pressure, temperature and volume in magmatic systems can produce unrest, which. 10 ... We calculate the chemical and physical evolution of magmas subjected to CO2-. 19 .... the impact of CO2 flushing on the evolution of the chemical and physical .... with their K2O content, a proxy for magma crystallisation, to identify the ...
1
Modulation of magmatic processes by CO2 flushing
2
Luca Caricchi1, Tom E. Sheldrake1, Jon Blundy2
3
1) Department of Earth Sciences, University of Geneva, Rue des Maraîchers 13, 1205, Geneva,
4
Switzerland
5
2) School of Earth Sciences, University of Bristol, Wills Memorial Building, BS8 1RJ, Bristol, UK
6 7
Abstract
8
Magmatic systems are the engines driving volcanic eruptions and the source of fluids
9
responsible for the formation of porphyry-type ore deposits. Sudden variations of
10
pressure, temperature and volume in magmatic systems can produce unrest, which
11
may culminate in a volcanic eruption and/or the abrupt release of ore-forming fluids.
12
Such variations of the conditions within magmatic systems are commonly ascribed to
13
the injection of new magma from depth. However, as magmas fractionating at depth
14
or rising to the upper crust release CO2-rich fluids, the interaction between carbonic
15
fluids and H2O-rich magmas stored in the upper crust (CO2 flushing), must also be a
16
common process affecting the evolution of subvolcanic magma reservoirs. Here, we
17
investigate the effect of gas injection on the stability and chemical evolution of
18
magmatic systems.
19
We calculate the chemical and physical evolution of magmas subjected to CO2-
20
flushing using rhyolite-MELTS. We compare the calculations with a set of melt
21
inclusion data for Mount St. Helens, Merapi, Etna, and Stromboli volcanoes. We
22
provide an approach that can be used to distinguish between melt inclusions trapped
23
during CO2 flushing, magma ascent and decompression, or those affected by post-
24
entrapment H2O-loss. Our results show that CO2 flushing is a widespread process in
25
both felsic and mafic magmatic systems. Depending upon initial magma crystallinity
26
and duration of CO2 input, flushing can either lead to volcanic eruption or fluid
1
27
release. We suggest that CO2 flushing is a fundamental process modulating the
28
behaviour and chemical evolution of crustal magmatic systems.
29 30
1. Introduction
31
Magmatic fluids are mainly composed of H2O and CO2, the latter being significantly
32
less soluble in silicate melts (Ghiorso and Gualda, 2015; Newman and Lowenstern,
33
2002). Crystallisation of silicate magma at depth and/or its ascent through the crust
34
leads to preferential release of CO2-rich fluids (Blundy et al., 2010; Newman and
35
Lowenstern, 2002; Papale et al., 2006), such that magmas stored in the upper crust
36
contain 40-60 times more H2O than CO2 (GEOROC database; Fig. 1a; Barsanti et al.,
37
2009). Thus, magmas stored at shallow depths are, in effect, high temperature
38
chemical reactors exposed to flushing of hot CO2-rich fluids released from deeper in
39
the magmatic system (Barsanti et al., 2009; Blundy et al., 2010; Evans et al., 2016;
40
Metrich et al., 2004; Moretti et al., 2013; Yoshimura and Nakamura, 2011; 2010).
41
Similar interactions between CO2-rich fluids and magmas can also occur in carbonate-
42
hosted magmatic systems (Blythe et al., 2015; Cross et al., 2013; Di Rocco et al.,
43
2012; Iacono Marziano et al., 2007; Jolis et al., 2013; Mason et al., 2017; Preece et
44
al., 2014).
45
An increase in the partial pressure of CO2 in the fluid phase in equilibrium with a
46
magma, produced either by CO2 flushing or by interaction with carbonate, leads to the
47
dissolution of modest amounts of CO2 in the melt and exsolution of several wt.% of
48
H2O (Ghiorso and Gualda, 2015; Newman and Lowenstern, 2002; Yoshimura and
49
Nakamura, 2011; Fig. 1a, b). The effects are a net increase of magma volume (i.e.
50
excess fluid volume fraction), progressive magma crystallisation due to reduced
51
weight fraction of dissolved H2O (Fig. 1a,b; Blundy and Cashman, 2008; Blundy et
2
52
al., 2010; Riker et al., 2015), and changes in crystallising assemblage and mineral
53
chemistry (Riker et al., 2015).
54
Previous studies have applied Rhyolite-MELTS (Gualda et al., 2012) to quantify the
55
effect of crystallisation and excess volatile exsolution on the evolution of the physical
56
properties of magma and fluid overpressure (Tramontano et al., 2017; Fowler and
57
Spera, 2008). Here, we use Rhyolite-MELTS to simulate interaction between CO2-
58
rich fluids and magmas stored in the upper crust and quantify the impact of this
59
process on the chemical and physical evolution of the system. We also simulate
60
simple, adiabatic decompression of volatile-bearing magmas to compare the effects of
61
these two contrasted processes on the chemical evolution of the residual melt (i.e.
62
melt in equilibrium with fluid phase and crystals). The results of our calculations
63
allow us to explore melt inclusion data collected in systems for which extensive
64
datasets exist (Mt. St. Helens, Merapi, Etna and Stromboli; GEOROC and Blundy et
65
al., 2010; Metrich et al., 2010; Preece et al., 2014). We propose a new approach that
66
can be applied to melt inclusions datasets to distinguish between magma-ascent,
67
interaction with CO2-rich fluids or post-entrapment H2O loss. The application of this
68
method to felsic and mafic magmatic systems (GEOROC and Blundy et al., 2010;
69
Metrich et al., 2010; Preece et al., 2014) show that melt inclusion data invariably
70
record interaction between magma and CO2-rich fluids. Consequently, we focus on
71
the impact of CO2 flushing on the evolution of the chemical and physical properties of
72
upper crustal magma reservoirs, and discuss the influence of CO2 flushing on the
73
ability of a magma reservoir to erupt magma and/or release magmatic volatiles.
74 75
2. Methods
76
2.1 Rhyolite-MELTS
3
77
We performed calculations using the thermodynamic software Rhyolite-MELTS
78
(version 1.1.0; Ghiorso and Gualda, 2015; Gualda et al., 2012) to simulate the effects
79
of CO2 flushing and adiabatic ascent on the chemical and physical evolution of
80
volatile-bearing magma. The chemical interaction between pure, hot CO2 and magma
81
is performed considering no subsequent heat loss from the system. CO2 is injected at
82
the liquidus temperature of the magma. In the calculations we consider either an
83
initially H2O-saturated (CO2-free) rhyolitic magma (composition equivalent to the
84
average melt inclusion composition of the early phase of the Bishop Tuff eruption:
85
Gualda et al., 2012; Ghiorso and Gualda, 2015) or a H2O-bearing basalt (Etna; melt
86
inclusion composition MI 33 of Metrich et al., 2004; 3 wt. % H2O; Tables 1, 2) at
87
their respective liquidus temperatures. Rhyolite-MELTS was chosen as it performs all
88
thermodynamic calculations with the latest available solubility data for magmas of a
89
wide variety of compositions (Ghiorso and Gualda, 2015). Simulations were carried
90
out at fixed confining pressures of 100, 200 and 300 MPa and constant oxygen
91
fugacity (fO2). Simulations for Bishop Tuff and Etna compositions were completed
92
with fO2 = NNO (Tables 1, 2). To check for the effect of redox state additional
93
simulations for the Etna composition were performed at fO2 = QFM (Table 1S). We
94
simulate isenthalpic assimilation of a fluid phase composed of pure CO2. CO2 is
95
added in steps of 0.01 grams up to a total of 10 grams of CO2 for 100 grams of
96
magma. Thermodynamic equilibrium is achieved at each step of the calculations
97
between melt, injected fluid and crystals (i.e. we consider crystallisation at
98
equilibrium).
99
A second set of simulations was performed to simulate adiabatic ascent of volatile-
100
bearing magma: 1) isentropic ascent with the same compositions as those used for the
101
CO2 flushing simulations, but with the magma initially containing some CO2 (Tables
4
102
1, 2); 2) ascent of magma initially containing excess fluids (Tables 1, 2). In both
103
cases, the starting composition was equilibrated at fO2 = NNO buffer, but left free to
104
vary during magma ascent and degassing.
105
In all simulations we trace the evolution of magma volume, temperature, crystallinity,
106
and residual melt chemistry (Figs. 2-5) to provide insights into how CO2 flushing and
107
adiabatic ascent of magmas influence respectively the chemical variability and
108
volatile content of melt inclusions. We consider closed-system (no release of excess
109
fluids) and thermodynamic equilibrium in our simulations. The release of the excess
110
volatile phase, if equilibrium between fluid and magma is maintained, does not affect
111
the chemical evolution of the system. However, as discussed more extensively in
112
Section 4.2, the release of a portion of the excess fluids can affect the total volume
113
increase of the system and magma compressibility, thus influencing the overpressure
114
produced. Whilst, closed-system assumption might be more conceivable for large
115
silicic systems in comparison to more dynamic basaltic open-conduit systems, we
116
focus on closed system conditions to understand the effects of CO2-flushing on
117
magmatic systems of varying chemistry.
118 119
2.2. Melt inclusions calculations
120
We calculated pressure of entrapment, and composition (xH2O) of the excess fluid
121
phase for melt inclusions (MI) assuming entrapment at fluid-saturated conditions. We
122
do not consider the individual composition of each melt inclusions when computing
123
entrapment pressure and the composition of the excess fluid phase, but consider
124
instead generic solubility models applicable to felsic and mafic compositions from
125
Ghiorso and Gualda (2015). Calculations for Mount. St. Helens and Merapi MIs were
126
performed with the solubility model provided for the Bishop Tuff in Ghiorso and
5
127
Gualda, (2015). For Etna and Stromboli MIs, we considered the H2O-CO2 solubility
128
model provided for MORB in Ghiorso and Gualda (2015). We calculate the
129
entrapment pressure and the composition of the excess fluid phase by determining the
130
isobar and fluid isopleth passing through the H2O and CO2 content of each MI. The
131
resulting entrapment pressure and composition of the excess fluid phase are combined
132
with their K2O content, a proxy for magma crystallisation, to identify the processes
133
responsible for MI entrapment and chemical composition. Note that K2O can be
134
substituted by any other element that is incompatible during crystallisation.
135
To test the reliability of our generic solubility approach, we calculated the entrapment
136
pressure for several pairs of melt inclusions (a couple from Merapi and Mt. St.
137
Helens, and another from Stromboli-Etna) considering explicitly their compositions
138
and using MagmaSat (Ghiorso and Gualda, 2015) to calculate pressure and fluid
139
chemistry. The comparison between the calculations performed with our approach
140
and MagmaSat, shows that the uncertainty is never larger than ±25 MPa and about
141
0.05 in xH2O (Figs. 6, 7).
142 143
3. Results
144
3.1. CO2 - magma interaction
145
The progressive addition of CO2 to H2O-saturated magma at constant confining
146
pressure leads to progressive dissolution of CO2 and concomitant exsolution of H2O
147
(Fig. 1) leading to an increase in the mass of excess fluid phase in equilibrium with
148
the magma (Figs. 1, 2a, 3a). The impact of CO2 flushing on melt fraction and
149
temperature differs between felsic and mafic magmatic systems (Fig. 2, 3). For felsic
150
magma, the molar fraction of H2O in the excess fluid phase (xH2Ofluid) decreases as a
151
consequence of the increasing amount of CO2 added to the system and the minor
6
152
amount dissolved in the residual melt phase (Fig. 2). The decrease of H2O
153
concentration in the residual melt phase drives magma crystallisation due to liquidus
154
temperature increase (Fig. 2b; Riker et al., 2015). As the calculations were performed
155
at isenthalpic conditions (no exchange of heat with the exterior), the release of latent
156
heat of crystallisation during CO2 flushing results in an increase of temperature,
157
which is more pronounced with decreasing confining pressure (Fig. 2c). In contrast, in
158
mafic magma CO2 flushing leads to progressive dissolution of CO2, but exsolution of
159
H2O is delayed until the system reaches fluid saturation (Fig. 3). The lower amount of
160
H2O initially dissolved in the mafic magma results in a smaller drop of the residual
161
melt fraction during progressive CO2 addition (Fig. 3b). In turn, this leads to the
162
progressive drop of temperature, as the latent heat of crystallisation is insufficient to
163
compensate for the heat required for excess fluid vaporisation (Fig. 3c).
164
Residual melt chemistry also changes with CO2 flushing because of progressive
165
magma crystallisation. We plot the results of the thermodynamic calculations
166
considering H2O, CO2 and K2O concentration in the residual melt combined with the
167
evolution of the molar fraction of xH2O in the excess volatile phase (Fig. 4). We have
168
selected these parameters because their evolution during CO2-flushing, magma ascent
169
(Blundy and Cashman, 2008) and post-entrapment H2O loss (Gaetani et al., 2012) is
170
different. This provides a chance to distinguish between these processes simply from
171
analysis of melt inclusion data. CO2 flushing at constant confining pressure produces
172
a decrease of H2O in the residual melt accompanied by an increase of CO2 and K2O
173
(Fig. 4). The variations occur together with a progressive decrease of xH2O in the
174
excess fluid phase (Fig. 4b). Magmas initially containing lower H2O (e.g. Etna with
175
respect to Bishop Tuff), exposed to CO2-flushing, experience less crystallisation than
176
magmas containing more H2O for comparable drops in xH2Ofluid (Fig. 4b, d). This
7
177
manifests graphically as a steeper drop of xH2Ofluid with respect to the decrease of
178
K2Omin/K2O in the residual melt (Fig. 4b, d; where K2Omin is the lowest content of
179
K2O in the set of melt inclusions of interest, i.e. the K2O of the less chemically
180
evolved MI).
181
Results of calculations performed for the Etna composition at fO2 = QFM buffer, are
182
almost identical to those performed at NNO (Tables 1, 2; Figs. 3, S1).
183 184
3.2. Variation of residual melt chemistry during magma ascent
185
We calculate the evolution of residual melt chemistry during magma ascent to identify
186
differences to the variations produced by CO2-flushing. Results, presented in H2O-
187
CO2 and K2Omin/K2O – xH2Ofluid space (Fig. 5), show that the increase of K2O content
188
of the residual melt phase occurs at relatively low pressure only after significant
189
amounts of H2O have been lost to the excess volatile phase (Fig. 5). This modest K2O
190
increase is associated with the limited crystallisation occurring at intermediate to high
191
pressure (500-200 MPa) during magma ascent (Fig. 5a). Upon decompression
192
xH2Ofluid increases while K2Omin/K2O of the residual melt only starts to increase at the
193
pressure at which significant H2O degassing occurs (Fig. 5b, d; Papale, 2005).
194
We tested also the effect of excess fluid on the evolution of residual melt chemistry
195
upon magma ascent using the Etna composition and considering an exsolved fluid
196
phase already present at depth (Fig. 5d; Table 1). The initial presence of excess CO2-
197
rich fluids promotes magma crystallisation and limits the increase of xH2Ofluid upon
198
ascent (Fig. 5e, f). As identified in previous melt inclusions studies, if sufficient
199
excess fluid is present initially, the volatile content of the ascending magma can
200
follow an isopleth corresponding to the molar ratio of CO2 and H2O in the excess
201
fluid phase (Metrich et al., 2004). However, we note that even in the presence of an
8
202
excess fluid phase, decompression between 500 and 200 MPa results in a modest
203
amount of crystallisation (Fig. 5e, f). Increasing the mass of excess fluid by a factor of
204
five with respect to the amount reported in Table 1 produced results similar to those
205
reported in Figure 5e and f.
206
In summary, thermodynamic calculations performed to simulate CO2-flushing and
207
magma ascent show that the residual melt chemistry changes in very distinctive ways
208
for these two processes. Therefore, the analysis of dissolved volatile content and
209
chemistry of melt inclusions remains a useful tool to probe processes accompanying
210
magma transfer and storage in the crust.
211 212
3.3 Melt inclusions
213
Once any effects of post-entrapment processes are accounted for, MIs can provide
214
information on magmatic processes taking place in crystallising portions of magmatic
215
systems (Metrich et al., 2004; Blundy and Cashman, 2001; Gaetani et al., 2012;
216
Papale, 2005). The relative chemical signals of magmatic and post-entrapment
217
processes can be deconvolved through the combined analysis of MI volatile content
218
and chemistry (inset Fig. 6b). Ascent of fluid-saturated magmas triggers exsolution of
219
volatiles, which are initially dominated by CO2 (Ghiorso and Gualda, 2015; Newman
220
and Lowenstern, 2002). As ascent continues, the amount of CO2 dissolved in magma
221
decreases and the proportion of H2O in the exsolved fluid phase increases (Fig. 5). If
222
ascent is accompanied by crystallisation (Blundy et al., 2010), MIs will record this
223
process with specific CO2-, H2O- and K2O trends (Fig. 5). Post-entrapment hydrogen
224
loss from MIs (Gaetani et al., 2012) leads to a simultaneous decrease in H2O and
225
increase in CO2 with respect to chemical trajectories produced by fluid-saturated
226
magma ascent but without variation of MI chemistry (inset Fig. 6a,b). CO2 flushing
9
227
leads consistently to a decrease of dissolved H2O and the combined increase of CO2
228
and incompatible elements such as K2O (Figs. 2-4, 6).
229
We consider MI data from four well-studied felsic and mafic volcanic systems
230
(Mount St. Helens, Merapi, Etna and Stromboli; GEOROC; Blundy et al., 2010;
231
Metrich et al., 2010; Preece et al., 2014). In all systems, MIs with the highest K2O
232
content contain the lowest H2O (Figs. 6a,c; 7a, c). Whereas ascent-induced
233
crystallisation can be responsible for the high K2O of MIs with the lowest H2O and
234
CO2 content, only CO2 flushing can produce water-poor and CO2 and K2O-rich melt
235
inclusions (Figs. 6a, c; 7a, c). The same conclusion is reached when projecting the
236
data in K2Omin/K2O – xH2Ofluid space (Figs. 6b, d; 7b, d). MI for Mount St. Helens and
237
Merapi seem dominated by the effects of CO2 flushing, while a minority can be
238
explained by ascent-induced crystallisation or H2O-loss (Fig. 6). This interpretation
239
differs from that of Blundy and Cashman (2005). In basaltic systems most of the melt
240
inclusions can be explained by CO2 flushing or ascent of magma initially saturated in
241
CO2-rich fluids (0.25 3.15]
0.4
0.6
K2Omin/K2O
0.8
1.0
[341 to 228] [455 to 341] [500 to 455]
decompression - fluid excess
0.2
0.10
0.20
0.30
0.0
Conf. Pressure (MPa)
0.40
K2O residual melt (wt.%)
dissolved H2O (wt.%)
8
(d)
1
6
0.8
4
Crystallisation
0.6
2
(c)
xH2Ofluid
0
2
4
6
dissolved H2O (wt.%)
(e) 8
Crystallisation 0
0.00 0.50
0
dissolved CO2 (wt.%)
0.00
0.2
[1.97 to1.98]
0.20
0.30
dissolved CO2 (wt.%)
Etna
0.0
(b)
1
dissolved H2O (wt.%)
0
6
0.8
4
Crystallisation
0.6
2
(a)
xH2Ofluid
0
0.50
0.00
0.6
[4.67 to 5.42]
[369 to 239]
xH2Ofluid
[4.63 to 4.67]
Conf. Pressure (MPa)
1
[108 to 10]
0.8
[4.61 to 4.62]
0.04
0.06
dissolved CO2 (wt.%)
0.08
K2O residual melt (wt.%)
0.10
Bishop Tuff
0.0
0.2
0.4
0.6
K2Omin/K2O
0.8
(f)
1.0
xH O 2 =0. 5
0.8
Ascent + initial fluids ng
hi
s flu
O2
C
5 =0 .7 xH O
0.4
2
0.1
200 MPa
Ascent
0.6
300 MPa
CO2 [wt.%]
1
[1.6 to 2.1] [2.1 to 2.3] [2.3 to 2.7] [2.7 to 2.9] [2.9 to 3.0] [3.0 to 6.3]
xH2O
0.15
K2O (wt.%)
H2O-loss
H2O-loss
xH O 2 =0.3
CO2-flushing
Ascent +/- initial fluid
0.2
Mt St Helens
typical uncertainty Pressure (MPa) [6 to 12]
8
0.0
0.2
0.4
0.6
K2Omin/K2O
0.8
(b)
1.0
1
Merapi [4.3 to 4.5] [4.5 to 4.7] [4.7 to 5.2] [5.2 to 5.4] [5.4 to 5.8] [5.8 to 6.0]
xH2O
0.15
Crystallisation
[179 to 400]
K2O (wt.%)
300 MPa
CO2 [wt.%]
[131 to 179]
0.8
3 xH O 2 =0.
0.2
6
[41 to 131]
0.6
4
H2O [wt.%] xH O 2 =0. 5
2
0.25
0
(a)
[12 to 41]
0.2
0.9
O= xH 2
0
100 MPa
0
0.05
±25 MPa
0.1
0.4
200 MPa
typical uncertainty Pressure (MPa)
0
[0 to 14]
100 MPa
2
4
H2O [wt.%]
6
0.2
xH
0.9
O=
2
(c)
[14 to 97] [97 to 179] [179 to 259] [259 to 400]
8
0
0
0.05
±25 MPa
0.0
0.2
0.4
0.6
K2Omin/K2O
0.8
(d)
1.0
1
K2O (wt.%)
xH O 2 =0.5
typical uncertainty
±25 MPa
0.6
0.8
[0 to 13] [13 to 23] [23 to 130] [130 to 206] [206 to 364] [364 to 500]
0.2 8
[1.2 to 1.5] [1.5 to 1.6] [1.6 to 1.7] [1.7 to 1.8] [1.8 to 3.7] [3.7 to 4.2]
0.4
0.6
K2Omin/K2O
0.8
O2
C
O= 2
1.0
[10 to 51] [51 to 114] [114 to 223] [233 to 276] [276 to 306] [306 to 380]
typical uncertainty
0.4
xH
0.8
±25 MPa
ng
hi
s flu
(b)
Pressure (MPa)
Ascent Ascent + initial fluids
0.7 5
0.15
0.2
1
Stromboli
0.1
2
200 MPa
4
H2O [wt.%]
xH
0.2
300 MPa
0.9
O= 2
(c) 6
8
0
0.05 0
100 MPa
0
0.0
0.6
xH O=0 2 .3
6
K2O (wt.%)
CO2 [wt.%]
0.2
4
H2O [wt.%]
H2O-loss
H2O-loss
Ascent +/- initial fluid
CO2-flushing
Crystallisation
(a)
xH2O
2
0.25
0
xH
300 MPa
xH O 2 =0.5
0
100 MPa
200 MPa
0.9
O= 2
0
0.1
0.4
xH
2
0.2
O= 0.7
5
CO2 [wt.%]
0.3
[0.5 to 0.9] [0.9 to 2.1] [2.1 to 2.7] [2.7 to 3.1] [3.1 to 3.4] [3.4 to 4.3]
Pressure (MPa)
xH2O
xH O=0 2 .3
0.4
Etna
0.0
(d) 0.2
0.4
0.6
K2Omin/K2O
0.8
1.0
100 MPa
0.7
0.6
0.5
0.4 0.70
0.76
0.68
0.63
100 MPa
0.58
0.55
0.52
0.49
0.47 1.5
0.65
0.8
1.3
Initial melt fraction 0.4 0.6
1.1
0.45
(a)
10
100 MPa
300 MPa 100 MPa
300 MPa
(b)
300 MPa
0.7
0.6
0
0.5
1.0
1.0
0.8
0.3
Solid xH2O fluid
Exsolved H2O (wt.%)
0
0
0 Fractions
0.2 0.4 0.6 0.8 1.0
2 4 6 8 Injected fluid ([wt.%] of flushed magma)
0.9
0.5
0.1
Melt fraction Bubble fraction
0
0.7
0.2
Solid
0.05
0.1
0.9
2 4 6 8 Injected fluid ([wt.%] of flushed magma)
300 MPa 100 MPa
0.82
0.77
0.73
300 MPa
0.69
0.67
0.64
100 MPa
(e)
0.62
0.60 2.5 2.3
0.8
0.8
0.3
0.30
2.1
0.25 0.1
0
0
0
2 4 6 8 Injected fluid ([wt.%] of flushed magma)
1.7 1.5
1.1 0.9 0.7 0.5
0.2
(c)
10
0
0.2
Solid
equilibrium
0.20
0.05
1.9
Initial melt fraction 0.4 0.6
Initial melt fraction 0.4 0.6
0.2
0.10
10
300 MPa
0.4 0.35
0.15
(d)
0
0.3
Solid
1.3
equilibrium
0.10
0.2
0
0.20
0.4 0.8 1.2 1.6 2.0
0.15
0.30
0.40
equilibrium
0.3
0.25
0.35
equilibrium
0.50
xH2O fluid
0.55
Exsolved H2O (wt.%)
0.2
Initial melt fraction 0.4 0.6
0.8
0.60
0.5 0.6 0.7 0.8
0.8
1.0
1.0
0.9
0 0.1
2 4 6 8 Injected fluid ([wt.%] of flushed magma)
(f)
10
3 10
log10 (time [yrs]) −1 0 1 2
-7
10
-6
1020 Pa s
Storage and fluid release 10
-5
Eruption
Viscous relaxation crust 1019 Pa s
10
-4
2 /s m / l 10 mo -3
10MPa < ΔPcrit