Vol48/04/2005def


775

ANNALS  OF  GEOPHYSICS,  VOL.  48,  N.  4/5,  August/October  2005

Key  words SiF4 – vapour crystallisation – silica –
degassing

1. Introduction

Silica minerals are frequently found as vesi-
cle fillings in recently erupted volcanic products,
and are commonly interpreted as late or post-
magmatic secondary phases. In a study of Mont-
serrat it was reported that groundmass silica
phases occur in slowly cooled (dome) lavas only
and not in pumice from explosive eruptions, in-
dicating that crystalline silica was produced by

vapour-phase crystallisation and devitrification
within the lava dome (Baxter et al., 1999). As
crystalline silica may pose a serious health threat
to humans if it enters the respiratory system, it
was suggested that volcanoes with lava domes
may pose a greater risk for silicosis than other
types of volcanoes (Baxter et al., 1999). Silica
enrichment in hydrothermal aureoles around
magma chambers is commonly interpreted as the
result of crystallisation from silica-saturated so-
lutions, and leaching of other elements by acidic
hydrothermal fluids (e.g., Giggenbach, 1984).
However, these processes are typical of post-in-
trusive alteration and it is unlikely that such
processes act in erupting lava flows and domes.
The importance of F species such as HF and SiF4
in mass transport during degassing was demon-
strated by White and Hochella (1992), who de-
scribe the depletion of weathering surfaces of
basaltic lava flows in SiO2 and enrichment in Ca,

Vapour-phase crystallisation of silica 
from SiF4-bearing volcanic gases

Jan C.M. de Hoog, Manfred J. van Bergen and Michel H.G. Jacobs
Department of Earth Sciences, Utrecht University, Utrecht, The Netherlands

Abstract
Thermodynamic modelling of magmatic gases shows that SiF4 may be an important F-bearing species at the high
pressures typical of magma reservoirs. Upon decompression during degassing, SiF4 will react with water vapour
to form HF and silica. Common magmatic gases of high-T fumaroles seem to contain too little SiF4 to be a sig-
nificant source of silica, except if extremely large amounts of gas percolate through a small volume of rock, as
is the case in lava domes. Only if fluorine contents of the gases exceed 1 mol% detectable amounts of silica may
be formed, but such high fluorine contents have not yet been observed in natural gases. Alternatively, silica may
be formed by heating of cool SiF4-rich gases circulating in cooling lava bodies. We suggest that these mecha-
nisms may be responsible for the deposition of crystalline silica, most probably cristobalite, observed in vesicles
in lavas from Lewotolo volcano (Eastern Sunda Arc, Indonesia). Silica occurs as vapour-crystallised patches in
vesicles, and is sometimes associated with F-phlogopite, which further supports F-rich conditions during depo-
sition. Because of the connection between F-rich conditions and high-K volcanism, we propose that late-stage
gaseous transport and deposition of silica may be more widespread in K-rich volcanoes than elsewhere, and
long-term exposure to ash from eruptions of such volcanoes could  therefore carry an increased risk for respira-
tory diseases. The dependence of SiF4/HF on temperature reported here differs from the current calibration used
for temperature measurements of fumarolic gases by remote sensing techniques, and we suggest an updated cal-
ibration.

Mailing address: Dr. Jan C.M. de Hoog, Department of
Earth Sciences, Gothenburg University, Guldhedsgatan 5A,
413 20 Göteborg, Sweden; e-mail: cees-jan@gvc.gu.se



776

Jan C.M. de Hoog, Manfred J. van Bergen and Michel H.G. Jacobs

Al, and Mg during post-eruptive degassing of
cooling magma bodies. 

Here we present thermodynamic calculations
that indicate that silica phases may form as a re-
sult of decompression of HF-rich gas during as-
cent of vapour-saturated magma, but also during
heating of SiF4-bearing gases. SiF4 is a highly
volatile compound at ambient pressures and tem-
peratures (sublimation point = − 86°C at atmos-
pheric pressure), which at elevated temperatures

readily reacts with water vapour to form HF and
silica. The reaction is very sensitive to tempera-
ture and is shifted in favour of SiF4 at low tem-
peratures (< 400ºC; Rosenberg, 1973), whereas in
high-temperature fumaroles HF is the dominant
fluorine-bearing species. The results will be used
to explain the presence of small patches of a crys-
talline silica phase, associated with fluorine-phl-
ogopite, in vesicles in lavas from Lewotolo vol-
cano (Eastern Sunda Arc, Indonesia). 

Fig. 1. Back-scatter electron images of a vesicle filling containing several patches of crystalline silica (cr), in-
ferred to be cristobalite. The dashed lines indicate the approximate outline of the vesicle and interconnecting
veins before the vesicle-filling assemblage was formed. Note the radial growth texture of the silica patches. The
small needles sticking into the open space are F-phlogopite. Inset shows detail of one of the silica patches.



777

Vapour-phase crystallisation of silica from SiF4-bearing volcanic gases

1.1. Geological background

Lewotolo volcano is located at Lomblen Is-
land (8°27′S, 123°59′E) in the Nusa Tenggara
(Lesser Sunda) Islands region of Indonesia.
Strong fumarolic action at the summit produces
a continuous outflow of hot (≤ 490°C) gases
rich in H2O, CO2, H2S, SO2, HCl and HF
(Poorter et al., 1991), pointing to the presence
of a degassing magma chamber at shallow
depth. Lewotolo has erupted porphyritic, high-
K calc-alkaline rocks over a wide range of sili-
ca contents from basalts to trachyandesites (47-
62 wt% SiO2; Stolz et al., 1990; Hoogewerff,
1999). High fluorine contents of bulk rock (up
to 800 ppm; Vroon, 1989), groundmass glass
(up to 6000 ppm), apatite and phlogopite (De
Hoog, 1995) indicate that the Lewotolo mag-
mas are relatively rich in fluorine. 

1.2. Petrography of vesicle fillings

A detailed description of the vesicle fillings
was presented in De Hoog and Van Bergen
(2000), in which the authors describe the pres-
ence of the rare mineral zirconolite in these fill-
ings. We will summarize the relevant parts here.

Vesicles in about half of the lavas from
Lewotolo volcano are partially to completely
filled with an assemblage of mainly alkali
feldspar, plagioclase and a silica phase. In addi-
tion the vesicles contain small amounts of cli-
no- and orthopyroxene and magnetite, and trace
amounts of zirconolite, ilmenite and apatite.
These partially filled former vesicles will be re-
ferred to as vesicle fillings (fig. 1). The fillings
have sieve-like textures and can be distin-
guished from the groundmass by a coarser
grain-size, higher porosity, and the absence of
glass. In samples with large amounts of vesicle
fillings the whole network of former vesicles is
interconnected. If the assemblage fills veinlets,
the texture is denser, with a lower porosity
close to that of the groundmass. 

A crystalline silica phase grew as isotropic
to low-birefringent patches with a radial struc-
ture into voids, or occupies interstitial spaces,
e.g., between crystals or near cracks (fig. 1). In-
dividual patches can reach sizes up to 200 µm;

the total volume in the rock is small, however,
and does not exceed 0.02 vol%. Optical mi-
croscopy did not conclusively distinguish be-
tween cristobalite and tridymite, but the com-
mon anisotropy and textures typical of shrink-
age features point to cristobalite. Thin whiskers
of phlogopite occasionally intrude into the left-
open space within the former vesicles. 

The vesicle fillings were probably deposited
from a migrating fluid phase, possibly derived
from deeper parts of the magma chamber, dur-
ing a very late stage of magmatic evolution (De
Hoog and Van Bergen, 2000). The morphology
of later-stage minerals (crystalline silica, phlo-
gopite) suggests that these may have crys-
tallised from a less dense vapour phase.

2. Thermodynamic modeling

We studied the formation of silica phases as
a result of decompression of SiF4-rich phases
using the equilibrium reaction

(2.1)

Although equilibrium between a vapour phase
and silica is a valid assumption for the forma-
tion of silica from a SiF4-bearing gas, reaction
(2.1) is only an approximation of equilibrium
between vapour and siliceous melt. In the un-
derlying paper we follow the deductions of
White and Hochella (1992) that the net free-en-
ergy differences between silica and siliceous
melt are negligible compared to temperature
and pressure effects.

In the calculations it is assumed that the rest
gases SO2, H2S, CO2 and HCl do not participate
in the equilibrium reaction given by reaction
(2.1). However H2S will react with H2O to form
SO2 and H2 following the reaction

(2.2)

The equilibrium of the reaction shifts to the
right with decreasing pressure, thus consuming
H2O. However, because the amount of S in vol-
canic gases is much smaller than the amount of
H2O, the effect of reaction (2.2) on reaction
(2.1) will be negligible.

( ) ( ) .g gO( ) ( )g gH S 2H SO 3H2 2 2 2D+ +

.( )gOH2+( ) ( ) ( )s g gSiO 4HF SiF2 4 2D+



778

Jan C.M. de Hoog, Manfred J. van Bergen and Michel H.G. Jacobs

Applying the equilibrium condition to reac-
tion (2.1) by setting the Gibbs energy of the re-
actants equal to that of the products results in

(2.3)

(2.4)

where P represents pressure, T thermodynamic
temperature, R the gas constant, K the equilibri-
um constant and the superscript «o» refers to the
reference pressure Po=1 bar and Pj represent the
partial pressures of the substances j = SiF4, H2O
and HF.

The Gibbs energies of the pure substances,
Gj

o(T), at 1 bar pressure are expressed in terms
of the enthalpy of formation, ∆Hf, jo , absolute
entropies, Sjo, and heat capacities CP, jo, as

(2.5)

The enthalpies of formation and absolute en-
tropies refer to a common reference tempera-
ture T0 = 298.15 K and are found in thermody-
namic tables (JANAF, 1985; Barin, 1989). The
heat capacity data in these tables were fitted
with polynomial functions with an accuracy
better than 0.1%.

The Gibbs energy of solid SiO2 at arbitrary
pressure and temperature conditions is calculat-
ed using

(2.6)

where j=SiO2 and Vj, the molar volume of
SiO2. The molar volume of a solid substance is
calculated using a Murnaghan equation of state

(2.7)
( )

( )V V
K
K

P P1
( )K

o
o

o
o

1
o

= + -
l l

: D

S +T $-+H∆=, T

+

( )G P C dT

T
C

dT V dP

, ,

,

j f j P j

T

T

j

P j

T

T

j

P

P

o o o

o

0

0
o

+

*

4

#

# #

.$T-+H∆ C dT=( )G T S
T

C
dT, ,

,

j f j P j

T

T

j

P j

T

T

o o o o

o

0 0

+* 4# #

+G-G4+( ,G P T)=( ,G P T T T

G T2

SiO HF
o

SiF
o

H O
o

2 4

2
-

∆ ) ] ]

]

g g

g

( )lnRT K=lnRT=( ,G P T
P

P P
P

HF
4

SiF H O
2

o4 2∆ ) c m

where Ko represents the isothermal bulk modu-
lus and (Ko)′ its pressure derivative at 1 bar
pressure. The molar volume at 1 bar pressure,
Vo, is calculated using the expression

(2.8)

where αo represents the thermal expansivity at
1 bar pressure. Polynomial expressions for Ko,
(Ko)′ and αo as function of temperature and the
molar volume of SiO2 were taken from the
work of Swamy et al. (1994).

Depending on pressure and temperature,
SiO2 may exist in the forms α-quartz, β-quartz,
tridymite, cristobalite, coesite and stishovite.
From these six forms of SiO2 we considered on-
ly α-quartz, β-quartz and tridymite because our
study was limited to a pressure range between
0-2.5 GPa and temperature range between
1100-1400 K. Note that cristobalite is not a sta-
ble phase at the pressure and temperature con-
ditions of our investigation, although it is often
observed to be the most common crystalline sil-
ica phase in volcanic rocks (e.g., Baxter et al.,
1999). The crystallisation of cristobalite, in-
stead of a stable silica form, has a negligible ef-
fect on the results of the thermodynamic calcu-
lations compared to effects associated with er-
rors in thermodynamic data. 

The calculations of the partial pressures of
HF, SiF4 and H2O and the partial pressure of the
rest gas at arbitrary pressure and temperature
conditions proceeds as follows. When a pressure
is applied to the system or when the temperature
of the system is changed, the partial pressures of
HF, H2O and SiF4 change and an amount of SiO2
is consumed. At the condition (Po,T), the number
of moles noHF, noSiF4 and n

o
H2O of each gas phase

species is calculated from the measured partial
pressures. According to eq. (2.1) at arbitrary con-
ditions (P, T′ ) an amount, m, of SiO2 reacts with
4m moles of HF to form m moles of SiF4 and 2m
moles of H2O. The amount of moles in the gas
phase changes from noHF + noSiF4 + n

o
H2O + n

o
restgas

to noHF + noSiF4 + n
o
H2O + n

o
restgas−m. The partial

pressures of the substances in the gas phase are

P
N

n m P
P

N
n m P4 2

HF

HF
o

H O

H O
o

2

2

=
-

=
+] ^g h

T T( ) ( ) ( )expV V T dT
T

T

o o
o

o

0

= af p#



779

Vapour-phase crystallisation of silica from SiF4-bearing volcanic gases

(2.9)

Inserting these relations into eqs. (2.3) and (2.4)
leads to

(2.10)

The Gibbs energy difference ∆G(P,T) is calcu-
lated by eq. (2.4). It is constrained by thermo-
dynamic data of the gas phase species and the
solid forms of SiO2. With noHF, noSiF4, n

o
H2O and

norestgas measured at the experimental condition
(Po,T), the amount m is calculated with eq.
(2.10), leading to calculated partial pressures of
the gas phase species at arbitrary pressure and
temperature conditions. The algorithm of the
computational procedure is included in the pro-
gram XiPT, which is a successor of the program
TXY-Calc (Jacobs et al., 1996; Jacobs and
Spencer, 1996).

( ,
.exp

n m n m n n n

n m

n m

P
P

RT
G P T

2

4

SiF
o

H O
o

HF
o

SiF
o

H O
o

HF
o

restgas
o

o

2

4

4 2 4 2
+ + + + +

-

+ -
=

∆ )

^ ^ ^

]
b

h h

g

h
l

.

P
N

n m P
P

N
n P

N n n n n m

SiF

SiF
o

restgas

restgas
o

HF
o

SiF
o

H O
o

restgas
o

4

4

4 2

=
+

=

= + + + -

^

^

h

h

The calculations were performed using the
compositions of hot (up to 490°C) summit fu-
maroles from Lewotolo volcano (Poorter et al.,
1991). Equilibrium compositions and tempera-
tures were recalculated by an iterative gas-
restoration procedure (Poorter et al., 1991) and
are given in table I. HCl, HF and H2 values are
in the normal range of volcanoes related to sub-
duction zones, but CO2 and SO2 values and
HF/HCl ratios are high compared to world-
wide high-temperature fumarole data given by
Symonds et al. (1994). We note that, because of
differences in solubility of the various volatiles,
the composition of gases at higher pressures is
potentially different from those at atmospheric
pressure, but these effects are difficult to quan-
tify and therefore not considered in the paper.

3. Results and discussion

3.1. Temperature dependence of HF/SiF4

The variation of HF/SiF4 with T for various
partial pressures of HF in the system SiO2-HF-
H2O-SiF4 is plotted in fig. 2. Note that the x-ax-
is in fig. 2 is a log scale, so that the increase in
HF/SiF4 between two markers (2 units) corre-
sponds to a 100-fold increase of HF/SiF4. The
temperature dependence of reaction (2.1) at 1
bar has been documented earlier by Rosenberg
(1973) and White and Hochella (1992), and
SiF4 and HF are also included in the GAS-
WORKS program used to calculate equilibrium
compositions of volcanic gases (Symonds et al.,
1992). As shown in fig. 2, our results closely
match those of Rosenberg (1973), except for a
slight deviation at lower temperatures. A few
data points published by Symonds et al. (1992)
are consistent with these trends. However, the
calculations of White and Hochella (1992) devi-
ate, showing a straight line with a slope which
is considerably less steep than the curves pre-
dicted by Rosenberg (1973) and our data. As
largely the same thermodynamic databases
were used in all cases, we hypothesize that the
trends of White and Hochella (1992) were
based on models or assumptions that are not
documented in their paper, which makes it dif-
ficult to verify their calculations. The agreement

Table I. Recalculated compositions of Lewotolo
gas samples in mol% (after Poorter et al., 1991).

Average 1σ
(n = 4)

H2O 73.5 4.4
CO2 15.9 4.4
SO2 7.76 2.5
H2S 0.86 0.51
HCl 0.169 0.106
HF 0.0465 0.0175
N2 1.39 1
H2 0.34 0.355

Equilibrium temperatures are 950 ± 50°C; sampling
temperatures were ∼ 490°C at 50 cm depth. Ar, He
and CH4 are present at ppm-level. Calculated oxygen
fugacities are ± 10−11 bar, slightly higher than the
QFM-buffer.



780

Jan C.M. de Hoog, Manfred J. van Bergen and Michel H.G. Jacobs

with the data of Rosenberg (1973) and Symonds
et al. (1992) supports that our model is valid for
1 bar, and can be used to predict equilibria at
higher pressures (see below).

The curves in fig. 2 show the variation of
HF/SiF4 with T for a fixed value of PHF and at a
total vapour pressure of 1 bar. As production of
SiF4 consumes HF, PHF is not constant, and the
curves do therefore not represent cooling paths.
However, HF consumption by production of
SiF4 is important only if PSiF4 is close to PHF, so
that in fig. 2 only the bottom part of the curves
would be noticeably influenced.

Reaction (2.1) is strongly dependent on
temperature and on partial HF pressure. For ex-
ample, a drop in temperature from 1400 to 800

K will decrease HF/SiF4 by over 3 log units
(∼1250 times). The amount of SiF4 will be
higher than HF below 500 K if PHF = 0.01 bar.
Therefore, at atmospheric pressure, SiF4 does
occur only in significant amounts at high partial
pressures of HF and low temperatures. Dilution
of HF-rich gas with HF-poor gas of the same
temperature will have a strong effect on SiF4,
e.g., diluting a gas 10 times will decrease PSiF4
by a factor 10000. Dilution of volcanic with at-
mospheric gases will lower T and PHF, two pa-
rameters with opposing effects on SiF4 produc-
tion. Modelling shows that conservative mixing
of high-T volcanic gases with air will increase
the relative amount of SiF4, because the tem-
perature effect outweighs the dilution effect.

Fig. 2. Variation of SiF4/HF (expressed as log[PHF/PSiF4]) with T (in K) at different partial pressures of HF, and
1 bar total vapour pressure (PHF+PSiF4+PH2O = 1). Continuous lines are from our own calculations, whereas dot-
ted lines are data from Rosenberg (1973), dashed lines are after White and Hochella (1993), and the two circles
are data points from Symonds et al. (1992) with PHF = 10−3.07 bar. Uncertainties in our data due to uncertainties
in heat of formation and enthalpy of SiO2, HF, H2O and SiF4 are about 25° and 0.2 log units. Note that the rel-
ative importance of SiF4 increases both with total amount of F in the system (i.e., PHF), and with decreasing tem-
perature. White and Hochella (1992) trends are clearly offset and predict HF/SiF4 variations with temperature
that are much larger than other models.



781

Vapour-phase crystallisation of silica from SiF4-bearing volcanic gases

3.2. Pressure dependence of HF/SiF4

Figure 3 illustrates the variation of PHF/PSiF4
with pressure at P*HF = 10−2.5 and 10−1.0 and tem-
peratures typical of magmatic gases (where P*HF
is the partial pressure of HF(PHF/Ptotal); P*HF =
=PHF at atmospheric pressure). The ratio PHF/
/PSiF4 increases strongly with decreasing pres-
sure, especially in the low-pressure part of the
system. The various curves that were calculated
for different temperatures between 1100 and
1400 K are close to parallel. Applying a pres-
sure of 860 bar to a system initially at atmos-
pheric pressure will shift the equilibrium by 3
log units, equivalent to a 10-fold increase in
PHF in fig. 2. Therefore, at this pressure the fig-
ure is similar to the one at 1 bar, except that the

continuous lines are valid at P*HF values that are
ten times lower than indicated at 1 bar. 

Even though the relative amount of SiF4 in-
creases at high pressure, the absolute quantities
are still very low in volcanic gases, as HF con-
centrations are generally low (PHF < 10−2) and
the total amount of SiF4 is even smaller than HF
(as PHF/PSiF4 > 10

5 at magmatic temperatures,
see fig. 2). However, if gases would be more
HF-rich, SiF4 would become a significant com-
ponent even at moderate pressures (fig. 3). For
example, if P*HF = 10−1 then SiF4 contains more
than 1% of the total amount of fluorine over the
whole temperature and pressure interval stud-
ied here, except at pressures lower than ∼100
bar, and at 1100 K it is the dominant fluorine
species at pressures higher than 500 bar. We

Fig. 3. Variation of log (PHF/PSiF4) with total pressure for various temperatures at P
*
HF = 10−2.5 and P*HF = 10−1.0,

respectively. The importance of SiF4 increases with decreasing temperature, increasing PHF, and increasing pres-
sure. The circles represent data at 1 bar, and are equivalent to the values in fig. 2 at the same temperature and P*HF .



782

Jan C.M. de Hoog, Manfred J. van Bergen and Michel H.G. Jacobs

note, however, that such high partial HF pres-
sures have never been observed in nature.

3.3. Formation of SiO2

3.3.1.  Pressure-release induced crystallisation
of SiO2

Because of the temperature and pressure de-
pendence of the SiF4/HF ratio, SiO2 can be pro-
duced in two ways: 1) by an increase of the
temperature of the system, or 2) by a decrease
of pressure. As pressure decrease is common
during magmatic degassing, whereby exsolved
vapour rises through a magma column, we will
discuss this process in more detail. 

When magma becomes volatile saturated,
e.g., in response to decompression or fractional
crystallisation, a separate fluid phase forms in
which volatile components, including HF, will
be concentrated. HF is a very reactive gas, and
it will react with the melt to form SiF4, follow-
ing reaction (2.1), and other less volatile
species, such as MgF2, CaF2 and AlF3 (White
and Hochella, 1992). This way, appreciable
amounts of silica may be entrained in the
volatile phase, the exact amount being deter-
mined by thermodynamic equilibrium and by
the solubility and concentration of fluorine in
the system. Here we assume that thermodynam-
ic properties during reaction with siliceous melt
are not significantly different from reaction
with pure SiO2 (White and Hochella, 1992; see
Methods). If the magma starts or continues to
rise, vapour pressure decreases and Reaction 1
will shift to the left resulting in precipitation
(sublimation) of silica from the vapour.

If we assume a magma which contains a
vapour phase with PHF = 10−2.6 bar, an isothermal
pressure decrease from 1000 to 1 bar at 1400 K
would cause a shift in logPSiF4 from −8.5 to 
−11.6, and would result in the production of
SiO2. However, as the total amount of SiF4 in
the system is very small, the amount of SiO2
formed during pressure release is limited. Un-
der the assumed conditions 1 mol SiO2 would
be formed from ∼794 kmol HF, or, expressed in
weight, ∼260 kg HF is needed to form 1 g sili-
ca. This amount of HF is equivalent to ∼105 000

kg magmatic vapour. A dynamic model in
which large amounts of gas percolate through a
relatively small rock volume is therefore need-
ed to deposit significant amounts of silica. In
our example from Lewotolo, a gas/rock ratio of
1300 is needed to deposit 0.2‰ silica by vol-
ume (the estimated amount of silica in some of
the samples). However, while such as scenario
might work for dome lavas, it is difficult to rec-
oncile with the presence of silica patches with-
in vesicles in lavas on the flanks of Lewotolo
volcano. If we apply the model to Mt. Etna
(Sicily, Italy), which is one of the larger gas-
producing volcanoes on Earth, with an average
HF production of 100-300 ton per day (Francis
et al., 1998; Pennisi and Le Cloarec, 1998), the
amount of silica formed during decompression
of SiF4-bearing gas could be about 1 kg per day,
depending on assumptions regarding P*HF at
depth. Clearly, these amounts are exceedingly
small.

3.3.2.  The effect of changing temperatures

In the above calculations we assumed that
the temperature of the magmatic gas phase is
constant. Because the HF/SiF4 ratio is strongly
temperature dependent, a temperature decrease
during degassing will counteract the formation
of SiO2 induced by pressure release. Thermody-
namic calculations on ascending magma-gas
mixtures (Mastin and Ghiorso, 1998) for water-
saturated magmas decompressing from 50 MPa
to 1 atm suggest that temperature changes may
vary between a 50°C decrease to an almost
10°C increase, depending on whether degassing
takes place below the fragmentation depth and
whether the magma is mafic or silicic. In the
case of passive degassing (no fragmentation)
the process is isenthalpic for silicic magmas
(temperature may increase) and isentropic for
basaltic magmas (temperature decreases).
However, the temperature effect is small com-
pared to the pressure effect (fig. 3). For exam-
ple, in order to balance PHF/PSiF4 in a 1400 K
vapour during decompression from 1 kbar to at-
mospheric pressure, a simultaneous tempera-
ture decrease of almost 600°C would be need-
ed. We conclude therefore that any tempera-



Vapour-phase crystallisation of silica from SiF4-bearing volcanic gases

tures variations during decompression and de-
gassing have an effect on the formation of SiO2
that is negligible compared to that of pressure
release. 

3.3.3.  The effect of PHF

As almost all SiF4 is converted to SiO2 dur-
ing decompression, the amount of SiO2 formed
is essentially limited by the amount of silicon
initially present in the gas phase. At partial
pressures of HF common in high-T fumaroles
the amount of SiO2 formed is extremely small.
Increasing the amount of SiF4 in the gas phase
in response to 1) a higher fluorine activity in the
vapour, or 2) lower gas temperatures would in-
crease the amount of SiO2 formed. A tempera-
ture decrease of about 350 K would result in a
ten times higher SiO2 production, which is still
rather insignificant. The effect of P*HF , howev-
er, is much stronger: an increase of P*HF from
0.001 to 0.01 results in a decrease of 3 log units
of PHF/PSiF4. In other words, while the amount
of HF increases by a factor of 10, the amount of
SiF4 increases by a factor of 10000. If we as-
sume an initial P*HF = 0.05 and T = 1250 K, the
amount of SiO2 formed during decompression
is about 2.5 mol% (or 7.5 wt%) of the amount
of HF present in the gas. The validity of such a
scenario is difficult to evaluate, as little is
known about the concentration of fluorine in
late-stage magmatic vapours at high pressures.
It is of interest to note that silica patches in the
vesicles in Lewotolo lavas are often associated
with thin whiskers of F-phlogopite that stick
out of the silica mass. This suggests a common
origin of both phases and a very high fluorine
activity during crystallisation (De Hoog and
Van Bergen, 2000). Although such F-rich com-
positions have yet to be observed in volcanic
gases, the co-existence with F-phlogopite
makes crystallisation of silica from F-rich
vapour an attractive hypothesis.

As an alternative, silica may have been de-
posited from within the lava flow, i.e., not in re-
sponse to depressurisation but as a result of the
large temperature gradients in the cooling lava.
Although in cooling gases SiF4/HF will in-
crease and therefore the lava surface may be-

come depleted in silica (White and Hochella,
1992), the heating of SiF4-rich gases is an effi-
cient way of producing silica because of the
large SiF4/HF ratios of these gases compared to
hot, pressurized magmatic vapour. For exam-
ple, SiF4/HF ratios of up to 0.2 and SiF4 fluxes
of up to 1 kton have been measured in the
plume of Popocatepetl (Love et al., 1998), and
because of the steep slope of SiF4/HF ratios
with temperature at low T (fig. 2), heating of
these gases by only 100°C would convert most
SiF4 to SiO2. In a similar manner, silica may be
deposited locally by gas circulating in cooling
lava bodies.

It is of interest to emphasize that Lewotolo
is a high-K volcano, as K-rich magmatic sys-
tems are considered to be often enriched in flu-
orine as well (e.g., Edgar et al., 1996; Strecher,
1998). Therefore, silica deposited during
vapour decompression might be relatively com-
mon in volcanoes with high-K characteristics.
We discovered silica patches (without F-phlo-
gopite) similar in texture and appearance to
those of Lewotolo, in lavas from Rinjani, a K-
rich volcano on Lombok, which supports this
supposition. Exposure to ash from eruptions of
high-K volcanoes could therefore pose an in-
creased risk for respiratory diseases, in a simi-
lar manner as eruptions from dome-forming
lavas raised concerns about adverse health ef-
fects because of their high contents of crys-
talline silica (Baxter et al., 1999). We note that
Lewotolo does not form lava domes, yet crys-
talline silica is present in many of its lavas. 

3.4. Implications for measurements of fumarole
temperatures by remote sensing

Figure 2 shows that the equilibrium temper-
ature can be calculated if PHF and PSiF4 are
known. This relation was used by Francis et al.
(1996) to remotely determine fumarole temper-
atures at Vulcano (Italy). The authors calculat-
ed that fumarole temperatures were 200-450°C,
which was in good agreement with directly
measured temperatures of 210-690°C.

Francis et al. (1996) used the thermodynam-
ic calculations of White and Hochella (1992),
which, as discussed above, are inconsistent with

783



784

Jan C.M. de Hoog, Manfred J. van Bergen and Michel H.G. Jacobs

the results of Rosenberg (1973), Symonds et al.
(1992) and the calculations presented here (fig.
2). Our thermodynamic model (cf. reaction (2.3))
predicts the temperature dependence of SiF4/HF
as follows: 

(3.1)

in which it is assumed that PH2O = 0.8 bar and
PHF, PSiF4 << PH2O , which are common condi-
tions in volcanic plumes and fumaroles.

Repeating the calculations of Francis et al.
(1996) with our thermodynamic data yields
equilibrium temperatures around 100°C, with
a maximum of ∼180°C. These values are much
lower than the findings of Francis et al. (1996)
and directly measured temperatures of 250-
600°C. In addition, the calculated range of
temperatures for the different fumaroles is
much smaller than observed. We hypothesize
that the initial match between remotely meas-
ured and observed fumarole temperatures was
fortuitous, and suggest that more work is need-
ed to confirm that the remote-sensing ap-
proach can be reliably used to estimate fuma-
role temperatures. The discrepancy might be
due to the absence of equilibrium between the
gas components (which is more likely for het-
erogeneous equilibria, in which case one of the
phases in the reaction is a solid, here silica).
Alternatively, in view of our inferred tempera-
ture of ∼100°C, the remote-sensing results
may represent the temperature of condensing
water vapour (steam), as water is the main
component of the gases. Irrespective of these
considerations, the strong influence of PHF on
calculated temperatures implies that errors in
remote-sensing thermometry will be large un-
less HF can be measured simultaneously with
SiF4. 

Acknowledgements

We thank Marie Edmonds and particularly
Mike Burton for insightful reviews, and Mike
Carroll and Roberto Moretti for editorial han-
dling of the manuscript. 

( )
.

T

P
P

logP
C

log 3 4 755

5574
273

SiF

HF
HF

4

=
+ -

-
-c

c m

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