PI
where PI and P2 are the density of the original magma and of the magma plus bubbles, respectively. The density of the magma plus bubbles can be estimated (Wilson and Head (1981) according to: 1
n
1- n
P
Pg
PI
-=-+-where n is the weight fraction of ex solved gas, and Pg and PI are the density of gas and liquid, respectively. The gas density can be estimated through the equation of state of the gas, assuming a good thermal contact between gas and magma:
RT
p=pg
m
R. Scandone / Journal of Volcanology and Geothermal Research 72 (J996) 7I -83
R = 8.3143 JK- 1 mol- 1 is the universal gas constant and m is the molecular weight for water (= 0.Q18 kg). The content of water for the mafic dacite of Mt. St. Helens has been estimated by Rutherford et al. (1985) at 4.6(± 1.0) wt.%. This estimate is between the value reported by Melson and Hopson (1982) of about 7%, and that of Eichelberger and Hays (1982) of 0.7-1.7%. Rutherford et ai. (1985) suggest that the magma was undersaturated in water; however, it is possible that a mixed volatile gas (C0 2 + H 20 + S02) was present (Johnson et aI., 1994). I will hypothesize, for modeling purposes, that the magma of Mt. St. Helens was gas saturated, and use the ex solution law for water provided by Burnham (1975) for rhyolite. The weight fraction of water dissolved in the melt is given by: n
=
4.11 X 10- 6 po.s (p in pascals)
A rhyolitic melt at a pressure of 185 MPa (depth of - 7 km) has a saturation content of 5.6 wt.% of water (the upper limit of Rutherford et aI., 1985). I use a total pressure decrease of 40 MPa (hypothesizing an overpressured magma chamber, for example 20 MPa, and a decrease of another 20 MPa below lithostatic pressure, corresponding to the strength of rocks). This decompression produces an ex solution of about 0.6 wt% of water, causing a density decrease from 2400 to 2282 kg/m3. The amount of magma necessary to be drained to produce such a pressure decrease is approximately 4.9% of the original volume. This value gives an estimate of 5.1-10.1 km 3 for the volume of the Mt. St. Helens magma chamber, using the volume of products erupted during the May 18 eruption (0.25-0.5 km 3 ) (Lipman et aI., 1981; Criswell, 1987). The value agrees well with the estimate of 10-20 km 3 made by Scandone and Malone (1985) for the earthquake-free zone considered to be the magma reservoir. The same method, applied to the eruption of Mt. Pinatubo in 1991, gives a result of a magma chamber of approximately 96 km 3 for an eruption of 5 km 3 of dacite (6.4% weight of water at 220 MPa and decompression of 40 MPa). The results is comparable with the 40-90 km 3 estimated from seismic evidence (Mori et aI., 1993). These rough estimates support the idea that highpressure vesiculation in the chamber and the amount
79
of saturated melt are important factors affecting the development of an eruption. 3.7. Deposit associated with the collapse
The collapse of the magma chamber tends to re-establish the lithostatic load. Collapse may not occur in a single phase as arching of the roof eventually determines a confining pressure less than the original one, but, as a rule, collapse is followed by a rapid waning in the rate of discharge. The violent earthquakes occurring around and immediately above the magma chamber possibly cause extensive fracturing and produce lithic breccia (Walker, 1985; Scandone, 1990; Varekamp, 1993). During the May 18 eruption of Mt. St. Helens, the emission of lithic breccia (lithic breccia zone of Criswell, 1987) between 15:00 and 16:30 occurs contemporaneously with the highest seismic energy release and peak magma discharge (Carey et aI., 1990). After this, there is a decrease in seismic energy release and magma discharge.
4. Discussion and conclusions Several other recent explosive eruptions closely monitored by seismic networks give support to the ideas proposed in this paper. The eruptions of Pinatubo in 1991, of Redoubt in 1989-1990 and of Mt. Spurr in 1992 provide further evidences of a close cause-effect relationship between the explosive emission of magma and the following swarms of deep earthquakes (Pinatubo Volcano Observatory Team, 1991; Wolfe, 1992; Power et aI., 1994; Alaska Volcano Observatory, 1993). The comparison of post-eruption seismicity observed at Mt. Pinatubo and Mt. St. Helens is shown in Fig. 5. The spacetime evolution of seismicity of the 1992 eruption of Mt. Spurr shows a close similarity with that of Mt. St. Helens shown in Fig. 2 (Alaska Volcano Observatory, 1993). Pre-eruption seismicity probably reflects either ascent of viscous magma because of its buoyancy (Mt. St. Helens and Pinatubo) and/or opening of fractures favoured by expansion of a vapour phase (as suggested for example by the occurrence of
80
R. Scan done / Journal of Volcanology and Geothermal Research 72 (1996) 71-83
low-frequency seismic events at Redoubt - Power et aI., 1994). Deep seismic activity around the magma chamber, which starts after the beginning of the eruption, reflects instead collapse of the chamber because of decompression of the magma and its readjustment to equilibrium conditions. This phenomenon can occur only if the walls of the chamber have an initial rigid behaviour. The basic ideas expressed in this paper suggest that the mechanical properties of the rocks surrounding the magma chamber control the mechanism of
the eruption. In some eruptions, especially basaltic ones, the driving pressure is furnished by the elastic strain of the medium surrounding the magma chamber. A purely elastic response produces a high pressure gradient, generally at the beginning of the eruption, and will result in a high initial discharge rate (in the case of constant conduit shape and water content of the magma). The gradual decrease of the overpressure does not permit extensive vesiculation within the magma chamber. The elastic mechanism is not permitted in the case
w
°
(f') I-
•
(1)
E
E
a
10
~\ • ,," l i"
-
..::o!
0
c:
15
L:
+-'
a. (1) o
20
0
o
°'-----'
I km
B
a°,?.c>#, o
0-""
• 0
<> a
.0
"
\l ' a
0
"
0
~
+-' (1)
E a
r
10
"
~
00
~ ":00 o '" ~ 0
+-J
a. (1) o
rIl a '0 OJ
L:
•• "
15 km
(1)
c:
.
10
5
Distance in kilometers
..::o!
15°10'
..
.f-'J
5
+-J (1)
E
0
t)
O("J 0 0 •
80 r,)Qo
., t)
20~--~~ ? --~-----L--~
o
5
10
15
Distance in kilometers
15°05'
if
c
t
15°00' Fig. 5. Comparison of post eruption seismicity of Mt. St. Helens (top) (Scandone and Malone, 1985) and Pinatubo (bottom) (Wolfe, 1992). Note that, in both cases, the collapse of rocks defined by the seismic zone around the seismic-free magma chamber occurred toward the end of the eruption.
R. Scandone / Journal of Volcanology and Geothermal Research 72 (J996) 71-83
of a rigid behaviour of the chamber walls (commonly during silicic eruptions). In this case, the pressure driving the eruption is furnished at the expense of the magma internal energy, mostly by an expansion of the volatile phase. The rate of bubble growth and expansion of the magma govern the rate at which magma is expelled by the chamber. Maximum eruption rate possibly occurs when the bubble growth rate attains high values. The high discharge favours the tapping of deeper levels in zoned magma chambers (Spera, 1984; Blake and Ivey, 1986) with possible eruption of magma with a lower gas content. These two factors greatly favour column collapse (Wilson et aI., 1980) with consequent pyroclastic flow activity, as observed at Mt. St. Helens (Carey et aI., 1990). The length of the eruption is controlled by the amount of volatiles in solution in the magma and by the mechanical stability of the reservoir. These two factors combine in controlling the amount of material that must be erupted before a decompression sufficient to cause a collapse of the wall rocks occurs. The work of Proussevitch et aI. (1993) suggests that the ambient pressure (depth) strongly controls the growth rate of bubbles, with higher pressures slowing the process. Within this frame, I suggest that deep chambers possibly favour a slow development of the eruption with a relatively long time span between the start of eruptions and their paroxysmal phases when there are maximum discharge rates and possible pyroclastic flow emplacement events. Shallower reservoirs would instead cause fast-developing eruptions because of higher bubble growth rates at lower confining pressures. Although this model makes a strong separation between the two modes of behaviour of surrounding rocks, it is probable that a mixed behaviour is the norm. However, the predominance of an elastic readjustment of the rocks seems more likely during basaltic eruptions; the rigid behaviour is instead favoured during the eruption of silicic magma.
Acknowledgements
My sincere appreciation goes to Christopher Kilburn, Lisetta Giacomelli and Grazia Giberti for critical reading and suggestions to an earlier version of
81
this manuscript. Steve McNutt, Elisabeth Parfitt, Steve Sparks and Lionel Wilson made thoughtful reviews of this paper; their comments and suggestions resulted in a substantial improvement of the manuscript. I am particularly grateful to Steve Sparks for drawing my attention to recent models of diffusive bubble growth. This work was made when the author was at Dipartimento di Geologia e Geofisica, Universita di Napoli. The financial support of grant MURST 60%, Universita di Napoli, is aknowledged.
References Aki, K., Fehler, M. and Das, S., 1977. Source mechanism of volcanic tremor: fluid-driven crack models and their application to the 1963 Kilauea eruption. J. Volcano!. Geotherm. Res., 2: 259-287. Alaska Volcano Observatory, 1993. Mt Spurr's 1992 eruptions. EOS, Trans. Am. Geophys. Union, 74: 217, 221-222. Anderson Jr., AT, 1991. Hourglass inclusions: theory and application to the Bishop Rhyolitic Tuff. Am. Minera!., 76: 530547. Anderson Jr., A.T., Newmann, S., Williams, S.N., Druitt, T.H., Skirius, C. and Stolper, E., 1989. H 2 0, CO, CI, and gas in plinian and ash flow Bishop rhyolite. Geology, 17: 221-225. Anderson, O.L. and Grew, P.C., 1977. Stress corrosion theory of crack propagation with applications to geophysics. Rev. Geophys. Space Phys., 15: 77-104. Banks, N.G., Tilling, R.I., Harlow, D.H. and Ewert, J.W., 1989. Volcano monitoring and short term forecasts. In: R. Tilling (Editor), Volcanic Hazards. (Short Course in Volcanology') Am. Geophys. Union, Washington, DC, pp. 51-80. Barker, S.E. and Malone, S.D., 1991. Magmatic system geometry at Mount St. Helens modeled from the stress field associated with posteruptive earthquakes. J. Geophys. Res., 96: 11,88311,894. Belkin, H.E. and De Vivo, B., 1993. Fluid inclusions study of ejected nodules from plinian eruptions of Mt SommaVesuvius. J. Volcanol. Geotherm. Res., 58: 89-100. Blake, S., 1981. Volcanism and the dynamics of open magma chamber. Nature, 289: 783-785. Blake, S., 1984. Volatile oversaturation during the evolution of silicic magma chambers as an eruption trigger. J. Geophys. Res., 89: 8237-8244. Blake, S. and Ivey, G.N., 1986. Density and viscosity gradients in zoned magma chambers, and their influence on withdrawal dynamics. J. Volcano!' Geotherm. Res., 30: 201-230. Burnham, C.W., 1972. The energy of explosive eruption. Earth Mineral. Sci. Pa. State Univ., 41: 69-70. Burnham, C.W., 1975. Water and magmas; a mixing mode!. Geochim. Cosmochim. Acta, 39: 1077-1084. Burnham, C.W., 1985. Energy release in subvolcanic Environments: Implications for breccia formation. Econ. Geol., 80: 1515-1522.
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