Modulation of magmatic processes by CO2 flushing

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


(e) 8

Crystallisation 0

0.00 0.50

0

dissolved CO2 (wt.%)

0.00

0.2

[1.97 to1.98]

0.20

0.30


Etna

0.0

(b)

1


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


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


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


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



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


0.8

0.60

0.5 0.6 0.7 0.8

0.8

1.0

1.0

0.9

0 0.1


(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