Interactions of Fluids in Geology

Prévia do material em texto

common ground between fluid
an outlet for research that considers the interactions of
recent advances, we are delighted to have been able to bring
together some excellent and wide-ranging new science
that continues in this tradition.
The first four articles all concern our fundamental theo-
retical and experimental understanding of essentially aque-
of water-rich fluid systems and how these are affected by
solutes, while noting the remaining limitations in the
experimental database. Dolejs and Manning present the
first comprehensive study to produce a more flexible alter-
native to the HKF model for aqueous electrolytes, better
suited to the range of compositions and conditions
encountered in nature, while Sherman shows how modern
computational power means that some fundamental prob-
lems in natural fluid chemistry can be addressed from first
principles using quantum chemistry and molecular dynam-
ics. In the final article in this section, Newton and Man-
ning review recent experimental results for lower crustal
conditions and present new data to quantify the impor-
tance of dissolved salts for the solubility of the major rock
–forming elements, Si and Al, and for a range of important
Ca-minerals.
The second group of articles relate to a specific geologi-
cal setting where fluid processes are of the highest impor-
tance: sedimentary basins. Parnell provides a concise review
of the use of hydrocarbon fluid inclusions to understand
the evolution of reservoirs through time and the relation-
ships between fluid stages and mineral cements. He shows
in particular how this approach has contributed to under-
standing the oil charge history of the North Sea and
UK Atlantic margin. Hanor and Mercer describe the
behaviour of saline waters and their distribution in the
Gulf of Mexico, and show how salinity differences arising
through salt dissolution can dictate flow patterns. They
also explore the likely impacts of salt on the potential of
the region as a source of methane hydrates. The article by
Manzocchi, Childs and Walsh reviews how faults affect the
flow of fluids, in particular hydrocarbons, in siliciclastic
basins, and also comment on the extent to which current
industry practice for evaluating the effects of faults is actu-
ally grounded in science.
A third group of article deals with fluid processes in oce-
anic settings. Saffer has modelled the lateral variations
along the Nankai margin of Japan and shown that large
scale variations along strike in the taper angle of the accre-
tionary wedge can be linked back to lithological variations
from more turbidite-rich sequences to mudrocks. The
lithology affects the development of fluid overpressure and
the draining of the subduction zone fault, which in turn
influences the overall geometry of the wedge. The interplay
between permeability, heat flow and discharge characteris-
tics at mid-ocean ridges is explored by Driesner. His results
support some findings from terrestrial geothermal systems:
high temperature discharges, and the highest fluid salini-
ties, may be associated with low fluid fluxes, while large
discharges at relatively low temperatures may in fact domi-
nate the removal of heat. Fisher and Harris take three spe-
cific examples of mid-ocean ridge settings to explore the
controls on heat loss. The relative importance of conduc-
tive heat loss is variable, and specific features of the base-
ment geology can serve to target fluid flow and hence heat
loss. Hydrothermal vents are also of likely significance for
both abiotic and metabolic organosynthesis and this is
explored by Shock and Canovas. Different patterns of mix-
ing of seawater with different hydrothermal fluids can lead
to different evolutionary paths, but in general, the mixing
favours formation of organic compounds from inorganic
reactants. Hence, microbes could produce components of
biomolecules simply by catalysis of reactions that are
already energetically favoured.
A fourth group of articles deals with the continental
crust. Ingebritsen and Manning present a crustal-scale
overview of permeability and argue that while there is a
power–law relation between permeability and depth in tec-
tonically active continental crust, some regions exhibit
markedly higher permeabilities, probably as transients,
while stable crust may decay to lower permeability. The
specific issue of the relationship of hydrologic response to
Frontiers in Geoflui , 1st edition. Edited by Bruce Yardley, Craig Manning and Grant Garven. © 2011 by Blackwell Publishing Ltd. ds
INTRODU
c
CTION
Frontiers in geofluids: introdu tion
to provide a comprehensive coverage of all the important
ous fluids. Liebscher provides an overview of the properties
This set of papers was originally published electronically
anniversary of the launch of the journal. For this volume, 
we sought to bring together a collection of papers span- 
ning a range of topics to which the role of fluids in the
chemical and physical processes. While we cannot pretend
Earth is central. Geofluids was founded to help empha-
sise the processes that
take place in different geological settings, and to provide
earthquake activity is discussed by Wang and Manga. They
as a special double issue of Geofluids to mark the tenth
demonstrate that, in the intermediate and far-field, changes
in groundwater flow are linked to changes in permeability
which arise in response to cyclic deformation and oscilla-
tory flow. The relationships between faulting and flow at
depth is explored by Cox, who shows how fluid pressure
and stress influence failure modes and hence the styles of
permeability enhancement and vein development, in both
mineralized and unmineralized systems. Fluid flow coupled
to deformation often introduces water into high grade
crystalline basement rocks which undergo retrogression.
Yardley, Harlov and Heinrich present the results of experi-
ments designed to measure the rate at which high grade
rocks undergo retrogression under mid- to lower-crustal
conditions, and conclude that water infiltrated along fine
cracks is likely to be rapidly consumed. The article by
Bucher and Stober addresses deep groundwaters found in
crystalline basement rocks today by tunnelling and drilling.
They argue that in areas of high relief such as the Alps,
such waters are of relatively low TDS because of flushing
by meteoric water, whereas much more saline brines may
evolve where the hydraulic gradients are less. Migration of
fluids can lead to mineralogical and chemical changes
(metasomatism) in a wide variety of crustal settings, and
Putnis and Austreheim explore some diverse examples of
metasomatism on a range of scales. They are able to dem-
onstrate that, while aqueous fluids partly act as a catalyst to
permit minerals to react, they can also influence the course
of the reaction through a thermodynamic role.
The final section comprises two interdisciplinary articles
that deal with ore deposits and draw on a range of aspects
of fluids. Boiron, Cathelineau and Richard review the
fluid systems that give rise to ore deposits near the uncon-
formity between sedimentary basins and their underlying
crystalline basement. They contrast the Proterozoic uncon-
formity uranium deposits with younger base metal deposits
that develop in similar settings, and conclude that there are
many similarities in both the nature of the fluids and the
flow patterns that give rise to mineralization. Kamenetsky
and Kamenetsky evaluate fluid processes at the other
temperature extreme of ore formation, associated with
magmatism. They present evidence from inclusions to doc-
ument the development of immiscibilityas magmas cool,
and evaluate the importance of immiscibility for magma
chamber processes, including degassing and the partition-
ing of metals.
Although we have grouped the articles for convenience,
we believe that the true value of the collection arises from
the basic new data presented, from the insight into the
interactions between physical and chemical processes, and
from the opportunity they provide to take ideas developed
in a field where particular types of observation or measure-
ment may be possible to understand processes in different
settings or at different times, where different types of data
may be available.
G. GARVEN1, C. E. MANNING 2 and
B. W. D. YARDLEY3
1Department of Geology, Tufts University, Medford, MA,
USA; 2Department of Earth and Space Sciences,
University of California, Los Angeles, CA, USA;
3School of Earth and Environment, University of Leeds,
Leeds, UK
Corresponding author: B. W. D. Yardley
School of Earth and Environment, University of Leeds,
Leeds LS2 9JT, UK.
Email: B. W. D. Yardley@leeds.ac.uk.
Tel: +44 113 343 5227. Fax: +44 113 343 5259.
2 C. Garven et al.
REVIEW
Aqueous fluids at elevated pressure and temperature
A. LIEBSCHER
Centre for CO2 Storage, Helmholtz Centre Potsdam, German Research Centre for Geosciences GFZ, Telegrafenberg, Potsdam,
Germany
ABSTRACT
The general major component composition of aqueous fluids at elevated pressure and temperature conditions can
be represented by H2O, different non-polar gases like CO2 and different dissolved metal halides like NaCl or
CaCl2. At high pressure, the mutual solubility of H2O and silicate melts increases and also silicates may form
essential components of aqueous fluids. Given the huge range of P–T–x regimes in crust and mantle, aqueous flu-
ids at elevated pressure and temperature are highly variable in composition and exhibit specific physicochemical
properties. This paper reviews principal phase relations in one- and two-component fluid systems, phase relations
and properties of binary and ternary fluid systems, properties of pure H2O at elevated P–T conditions, and aque-
ous fluids in H2O–silicate systems at high pressure and temperature. At metamorphic conditions, even the physi-
cochemical properties of pure water substantially differ from those at ambient conditions. Under typical mid- to
lower-crustal metamorphic conditions, the density of pure H2O is qH2O ¼ 0:6�1:0 g cm�3, the ion product
Kw = 10
)7.5 to approximately 10)12.5, the dielectric constant e = 8–25, and the viscosity g = 0.0001–
0.0002 Pa sec compared to qH2O ¼ 1:0 g cm�3, Kw = 10)14, e = 78 and g = 0.001 Pa sec at ambient conditions.
Adding dissolved metal halides and non-polar gases to H2O significantly enlarges the pressure–temperature
range, where different aqueous fluids may co-exist and leads to potential two-phase fluid conditions under must
mid- to lower-crustal P–T conditions. As a result of the increased mutual solubility between aqueous fluids and
silicate melts at high pressure, the differences between fluid and melt vanishes and the distinction between fluid
and melt becomes obsolete. Both are completely miscible at pressures above the respective critical curve giving
rise to so-called supercritical fluids. These supercritical fluids combine comparably low viscosity with high solute
contents and are very effective metasomatising agents within the mantle wedge above subduction zones.
Key words: fluid–fluid interactions, fluid phase relations, fluid properties, fluid systems, metamorphic fluids,
supercritical melts ⁄ fluids
Received 23 July 2009; accepted 14 March 2010
Corresponding author: Axel Liebscher, Centre for CO2 Storage, Helmholtz Centre Potsdam, German Research
Centre for Geosciences GFZ, Telegrafenberg, D-14473 Potsdam, Germany.
Email: alieb@gfz-potsdam.de. Tel: +49 (0)331 288 1553. Fax: +49 (0)331 288 1502.
Geofluids (2010) 10, 3–19
INTRODUCTION
Aqueous fluids play a fundamental role in the geochemical
evolution of the Earth. By transport and recycling of vola-
tile components and different solutes from the atmosphere
and hydrosphere through the solid interior of the Earth
and back to the surface they chemically link the different
spheres of the Earth on the local, regional and also global
scale. As a mobile phase, they can transport heat very effi-
ciently by convection and contribute to the local and
regional heat budget and heat distribution. The highly var-
iable P–T–x regimes at the surface, in the crust and in the
mantle generate equally variable, characteristic fluids with
specific compositions and specific physical properties like
density and viscosity. Aqueous fluids at elevated pressure
and temperature conditions may form by infiltration of
meteoric waters in geothermal and basinal systems (e.g.,
Hanor 1994; Arno´rsson et al. 2007), by diagenetic reac-
Frontiers in Geoflui , 1st edition. Edited by Bruce Yardley, Craig Manning and Grant Garven. © 2011 by Blackwell Publishing Ltd. ds
tions with connate fluids (Hanor 1994), by infiltration of
seawater in oceanic hydrothermal systems (e.g., German &
Von Damm 2003; Foustoukos & Seyfried 2007), by pro-
grade metamorphic dehydration and decarbonation reac-
tions as response to changes in P–T–x conditions (e.g.,
Yardley & Graham 2002; W. Heinrich 2007), and by liber-
ation from crystallizing magmas (e.g., Cline & Bodnar
1994; C.A. Heinrich 2007). Due to the different fluid
sources and different geologic environments with notably
different P–T–x regimes, the composition of these aqueous
fluids is highly variable. Meteoric fluids are almost pure
water, which may gain considerable amounts of dissolved
salts during diagenetic reactions. The salinity of aqueous
basinal fluids ranges over five orders of magnitude from a
few ppm in shallow meteoric regimes to more than approx-
imately 40 wt% in evaporite-rich basins (Hanor 1994). In
oceanic hydrothermal systems, the dominant fluid source is
seawater, which can be modelled as a 3.3 wt% aqueous
NaCl equivalent solution (NaCleq; Bischoff & Rosenbauer
1984). However, by interaction with the oceanic crust, the
salinity of the generated hydrothermal fluids may range
from only approximately 0.2 up to approximately 7.3 wt%
NaCl (e.g., Oosting & Von Damm 1996; Lu¨ders et al.
2002). Prograde metamorphic dehydration reactions form
highly variable saline fluids that may even reach salt satura-
tion (Fig. 1B; Yardley & Graham 2002). However, the
salinity of metamorphic fluids depends only little on meta-
morphic grade but is strongly linked to the protolith’s ori-
ginal setting. Metamorphism of oceanic or accretionary
prism protoliths generally forms fluids with salinity below
approximately 6 wt% NaCleq, whereas metamorphism of
rocks from shallow marine or continental margin origin
forms fluids that span the complete range from almost salt
free up to approximately 60 wt% NaCleq (Fig. 1B; most
salinity data for metamorphic fluids come from fluid inclu-
sions and are estimated from the melting temperature of
ice. They are usually given as ‘equivalent’ concentration of
NaCl or NaCleq, referring to the salinity of a NaCl solution
that would yield the same melting temperature of ice as
the measured fluid inclusion). Metamorphic decarbonation
reactions, triggerd by changes in pressure and temperature
or by an infiltrating fluid itself, potentially supply CO2 to
the system, force the fluids to un-mix and generate aquo-
carbonic fluids co-existing with variable saline H2O–salt
fluids (e.g., Trommsdorff et al. 1985; Skippen & Trom-
msdorff 1986; Trommsdorff & Skippen 1986; W. Heinrich
1993, 2007). H2O–salt fluids are also reported from fluid
inclusions in diamonds from the uppermantle (Navon
et al. 1988; Izraeli et al. 2001). But at high pressure, also
the solubility of silicates in aqueous fluids notably increases
and silicates may form important solutes at these condi-
tions. This may even lead to complete miscibility between
aqueous fluids and silicate melts (e.g., Shen & Keppler
1997; Hack et al. 2007) and individual components
continuously change their character from low, typical
solute-like concentrations to high, typical solvent-like
concentrations; finally, any major element may become an
essential phase component of the fluid.
This review aims at presenting some of the fundamen-
tal topics of aqueous fluids at elevated pressure and tem-
perature. Given the vast number of different geological
environments in which fluids evolve, each of which is
unique in space, time and physicochemical properties,
such a review is necessarily incomplete and the reader is
kindly referred to the references given for more and
detailed information on specific aspects of aqueous fluids.
The review first describes the principal phase relations
in one- and two-component fluid systems. Then, phase
Volcanic systems
Meteoric
systems
Mantle
Mid ocean
ridge
Hydrated
 oceanic crust
Oceanic
systems
Continuous
 dehydration
 of
 
Unde
rplatin
g
High-P metamorphism
subducting slab
Mantle
melting
Magmatic
hydrothermal
systems
Metamorphic systems
Basinal
 systems
Mantle
Geothermal
systems
Te
m
pe
ra
tu
re
 (°
C)
200
300
400
500
600
700
0 605040302010
Ap
pro
x.
sa
lt s
atu
rat
ion
Salinity (eq. wt% NaCl)
0
100 Basinal
Magmatic
Oceanic
Metamorphic
(A) (B)
Fig. 1. (A) Schematic drawing showing the different geological fluid systems in crust and mantle. Modified from Liebscher & Heinrich (2007). (B) Compilation
of salinity data for metamorphic fluids. Dark grey fields represent protoliths of shallow marine or continental margin origin, white boxes those from oceanic
or accretionary prism origin. Hatched boxes indicate high-pressure metamorphic rocks. Redrawn and modified from Yardley & Graham (2002). Lighter grey
fields indicate compositional ranges for fluids from magmatic, metamorphic, oceanic and basinal systems as compiled by Kesler (2005).
4 A. LIEBSCHER
relations and properties of binary and ternary fluid sys-
tems, mostly based on experimental studies, are summa-
rized followed by a review of the properties of pure H2O
at elevated P–T conditions. Finally, aqueous fluids in
H2O–silicate systems at high pressure and temperature
are discussed.
PRINCIPAL PHASE RELATIONS IN ONE- AND
TWO-COMPONENT FLUID SYSTEMS
Aqueous geological fluids are rarely pure water at elevated
P–T conditions and represent mixtures with several addi-
tional major components like salts (here used in the
restricted sense of dissolved metal halides), non-polar gases
and rock components like silica. In these multicomponent
systems, phase relations become more complicated, fluid
miscibility and immiscibility play an important role and
concentrations of individual components may continuously
change from solute-like to solvent-like or vice versa. In the
following, the principal phase relations in two-component
model fluid systems are described starting from the one-
component case as exemplified for H2O (Fig. 2A).
In any one-component system, the three phase states
solid (in case of H2O ‘ice’), liquid and vapour (in case of
H2O ‘steam’) have identical and fixed composition. All
three co-exist at the invariant triple point from which the
three univariant solid–liquid, solid–vapour and liquid–
vapour equilibria emanate (Fig. 2A). In case of H2O, the
triple point is at 0.00061 MPa ⁄0.01�C. In one-component
systems, co-existence of different fluid phases is exclusively
restricted to the liquid–vapour equilibrium. However,
although identical in composition, liquid and vapour differ
in their physical properties like density, viscosity and elec-
tric permittivity. These properties show abrupt, discontinu-
ous changes at the first-order liquid–vapour phase
transition along the liquid–vapour equilibrium. The differ-
ences in physical properties, however, diminish along the
liquid–vapour equilibrium towards higher temperature and
pressure and finally disappear at the critical point. The criti-
cal point of H2O is at 373.95�C ⁄22.06 MPa (Wagner &
Pruß 2002). Above critical temperature and pressure, nei-
ther changes in temperature nor pressure induce any phase
transition and the physical properties of the homogeneous
single phase fluid continuously change in response to
changes in P and T.
Addition of a second component to a one-component
system adds composition as an additional degree of free-
dom and co-existing fluids not only differ in their physical
properties but also in composition. Consequently, the
invariant triple point, the univariant liquid–vapour equilib-
rium and the critical point of the one-component system
turn into a univariant solid–liquid–vapour equilibrium, a
divariant liquid–vapour field and a critical curve in the two-
component system. The P–T range in which two fluids co-
exist may thereby greatly expand (Fig. 2B). The principal
phase relations of two-component fluid systems and how
they apply to different H2O–salt systems are shown in
Fig. 3. For an in-depth presentation and discussion of the
different phase topologies the reader is referred to Ravich
(1974) and Valyashko (1990, 2004) and references given
therein. A discussion of the principal phase topologies in
the context of liquid–liquid immiscibility and the mag-
matic-hydrothermal transition is given by Veksler (2004).
The overall phase topology of H2O–salt systems is deter-
mined by the pressure and temperature conditions of the
univariant salt–liquid–vapour equilibrium, which emanates
from the triple point of the salt endmember systems, rela-
tive to the pressure–temperature conditions of the critical
curve. In so-called Type 1-systems, with a high and pro-
grade solubility of the salt in H2O, the univariant salt–
liquid–vapour equilibrium does not intersect the critical
curve, which is continuous over the entire P–T–x space
and connects the critical points of the pure systems
(Fig. 3A, B). Between the univariant salt–liquid–vapour
Pr
es
su
re
 (M
Pa
)
Temperature (°C)
Liquid
Critical point
Triple points 
+ v
s
 +
 l
374
22.1
0.00061
0.01
CO2 CH4
0
50
100
Temperature (°C)
Pr
es
su
re
 (M
Pa
)
Critical point
H2O
Liquid-vap
our 
H2O
 
NaCl
CaCl2
200 500400300
Critical isochore
(0.322 g cm–3)
H2O–
H2O–
(A) (B)
Fig. 2. (A) Phase diagram for pure H2O showing the stability fields of the different phase states of H2O. Co-existing fluid phases are restricted to the
liquid–vapour (‘l + v’) curve. Above the critical point, only one homogeneous single phase fluid exists. (B) P–T projection of the critical curves for the systems
H2O–CO2, H2O–CH4, H2O–CaCl2 and H2O–NaCl. Hatched sides for of the critical curves mark direction of fluid un-mixing. Based on data by Takenouchi &
Kennedy (1964) for H2O–CO2, Welsch (1973) for H2O–CH4, Shmulovich et al. (1995), Bischoff et al. (1996) and Zhang & Frantz (1989) for H2O–CaCl2,
and Driesner & Heinrich (2007) for H2O–NaCl. Liquid–vapour curve for H2O from Wagner & Pruß (2002).
Aqueous fluids at elevated pressure and temperature 5
Pr
es
su
re
Composition
Tempe
rature
H2O
Salt
v + Saltv
l + SaltSteam
Ice
c.p.
(H2O) c.p.
(salt)
t.p. (salt)
Critical curve
t.p.
(H2O)
Ice
Water
Steam
Water
v + l
Pr
es
su
reComposition
Tempe
rature
H2O
Salt
v + Salt
v
l + Salt
Steam
Ice
c.p.
(H2O) c.p.(salt)
t.p. (salt)
Critical curve
t.p.
(H2O)
Ice
Water
Steam
Water
v + l
Lower
critical
endpoint
Upper
critical
endpoint
c.p.
c.p.
t.p.
Cri
tic
al 
cur
ve v
 =
 l
t.p.Ice
Liquid
Steam
Solid salt
Liquid
Vapour
Pr
es
su
re
Temperature
c.p.
c.p.
t.p.
Critical
 
 
 
 
 
 Curve
 v
 =
 l
t.p.
Ice
Liquid
Steam
Solid salt
Liquid
Vapour
Critical
endpoints
Lower
Upper
Pr
es
su
re
Temperature
Type 1a
Type 2a
salt = NaCl, KCl, LiCl, MgCl2, CaCl2, MnCl2, NaOH, K2CO3 
Pr
es
su
re
Composition
Tempe
rature
H2O
Salt
v + Salt
v
l + Salt
Steam
Ice
c.p.
(H2O) c.p.(salt)
t.p. (salt)
Critical curve
t.p.
(H2O)
Ice
Water
Steam
Water
v + l
v + lv + l
l1 + l2
l1 + l2
v + l1
N
R c.p.
c.p.
t.p.
Cri
tical
 
curve l1 = l2
t.p.
Ice
Liquid
Steam
Solid salt
Liquid
Vapour
Pr
es
su
re
Temperature
N
R
v +
 
l1 + 
l2
v = l1
Type 1d
Salt = KH2PO4, Na2B4O7, Na2HPO4, UO2SO4 
Salt = Al2O3, SiO2, BaCl2, BaSO4, K2SO4, Na2SO4, Na2CO3 
v + l
v + l
v + l
Ice +
v + l
salt +
v + l
Salt +
v + l
I
H2O Salt
Steam
Ice
Ice
Water
Ice + l
v
Ice
+ v
v + Salt
v + l
l + Salt
l
H2O Salt
steam
water
v
v + Salt
v + l
l + Salt
l
H2O Salt
v
v + Salt
v + l
l + Saltl
Single phase
fluid
Pr
es
su
re
Pr
es
su
re
Pr
es
su
re
TI < T1 < Tt.p. (H2O) Tt.p. (H2O) < T2 < Tc.p. (H2O) Tc.p. (H2O) < T3 < Tt.p. (salt)
T1
T3T2
Salt + l + v
Salt + l + v
salt
 +
 l
 +
 v
Sal
t +
 
l +
 
v
(A)
(B)
(C)
Fig. 3. Schematic drawings of phase relations in water–salt systems of (A) Type 1a, (B) Type 1d and (C) Type 2a as P–T–x presentation and projections onto the
P–T and P–x planes. Assignment of different salts to the different systems after Valyashko (2004). Drawn and modified according to Valyashko (1990, 2004).
6 A. LIEBSCHER
equilibrium and the critical curve a two-phase fluid
volume forms in P–T–x space. The univariant salt–liquid–
vapour equilibrium may intersect the univariant ice–liquid–
vapour equilibrium, which emanates from the triple point
of pure H2O, forming an invariant point where liquid,
vapour, ice, and salt co-exist (Fig. 3A). In so-called Type
2-systems, with a low solubility of the salt in H2O, the
univariant salt–liquid–vapour equilibrium intersects the
critical curve at the lower and upper critical endpoints
(Fig. 3C). These general Type 1 and 2 phase topologies
may get more complicated due to liquid–liquid immiscibil-
ity. Such liquid–liquid immiscibility gives rise to a number
of principal phase topologies, which are presented and dis-
cussed in detail by Valyashko (1990, 2004). Not all of
these principal phase topologies, however, have yet been
found or are relevant for geological fluid systems. The
most simple Type 1a system represents the binaries
between H2O and the highly soluble salts like NaCl (e.g.,
Sourirajan & Kennedy 1962), KCl (Tkachenko & Shmulo-
vich 1992; Dubois et al. 1994), CaCl2 (Tkachenko &
Shmulovich 1992; Bischoff et al. 1996) but also NaOH
(Urusova 1974) (Fig. 3A). Here, the critical curve is con-
tinuous over the entire P–T–x space and no liquid–liquid
immiscibility occurs. The effects of liquid–liquid immisci-
bility are exemplified by Type 1d (Fig. 3B). Here, two sep-
arate critical curves form. One critical curve emanates from
the critical point of the more volatile component (H2O in
Fig. 3B) and shows critical behaviour between vapour and
liquid 1, the other critical curve emanates from the critical
point of the less volatile component (‘salt’ in Fig. 3B) and
shows critical behaviour between liquid 1 and 2. H2O–salt
systems of Type 1d include exotic salts like KH2PO4 (Mar-
shall et al. 1981) and are not relevant for geological fluids.
The simple Type 2a topology without liquid–liquid immis-
cibility is representative for binaries between H2O and
sparingly soluble second components like SiO2, Al2O3 and
other silicates, different sulphates and carbonates (Fig. 3C).
Owing to the intersection of the univariant liquid–vapour–
salt equilibrium with the critical curve at the lower and
upper critical endpoints, lower and upper segments of the
two-phase fluid volume form. In systems in which silicates
are the additional second component, the high tempera-
ture part of the liquid–vapour–salt equilibrium is identical
to the wet solidus and describes the silicate’s water satu-
rated melting behaviour. As the upper liquid–vapour–salt
equilibrium or wet solidus terminates at the upper critical
endpoint, no discrete melting reaction is possible at pres-
sure above the upper critical endpoint (see below).
BINARY AND TERNARY FLUID SYSTEMS
The phase relations in binary and ternary fluid systems are
simplified representations of natural fluids but form the
framework for the more realistic higher component sys-
tems. The most important binary systems for aqueous flu-
ids are the H2O–non-polar gas systems H2O–CO2 and
H2O–CH4 and the H2O–salt systems H2O–NaCl,
H2O–CaCl2 and H2O–KCl. Other components like the
non-polar gas N2 and salts like LiCl, SrCl2 or MgCl2 are
normally present only at the minor or even trace element
level. Combining these binary systems yields the important
ternary fluid systems H2O–NaCl–CO2, H2O–CaCl2–CO2
and H2O–NaCl–CH4.
The principal phase relations in the systems H2O–CO2
and H2O–CH4 resemble each other and can be taken as
proxies for other H2O–non-polar gas systems (Fig. 4A, B).
Starting from the critical point of pure water, their critical
curves initially extend towards lower temperature with only
minor pressure dependence, then pass through temperature
minima and extend to high pressure with only minor
temperature increase. The temperature minima occur at
155–190 MPa and approximately 265�C in the H2O–CO2
system (Takenouchi & Kennedy 1964) and approximately
100 MPa and 353�C in the H2O–CH4 system (Welsch
1973). At temperatures below the critical curves, both
systems show pronounced opening of the two-phase fluid
region towards lower temperature, reflecting the very low
solubility of non-polar solutes in water at low tempera-
tures. A detailed and thorough discussion of phase rela-
tions within the H2O–CO2 system that also treats the
important formation of clathrates at low temperatures is
given by Diamond (2001).
Most aqueous fluids contain important amounts of dis-
solved salts (see above), and water–salt systems therefore
have attracted much interest. The H2O–NaCl system has
been studied among others by Keevil (1942), Sourirajan &
Kennedy (1962), Khaibullin & Borisov (1966), Urusova
(1975), Bischoff & Rosenbauer (1984), Bodnar et al.
(1985), Chou (1987), Bischoff et al. (1986), Rosenbauer
& Bischoff (1987), Bischoff & Rosenbauer (1988), Bisc-
hoff & Pitzer (1989), Bischoff (1991) and Shmulovich
et al. (1995). Based on the available experimental data,
Driesner & Heinrich (2007) derived the most recent repre-
sentation of the phase topology in the H2O–NaCl system
up to 1000�C ⁄220 MPa, and xNaCl = 0–1.0, where x is
mole fraction. Their data form the basis for the H2O–NaCl
phase relations shown in Fig. 4C. The critical curve in the
H2O–NaCl system monotonously extends from the critical
point of pure H2O towards higher pressure and salinity
with increasing temperature, at least within reasonable
crustal P–T conditions. Beside salinity also the density
increases along thecritical curve with increasing pressure
and temperature (Fig. 4D; Urusova 1975; Chou 1987;
Bischoff 1991). The data show that at elevated P–T condi-
tions even the vapour in H2O–salt systems is considerably
dense and notably differ from vapour at ambient condi-
tions. For the role of dense vapour as an ore-forming fluid
of its own the reader is referred to C.A. Heinrich (2007;
Aqueous fluids at elevated pressure and temperature 7
and references given therein). The system H2O–KCl has
been studied by Keevil (1942), Hovey et al. (1990),
Dubois et al. (1994) and Shmulovich et al. (1995). The
critical curve in the H2O–KCl system as well as the corre-
sponding phase relations closely resemble those of the
H2O–NaCl system and the phase relations of natural KCl
dominated fluids may be approximated by the H2O–NaCl
system (Fig. 4E). However, the liquid–vapour–salt equilib-
rium is at slightly lower pressure in the H2O–KCl than in
the H2O–NaCl system. H2O–alkaline earth salt systems
were studied among others by Zhang & Frantz (1989),
Tkachenko & Shmulovich (1992), Shmulovich et al.
(1995), and Bischoff et al. (1996) for H2O–CaCl2 and
Urusova & Valyashko (1983, 1984) and Shmulovich et al.
(1995) for H2O–MgCl2. The data indicate that the critical
curves in H2O-alkaline earth salt systems are at higher
pressure than in the H2O-alkaline salt systems, whereas the
liquid–vapour–salt equilibrium is at lower pressure. In
H2O-alkaline earth salt systems the P–T range where a low
salinity vapour co-exists with high salinity liquid therefore
expands compared to H2O-alkaline salt systems.
Fluid un-mixing into low salinity vapour and high salin-
ity liquid in water–salt systems not only influences salinity
and density of the co-existing fluid phases but may also
change the availability of ligands and the acid or basic char-
acter of the co-existing fluids. In H2O–salt systems hydro-
lysis occurs according to the equilibrium
xH2OþMxþCl�x ¼ xHClþMxþðOHÞ�x :
At ambient conditions, the equilibrium constant is such
that reactant activities are significantly greater than those
of the products and the amounts of HCl and MxþðOHÞ�x
in the fluid are negligible. At elevated P–T conditions,
however, the equilibrium constant may change in such a
H2O–CO2
0 0.2 0.4 0.6 0.8 1
XCH4
0
50
100
150
200
250
300
Pr
es
su
re
 (M
Pa
)
H2O–CH4
36
0
35
3
35
0
33
0
30
0
28
0
350
330
280
360
35
3
Critical
curve
0 0.2 0.4 0.6 0.8 1
XCO2
Critical
curve
0
50
100
150
200
250
300
Pr
es
su
re
 (M
Pa
)
325 300
275
270 250 200200
250 265
265
264 264
0
10
20
30
40
50
60
70
80
90
0 0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6
Pr
es
su
re
 (M
Pa
)
Density (g cm–3)
Critical curve
c.p. H2O
c.p. H2O c.p. H2O
H2O–NaCl H2O–NaCl
CaCl2, Ca2+, Cl– (eq per kg)
Pr
es
su
re
 (M
Pa
)
CaCl2
Cl
–Ca
2+
0
10
20
30
40
50
60
70
80
90
0.000001 0.0001 0.01 1 100
500°C
Onset
of hydrolysis
LiquidVa
po
ur
0
20
40
60
80
100
120
140
0.000001 0.0001 0.01 1
550
700
650
600
400
450
500
350
XNaCl
Pr
es
su
re
 (M
Pa
)
Critical curve
c.p. H2O
H2O–CaCl2
500
430
475
400
550
350
310
Ion imbalanceNaCl
KClCaCl20
50
100
150
200 300 400 500 600 700 800
Temperature (°C)
Pr
es
su
re
 (M
Pa
)
c.p. H2O
200
NaCl
KCl
CaCl2
(A) (B)
(C) (D)
(E) (F)
Fig. 4. Phase relations in the binary systems
(A) H2O–CO2, (B) H2O–CH4, and (C) H2O–
NaCl; (D) densities in the system H2O–NaCl;
(E) P–T projections of the critical curves and
liquid + vapour + salt equilibria in the systems
H2O–NaCl, H2O–CaCl2 and H2O–KCl showing
the limits for co-existing fluids; and (F) hydrolysis
reaction in the system H2O–CaCl2. Thin lines in
(A), (B), (C), and (E) are isotherms with tempera-
ture given in �C. Based on data in (A) from
Takenouchi & Kennedy (1964), in (B) from
Welsch (1973), in (C) from Shmulovich et al.
(1995) and Driesner & Heinrich (2007), in (D)
from Urusova (1975), Chou (1987) and Bischoff
(1991), in (E) from Keevil (1942), Sourirajan &
Kennedy (1962), Ketsko et al. (1984), Chou
(1987), Zhang & Frantz (1989), Hovey et al.
(1990), Shmulovich et al. (1995), Bischoff et al.
(1996), and in (F) from Bischoff et al. (1996).
8 A. LIEBSCHER
way that product activities are significantly greater than
those of the reactants leading to notable amounts of
HCl and MxþðOHÞ�x . In case of co-existing fluids, HCl
preferentially fractionates into the vapour whereas
MxþðOHÞ�x fractionates into the liquid. This gives rise to
an HCl-enriched, potentially acidic vapour and a co-exist-
ing MxþðOHÞ�x -enriched, potentially basic liquid. This
effect of fluid un-mixing has been experimentally studied
by Bischoff et al. (1996), Vakulenko et al. (1989) and
Shmulovich et al. (2002). Bischoff et al. (1996) analysed
experimentally phase separated fluids in the system H2O–
CaCl2 for Ca and Cl. Below 25.0 MPa at 400�C and
58.0 MPa at 500�C they observed an increasing ion imbal-
ance between Ca2+ and Cl) in the vapour with Cl) concen-
trations being notably higher than Ca2+ concentrations
(Fig. 4F). This observation reflects formation and preferen-
tial fractionation of HCl into the vapour. The amount of
HCl produced by this reaction is remarkable and may
reach 0.1 mol kg)1 in the vapour.
The phase relations of the two binary H2O–non-polar
gas and H2O–salt subsystems form the framework for the
phase relations within the ternary H2O–salt–non-polar gas
systems. The third binary subsystem salt–non-polar gas
can safely be assumed as immiscible at all geologic rele-
vant P–T conditions. All ternary H2O–salt–non-polar gas
systems share some common features. Increasing salt con-
centrations decrease the solubility of the non-polar gas in
H2O–salt mixtures (‘salting out’; Fig. 5A) and salt concen-
trations in H2O–non-polar gas mixtures are generally low.
The P–T–x range of fluid immiscibility therefore greatly
expands and extends to conditions where the binary sub-
systems may already be completely miscible. Fluid immisci-
bility may therefore prevail over the entire crustal P–T
range (Heinrich et al. 2004). The H2O–NaCl–CO2 system
is the most relevant ternary system and has been studied,
among others, by Takenouchi & Kennedy (1965), Gehrig
et al. (1986), Kotel’nikov & Kotel’nikova (1991), Joyce &
Holloway (1993), Shmulovich & Plyasunova (1993),
Gibert et al. (1998), Shmulovich & Graham (1999, 2004),
Schmidt & Bodnar (2000) and Anovitz et al. (2004). The
H2O–CaCl2–CO2 ternary has been experimentally deter-
mined by Zhang & Frantz (1989), Plyasunova & Shmulo-
vich (1991), Shmulovich & Plyasunova (1993), and
Shmulovich & Graham (2004) and the H2O–NaCl–CH4
ternary by Krader (1985) and Lamb et al. (1996, 2002).
500°C/100 MPa
H2O 0 0.40.2
0
0.2
CO2
NaCl, CaCl2Salt saturation
in H2O–NaCl–CO2
xCO2
x N
aC
l; 
x C
aC
l 2
Vapour limb
Li
qu
id 
lim
b
Two-phase
fluid
Single-phase fluid
H2O–NaCl–CO2 H2O–CaCl2–CO2
l +
 sa
lt
l + v + Salt v + Salt
(C)
0
20
40
60
80
100
120
140
160
0 0.04 0.08 0.12 0.16
XCO2
Pr
es
su
re
 (M
Pa
)
nNaCl/(nNaCl + nH2O)
150°C
250°C
= 0.0715 = 0.0
 Increasing salt(A)
H2O
NaCl
CH4
Te
m
pe
ra
tu
re
400°C
450°C
500°C
550°C
350°C
50 MPa
X
NaCl
XCH4
l + v
l + v
Critical point
H2O-CH4
Critical point
H2O-NaCl
Single-phase fluid
Liquid + vapour
(D)
500°C/50 MPa
H2O CO2
XNaClXCO2
0 0.60.40.2 1.00.8
0
0.6
0.4
0.2
0
0.2
l
v
l +
 v
 +
 N
aC
l
Single-phase
Liquid or vapour
l + v
v + NaCl
l +
 N
aC
l
(B)
Fig. 5. (A) Salting-out effect in the system H2O–NaCl–CO2 at 150 and
250�C, and phase relations in the ternary systems (B) H2O–NaCl–CO2, (C)
H2O–NaCl–CO2 and H2O–CaCl2–CO2 and (D) H2O–NaCl–CH4. Thin lines
in (B) and dashed lines in (C) are only schematic to indicate principal phase
relations. Data in (A) from Takenouchi & Kennedy (1964, 1965), in (B)
from Anovitz et al. (2004), and in (C) from Kotel’nikov & Kotel’nikova
(1991) and Zhang & Frantz (1989). (D) Redrawn and modified after Krader
(1985).
Aqueous fluids at elevated pressure and temperature 9
The principal phase relations that occur in the H2O–NaCl–
CO2 but also in the other H2O–salt–non-polar gas systems
are exemplified in Fig. 5B based on the data by Anovitz
et al. (2004) for 500�C ⁄50 MPa. At these conditions, the
H2O–NaCl binary shows liquid–vapour immiscibility and
six principal stability fields can be distinguished: (i) At low
NaCl concentrations, single phase H2O–CO2 vapour is sta-
ble; (ii) at H2O rich compositions, the NaCl poor H2O–
CO2 vapour co-exists with a CO2 poor H2O–NaCl liquid,
defining a vapour + liquid two-phase field; (iii) at higher
NaCl concentrations, a single phase field of H2O–NaCl
liquid with only minor amounts of CO2 forms; (iv)–(vi) at
high NaCl concentrations the fluids are salt saturated and,
depending on the H2O ⁄CO2 ratio, liquid + salt, liquid +
vapour + salt or vapour + salt co-exist. With increasing
pressure, the miscibility gap within the H2O–NaCl binary
closes and the single phase fluid field extend from salt-satu-
rated CO2–poor H2O–NaCl fluids through ternary fluids
at H2O-dominated conditions to NaCl poor H2O–CO2
fluids (Fig. 5C). The available data clearly show that the
solubility of NaCl in H2O–CO2 vapour is very low and
substantially decreases with increasing pressure and temper-
ature. For instance, at 500�C, e.g., xNaCl in H2O–CO2
vapour decreases from about 0.074 at 50 MPa to <0.025
at 500 MPa (Shmulovich & Plyasunova 1993; Anovitz
et al. 2004), whereas at 200 MPa a temperature increase
from 600 to 800�C decreases xNaCl from about 0.052 to
<0.032 (Kotel’nikov & Kotel’nikova 1991). In the ternary
systems H2O–NaCl–CH4 and H2O–CaCl2–CO2 the solu-
bility of the salt in H2O–non-polar gas vapour is even less
and the field of co-existing fluids greatly expands (Fig. 5C;
see also Liebscher 2007). The progression of the phase
relations in ternary systems as function of T and x and the
effects of opening and closure of two-phase fluid fields are
shown for the system H2O–NaCl–CH4 at 50 MPa in
Fig. 5D based on the data by Krader (1985). Remember-
ing that increasing pressure favours fluid immiscibility in
H2O–salt–non-polar gas systems, it is clear, that single
phase fluids are restricted to very narrow P–T–x regimes
and that most parts of the geological relevant P–T–x space
are characterized by co-existing fluids with different physi-
cochemical properties.
Numerous equations of state (EOS) exist to describe the
pressure–temperature–volume relations in the different
fluid systems at elevated pressure–temperature conditions.
For a thorough review of the different EOS and the under-
lying fundamentals, the reader is referred to the reviews by
Ferry & Baumgartner (1987) and Gottschalk (2007) and
to Prausnitz et al. (1999) and Sengers et al. (2000). EOS
for H2O–non-polar gas systems are given among others by
Kerrick & Jacobs (1981) and Duan & Zhang (2006) for
H2O–CO2, Jacobs & Kerrick (1981) and Duan et al.
(1992a,b) for H2O–CO2–CH4, and Saxena & Fei (1987,
1988), Belonoshko & Saxena (1992), Duan et al. (1992c,
1996), Shi & Saxena (1992) and Churakov & Gottschalk
(2003a,b) for H2O–CO2–CH4 plus additional phase com-
ponents. H2O–salt systems are covered by Anderko & Pit-
zer (1993a,b) for H2O–NaCl–KCl, Jiang & Pitzer (1996)
for H2O–CaCl2 and Duan et al. (2006) for H2O–CaCl2
and H2O–MgCl2. EOS for ternary H2O–salt–non-polar
gas systems are restricted to those by Duan et al. (1995,
2003) for H2O–NaCl–CO2 and H2O–NaCl–CO2–CH4.
According to the authors, some of the EOS for H2O–non-
polar gas systems are valid to rather high pressure and tem-
perature conditions, e.g., Duan & Zhang (2006) up to
2300�C ⁄10 GPa and Churakov & Gottschalk (2003a,b) up
to 1730�C ⁄10 GPa. Unfortunately, those that include salts
are applicable only up to mid-crustal pressure conditions of
maximum 500 MPa (e.g., Duan et al. 1995, 2003).
PROPERTIES OF PURE H2O AT ELEVATED
P–T CONDITIONS
For the evaluation of the properties of aqueous fluids at
elevated pressure and temperature conditions, knowledge
of the fundamental properties density, dissociation behav-
iour, static dielectric constant and dynamic viscosity of
pure water at elevated pressure and temperature conditions
is of primary importance. The available experimental data
as well as formulations for the calculation of these proper-
ties at elevated pressure and temperature are summarized
below and shown in Fig. 6. Emphasis is given to meta-
morphic P–T conditions (grey fields in Fig. 6), which are
here somewhat arbitrarily defined as P > 100 MPa and T >
150�C. Metamorphic conditions are limited towards low
temperature and high pressure by the lowest possible geo-
therm, which was assumed to 5 MPa ⁄ �C. The shown
boundaries of the different metamorphic facies are from
Spear (1993). Some of the reviewed studies report the
respective properties as a function of temperature and den-
sity while others use temperature and pressure. In Fig. 6,
all properties are plotted as a function of temperature with
isobars showing the pressure dependence. For data that
were presented as function of temperature and density in
the original papers density was converted to pressure by
applying the IAPWS formulation for H2O by Wagner &
Pruß (2002). For near-critical behaviour of H2O and aque-
ous systems, see Anisimov et al. (2004).
Density
At ambient conditions, liquid water has a density of
approximately 1 g cm)3. With increasing pressure and tem-
perature along the liquid–vapour equilibrium, the density
contrast between liquid water and steam diminishes and at
the critical point both approach the critical density of
0.322 g cm)3. Numerous EOS and molecular dynamics
simulations describe the PVT properties of pure H2O.
10 A. LIEBSCHER
Among others, these include those by Kerrick & Jacobs
(1981), Halbach & Chatterjee (1982), Saul & Wagner
(1989), Belonoshko & Saxena (1991) and Wagner & Pruß
(2002). At elevated P–T conditions, the density of H2O
ranges from approximately 0.3 g cm)3 up to >1.2 g cm)3
(Fig. 6A; calculated with the IAPWS formulation by
Wagner & Pruß 2002). Only at low pressure the density
may be < 0.3 g cm)3. Increasing temperature generally
decreases density whereas increasing pressure has the oppo-
site effect. In low-P ⁄high-T settings as they occur in
oceanic hydrothermal systems, volcanic systems and meta-
morphic systems with very high geothermal gradients H2O
densities are below approximately 0.6 g cm)3. Under typi-
cal crustal metamorphic conditions, however, the density
of H2O ranges between approximately 0.6 and 1.0 g cm
)3.
Only at rather high pressures, as in subduction zones or
within the upper mantle, are conditions such that H2O
densities exceed 1.2 g cm)3.
Self-dissociation
The dissociation equilibrium of water
1100
300
400
500
600
700
800
900
1000
200
5 25201510
500
Isobars (MPa)
A
G
SGr
Gr
B
EA
Te
m
pe
ra
tu
re
 (°
C)
-log ion product KW
30
1000
Liquid + Vapour
Critical point
100 2550200
Liquid + Vapour
Critical point N
ot
 re
al
ize
d 
on
 E
ar
th
100 250 500 1000 2000
Isobars (MPa)
A
G
SGr
Gr
EA
E
300
400
500
600
700
800
900
1000
200
0 1.20.80.4
Te
m
pe
ra
tu
re
 (°
C)
1100
0.2 1.41.00.6
Density (g cm–3)
B
50
10
100
0
1000
Not realized
on Earth
300
400
500
600
700
800
900
1100
200
0 0.00040.00030.00020.0001
Te
m
pe
ra
tu
re
 (°
C)
Dynamic viscosity (Pa s)
Liquid + Vapour
Critical
point
100
10 50
Isobars (MPa)
0.0005
0
100
1000
15001000500
N
ot
 realized
 on
 Earth
200
300
350
G
A
Gr
B
SGr
Te
m
pe
ra
tu
re
 (°
C)
Liquid + Vapour
Critical
point
100
250
500
1000
Isobars (MPa)
Static dielectric constant
0
100
200
300
400
500
600
700
800
900
1000
1100
0 30 40 60 80 90
50
10
Gr
B
SGr
G
A
EA
Not
 realized
 on
 Earth
Heger (1969) Deul (1984)
50 MPa
500 MPa
250 MPa
100 MPa
Shock et al. (1992)
Fernández et al. (1997)
70502010
E
Abramson (2007)
Sengers & 
Kamgar-Parsi (1984)
Marshall & Franck (1981)
Tanger & Pitzer (1989)
Quist (1970)
50 100 200 MPa
Franck et al. (1990a,b)
255 MPa
~ 500 MPa
1.55 GPa
Pitzer (1983)
229 MPa
885 MPa
E
Franck (1956)
350 670 970 MPa
(C) (D)
(A) (B)
Fig. 6. Density (A), ion product (B), static dielectric constant (C) and dynamic viscosity (D) of pure water as function of temperature. Thin lines are isobars
with pressure given in MPa. Shadowed fields represent range of metamorphic conditions with 100 MPa arbitrarily chosen as lower P-limit. Thick grey lines
define the different metamorphic facies where SGr = sub-greenschist, Gr = greenschist, B = blueschist, EA = epidote-amphibolite, A = amphibolites,
G = granulite and E = eclogite facies (after Spear 1993). Data in (A) from Wagner & Pruß (2002), other data sources given in the Figure. Stippled lines in (D)
are eye-drawn through the data points of Abramson (2007).
Aqueous fluids at elevated pressure and temperature 11
H2O ¼ HþþOH�
is described by the equilibrium constant
K ¼ aHþ � aOH�
aH2O
or by the ion product
Kw ¼ aHþ � aOH� :
At ambient conditions, the dissociation equilibrium of
liquid water is strongly shifted to the left hand side, con-
centrations and activities of H+ and OH) are low and
Kw = 1 · 10)14. With pH ¼ � log aHþ this turns into the
well known value of pH = 7 for neutral water. At ele-
vated P–T conditions, however, Kw notably changes. The
ion product along the liquid–vapour equilibrium has been
experimentally determined among others by Bignold et al.
(1971), Fischer & Barnes (1972), MacDonald et al.
(1973) and Sweeton et al. (1974). Measurements at ele-
vated pressure and temperature conditions above the criti-
cal point of water were performed by Franck (1956) at
350–970 MPa ⁄500–1000�C and by Quist (1970) up to
555 MPa and 800�C. Extreme conditions have been
studied by Holzapfel & Franck (1966) at 4.5–
10.0 GPa ⁄600–1000�C. Formulations for the change of
Kw with pressure and temperature are given by Marshall
& Franck (1981) up to 1000�C ⁄1000 MPa, Pitzer
(1983) for low densities up to approximately 0.1 g cm)3,
Tanger & Pitzer (1989) up to 500 MPa ⁄2000�C, and
Bandura & Lvov (2006) up to 1.0 GPa ⁄1000�C. Kw as
function of pressure and temperature calculated by the
formulations of Marshall & Franck (1981) and Tanger &
Pitzer (1989) together with selected data by Franck
(1956) and Quist (1970) are shown in Fig. 6B. The
Marshall & Franck (1981) and Tanger & Pitzer (1989)
data agree well for the pressure range between 100 and
500 MPa and reproduce the available experimental data
by Quist (1970). However, at pressures below approxi-
mately 100 MPa, both formulations diverge and the for-
mulation by Marshall & Franck (1981) predicts notably
higher values for Kw than the Pitzer (1983) and Tanger
& Pitzer (1989) formulations. In this low pressure, low
density region the formulations of Pitzer (1983) and
Tanger & Pitzer (1989) are probably more reliable (see
discussion in Tanger & Pitzer 1989). Within the pressure
range 200–1000 MPa, temperature has only a minor
effect on Kw, whereas Kw generally increases with increas-
ing pressure. Under normal mid- to lower-crustal P–T
conditions, Kw ranges between approximately 10
)7.5 and
approximately 10)12.5. However, at very high pressures of
the eclogite facies, Kw may even be substantially higher
than 10)7.5. In line with this, Holzapfel & Franck (1966)
determined Kw = 10
)2.8 to 10)1.2 at approximately 7–9
GPa ⁄500–1000�C.
Dielectric constant
The dielectric constant is of primary importance for the
evaluation of the solvent properties of water. The attractive
or repulsive force between any point charges q1 and q2 in a
medium, e.g., ions in aqueous solution, that are separated
by distance r is given by
F ¼ q1q2
4p"r2
with e being the electric permittivity of the medium. The
electric permittivity of the vacuum is typically designated e0
and the dimensionless ratio e ⁄ e0 then yields the static rela-
tive permittivity or static dielectric constant for any med-
ium. A high dielectric constant corresponds to a high
resistance of the medium to the transmission of an electric
field. At ambient pressure and 25�C liquid water has a
notably high static dielectric constant of 78.4 (Ferna´ndez
et al. 1995). This high static dielectric constant makes
liquid water a good solvent for charged species at ambient
conditions as it minimizes the electrostatic forces between
the dissolved ions and prevents them to combine to crys-
tals. However, with increasing pressure and temperature,
the static dielectric constant of liquid water substantially
decreases. Ferna´ndez et al. (1995) presented a database for
the static dielectric constant of water and steam. It extends
up to a temperature of 873 K and a pressure of 1189 MPa
and covers the data available at that time. The static dielec-
tric constant along the liquid–vapour equilibrium has been
studied by Oshry (1949), Svistunov (1975), Lukashov
(1981), Muchailov (1988) and Mulev et al. (1994). Along
the liquid–vapour equilibrium, e of liquid water continu-
ously decreases from approximately 55 at 100�C to only
approximately 9.7 at 370�C, whereas e of co-existing steam
only slightly increases from approximately 1.03 at 150�C
to approximately 3.02 at 370�C. Above the critical point, e
has been measured by Heger (1969) and Heger et al.
(1980) at 25–500 MPa ⁄400–550�C, Lukashov et al. (1975)
at 23–58 MPa ⁄400–600�C, Golubev (1978) at 23–39 MPa ⁄
420–510�C, Lukashov (1981) at 24–580 MPa ⁄400–600�C,
and Deul (1984) at 30–300 MPa ⁄400�C. Based on the
available experimental data, formulations and calculated
values for e up to high P–T conditions are given by
Quist & Marshall (1965) up to 1.55 GPa ⁄800�C, Bradley
& Pitzer (1979) up to 100 MPa ⁄350�C, Pitzer (1983) up
to 880 MPa ⁄927�C, Franck et al. (1990a,b) up to 1.55
GPa ⁄1000�C, Shock et al. (1992) up to 500 MPa ⁄
1000�C, Wasserman et al. (1995), and Ferna´ndez et al.
(1997) up to 1.2 GPa ⁄600�C. Fig. 6C reviews the experi-
mental data by Heger (1969) and Deul (1984) together
with isobars calculated by formulations of Shock et al.
(1992) and Ferna´ndez et al. (1997). Also shown are values
for the static dielectric constant for selected P–T conditions
as calculated by formulations of Pitzer (1983) and
12 A. LIEBSCHER
Franck et al. (1990a,b). For temperatures up to 1000�C
and pressures up to 500 MPa, i.e. the P-range largely
covered by experimental data, the differentformulations
agree quite well. However, at pressures above 500 MPa,
the different formulations diverge. At 800�C, e.g.,
Ferna´ndez et al. (1997) predict a static dielectric constant
of approximately 14 at 1.0 GPa, whereas Franck et al.
(1990b) predict e � 14 only at notably higher pressure of
1.55 GPa. At slightly higher temperature of 930�C, Pitzer
(1983) predicts e � 11 at 885 MPa, roughly 120 MPa
below the prediction of Ferna´ndez et al. (1997). This indi-
cates a higher pressure dependence of e in Pitzer (1983)
compared to Ferna´ndez et al. (1997) and Franck et al.
(1990a,b). Despite these differences at high P conditions,
the static dielectric constant of H2O generally decreases
with increasing temperature but increases only slightly with
increasing pressure. Because the effect of temperature is
more pronounced than that of pressure, e decreases with
increasing P–T conditions and range between approxi-
mately 8 and 25 at normal crustal P–T conditions. Only at
low temperature of greenschist to sub-greenschist facies
conditions or notably high pressure of blueschist to eclog-
ite facies conditions e may exceed 25. But given the
restricted P–T range covered by experimental data, any
calculation of the static dielectric constant at pressures
notably in excess of 500 MPa has to be done with great
caution.
Viscosity
The transport properties of water strongly depend on its
dynamic viscosity g, which also plays an important role for
diffusion. Several studies have addressed the dynamic vis-
cosity of water at low to moderate temperature and pres-
sure (see Watson et al. (1980) for a review of available
experiments at that time). The dynamic viscosity of water
at elevated P–T conditions has been studied by Dudziak &
Franck (1966) up to 350 MPa and 560�C and by Abram-
son (2007) up to 6 GPa and 300�C. Based on the available
data, Watson et al. (1980) and Sengers & Kamgar-Parsi
(1984) derived representative equations for g (Fig. 6D).
At ambient conditions, liquid water has a dynamic viscosity
of g = 0.001 Pa sec (Sengers & Kamgar-Parsi 1984).
Along the liquid–vapour equilibrium, the viscosity of liquid
water substantially decreases whereas that of steam only
slightly increases, so that the critical dynamic viscosity of
H2O is approximately 3.9 · 10)5 Pa sec. At temperature
conditions above approximately 400�C, the data suggest
that at isobaric conditions the temperature effect on the
dynamic viscosity is only minor whereas increasing pressure
increases g. Except for low temperatures of sub-greenschist
facies conditions, H2O has dynamic viscosities at normal
crustal P–T conditions between 0.0001 and 0.0002 Pa sec.
Although no data are available for upper blueschist and
eclogite facies, the data suggest that at these high pressure
conditions the dynamic viscosity of H2O may range up to
0.0003 Pa sec (Fig. 6D).
H2O–SILICATE SYSTEMS AT HIGH P–T
CONDITIONS
Aqueous fluids and silicate melts are the two most impor-
tant mobile phases in the Earth crust and mantle. H2O–
silicate systems typically belong to Type 2 systems (see
above), which are characterized by a discontinuous critical
curve and lower and upper segments of the two-phase fluid
volume that terminate at the lower and upper critical end-
points (see Fig. 7C). Up to moderate pressure and temper-
ature conditions the solubility of silicates in aqueous fluids
and of H2O in silicate melts is typically low and both
phases fundamentally differ in their physicochemical prop-
erties. With increasing pressure, however, the solubility of
H2O in silicate melts and of silicates in aqueous fluids
increases and the differences between both phases finally
vanish at the critical curve above which aqueous fluids and
silicate melts are completely miscible giving rise to so-
called ‘supercritical’ fluids. Here, only some key aspects of
H2O–silicate systems at high pressure and temperature
conditions are reviewed. A thorough review of this topic is
given by Hack et al. (2007). Kennedy et al. (1962) pro-
vided the first experimental evidence for complete miscibil-
ity between aqueous fluids and silicate melts at high
pressure and temperature within the system H2O–SiO2
and proposed an upper critical endpoint at 970 MPa ⁄
1080�C with a composition of approximately 25 wt%
H2O ⁄75 wt% SiO2. This has been confirmed by Newton
& Manning (2008). Shen & Keppler (1997) directly
observed complete miscibility between aqueous fluids and
silicate melts in the system H2O–albite and determined the
location of the critical curve (Fig. 7A). Further experiments
by Stalder et al. (2000) at near wet solidus conditions then
located the upper critical endpoint in the system H2O–
albite at approximately 1.5 GPa ⁄700�C (Fig. 7A). The
resulting phase relations within the system H2O–albite up
to 1.7 GPa and T > 500�C based on the data by Paillat
et al. (1992), Shen & Keppler (1997) and Stalder et al.
(2000) are shown in Fig. 7B. At pressure below approxi-
mately 1.5 GPa, the critical curve is at higher temperature
than the wet solidus, which is defined as onset of melting
by the reaction albite + vapour = liquid. However, with
increasing pressure, the critical curve shifts to lower tem-
perature, the liquid + vapour two-phase field narrows and
at approximately 1.5 GPa the critical curve intersects the
wet solidus at about 700�C. Above approximately 1.5 GPa
no discrete melting occurs but albite continuously dissolves
in the fluid with increasing temperature giving rise to a
continuous range from solute poor, almost pure H2O at
low temperature to H2O poor silicate dominated liquids at
Aqueous fluids at elevated pressure and temperature 13
high temperature. Addition of fluorine, boron and sodium
to the H2O–albite system shifts the location of the critical
curve to even lower pressure and temperature (Sowerby &
Keppler 2002). Critical curves for the systems H2O–nepheline
and H2O–jadeite have then been determined by Bureau &
Keppler (1999). Termination of the wet solidus at the
upper critical endpoint, however, not only occurs in simple
H2O–mineral systems but is also observed in H2O–rock
systems. Kessel et al. (2005a,b) studied the system H2O–
potassium-free basalt at 4–6 GPa ⁄700–1400�C. They
observed eutectic and peritectic melting at pressures of
4 and 5 GPa, respectively, but no discrete melting at
6 GPa (Fig. 7C). At 6 GPa, the aqueous fluid continuously
increases its solute content with increasing temperature
suggesting a critical endpoint in this system between 5 and
6 GPa ⁄1050�C. Complete miscibility between aqueous
fluids and silicate melts in the systems H2O–Ca-bearing
granite and H2O–haplogranite was observed by Bureau &
Keppler (1999). The corresponding critical curves are
located at 1.34 GPa ⁄1003�C to 2.04 GPa ⁄735�C in the
H2O–haplogranite system and at slightly higher pressure
and temperature conditions of 1.61 GPa ⁄898�C to
2.08 GPa ⁄820�C in the H2O–Ca-bearing granite system.
However, no critical endpoints, which determine the ter-
 T (°C)
1400
1200
1000
800
1600
 T (°C)
2000
1500
1000
500
2500
0.1
1.7
1.5
1.2
5
1.0
0.7
5
0.5
0.2
5
2000
1500
2500
755025
1500
1000
500
P (GPa)
1000
NaAlSi3O8 H2O
Upper
critical
endpoint
Critical
 curve
l + v
ab + v
ab
 +
 lW
et
 solidus
Dry
 solidus
ab + scf
4.0
6.0
5.0
604020
P (GP
a)
mol % H2O
Upper
critical
endpoint
1300900500
0
0.5
1.0
1.5
Temperature (°C)
Pr
es
su
re
 (G
Pa
)
Upper
critical
endpoint
W
et
 solidus
Dr
y 
so
lid
u
s
Liq
uid
Albite
 +
 vapour
Li
qu
id
Al
bi
te
1400
12001000
800
1600
80
Critical curve
1000
800
Rock
 + v
Rock + l
Rock + scf
l + v
D
ry
 solidus
W
et
solidus
H2OEclogite
6 1412108
–4
–3
–2
–1
0
1
2
3
93 mol % H2O
81 mol % H2O
72 mol % H2O
57 mol % H2O
33 mol % H2O
Temperature (10 000/K)
Lo
g 
vis
co
sit
y 
(P
a
 s
)
mol % H2O
(A) (B)
(C) (D)
Fig. 7. Water–silicate systems at high pressure and temperature. (A) P–T diagram showing location of the wet and dry solidus, critical curve and upper critical
endpoint in the system H2O–albite. Wet solidus, filled and empty circles from Stalder et al. (2000), critical curve and diamonds from Shen & Keppler (1997)
and dry solidus from Boyd & England (1963). (B) Phase relations within the system albite–H2O up to 1.7 GPa and T > 500�C showing the termination of the
water saturated solidus and the critical curve at the upper critical endpoint. Drawn and modified after Paillat et al. (1992), Shen & Keppler (1997) and Stalder
et al. (2000) (ab = albite, v = vapour, l = liquid, scf = supercritical fluid). (C) Phase relations within the system eclogite–H2O at 4–6 GPa and T > 600�C
based on the data of Kessel et al. (2005a). Critical curve, dashed and stippled lines are hypothetical additions to the data given in Kessel et al. (2005a) to
clarify phase relations (v = vapour, l = liquid, scf = supercritical fluid). (D) Arrhenius plot showing the viscosity of H2O–albite solutions as function of temper-
ature and water content at high pressures >1.0 GPa. Data from Aude´tat & Keppler (2004), stippled lines calculated according to their Eqn (1).
14 A. LIEBSCHER
mination of the wet solidus, are given for the systems
H2O–Ca-bearing granite and H2O–haplogranite in Bureau
& Keppler (1999). One important aspect of H2O–silicate
systems at pressure conditions above the critical endpoint
is the combination of high solute contents in the fluids
with comparably low viscosities. Aude´tat & Keppler (2004)
performed a seminal experimental study on the viscosity of
high-pressure silicate rich aqueous fluids in the systems
H2O–albite, H2O–nepheline and H2O–pectolite. In the
H2O–albite system, fluid viscosity linearly increases with
increasing solute content from approximately 3.5 · 10)4
Pa sec for 7 mol% solutes to approximately 4 · 100 Pa sec
for 67 mol% solutes (Fig. 7D). The viscosity of dry and
hydrous albite melts with up to 3 wt% H2O is notably
higher and ranges between approximately 109 and
1011.5 Pa sec (Romano et al. 2001). The data therefore
indicate an exponentially decrease of viscosity with increas-
ing H2O content for <20 wt% H2O (Aude´tat & Keppler
2004). Given their comparably low viscosity with high
solute content, aqueous fluids above the upper critical end-
point potentially form very effective metasomatising agents
within the mantle wedge.
SUMMARY
This paper reviews some of the fundamental physical and
chemical properties of aqueous fluids at elevated pressure
and temperature conditions. For certain geological environ-
ments these physicochemical fluid properties are fairly well
known. This is, at least partly, true for low pressure and
temperature systems like geothermal systems, basinal sys-
tems, and oceanic hydrothermal systems, in which fluids
can be sampled and studied directly or which are easily
investigated by experimental methods. At higher, typical
metamorphic pressure and temperature conditions, the
principal phase relations in model fluid systems are known.
However, there are important gaps in knowledge of aque-
ous fluids at elevated pressure and temperature. First, even
for pure H2O, most data for important properties like
ion product, static dielectric constant, and viscosity are
restricted to pressures below 1.0 GPa. Also, some of the
available formulations for these properties substantially
diverge outside the pressure–temperature range covered by
the experimental data. Second, while the principal phase
relations in binary and ternary aqueous fluid systems are
well known, experimental data precisely addressing the
composition of co-existing fluid phases are rare. In addition,
data on trace element and isotope fractionation behaviour
between co-existing aqueous fluids are rare. Knowledge of
these fractionation behaviours are a prerequisite for model-
ling geochemical cycles in fluid mediated or fluid dominated
systems. Finally, for high pressure and temperature of the
lower crust and upper mantle, the knowledge of physical
fluid properties like viscosity and density in ‘supercritical’
H2O–silicate systems is rather limited. However, for model-
ling the role of these ‘supercritical’ H2O–silicate fluids as
metasomatizing agent for crust–mantle interactions within
subduction zones, these data are essential.
ACKNOWLEDGMENTS
I thank the editors for inviting me to contribute to this
Geofluids anniversary volume. This paper greatly benefitted
from careful and critical reviews by two anonymous review-
ers. Helpful editorial handling by C. Manning is gratefully
acknowledged.
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