the chemistry of metal transport and deposition by ore

Tre

13.2 The Chemistry of Metal Transport and Deposition by Ore-FormingHydrothermal FluidsTM Seward, Victoria University of Wellington, Wellington, New ZealandAE Williams-Jones and AA Migdisov, McGill University, Montreal, QC, Canada

ã 2014 Elsevier Ltd. All rights reserved.

13.2.1 Introduction 2913.2.1.1 Compositions of Ore-Forming Fluids 3013.2.2 Hydrothermal Ore Solution Chemistry – The Main Dissolved Components 3313.2.2.1 Water Solvent at Hydrothermal Conditions 3313.2.2.2 NaCl – The Main Dissolved Electrolyte Component 3513.2.2.3 Ion Hydration, Association, and Water Activity 3513.2.2.4 Weak Acid/Base Equilibria in Hydrothermal Systems 3613.2.3 Mineral Solubility in Water and Salt Solutions at High Temperature and Pressure 3713.2.4 Ore Metal Transport and Deposition 4013.2.4.1 Ore Fluids with Liquid-Like Densities 4013.2.4.1.1 Ligands in hydrothermal ore solutions 4013.2.4.1.2 Metal chloride complexing 4113.2.4.1.3 Complexing with other halide ligands 4213.2.4.1.4 Metal complexes with hydroxide and other oxygen electron donor ligands 4313.2.4.1.5 Complexing with hydrosulfide/sulfide ligands 4513.2.4.1.6 Thioanions 4613.2.4.1.7 Complexing with other sulfur-containing ligands 4713.2.4.1.8 Other complexing ligands 4813.2.4.1.9 Ore fluids with gas-like density 4813.2.5 Epilogue 50Acknowledgments 50References 50

13.2.1 Introduction

Economically exploitable deposits of metallic minerals (ore

deposits) form in the Earth’s crust through a variety of geologic

processes. These involve extraction of metals at low

concentration and transport of these metals to sites of deposi-

tion where they accumulate in very much higher concentra-

tion. Hydrothermal fluids are by far the most important agents

of metal transport. Theymay be formational waters of meteoric

or seawater origin, metamorphic fluids produced during devo-

latilization of hydrous minerals, or magmatic fluids released

during decompression and/or crystallization of magmas. Their

physical state may be that of liquid, vapor, or a supercritical

fluid, and chemically, they have highly variable concentrations

of dissolved components, including charged and uncharged

species containing elements such as Na, K, Ca, Fe, Si, Cl, C,

H, O, and S and, of course, the ore metals. The challenge, if

we are to understand the processes involving the hydrothermal

extraction, transport, and deposition of metals, is to determine

the nature of these fluids, the properties that allow them to

dissolve metals, and the conditions under which dissolution

and deposition are optimized.

Active hydrothermal systems are an important source of

information on the composition of ore fluids either because

they are forming deposits analogous to those that are mined or

there is strong evidence linking them to such deposits. For

atise on Geochemistry 2nd Edition http://dx.doi.org/10.1016/B978-0-08-095975

example, the fluids venting through chimneys at spreading

centers and subduction zones are currently forming deposits

with the characteristics of volcanogenic massive sulfide (VMS)

base metal deposits (Scott, 1997). There is strong evidence

that, in many cases, geothermal fluids of the type employed

in energy generation are responsible for the formation of low-

sulfidation epithermal precious metal deposits (Barnes and

Seward, 1997; Clark and Williams-Jones, 1990; Krupp and

Seward, 1987, 1990; Weissberg et al., 1979). The same is true

for oil-field brines, which are analogous to the basinal brines

interpreted to form Mississippi Valley-type (MVT) lead–zinc

deposits. Although they are not ore fluids, per se, volcanic

gases are representative of the fluids that at greater pressure

may form high-sulfidation epithermal precious metal and por-

phyry copper deposits. Indeed, the ore fluids forming these

deposits actually may be gases, albeit of greater density than

volcanic gases (Williams-Jones and Heinrich, 2005). The fluids

from these active hydrothermal systems have been sampled

extensively during the past 50 years and have provided

people with a wealth of data on the chemistry of ore-forming

fluids and, in some cases (e.g., VMS and low-sulfidation

epithermal deposits), the pressure–temperature conditions of

potential ore formation. The other important sources of infor-

mation on the composition of ore-forming fluids and the

conditions of ore formation are fluid inclusions. With the

recent development of tools capable of analyzing their element

-7.01102-5 29

http://dx.doi.org/10.1016/B978-0-08-095975-7.01102-5

30 The Chemistry of Metal Transport and Deposition by Ore-Forming Hydrothermal Fluids

concentrations down to ppm levels, fluid inclusions are now

starting to provide chemical data comparable in quality to

those available from active hydrothermal systems. They are

also the principal source of data on the pressure–temperature

conditions of ore formation.

13.2.1.1 Compositions of Ore-Forming Fluids

Hydrothermal fluids venting on the seafloor were first discov-

ered in 1977 (Weiss, et al., 1977). Since then, large numbers of

studies, mainly along mid-ocean ridges, have documented the

chemistry of these fluids (Table 1) and shown that those at

T �350 �C are forming massive sulfide base metal (and gold)

deposits similar to VMS deposits mined in terrestrial settings

(Scott, 1997; Tivey, 2005). Studies of mid-ocean ridge systems

have established that the composition of the fluids (for the

most part, hydrothermal liquids) is largely the result of inter-

action of seawater with basalt; there are alsominor additions of

magmatic fluids (Shanks, 2001; Yang and Scott, 1996, 2006).

During this interaction, there is a sharp decrease in ex situ pH

(from �7.8 to between 3 and 4) due to removal of magnesium

to form secondary silicate minerals (Seyfried, 1987). Chlorine

is conserved, and consequently, the chlorinity of the hydro-

thermal fluids is generally very similar to that of seawater, that

is,�0.5 m (German and Von Damm, 2003). Sodium, which is

the overwhelmingly dominant cation in the hydrothermal

fluids, is also largely conserved, in most cases, retaining its

seawater concentration of �0.45 m; some Na may be lost to

albitization. Among the major elements, only Mg and S are lost

from the seawater, Mg almost completely so. The sulfur con-

tent drops from nearly 30 mm in seawater to <10 mm in the

vent fluid. However, the major change is in its oxidation state.

Whereas in seawater, sulfur occurs as sulfate, in the vent fluid it

is present dominantly as H2S (Shanks, 2001). The carbon

content, mainly in the form of CO2, is typically two to five

times its seawater content of �2.3 mm. The other major ele-

ments, that is, K and Ca, are also added (Ca through the

breakdown of the anorthite component in plagioclase), typi-

cally reaching concentrations two to three times their seawater

concentration of �0.01 m. Silicon, which is a minor constitu-

ent of seawater, becomes a major element with a concentration

of �0.02 m. Iron, which is present in seawater at very low

concentrations (<1 nm), may also become a major element

in the hydrothermal fluid; in some cases, its concentration may

exceed 0.1 m. The concentrations of the main ore elements

(Cu, Zn, and Pb) are highly variable, generally ranging from

a few tens to a few thousand ppb, but are thousands to tens of

thousands of times higher than in seawater; gold concentra-

tions have not been measured (Table 1).

Although most of the information on the composition of

seafloor hydrothermal fluids comes from systems at mid-ocean

ridges, data on the compositions of hydrothermal fluids in arcs

and back-arc basins are being increasingly reported (Table 1).

The back-arc fluids are very similar in composition to those of

mid-ocean ridges and are also mainly products of seawater–

basalt interaction (Reeves et al., 2011). However, in the arc and

near-arc environment, where the magmas are much more si-

licic, their compositions differ considerably from those of hy-

drothermal fluids at mid-ocean ridges. The main differences

are significantly lower pH (2–3 vs. 3–4), very much higher CO2

(0.25 m) and F contents (0.4 mm), and higher concentrations

of K (up to 0.1 m). These differences are attributed to the

addition of a large magmatic component to the hydrothermal

fluid, possibly as much as 25% (Reeves et al., 2011).

The only ligands likely to be present in concentrations

sufficient to form the complexes necessary for metal transport

are Cl�, CO2 species, reduced sulfur species, and F� (in arc

environments) (Table 1). Ammonium ions are generally not

detected, except in sediment-covered settings (German and

Von Damm, 2003), and in these settings, it seems likely that

they are incorporated at relatively shallow depths of several

hundred meters (Lizarralde et al., 2010) and thus are not likely

to play a significant role in metal extraction, which is thought

to mainly take place at greater depth (Lowell et al., 1995;

Tivey, 2005).

Geothermal well fluids have been extensively sampled and

analyzed for over 50 years. Until recently, this sampling had

been restricted to the wellhead, and because the fluid generally

boils (a requirement for conventional geothermal energy pro-

duction) and precipitates minerals before reaching the surface,

the composition of the fluid differs considerably from that of

the reservoir fluid. Geothermal fluids display a broad range of

compositions, reflecting the diverse geology of terrestrial set-

tings (Table 1). For example, the chlorinity of the high-

temperature fluids (>250 �C) in the Taupo volcanic zone, a

young rifted arc dominated by andesites and rhyolites, ranges

between 0.02 and 0.05 m, anorder ofmagnitude lower than that

of seawater (Ellis, 1979a,b; Ellis andMahon, 1977). At the other

extreme, the high-temperature fluids (>250 �C) in the SaltonSeageothermal system, which is located in a closed evaporitic,

sedimentary basin (Imperial Valley), commonly have a chlorin-

ity between 3 and 4 m; immediately to the south, in the same

basin, geothermal fluids at Cerro Prieto have a chlorinity of 0.2

to 0.5 m (Mercado andHurtado, 1992;Williams andMcKibben,

1989). In both the Taupo volcanic zone and the Imperial Valley,

the fluid is predominantly meteoric water with a small contri-

bution of magmatic water. Geothermal systems may also be

dominated by seawater, as is the case for the Reykjanes field in

Iceland where the chlorinity is very similar to that of seawater

(Arnorsson, 1978). The other major components of geothermal

fluids, for example,Na, K, Ca, andCO2, also vary considerably in

concentration; the seawater-dominated, basalt-hosted Reykjanes

fluids are compositionally very similar to seafloor vent fluids

(Table 1). Until relatively recently, it has not been possible to

directly measure the metal concentrations of preboiled geo-

thermal fluids; however, such data are now starting to become

available through the development of ‘downhole’ devices that

can sample the reservoir fluids directly (e.g., Simmons and

Brown, 2006). For example, in the case of the geothermal fields

of the Taupo volcanic zone, these data show that the reservoir

fluid contains up to 23 ppb Au, 2400 ppb Ag, and 4850 ppb As;

the H2S content ranges up to 0.007 m (Ellis, 1979a,b; Seward,

1989). Lower concentrations of these metals were measured

in the Reykjanes reservoir fluid using the same method, that

is, up to 6 ppb Au, 34 ppb Ag, and <1 ppb As; concentrations

for Cu, Zn, and Pb were 16, 26, and <1 ppb, respectively

(Hardardottir et al., 2009). By comparison, concentrations of

Cu, Zn, and Pb reported for the very high-salinity Salton Sea

geothermal fluid, although sampled at the wellhead, are up to 7,

518, and 107 ppm, respectively; the H2S concentration is

Table 1 Selected analyses of the composition of fluids in active hydrothermal systems

Seafloor vent fluids Geothermal fluids Oil-field brines

Seawater Spreadingcenter

Back-arc basin Arc Rotokawa Salton Sea Cerro Prieto Reykjanes Ladolam golddeposit

Mississippi

Temp (�C) 2 350 282 268 320 330 337 296 275 107pH (25 �C) 7.8 3.400 4.1 2.6 6.03 5.1 8 – 8.01 5.650Na (mm) 469.003 436.334 540.637 449.600 12.139 2382.609 419.980 393.000 1154.697 76.257K (mm) 9.804 23.021 24.023 86.290 2.046 453.846 77.064 38.000 125.641 17.686Ca (mm) 10.204 16.010 82.270 15.009 0.012 712.500 10.059 47.000 0.227 774.343Mg (mm) 52.768 – 0.000 0.000 0.000 2.042 0.002 0.390 0.004Mn (mm) – 0.960 0.348 3.117 – 27.273 0.010 52 000.000 – 1.163Fe (mm) – 1.664 0.109 2.404 – 30.536 0.004 0.430 – 6.198Si (mm) 0.200 18.010 15.007 16.008 23.409 9.800 48.169 10.000 19.238 0.712Ba (mm) 0.140 8.000 17.000 91.001 – 2576.642 – 71.000 – 1056.004Cu (ppm) – 2.224 0.127 2.288 – 6.800 0.005 16.586 4.450 0.020Zn (ppm) – 6.931 0.654 7.520 – 507.000 0.006 24.987 0.185 222.000Pb (ppm) – 63.818 0.829 1450.400 – 102.000 0.005 0.001 0.023 53.200Mo (ppm) – – – – – – – 0.015 0.043Sn (ppm) – – – – – – – 0.001 0.840Cd (ppm) – 0.180 – – – 2.300 – 0.135 – 0.830Ag (ppb) – 0.037 – – 1100.000 1400.000 4.000 34.626 6.000Au (ppb) – – – – 7.800 – 4.000 6.106 16.000As (ppm) – – – – 5.400 – 2.000 0.112 17.000Sb (ppm) – – – – 1.200 – 0.400 0.025 0.004Cl (mm) 551.626 497.669 730.500 583.889 14.620 4500.000 526.019 524.000 613.923 5614.690Br (mm) 0.808 0.855 1.030 1.000 – 1.388 – 0.780 0.451 13.029F (mm) 0.064 – 0.023 0.116 0.300 – – – – 0.042B (mm) 0.426 0.548 0.240 1.620 9.000 24.636 2.220 0.709 12.213 8.049SO4 (mm) 22.366 0.400 1.880 0.420 0.041 0.552 – – 410.375 0.187H2S (mm) 0.000 7.303 1.800 6.803 7.310 0.294 22.617 0.900 –HCO3 (mm) 2.931 7.266 7.649 40.854 1.072 35.909 1.072 – 48.261References Seawater

(Hanningtonet al., 2005)

EPR 21�N(Hanningtonet al., 2005)

Vienna Woods(Hanningtonet al., 2005)

PACMANUSbasin(Hanningtonet al., 2005)

Rotokawa(Weissberget al., 1979;Krupp andSeward,1990;Simmonsand Brown,2007)

Salton Sea(WilliamsandMcKibben,1989)

Cerro Prieto(Weissberget al., 1979;Clark andWilliams-Jones, 1990)

Reykjanes(Hardardottiret al., 2009)

Ladolam golddeposit(Simmonsand Brown,2007)

Oil-field brines(Kharaka et al.,1987)

Table 2 Compositions of selected fluid inclusions from a variety of ore deposit types

Deposit type Sn granite Sn granite Porphyry Cu Porphyry Cu Porphyry Cu Porphyry Cu Zn skarn REE granite MVT MVT MVT

Fluid type Brine Vapor Brine Brine Brine Vapor Brine Brine Brine Brine BrineSalinity 35.20 4.00 31.20 33.10 49.90 9.50 72.50 19.30 23.70 19.70Temp (�C) 453 650 720 471 395 372 595 112 114 110Na (ppm) 74000 14000 75000 67900 80000 30000 95461 222600 60000 74000 67000K (ppm) 57000 3300 61300 58300 99000 16000 41530 92600 <1040 1400 7200Ca (ppm) – – – – – – – – 23000 – 14000Mg (ppm) – – – – – – – – 2110 – 1400Mn (ppm) 22000 1600 n.a. 14200 23000 1500 – 17900 <32.7 – –Fe (ppm) 73000 4100 54700 74700 130000 29000 – 43700 – –Cu (ppm) 2300 4600 2200 2300 5500 30000 – 230 50 6 –Zn (ppm) 3600 680 n.a. 3400 9900 6500 5930 2500 n.d. 13 –Pb (ppm) 3400 190 780 800 2400 620 4350 470 240 – 180Mo (ppm) – – n.a. <120 90 n.a. – 81 – – –Sn (ppm) 390 – n.a. <350 n.a. n.a. – 63 – – –W (ppm) 56 – 57 <67 <80 <50 – 30 – – –Ag (ppm) 290 5 7 <67 n.a. <70 50 3 – – –As (ppm) 120 34 8 <400 50 n.a. 248 25 – – –Sb (ppm) 110 31 – – – – 365 – – –Ce (ppm) 2 – 6 15 210 <2 – 300 – – –Bi (ppm) 10 3 n.a. n.a. – – – 5 – – –Cl (ppm) – – – – – – – – 117000 140000 120000References Mole granite

(Audetatet al.,2000)

Mole granite(Audetatet al., 2000)

Santa Ritaporphyry Cudeposit(Audetatet al., 2008)

Santa Ritaporphyry Cudeposit(Audetatet al., 2008)

Bajo de laAlumbreraporphyry Cu–Audeposit (Ulrichet al., 2002)

Bajo de laAlumbreraporphyryCu–Au deposit(Ulrich et al.,2002)

El MochitoZn–Agskarn(Samsonet al.,2008)

Capitan REEgranite(Bankset al.,1994)

Ozark MVT Pb–Zn deposits(Wilkinsonet al., 2009)

NorthArkansasMVT Pb–Zndeposits(Stoffellet al., 2004)

Tri-State MVTPb–Zndeposits(Stoffellet al., 2004)

The Chemistry of Metal Transport and Deposition by Ore-Forming Hydrothermal Fluids 33

�0.0005 m (Williams and McKibben, 1989). The giant

Ladolam gold deposit (a low-sulfidation epithermal deposit) is

an actively forming ore deposit on Lihir Island, Papua New

Guinea, which contains 1300 t of gold. The fluids forming

this deposit contain up to 16 ppb Au and 4450 ppb Cu; the

chloride content is 0.6 m, the CO2 content 0.2 m, and the H2S

content 0.0006 m (Simmons and Brown, 2006).

The ligands that could be important for metal transport in

geothermal fluids vary considerably with the environment in

which the geothermal system is located. In systems like the

Salton Sea, Cerro Prieto, Reykjanes, and Lihir Island, chloride is

clearly an important complexing ligand for some metals such as

Pb. However, in systems like those of the Taupo volcanic zone,

the chloride concentrations are typically low (e.g., 0.02–0.06 m),

although some of these systems have quite elevated reduced

sulfur concentrations, such as at Ohaaki (0.002 m) and Roto-

kawa (0.007 m) (Krupp and Seward, 1987, 1990; Seward, 1989).

Other species reported to be present in active geothermal systems

that might be important for metal transport are ammonia and

sulfate, the latter reaching 0.02 m at Salton Sea and >0.3 m at

Lihir Island (Simmons and Brown, 2006; Williams and Mckib-

ben, 1989). Ammonia (i.e., NH3þNH4þ) occurs ubiquitously in

geothermal fluids at concentrations up to and/or>0.002 m (see

Ellis, 1979a,b) and forms as an expected consequence of the

redox equilibria involving N2 and H2 (Seward, 1974a,b).

Oil-field brines, as noted earlier, have long been considered

to be modern analogues of the fluids interpreted to form MVT

lead–zinc deposits (Hall and Friedman, 1963). Although the

origin of these basinal brines is still a matter of conjecture, for

example, connate or evaporated seawater, variably modified by

mixing with meteoric water and diagenetic reaction (Wilson

and Long, 1993), they are able to transport significant concen-

trations of metals at relatively low temperature (85–150 �C).Their chlorinity ranges from about 2 to nearly 6 m, and Naþ is

the dominant cation (up to 3.5 m), followed by Ca2þ (up to

2 m; Table 1). The sulfate content ranges up to 0.004 m and that

of NH4þ to 0.01 m, but H2S and inorganic carbon concentra-

tions are very low, <1�10�7 and <5�10�4 m, respectively

(Carpenter et al., 1974; Kharaka et al., 1987). In addition, oil-

field brines contain significant proportions of dissolved organic

species, such as acetate (Table 1). The principal ore metals are

Zn and Pb, which can reach concentrations of up to 350 ppm

and 100 ppm, respectively (Carpenter et al., 1974; Kharaka

et al., 1987). Copper concentrations are typically <200 ppb.

On the other hand, gold concentrations as high as 18 ppb

have been reported (Saunders and Swann, 1990).

Volcanic gases are the surface representatives of magmatic

hydrothermal fluids, and although they are considerably less

dense than the vapors that are believed to be ore fluids in some

settings, they do provide information about the compositions

of these fluids. In addition to water vapor, volcanic gases can

contain >30 mol% CO2, >20 mol% SO2, >5 mol% H2S,

>2mol% HCl, and >0.2 mol% HF; the average contents of

these gases for 19 representative volcanoes are 10.6, 5.2, 1.1,

0.8, and 0.1 mol%, respectively (Halmer, 2002). Metal con-

tents of volcanic gases range up to 6 ppm Cu, 12 ppm Pb,

11 ppm Zn, 7 ppm Sn, 250 ppb Ag, and 24 ppb Au (e.g.,

Symonds et al., 1987; Wahrenberger et al., 2002; Williams-

Jones et al., 2002), and the total metal content of discharging

volcanic gases from a single volcano may constitute an

appreciable metallic flux on the order of thousands of tonnes

per annum to the atmosphere (Calabrese et al., 2011).

Fluid inclusions provide the most direct source of data on

the compositions of ore fluids in extinct hydrothermal systems,

and analyses of individual fluid inclusions (e.g., using LA-

ICPMS) are the most reliable sources for these data (Table 2).

With a few exceptions, studies reporting compositions of indi-

vidual inclusions have focused on magmatic hydrothermal ore

deposits, and data are available for both the single-phase fluids

that exsolved from the magma and the later brine and vapor

that were the products of phase separation (Klemm et al., 2007;

Samson et al., 2008). These data show that supercritical mag-

matic fluid inclusions typically contain 0.1–0.8 m Cl, 0.1–

0.6 m Na, up to 0.3 m K, up to 0.1 m Fe, on average

3500 ppm Cu, 200 ppm Zn, 40 ppm Pb, and 10 ppm Mo.

Gold concentrations are typically at high ppb levels. The com-

positions of magmatic vapor inclusions are similar, except that

Cu concentrations can be very much higher, over 1 wt%, and

gold concentrations up to 10 ppm have been reported (Ulrich

et al., 1999). Magmatic brine inclusions have much higher

concentrations of salts, commonly >6 m Cl, >3 m Na,

>1.5 m K, >1.5 m Fe, and on average 1500 ppm Cu,

3500 ppm Zn, 2000 ppm Pb, and 80 ppm Mo, and up to

3000 ppb Au. Some compositional data have been reported by

Stoffell et al. (2004) for fluid inclusions from MVT deposits.

These show that the fluids contain an average of 3.4 m Cl, 2.6 m

Na, 0.2 m K, 0.6 m Ca, 212 ppm Zn, and 94 ppm Pb.

13.2.2 Hydrothermal Ore Solution Chemistry –The Main Dissolved Components

13.2.2.1 Water Solvent at Hydrothermal Conditions

The review of the chemistry of ore-forming fluids has shown

clearly that they are multicomponent aqueous electrolyte solu-

tions containing both simple ionic and more complex molec-

ular species. In this section, the properties of water solvent that

enable it to be an effective agent for the transport of ore metals

and associated elements are considered.

At ambient temperature and pressure, water is the archetype

protic solvent that exhibits various unusual properties that may

be attributed to the large dipole moment of water and the

associated hydrogen bonding among neighboring molecules.

Dissolved ionic and molecular species are ‘hydrated’ and inter-

act with water solvent (Figure 1). As temperature and pressure

increase at conditions defined by the two-phase curve (i.e.,

liquid–vapor equilibrium), liquid water expands as the posi-

tional and orientational constraints alter and the extent of

hydrogen bonding diminishes. The increasing molar volume

(decreasing density) of water liquid as one proceeds to the

critical point is thus accompanied by concomitant decreases

in its dielectric permittivity and other properties, such as vis-

cosity. The decrease in the dielectric constant from 80.10 to

7.22 for liquid water as temperature increases from 20 to

373 �C (at the equilibrium vapor pressure) (Fernandez et al.,

1997; Uematsu and Franck, 1980) gives some premonition of

the changing nature of water dipole–ion interaction, which

results in enhanced ion pairing/association and metal complex

formation. Increasing pressure acts in an opposing way, caus-

ing an increase in the static permittivity (Fernandez et al.,

2.952 Å48.5 �

-

49.9 �

+

Figure 1 Structure of water in the vicinity of a metal cation. Watermolecules immediately adjacent to the ion are arranged so that theirnegatively charged oxygen atoms are directed inward to form an innerhydration shell; these molecules are surrounded by an outer shell ofwater molecules that are similarly directed but with less consistency andlocally share hydrogen bonds with the bulk water solvent. Thedevelopment of hydration shells facilitates dissolution by separatingcations from anions.


1997), thus encouraging the stability of some ionic species

under supercritical conditions. The self-ionization of liquid

water as manifested by the ion product constant, Kw, changes

as a function of temperature and pressure as the energetics of

deprotonation of water molecules changes in response to re-

lated changes in the hydrogen bonding array. Steam is domi-

nated by the presence of water clusters that increasingly

interact with each other via hydrogen bonding as the density

of the vapor phase increases to liquid-like densities as the

critical point is approached. As expected, therefore, the viscos-

ity of steam increases with increasing temperature at liquid–

vapor equilibrium.

The molecular structure and nature of the hydrogen bond-

ing of liquid water and supercritical water at liquid-like densi-

ties continues to be an important and ongoing research area of

fundamental importance to hydrothermal geochemistry and

ore solution chemistry. It has been generally considered that

liquid water exhibits a short-range tetrahedral ordering with

respect to the nearest neighbor water molecules, apparently

inheriting some remnant configurational aspects from ice.

Above about 150 �C, the ‘evidence’ from X-ray diffraction

and neutron diffraction isotopic dilution (NDIS) (Gorbaty

and Kalinichev, 1995; Soper, 2000) for the nearest neighbor

ordering diminishes, suggesting that water might be con-

sidered as a hydrogen-bonded continuum. But recent X-ray

Raman scattering data (Wernet et al., 2005) and NDIS mea-

surements and simulations (Bernabei et al., 2008) of supercrit-

ical water indicate that such a view is too simplistic. The X-ray

Raman spectra are consistent with a structural model for su-

percritical water comprising small domains of water molecules

in various tetrahedral configurations in a gas-phase-like mo-

lecular continuum. Extensive hydrogen bonding persists with

increasing temperature and pressure, as demonstrated by com-

putational/theoretical and spectroscopic (Raman and NMR)

studies (e.g., Chialvo and Cummings, 1994; Hoffmann and

Conradi, 1997; Kalinichev et al., 1999; Walrafen et al., 1996).

For example, the NMR data of Hoffmann and Conradi (1997)

indicate that at 400 �C and 400 bar, 29% of the hydrogen

bonding still exists, compared to ambient conditions. Never-

theless, the detailed interpretation of such results on a molec-

ular level is awaited. The current perspective on the structure

of liquid water from X-ray scattering, spectroscopic and dif-

fraction data, as well as molecular dynamics (MD) and ab

initio approaches, is elegantly summarized by Nilsson and

Petterssen (2011).

Self-ionization is an important property of water and aque-

ous solutions, which involves the loss of a proton from the

water molecule, leaving behind a hydroxide ion (OH�). Spon-taneous reaction of the hydrogen nucleus with another water

molecule produces the hydronium ion, H3Oþ. The value of

pKw (i.e., � log Kw of the equilibrium ion product constant)

depends on temperature, pressure, and, in the case of an aque-

ous solution, ionic strength. It decreases with increasing tem-

perature to a minimum at between 250 and 350 �C(depending on pressure), above which it increases. For exam-

ple, the pKw at 250 �C and saturated water vapor pressure

(swvp) is 11, whereas it is 14 at 25 �C and swvp. The equilib-

rium pH therefore drops from 7 at 25 �C to 5.5 at 250 �C at

swvp. Increasing pressure decreases the pKw (due to the higher

stability of ionized species) and shifts the temperature mini-

mum to higher temperature (Marshall and Franck, 1981).

Addition of salts to the water may either increase or decrease

the ion product, depending on the nature and concentration of

the salt present in the solution (Kron et al., 1995). For example,

measurements of the ion product constant of water in solu-

tions of NaCl, KCl, KNO3, and NaNO3 at 25 �C and at swvp

demonstrate that addition of these electrolytes progressively

decreases Kw of water to a minimum value at an ionic strength

characteristic for the particular electrolyte; further addition of

these electrolytes increases the ion product, potentially to

values well above 14 (Figure 2). Although the pKw minima

for the salts in question are at concentrations <1 m, those for

other salts may occur at significantly higher concentrations

(Kron et al., 1995). As most interactions of water with rock

(i.e., with silicate minerals) occur through pH-dependent re-

actions and concentrations of many aqueous species also de-

pend on pH, pKw is one of the major parameters controlling

speciation in hydrothermal solutions. For example, the species

AlOH2þ forms via the reaction Al3þþOH�⇆AlOH2þ, and its

stability therefore depends on the concentration of the hy-

droxyl ion and, in turn, the solvent self-ionization.

As discussed earlier, the ability of water to stabilize ionized

(charged) species is controlled by the development of

Figure 1

14.6

14.4

14.2

pK

w

14.0

13.8

13.60 1 2 3

NaCl (mol kg-1)

4 5 6 7

Figure 2 The effect of salinity on the ionization constant of water (pKw)at 25 �C; the experimental points are shown by the open circles andcalculated values by the solid line. Modified from Kron I, Marshall SL,May PM, Hefter G, and Konigsberger E (1995) The ionic product of waterin highly concentrated aqueous electrolyte solutions. Monatshefte furChemie 126: 819–837. With permission from Springer.


hydrogen-bonded networks of water molecules. The expansion

associated with increasing temperature distorts these networks

and leads to a decrease in the relative permittivity and associ-

ated destabilization of ionized species. In the extreme, water

can expand to a fluid with gas-like density or separate a vapor,

and in this state, there is complete destruction of the network.

Water molecules cease to be part of a continuum but instead

form isolated clusters with small numbers of molecules

(Kalinichev and Churakov, 1999). Consequently, the ability

of gas-like aqueous fluids to stabilize ionized species is lower

(see the succeeding text), and metal solubility in these fluids is

predominantly, but not exclusively, molecular (i.e., dissolved

species have no charge). Nevertheless, these species commonly

disobey the rules for simple gaseous mixtures. In particular,

hydration still occurs and neutrally charged species may be

incorporated into water clusters. Moreover, there is evidence

that the clusters can even support charged species (e.g.,

Likholyot et al., 2005, 2007).

13.2.2.2 NaCl – The Main Dissolved Electrolyte Component

Hydrothermal ore fluids are multicomponent electrolyte solu-

tions in which NaCl is generally the dominant salt component.

Of course, natural hydrothermal systems are more complex

with numerous interrelated heterogeneous equilibria deter-

mining the hydrolytic (pH) and redox state of the fluids that

undergo phase separation as they migrate and ascend buoy-

antly though the Earth’s crust. Sodium chloride is the classic

strong electrolyte whose aqueous solutions are considered to

be extensively ionized into hydrated Naþ and Cl� ions at

ambient temperature and pressure, although concentrated so-

lutions (saturation¼6.013 m at 20 �C) may exhibit some as-

sociation. However, with increasing temperature and at lower

pressures (e.g., P�500 bar), sodium chloride solutions be-

come progressively more associated (to form simple NaCl0

ion pairs), as was elegantly demonstrated by the earlier con-

ductance study of Quist and Marshall (1968). For example, a

1.0 m NaCl solution at 300 �C and at the equilibrium vapor

pressure is about 50% associated. They also showed that in-

creasing pressure up to 4000 bar has the opposite effect and

causes the disproportionation of associated species, thus favor-

ing the formation of the simple hydrated ions, Naþ and Cl�,whose partial molar volumes are electrostrictively enhanced

relative to the associated moieties.

A number of recent MD and ab initio/MD studies have also

shed light on the molecular nature of sodium chloride solu-

tions at elevated temperatures and pressures (e.g., Dong et al.,

2008; Driesner et al., 1998; Guardia et al., 2006a,b; Sherman

and Collings, 2002; Yui et al., 2010). Driesner et al. (1998)

showed that for a 1 m NaCl solution at 380 �C and solution

density of 0.55 g cm�3, the simple hydrated ions (i.e., Naþ and

Cl�) comprise only about 13% of the species in solution, with

the remainder of the solute consisting of associated species such

as NaCl0, Na2Clþ, NaCl2

�, and Na2Cl2 but with the dominant

associated species being the simple NaCl0 ion pair. Some

short-lived, larger associated species, such as Na3Cl4�, were

also observed in the simulations but have lifetimes of �2 ps.

In summary then, the main electrolyte salt component of

hydrothermal fluids in the Earth’s crust is, with few exceptions,

NaCl. The bulk electrolyte salt association chemistry at high

temperatures and pressures plays a fundamental role in the

complexing and transport/deposition chemistry of many tran-

sition metals in hydrothermal ore-forming systems. Thus, in a

cooling hydrothermal solution ascending buoyantly through

the Earth’s crust at lower pressures (e.g., P�500 bar), more

chloride ligands become available for metal complexing as

associated sodium chloride solute species disproportionate

with decreasing temperature. Conversely, a hydrothermal ore

solution decompressing isothermally will lead to instability of

some metal complexes as chloride ligands are consumed by

association of the bulk electrolyte.

13.2.2.3 Ion Hydration, Association, and Water Activity

Of particular interest, as well, is the changing nature of ion

hydration in hydrothermal fluids as a function of temperature

and pressure. Insight into ion–solvent interaction at elevated

temperatures and pressures comprising ion hydration coordi-

nation numbers, ion–water bond lengths, and solvated water

residence times has come relatively recently from spectro-

scopic, neutron, and X-ray diffraction and computational stud-

ies, many of which are summarized and discussed by Seward

and Driesner (2004). More recently, computational studies

(ab initio, MD, and ab initio/MD) have added significantly

to our knowledge of ion hydration under hydrothermal con-

ditions, especially with respect to NaCl solutions (e.g.,

Bondarenko et al., 2006; Chialvo et al., 2009; Dong et al.,

2008; Fedotova, 2008; Guardia et al., 2006a,b; Nahtigal and

Svishchev, 2009). Driesner et al. (1998) studied (using MD)

Naþ ion hydration for both cases of infinite dilution and 1.0 m

from 30 to 590 K and at several densities. They showed that

there is a contraction (i.e., 2.26 to 2.23 A) of first-shell waters

around Naþ, which is also accompanied by a decrease in the

Figure 2


number of coordinated water molecules with increasing tem-

perature. Moreover, the residence times of first-shell waters

around Naþ and Cl� (in a 1.0 m solution) at 298 K are found

to be 22.5 and 12.0 ps, respectively, but decrease by about an

order of magnitude to 2.2 and 2.4 ps with increasing temper-

ature to 650 K (Guardia et al., 2006b).

In addition, a number of systematic X-ray absorption spec-

troscopic studies (i.e., EXAFS (extended X-ray absorption fine

structure)) of Agþ and Sr2þ hydration up to 350 �C and at

equilibrium saturated vapor pressures (Seward et al., 1996,

1999) have also demonstrated a contraction of the cation–

oxygen (water) distances with increasing temperature as well

as a decrease in the number of the first-shell water molecules. In

the case of Agþ, for example, the Agþ–oxygen (water) distance

undergoes an appreciable contraction of 0.1 A as temperature

increases from 25 to 350 �C, with an accompanying decrease in

the number of first-shell waters from 4 to 3. However, data for

Cd2þ hydration up to 300 �C and at equilibrium vapor pres-

sures indicate a decrease in the Cd2þ–oxygen (water) distance

with increasing temperature but with no change in the number

of first-shell waters (Seward et al., 2013). In the case of the

trivalent cation, In3þ, the hydration environment remains

unchanged (six first-shell waters bound to the In3þ at a distance

of 2.14 A) over the same temperature range (i.e., 25 to 300 �C)at saturated vapor pressures (Seward et al., 2000).

The nature of ion hydration (i.e., ion–solvent interaction) is

complex and varies with the changing dielectric properties of

water solvent, which depend fundamentally on the nature

and extent of hydrogen bonding, which is a function of temper-

ature and pressure. However, NaCl is the most concentrated

salt component in hydrothermal fluids in the Earth’s crust, and

its hydration behavior is therefore of importance in determining

the water activity in hydrothermal ore solutions. Changes in

the hydration of Naþ and Cl� ions as well as the formation

of associated species, such as NaCl0 (weakly hydrated), can

fundamentally affect water activity and hence the solubility of

minerals. Thus, a hydrothermal ore solution migrating through

the Earth’s crust will experience a range of physical and chemical

phenomena as it cools and decompresses (e.g., Heinrich

et al., 2004), which include element partitioning, mineral

precipitation, and changing metal complex stabilities, which

may, in some cases, be temporarily enhanced as a result of the

myriad of changes occurring in a complex multicomponent

system. As shown, for example, by the solubility of a simple

gangue mineral such as quartz in sodium chloride and other

electrolyte salt solutions (Newton and Manning, 2000, 2010;

Shmulovich et al., 2006), the dissolved silica concentration var-

ies not only with temperature and pressure but also with the salt

concentration and hence the extent and molecular nature of

association (ion pairing) and ion–solvent interaction with con-

sequent effects on water solvent activity. The changing value of

pK1 of silicic acid as a function of temperature and pressure as

well as bulk salt hydrolysis may also be of importance.

13.2.2.4 Weak Acid/Base Equilibria in HydrothermalSystems

In addition to the main electrolyte salt, NaCl, there are other

important solute components in hydrothermal ore solutions

in the Earth’s crust such as CO2, H2S, and NH3, which

are involved in various, weak acid hydrolytic equilibria and

play a role in metal transport by fluids in the Earth’s crust. CO2

may occur in high concentrations greater than 20 mol%

in H2OþCO2þNaCl fluids associated with metamorphic

reactions; however, in the majority of hydrothermal ore fluids

in the upper crust, the CO2 (as H2CO3þHCO3�þCO3

2�)would be in the range from 0.01 to 1.0 m. CO2 being a volatile

component means that carbonate equilibria and pH are sensi-

tive to phase separation and boiling in hydrothermal systems.

In active geothermal/hydrothermal systems, it is well known

that CO2 partitioning into the vapor (steam) phase during

boiling leads to an increase in the residual liquid-phase pH

by as much as two pH units (e.g., from, say, 6 to 8). This affects

the stability of metal carbonate complexes as well as the stabil-

ity of any other pH-dependent metal complex equilibria, such

as hydroxide and/or hydrosulfide complexes.

Of particular importance is H2S and its first deprotonation

equilibrium in aqueous media,

H2S ¼ Hþ þHS�

the equilibrium constant, K1, for which has been accurately

determined up to 360 �C at the equilibrium vapor pressure

(Suleimenov and Seward, 1997) as well as at 500 bar and up

to 500 �C (Stefansson and Seward, 2004). Note that the equi-

librium constant, K2, for the second ionization of H2S,

HS� ¼ Hþ þ S2�

is small (i.e., pK2>17.1 at 24 �C and >12.5 at 200 to 270 �C;Giggenbach, 1971a), and hence, the sulfide ion, S2�, is presentat negligible concentrations in natural hydrothermal solutions.

However, the hydrosulfide ligand, HS�, plays a fundamental

role in the transport and deposition chemistry of some

elements, such as gold in hydrothermal ore solutions, and

the availability of reduced sulfur is crucial for the precipitation

of sulfide ore minerals. During boiling and/or phase separa-

tion, H2S and volatile components, such as CO2, NH3, and H2,

partition into the vapor (steam) or less dense, volatile-rich

phase, which affects the stability of various metal complexes

with respect to ligand availability (e.g., H2S loss) and increas-

ing redox potential of the residual liquid (H2 loss) or more

dense supercritical fluid.

The volatile species NH3 and N2 and H2 are involved in

heterogeneous redox equilibria and ammonium ion stability

through the reactions

N2 gð Þ þH2 gð Þ ¼ NH3 gð Þ

and

NH3 aqð Þ þHþ ¼ NH4þ

The stability of simple amino complexes (involving the NH3

ligand) of transition metals under hydrothermal conditions is

still poorly known. The ammonium ion also participates in

wall rock alteration reactions in hydrothermal systems, giving

rise to the formation of ammonium-containing silicate min-

erals such as buddingtonite and tobelite.

Thus, it is against this background of changing water prop-

erties (e.g., hydrogen bonding) with associated changes in ion–

solvent interaction and ion pairing/association and changes in

water activity in the presence of ligand-supplying weak acid/

30 High-salinity

20 MPa


base systems that metal complex equilibria, metal transport

mechanisms, and ore mineral precipitation chemistry at ele-

vated temperatures and pressures relevant to hydrothermal

ore-depositing systems must be considered.

(a)

(b)

20

10

Single-phase fluid

Single-phase fluid

liquid

High-salinityliquid

50 MPa

Low-salinityvapor

Low-salinityvapor

Imm

isci

bili

ty

Imm

isci

bili

ty

Ret

rogr

ade

solu

bili

ty

3.2

10

0

10

3.2

0

wt% NaCl

wt% NaCl

50

Qua

rtz

solu

bili

ty (m

mol

kg-

1 )

Temperature (�C)

40

30

20

10

00 100 200 300 400 500

Figure 3 Isobaric quartz solubility as a function of temperature in pureH2O, H2Oþ3.2 wt% NaCl, and H2Oþ10 wt% NaCl fluids, (a) at 20 MPa(bottom) and (b) at 50 MPa; the shaded areas represent regions ofretrograde solubility. Modified from Steele-MacInnis M, Han L,Lowell RP, Rimstidt JD, and Bodnar RJ (2012) The role of fluid phaseimmiscibility in quartz dissolution and precipitation in sub-seafloorhydrothermal systems. Earth and Planetary Science Letters321/322: 139–151.

13.2.3 Mineral Solubility in Water and SaltSolutions at High Temperature and Pressure

The main interests of geochemists investigating hydrothermal

processes, including those of ore formation, are the mecha-

nisms that control the transport of elements in aqueous solu-

tions and the saturation of the solutions with minerals

containing them. At any given set of physicochemical condi-

tions, the concentration of an element in an aqueous solution

is limited by the solubility of the least soluble mineral contain-

ing this element. For example, in the system Na–Ca–F–H2O,

the concentration of fluorine is determined by the solubility of

fluorite (CaF2). Consequently, the mineral villiaumite (NaF),

which is much more soluble than fluorite, will not form unless

the buffering capacity of fluorite is exhausted, that is, the con-

centration of Ca is so low that the solubility product for the ions

comprising fluorite is less than that for villiaumite. In this

section, a brief review is provided of the data available on the

solubility of the major gangue (non-ore) minerals commonly

encountered in hydrothermal ore deposits.

Hydrothermal gangue minerals can be subdivided into two

groups, namely, those that display prograde solubility, that is,

their solubility increases with increasing temperature, and

those that display retrograde solubility, that is, decrease in

solubility with increasing temperature. Most chloride, oxide,

sulfide, and silicate minerals exhibit prograde solubility,

whereas sulfates, carbonates, and phosphates commonly

show retrograde solubility. These minerals can be further clas-

sified according to whether their dissolution is congruent or

incongruent. In the case of the former, all products of the

dissolution reaction are transferred stoichiometrically to the

solution. By contrast, incongruent dissolution involves forma-

tion of both aqueous species and new solids. Simple com-

pounds such as chlorides and oxides generally display

congruent dissolution, whereas complex silicates and sulfides

typically undergo incongruent dissolution.

Quartz is without question the gangue mineral most com-

monly encountered in ore-forming hydrothermal systems and

has been the subject of large numbers of studies by both

experimentalists and theoreticians. In pure water, the solubility

of quartz increases systematically with temperature up to

�370 �C, but at higher temperature, its solubility is retrograde;

quartz solubility increases with pressure at all temperatures.

The presence of alkali salts, such as NaCl and KCl, reduces the

solubility of quartz (salting out) at low temperature but in-

creases it at temperatures above 250 �C (Figure 3). However,

an understanding at a molecular level of how the presence of

an electrolyte salt, such as NaCl, affects quartz solubility over a

wide range of temperature and pressure, as illustrated by the

experiments of Newton and Manning (2000), eludes us at

present. It will necessarily involve a complex interplay of sol-

vent interaction with both ionic and molecular species as well

as the effect of temperature and pressure on ion pairing/asso-

ciation, salt hydrolysis, and the ionization of silicic acid.

The solubility of quartz in pure water is controlled by a

number of hydrated ‘SiO2’ species of which silicic acid,

H4SiO40 (�Si(OH)4

0), and its corresponding anions, H3SiO4�

and H2SiO42�, are the most important, although H2SiO4

2� is a

negligible species in hydrothermal fluids in the Earth’s crust.

Neutrally charged species predominate at pH<9, and therefore,

the solubility of quartz is independent of pH, except at condi-

tions of extreme alkalinity. At these conditions, the formation of

ionized species of silicic acid causes quartz solubility to increase

with increasing pH and, in salt solutions, this may be further

enhanced by the formation of ion pairs such as NaH3SiO40

(Seward, 1974a,b). Aqueous silica is also known to form poly-

mers (Chan, 1989; Iler, 1979), as illustrated by the following

stepwise reactions (e.g., Newton and Manning, 2010):

SiO2 þ 4H2Oquartz

¼ Si OHð Þ4 aqð Þsilica monomer

and

2Si OH4ð Þ aqð Þsilica monomer

¼ Si2O OHð Þ6 aqð Þ þH2Osilica dimer

Figure 3


An understanding of the nature and kinetics of silica poly-

merization as a function of temperature, pressure, pH, and salt

concentration (ionic strength) is of considerable importance in

the exploitation of active geothermal systems for energy use

(Gunnarsson and Arnorsson, 2005; Rothbaum and Rhode,

1979), particularly in managing the reservoir environment

when hot, ‘flashed,’ fluids (oversaturated with respect to amor-

phous silica) are reinjected. However, the chemistry of silica

polymerization, especially in the presence of aluminum and

salts (e.g., NaCl and KCl), under more extreme hydrothermal

conditions is still poorly known.

Another very common gangue mineral is calcite. Studies

of its solubility in hydrothermal fluids date back to the

publication of Ellis (1963) and have continued to the present

(e.g., Duan and Li, 2008a). In all cases and for most crustal

conditions, these studies have shown that calcite exhibits ret-

rograde solubility, that is, its solubility decreases in pure water

with increasing temperature. However, its solubility increases

with increasing pressure. Calcite solubility also increases with

increasing NaCl concentration up to a molality of �1–1.5

(Figure 4); however, it decreases at higher NaCl concentration

(Duan and Li, 2008; Newton and Manning, 2002). Neverthe-

less, at deep crustal conditions (i.e., at very high temperature

and pressure), the solubility of calcite is prograde, even at

relatively high salinity (Newton and Manning, 2002). The

aqueous species controlling the solubility of calcite are the

aquated ion and the ion pairs, Ca2þ, CaHCO3þ, and CaCO3

0

(the list also includes CaClþ and CaCl20 in chloride-containing

solutions), and the carbonic acid species, CO2(aq), HCO3�,

and CO32�. The hydrolytic species, CaOHþ, may also form,

but its contribution to calcite solubility is likely to be

insignificant in CO2-saturated solutions.

The other important carbonate mineral in ore-forming hy-

drothermal systems, particularly in systems forming MVT de-

posits, is dolomite. The behavior of this mineral at elevated

temperature, however, has attracted considerably less attention

from experimentalists than has calcite. A number of studies

were conducted on the solubility of dolomite at ambient tem-

perature during the 1950s and 1960s (e.g., Garrels et al., 1960)

but reported values for its solubility product differed by up to

three orders of magnitude. The probable reason for this has

since been revealed by experimental studies of the kinetics of

dolomite dissolution (Arvidson and Mackenzie, 1999;

6.0

5.0

4.0

3.0

Ca

(mm

ol k

g-1 )

NaCl (mol kg-1)

2.0

1.0

0.00 0.2

300 �C, 500 bar

0.4 0.6 0.8

Fluorite

Calcite

Anhydrite

1

Figure 4 The solubility of anhydrite, calcite, and fluorite as a function ofNaCl molality at 300 �C and 500 bar; calculated using thermodynamicdata from Holland and Powell (1998) and Shock et al. (1997).

Sherman and Barak, 2000). These studies showed that the

CaCO3 and MgCO3 components of dolomite dissolve at con-

siderably different rates. Consequently, evaluations of the

solubility of dolomite in aqueous solutions are now based on

calorimetric determinations of its enthalpy, entropy, and heat

capacity (e.g., Hemingway and Robie, 1994; Navrotsky and

Capobianco, 1987).

Sulfate minerals, such as anhydrite (CaSO4) and its hy-

drated form, gypsum (CaSO42H2O), and, to a lesser extent,

barite (BaSO4), are also important in some types of hydrother-

mal ore deposits, notably porphyry copper and molybdenum

deposits (anhydrite) and VMS deposits (anhydrite and barite).

The solubility of gypsum, which is significantly lower than that

of calcite, increases to a maximum of 0.015 m at �40 �C in

pure water and then decreases with further increases in tem-

perature (Booth and Bidwell, 1950). At 97 �C, gypsum dehy-

drates to anhydrite, generally via formation of an intermediate

hemihydrate phase, CaSO4�0.5H2O; anhydrite displays ret-

rograde solubility. As is the case for calcite, the solubility of

gypsum and anhydrite increases with increasing pressure. Sim-

ilarly, their solubility increases with addition of NaCl up to a

concentration of �1–2 m (Figure 4) and decreases at higher

NaCl concentration (Blount and Dickson, 1973; Raju and

Atkinson, 1990). Anhydrite also displays pronounced pro-

grade solubility in saline fluids at high temperature and pres-

sure. The main aqueous species controlling the solubility of

gypsum and anhydrite are Ca2þ, CaSO40, CaHSO4

þ, CaClþ,CaCl2

0, HSO4�, and SO4

2�.The behavior of barite (BaSO4) is similar to that of gypsum,

except that it is considerably less soluble. The solubility of

barite in pure water increases with temperature to a maximum

of 0.00002 m at �150 �C and, thereafter, decreases with in-

creasing temperature. With the addition of NaCl and/or KCl,

this maximum shifts to higher temperature. Pressure has a

similar effect, increasing the solubility of barite by a factor of

3 from 100 to 1000 bar (Blount, 1977). The speciation con-

trolling barite solubility is effectively the same as that for

gypsum and anhydrite, except that Ca is replaced by Ba.

Fluorite (CaF2) is by far the most commonly occurring

fluorine-containing mineral and is an important gangue min-

eral in MVT deposits, as well as in rare earth element (REE)

deposits, where it shows a clear spatial association with the

REE mineralization (Williams-Jones et al., 2000). In pure

water, fluorite is relatively insoluble, with a maximum solubil-

ity of 0.0002 m, which is reached at a temperature of 100 �C at

swvp (Strubel, 1965). Thus, like gypsum, its solubility is pro-

grade at low temperature and retrograde at higher temperature;

however, its maximum solubility is only about 10% of that of

gypsum. With the addition of alkali chlorides, fluorite solubil-

ity increases (Figure 4), and in a 2 m NaCl solution, its solu-

bility is prograde at all temperatures. The solubility of fluorite,

like that of the carbonate and sulfate minerals, also increases

with pressure (Tropper and Manning, 2007). Species control-

ling the solubility of fluorite comprise Ca2þ, CaClþ, CaCl20,

HF, and F�.From the preceding paragraphs, it is apparent that the

effects of temperature, pressure, and salinity on the solubility

of calcite, gypsum, anhydrite, and fluorite are quite similar,

although the absolute solubility of these minerals varies con-

siderably at comparable conditions. All these minerals display

Figure 4

250 �C, 1 m NaCl,pH 4.5

Pyrite

Magnetite

Pyrrhotite

log aH2S aq

log

a H2

aq

Hem

atite

-8-6-4

-4

-3

-2

-2

-2-4-60

-2

-4

-6

-8

-10

0

-3-4

Figure 5 Solubility contours (logm SFe) for pyrite, pyrrhotite,magnetite, and hematite as a function of log aH2S aq and log aH2 aq in a 1 mSCl aqueous solution at pH 4.5, a temperature of 250 �C, and vapor-saturated water pressure; calculated using thermodynamic data fromHolland and Powell (1998) and Shock et al. (1997).


retrograde solubility in dilute aqueous solutions, albeit over

quite different temperature intervals. They also all revert to

prograde solubility with the addition of alkali chlorides, and

they all increase in solubility within increasing pressure. The

main reason for their similar behavior is that these minerals

only differ in the nature of the ligand that binds with Ca2þ (i.e.,

CO32�, SO4

2�, or F�). These are all hard ligands having high

charge/radius ratios and are strong electron acceptors, and

therefore, their bonding with Ca2þ is dominantly ionic. More-

over, as the relative order of hardness for these ligands is F� >

CO32�>SO4

2�, then Ca2þ should form its strongest bonds

with these ligands in the same order, and consequently, at a

given set of conditions, anhydrite should have the highest

solubility followed, in turn, by calcite and fluorite. This is

what is observed (Figure 4).

Pyrite (FeS2) is the most important sulfide gangue mineral

and occurs in virtually all major hydrothermal base and pre-

cious metal mineral deposits, except those subjected to high-

grade metamorphism, in which pyrrhotite has commonly

replaced pyrite. Furthermore, pyrite is the principal sulfide

mineral in most of these deposits, and its concentration is

generally much greater than that of the sulfide minerals con-

taining the ore metals of interest. Numerous studies have

investigated the dissolution of pyrite and its deposition from

aqueous solutions (Ohmoto et al., 1994; Schoonen and

Barnes, 1991a,b). These studies have shown that the solubility

of pyrite is prograde and depends strongly on pH.

Formally, the dissolution of sulfide minerals can be de-

scribed by reactions of the type

MeS ¼ Me2þ þ S2�

but as hydrothermal solutions invariably contain insignificant

concentrations of S2� (Ellis and Giggenbach, 1971; Migdisov

et al., 2002), the dissolution of these minerals is more conve-

niently described by reactions involving H2S and HS�, thedominant sulfide species in solution, at acidic and near-neutral

to alkaline conditions, respectively. As mentioned earlier, pK2

for H2S is �17 at 25 �C. However, the sulfur in most sulfide

minerals is in the �2 state, whereas in pyrite, the sulfur occurs

in both the zerovalent and�2 state as the S0.S2� ion (i.e., as the

simple S22� polysulfide ion). The solubility of pyrite is there-

fore most appropriately described by a reaction involving the

polysulfane, H2S2, but as the stability of polysulfanes and

especially their nondissociated forms is poorly known at ele-

vated temperature, the dissolution of pyrite is normally

described by redox reactions involving H2S and HS�, that is,

FeS2 þ 2Hþ þH2O ¼ Fe2þ þ 2H2Sþ 0:5O2

or

FeS2 þH2O ¼ Fe2þ þ 2HS� þ 0:5O2

The dissolution of pyrite is incongruent and, depending on

the system, may proceed via complex reactions involving for-

mation of elemental sulfur or iron oxides (e.g., Caldeira et al.,

2010; Rimstidt, 2003). The process of pyrite precipitation is

also complex and, below 300 �C (and perhaps at higher

temperature), proceeds via intermediate steps involving forma-

tion of precursor iron monosulfide, mackinawite (FeS), and

the thiospinel, greigite (Fe3S4). In the presence of elemental

sulfur, thiosulfate, or polysulfides, the rate of conversion of

these intermediate phases to pyrite is extremely fast and may

occur within minutes, whereas with only hydrogen sulfide

or bisulfide present, the conversion rate is much slower

(Schoonen and Barnes, 1991a).

One of the important roles of pyrite in hydrothermal sys-

tems is to buffer parameters such as fO2 (or fH2) and pH

through reactions with minerals like pyrrhotite, magnetite,

and hematite. For example, the assemblage pyrite–pyrrhotite

constrains oxygen fugacity under acidic conditions through the

reaction

1� xð ÞFeS2 þ 1� xð ÞH2O ¼ Fe1�xSþ 1� xð ÞH2S

þ 0:5 1� xÞO2ðand the assemblage pyrite–pyrrhotite–magnetite buffers both

fO2 and pH, in both cases for a given total concentration of

aqueous sulfur (Figure 5).

The solubility of most silicate minerals increases with tem-

perature in aqueous solutions, that is, it is prograde, and de-

pends strongly on pH. In most cases, these minerals dissolve

incongruently, and the nature of the phase formed depends

strongly on the composition of the solution, including its pH.

For example, at relatively low pH and/or low activity of Kþ, thedissolution of K-feldspar may involve formation of kaolinite

(or another aluminosilicate phase) through the reaction

2KAlSi3O8 þH2Oþ 2Hþ ¼ Al2Si2O5 OHð Þ4þ 4SiO2 aqð Þ þ 2Kþ

Figure 5

100 �C, swp

K-feldspar

Muscovite

Kaolinite

log aSiO2

log

a K+

a H+

6543210-1-2

-4 -3 -2 -1 0 1

Figure 6 Stability relationships among muscovite, kaolinite, andK-feldspar as a function of log aSiO2 aq and log aKþ=aHþ

� �at 100 �C and

vapor-saturated water pressure; calculated using thermodynamic datafrom Holland and Powell (1998) and Shock et al. (1997).

Table 3 Ligands occurring in hydrothermal ore solutions

Hydroxide OH�

Halide ions F�, Cl�, Br�, I�

Sulfur species HS�, Sn�, SnS

2�, SO32�, S2O3

2�, SO42�

Ammonia (ammine) NH3

Oxyanions CO32�, PO4

3�, AsO33�, SbO3

3�, MoO42�,

WO42�, SiO4

4�

Thioanions AsS33�, SbS3

3�, MoS42�, WS4

2�

Carboxylates CH3COO� (acetate), C2H5COO

� (propionate),CH2(COO)2

2� (malonate), (COO)22� (oxalate)

Miscellaneousligands of possibleinterest

HTe�, Te22�, CN�, SCN�


By contrast, at higher pH and/or higher activity of Kþ, the disso-lution reaction leads to the formation of muscovite (or illite):

3KAlSi3O8 þ 2Hþ ¼ KAl3Si3O10 OHð Þ2 þ 6SiO2 aqð Þ þ 2Kþ

Both reactions also depend on the activity of aqueous silica

(Figure 6).

Because of their importance as rock-forming minerals,

many of these silicates play an important role in buffering

fluid composition, and analysis of hydrothermal equilibria

involving them may help in evaluating compositional param-

eters such as pH. As hydrothermal fluids commonly saturate

with respect to quartz, aqueous silica activity is typically buff-

ered, and reactions involving K-feldspar, muscovite, and an

aluminosilicate mineral can be rewritten as follows:

3K-feldspar þ 2Hþ ¼ Muscoviteþ 6Quartzþ 2Kþ

2Muscoviteþ 3H2Oþ 2Hþ ¼ 3Kaoliniteþ 2Kþ

Thus, the coexistence of K-feldspar with muscovite in the

presence of quartz or muscovite with kaolinite buffers the Kþ/Hþ activity ratio and, if albite is also present, the Kþ activity can

be deduced from fluid inclusion salinity and the Kþ/Naþ ratio

corresponding to equilibrium between K-feldspar and albite,

thereby permitting evaluation of pH (see also the discussion in

Chapter 13.1). At given concentrations of Kþ or Naþ, hydro-lysis of alkali feldspar tomuscovite and paragonite, or the latter

minerals to an aluminosilicate mineral, buffers pH to a fixed

value that depends only on temperature and pressure (e.g.,

Hemley et al., 1992).

13.2.4 Ore Metal Transport and Deposition

13.2.4.1 Ore Fluids with Liquid-Like Densities

13.2.4.1.1 Ligands in hydrothermal ore solutionsOre-forming hydrothermal fluids migrate through the Earth’s

crust transporting metals as complex ions and molecules.

Focused deposition of metals to form an ore deposit occurs

in response to physicochemical changes in the local environ-

ment. In essence then, ore deposition as a manifestation of the

stability/instability or disproportionation of metal complexes

at a given temperature and pressure (or T–P range) may be

considered. A detailed knowledge of the stability and

stoichiometry of metal complexes with available ligands as a

function of temperature, pressure, pH, and redox potential

comprises the fundamental basis for understanding the pro-

cesses of hydrothermal ore formation.

What ligands are present in hydrothermal ore solutions?

The three most important are Cl�, HS�, and OH�, althoughvarious other electron donors are also present (Table 3) and

may be important in the complexing and transport of some

metals and under certain conditions. In natural systems, ligand

availability is important. For example, chloride is always pre-

sent at much higher concentrations than, say, iodide, and

hence, in a 1.0 m chloride ore fluid at 500 �C and 500 bar,

chloride would be the more important complexing ligand for

gold transport even though the gold(I) iodide complex, AuI2�,

is more stable than AuCl2�.

Ligand availability and stability may be affected by various

homogeneous reactions occurring within the hydrothermal

ore solution itself. The activity of free chloride ion in an aque-

ous NaCl solution is much diminished with increasing tem-

perature due to ion pairing as discussed earlier. A 1.0 m NaCl

solution at 400 �C and 300 bar will be >99% associated,

which will determine the number of coordinated chlorides in

metal complexes. In addition, phase separation and boiling

may dramatically decrease the activity of compounds such as

H2S, H2Te, H2Se, NH3, and CO2 in the residual liquid (or

denser supercritical) phase, effectively removing the corre-

sponding weak acid/base ligands. Changes in the redox envi-

ronment that accompany boiling and/or mixing with

oxygenated fluids in the upper parts of geothermal/hydrother-

mal systems also lead to ligand destruction by oxidation as in

the case of H2S/HS�, which reacts to sulfate and other inter-

mediate oxidation state species of sulfur. Mineral deposition

(e.g., sulfide precipitation) also causes H2S/HS� loss, especially

in the ore-depositing environment where a cascade of interre-

lated chemical events leads to the almost complete deposition

of various metals from the ore fluid. In addition, some ligands

are not thermally stable. Acetate (CH3COO�) is well known to

decarboxylate at elevated temperatures (Bell and Palmer,

1994), and thiosulfate (S2O32�) disproportionates at elevated

temperatures Giggenbach, 1974b).

Finally, it is noted that complex stabilities differ signifi-

cantly because of the differing metal–ligand bonds. The differ-

ing complex stabilities can be predicted using hard/soft Lewis

acid/base theory (Ahrland, 1968; Pearson, 1963) as discussed

by Seward and Barnes (1997) with reference to hydrothermal

Figure 6


metal complexing. Thus, a soft electron donor such as HS�

would be predicted to form very stable complexes with Auþ (a

soft electron acceptor), whereas the fluoride (hard Lewis base/

electron donor) complexes of Auþ are much less stable and the

first isolatable Au(I) fluoride complex has only recently been

synthesized (Laitar et al., 2005), although fluoride complexes

with higher oxidation states of gold are well documented

(Mohr, 2004). This approach is simplistic compared to the

modern quantum chemical assessment of metal complex ge-

ometries and metal–ligand bond energies of complexes of

geochemical interest (e.g., Boily and Seward, 2005; Sherman,

2010) but facilitates an overview of the important metal

transport mechanisms in the multicomponent electrolyte

solutions, which comprise the diversity of hydrothermal ore

solutions.

13.2.4.1.2 Metal chloride complexingSodium chloride is the ubiquitous salt in hydrothermal ore

solutions, and hence, the stability of chloride complexes of

many elements at high temperatures and pressures plays a

fundamental role in the transport of these elements by migrat-

ing ore fluids throughout the Earth’s crust and their subsequent

precipitation as ore minerals in an ore-depositing environment

at a given temperature and pressure. It has long been known

that chloride may form a stepwise series of complexes with

many metal cations at ambient temperature as defined by

Mnþ þmCl� ¼ MClmn�m

0

100

80

60

% S

n

log Cltotal / mol kg-1

40

20

0-3 -2 -1 0

1

2-o

4

0

1

100

80

60

% C

o

Cltotal / mol kg-1

40

20

00 21 3 4 5 6

12 3 4

25 �C

35 �C

Figure 7 Stepwise formation of chloridotin(II) and chloridocobalt(II) complecurves are labeled according to the number of chloride ligands bound to eachcomplexes, 2-o refers to octahedrally coordinated CoCl2(H2O)4 and 2-t refersSeward TM (2001) Spectrophotometric determination of the stability of tin(IICosmochimica Acta 65: 4187–4199; Liu W, Borg SJ, Testemale D, Etschmannproperties for cobalt chloride complexes in hydrothermal fluids at 35–440 �C1227–1248.

but, until relatively recently, few studies extended to higher

temperatures. Most earlier studies have been discussed/sum-

marized by Seward (1981) and Seward and Barnes (1997).

Sverjensky et al. (1997) tried to overcome the paucity of reli-

able, experimentally based thermodynamic data for metal

complex equilibria at extreme conditions with computational

predictions extending to 1000 �C and 5000 bar. Studies of

chloridosilver(I) and chloridolead(II) stepwise complex for-

mation up to 350 �C and 300 �C, respectively (Seward, 1976,

1984), demonstrated that the higher coordination number

species, such as AgCl43� and PbCl4

2�, are not observed at

elevated temperatures (e.g., 350 �C) and equilibrium saturated

vapor pressures, mainly because of the low chloride ion activity

due to NaCl association. Several recent studies (Figure 7) reaf-

firm earlier observations concerning stepwise complex forma-

tion as well as demonstrating the coordination geometry

changes that may occur at elevated temperatures, as discussed

in some detail by Seward (1981) and Seward et al. (1996). In

the case of complexing of tin(II) with chloride (Muller and

Seward, 2001; Figure 7), the distribution of species changes

and the SnCl42� species is not observed in the more concen-

trated chloride solutions up to 3.0 m at 300 �C at the equilib-

rium saturated vapor pressure.

Coordination geometry changes attending stepwise com-

plex formation and changes in temperature and pressure

have long been known (Seward, 1981), especially in the case

of cobalt(II) and nickel(II) species in aqueous solution (e.g.,

Ludemann and Franck, 1967). Recently, the stepwise formation

1

2-t

4

100

80

60

% S

n

log Cltotal / mol kg-1

40

20

0-3 -2 -1 0 1

100

80

60

% C

o

Cltotal / mol kg-1

40

20

00 21 3 4 5 6

350 �C

300 �C0

1

2

3

xes at ambient and elevated temperatures at saturated vapor pressure; themetal cation (i.e., Sn2þ or Co2þ). In the case of the cobalt(II) chlorideto tetrahedrally coordinated CoCl2(H2O)2. Modified from Muller B and) chloride complexes in aqueous solution up to 300 �C. Geochimica etB, Hazemann J-L, and Brugger J (2011) Speciation and thermodynamic

and 600 bar: An in-situ XAS study. Geochimica et Cosmochimica Acta 75:

Figure 7


of cobalt(II) complexes and their octahedral/tetrahedral con-

figurational changes have been revisited by Liu et al. (2011)

(Figure 7), who demonstrated the predominance of tetrahe-

drally coordinated moieties, including CoCl42�, in expanded

high-temperature water. Of particular interest as well is the

nature of complex geometry and the manner in which the

solvent molecules interact with metal complex ions/molecules

to affect metal–ligand bonds and complex stability. In the case

of stepwise complex formation between Cd2þ and Cl�, forexample, solvent molecules (i.e., H2O) contribute fundamen-

tally to the stability of the complex (Seward et al., 2013). As

shown in Figure 8, the Cd–O (water) and Cd–Cl distances

obtained from EXAFS measurements are 2.30 and 2.50 A,

respectively, for the octahedral CdCl(H2O)5þ complex, but

these values differ from the quantum chemically calculated

values with only five first-shell water molecules bound to the

Cd2þ. However, with stepwise addition of solvatedwaters to the

second shell, one observes that the calculated Cd–O distance

asymptotically approaches the valuemeasured by X-ray absorp-

tion spectroscopy when n�4. However, at least eight second-

shell waters are necessary to reproduce the measured Cd–Cl

distance. These concepts are of considerable importance in

understanding the partitioning ofmetal complex ions andmol-

ecules to a steam-phase or low-density supercritical fluid in

hydrothermal systems. Metal complexes in such lower-density

fluids are also hydrated moieties, which comprise molecular

aggregates or clusters, as discussed in the succeeding text.

In recent years, further advances in our understanding of

metal chloride complexing at hydrothermal conditions have

come from solubility measurements, UV/Vis and X-ray absorp-

tion spectroscopy, and ab initio and ab initio/MD calculations.

[CdCl(H2O)5(H2O)2]+ [CdCl(H2O)5(H2O)3]+[CdCl(H2O)5(H2O)1]+

[CdCl(H2O)5(H2O)5]+

MP2 optimized geometries

Stepwise hydration of

[CdCl(H2O)5(H2O)6]+[CdCl(H2O)5(H2O)4]+

Figure 8 EXAFS and ab initio data for cadmium(II) chloride and oxygen (wat(H2O)5

þ; the EXAFS-determined Cd–O (water) and Cd–Cl distances are givendistances for the optimized geometries resulting from the addition of each secoHenderson CMB, and Charnock JM (2013) An X-ray absorption spectroscopichlorocadmium(II) complexing in hydrothermal solutions. Chemical Geology

These include studies of the chloride complexes of Ti4þ

(Ryzhenko et al., 2006), Mn2þ (Suleimenov and Seward,

2000), Fe2þ and Fe3þ (Liu et al., 2007; Stefansson et al.,

2008; Testemale et al., 2009), Co2þ (Liu et al., 2011; Migdisov

et al., 2011a,b), Ni2þ (Tian et al., 2012), Cuþ and Cu2þ

(Etschmann et al., 2010; Fulton et al., 2000; Liu et al., 2002,

2008; Sherman, 2007; Xiao et al., 1998), Zn2þ (Anderson et al.,

2000; Harris et al., 2003; Trevani et al., 2009), Cd2þ (Bazarkina

et al., 2010; Palmer et al., 2000; Seward and Driesner, 2004;

Seward et al., 2013), In3þ (Seward et al., 2000), Sn2þ and Sn4þ

(Kovalenko and Ryzhenko, 1997; Muller and Seward, 2001;

Sherman et al., 2000a,b; Uchida et al., 2002), Sb3þ (Oelkers

et al., 1998; Pokrovski et al., 2006), Pd2þ (Boily and Seward,

2005; Seward and Driesner, 2004), Auþ and Au3þ (Gammons

et al., 1997; Stefansson and Seward, 2003a,b,c; Usher et al.,

2009), Tlþ (Bebie et al., 1998), rare earth(III) elements except

Pm3þ (Gammons et al., 2002; Mayanovic et al., 2002, 2007,

2009; Migdisov and Williams-Jones, 2002, 2006; Migdisov

et al., 2008, 2009; Stepanchikova and Kolonin, 2005), Y3þ

(Ragnarsdottir et al., 1998), and U4þ (Kovalenko et al.,

2011). All these studies provide thermodynamic data and/or

molecular insight into complex energetics, geometries, and

metal–ligand bond lengths, the aim being to understand fun-

damental aspects of metal transport, metal (complex) parti-

tioning during fluid phase separation, and ore mineral

precipitation in ore-forming hydrothermal systems.

13.2.4.1.3 Complexing with other halide ligandsWith the exception of Gammons and Yu (1997) and Liu et al.’s

(2012) studies of bromide and iodide complexes of silver and

zinc at elevated temperatures, there have been few studies of

2.65

0 1 2 3

n, Solvation number

Cd–OEXAFS bulk, Cd–O, 298 K

EXAFS bulk, Cd–Cl, 298 K

[CdCl(H2O)5(H2O)n]+

CdCl(H2O)5+

Cd

–Cl (

Å)

Cd

–O (Å

)

Cd–Cl

HF/MSI*, Butterworth92

MP2/aug-cc-pVDZ(PP), this studyMP2/aug-cc-pVTZ(PP), this study

4 5 6

2.60

2.55

2.50

2.45

2.40

2.35

2.30

2.25

2.20

er) distances for the octahedral monochloridocadmium(II) complex, CdClby the horizontal dashed lines. The quantum chemically calculatednd-shell water are also shown. Reproduced from Seward TM, Lemke KH,c and ab initio computational study of the Cd(II) aquated ion and(in review).

Figure 8


metal bromide and iodide complexes under hydrothermal

conditions from which thermodynamic data have been

obtained. In near-magmatic to epithermal hydrothermal

fluids, bromide and iodide complexes are not considered to

play a significant role in ore transport. Nevertheless, bromide

and iodide complexes may play a role in the volatile metal

complex chemistry in magmatic/volcanic gas systems and tran-

sition metal transport chemistry. Various sulfide/halide subli-

mate minerals containing bromide and iodide occur in

fumarolic deposits of active volcanoes, such as at Vulcano in

Italy (Demartin et al., 2010). Bromide and iodide complexes

may play a role inmetal transport in sedimentary basinal brines,

but this has been little studied. Bromide interaction with some

metal cations has been studied at hydrothermal conditions

using X-ray absorption spectroscopy (Ni2þ: Hoffmann et al.,

1999; Wallen et al., 1998; Zn2þ: Anderson et al., 2000; Liu

et al., 2012; Simonet et al., 2002). These latter studies emphasize

the tendency toward tetrahedrally coordinated species, such as

ZnBr2(aq) (i.e., ZnBr2(H2O)20) and ZnBr4

2�, with increasing

temperature, as shown also by the high-temperature (up to

500 �C) Raman data of Mibe et al. (2009).

However, fluoride plays an important role in the transport

of some metals/elements such as titanium, tin, rare earths,

zirconium, and uranium in hydrothermal ore solutions. In

the case of titanium, for example, the solubility of rutile in

high-temperature hydrothermal fluids is much enhanced by

the presence of fluoride (Rapp et al., 2010) due to the complex-

ing of Ti4þ by F� (Ryzhenko et al., 2006) and the formation of

mixed hydroxofluoro -complexes, such as TiF(OH)30,

TiF2(OH)20, and TiF(OH)4

�. Kovalenko et al. (1992) and

Kovalenko and Ryzhenko (1997) have demonstrated the im-

portance of tin(II) fluoride complexes, such as SnF(OH)2� and

SnFCl0, in hydrothermal fluoride solutions at 500 �C and

1 kbar, but more experimentally based thermodynamic data

are required in order to adequately understand the hydrother-

mal transport and deposition chemistry of tin by fluids in the

Earth’s crust. Recent experimental studies (Migdisov et al.,

2011a,b; Prisyagina et al., 2008; Ryzhenko et al., 2008) have

also demonstrated that the zirconium(IV) hydroxofluoride

complexes are important in determining the transport (i.e.,

the mobility) of zirconium (and the solubility of zircon and

baddeleyite) in hydrothermal fluids in the Earth’s crust. The

fluoride complexes of the rare earth(III) elements (except

Pm3þ) have been studied using solubility and spectrophoto-

metric methods by Migdisov and Williams-Jones (2007) and

Migdisov et al. (2009) up to 300 �C at the equilibrium vapor

pressure. Previous theoretical predictions had overestimated

the stability at elevated temperatures of species such as NdF2þ

and hence its importance in hydrothermal fluids relative to the

chloride complexes, NdCl2þ and NdCl2þ. The values of the

equilibrium constants for the formation of the monofluorido-

and monochlorido-complexes of the rare earth(III) elements

are summarized in Figure 9. More data are required to under-

stand REE transport and deposition chemistry by hydrother-

mal fluids at T>350 �C. Kovalenko et al. (2012) have

measured the solubility of uraninite (UO2) in HF solutions at

500 �C and 1000 bar and identified the hydroxofluoride spe-

cies, U(OH)3F0 and U(OH)2F2

0, thus emphasizing the proba-

ble importance of fluoride complexing of U4þ in hydrothermal

transport, but more experimental studies are required before

an adequate understanding of the hydrothermal transport and

deposition chemistry of uranium is available. The uranyl ion,

UO22þ, also plays a role in hydrothermal uranium transport in

environments of intermediate redox potential.

13.2.4.1.4 Metal complexes with hydroxide and otheroxygen electron donor ligandsMany metal ions hydrolyze in high-temperature aqueous solu-

tions in the pH range encompassed by natural hydrothermal

fluids in the Earth’s crust, and hence, the hydroxide complexes

can play an important role in the chemistry of hydrothermal

ore transport. A detailed overview of metal ion hydrolysis

speciation is given by Baes and Mesmer (1976), and

Wesolowski et al. (2004) provide an erudite summary of

metal ion hydrolysis and metal oxide solubility measurements

at elevated temperatures. It is well known that the solubility of

metal oxides in aqueous media over wide ranges of tempera-

ture and pressure is pH-dependent (Wesolowski et al., 2004),

and as in the case of rutile solubility at 300 �C, for example, the

minimum solubility in the near-neutral region and the overall

shape of the solubility curve is due to the hydrolysis of Ti4þ

with the formation of various Ti(OH)n4�n species (Knauss

et al., 2001) according to the reaction

TiO2 rutileð Þ þ 4� nð ÞHþ þ n� 2ð ÞH2O ¼ Ti OHð Þ44�n

The elevated solubility of rutile in supercritical water (up to

1000 �C and 2.3 GPa) in the presence of albite (Audetat and

Keppler, 2005; Hayden and Manning, 2011; Manning et al.,

2008) indicates that dissolved Ti(IV) forms some sorts of ad-

ducts with dissolved aluminum silicate species in solution.

Iron occurs ubiquitously in hydrothermal ore deposits as

sulfide, oxide, silicate, and carbonate minerals. The hydrother-

mal ore transport chemistry of Fe(II) and Fe(III) is dominated

by chloride complexing (see the preceding text) and by hydro-

lysis equilibria, which have been extensively studied up to

300 �C (Wesolowski et al., 2004). Recently, the formation

of Fe(OH)2þ has been studied spectrophotometrically by

Stefansson et al. (2008) up to 300 �C, and their thermody-

namic data are in good agreement with the few earlier reported

data. Unfortunately, there are essentially no experimentally

derived thermodynamic data for iron complexing and hydro-

lysis at higher temperatures from 350 �C to supercritical

conditions, which limits our ability to model the transport

and precipitation chemistry of iron in many ore-depositing

environments in the Earth’s crust.

Significant gallium enrichments occur in the metal-

enriched surface precipitates of some active geothermal sys-

tems (Krupp and Seward, 1987), but the aqueous chemistry

of gallium (predominantly as Ga3þ) has been little studied at

elevated temperatures and pressures, and the transport chem-

istry by hydrothermal fluids in the Earth’s crust is poorly

known. Benezeth et al. (1997) measured the solubility of

a-GaOOH from 150 to 250 �C and estimated the equilibrium

formation constants for Ga(OH)n3�n from 25 to 300 �C.

The hydrolysis Nd3þ has been studied by solubility and

potentiometric methods up to 290 �C by Wood et al. (2000),

but otherwise, there are few experimentally based data available

on REE hydroxide complexes under hydrothermal conditions.

9

8

7

6

Log

bF 1Lo

g bC

l 15

4La Ce Pr Nd Pm Sm Eu

150 �C

150 �C

200 �C

200 �C

250 �C

250 �C

250 �C

200 �C

150 �C

Gd Tb Dy Ho Er Tm Yb Lu

La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

4

3.5

2.5

1.5

0.5

3

2

1

Figure 9 The equilibrium cumulative formation constants for the monofluorido- and monochlorido-complexes of the rare earth elements inaqueous solution up to 250 �C and at the saturated vapor pressure. Modified from Migdisov AA, Williams-Jones AE, and Wagner T (2009)An experimental study of the solubility and speciation of rare earth(III) elements in fluoride and chloride bearing aqueous solutions at temperaturesup to 300�C. Geochimica et Cosmochimica Acta 73: 7087–7109.


Hydrothermal molybdate chemistry has been studied by

Ulrich and Mavrogenes (2008) who measured the solubility

of Mo and MoO3 in water and KCl solutions from 500 to

800 �C and to 3000 bar and suggested (with the aid of

XANES spectra) that solubility was due to the simple molybdic

acid species as well as to chloromolybdate species (e.g.,

MoOmCln2�) (see also the thioaniondiscussion in the succeeding

text). Minubayeva and Seward (2010) used ultraviolet spectros-

copy to obtain the first thermodynamic data for molybdic acid

ionization up to 300 �C at the equilibrium saturated vapor

pressure, that is,

H2MoO4 ¼ HMoO4� þHþand

HMoO4� ¼ MoO4

2� þHþ

and Yan et al. (2011) have confirmed the MoO4 stoichiometry

(using EXAFS) in hydrothermal solutions up to 600 �C.In addition, Borg et al. (2012) have confirmed molybdate

stoichiometry and geometry up to 385 �C and 600 bar using

X-ray absorption spectroscopy and also identified chloromo-

lybdate species in concentrated HCl solutions. There are few

studies of hydrothermal tungstate chemistry, and the earlier

potentiometric study by Wesolowski et al. (1984) still com-

prises the benchmark work on aqueous tungstate chemistry

(i.e., tungstic acid ionization and polynuclear tungstate forma-

tion) up to 300 �C. Recently, Minubayeva and Seward (2013)

have also studied (spectrophotometrically) the deprotonation

of simple monomeric H4WO4 up to 300 �C and reported

similar results to those of Wesolowski et al. (1984).

Arsenic and antimony are fellow travelers in hydrothermal

ore fluids and are often involved in ore mineral deposition,

and hence, an understanding of their hydrothermal chemistry

will provide additional insight into ore formation. The equilib-

rium constants for arsenous and antimonous acid ionization

have been determined up to 300 �C at saturated vapor pressures

by Zakaznova-Herzog et al. (2006) and Zakaznova-Herzog and

Seward (2006). Their data permit the modeling of arsenic(III)

and antimony(III) transport and deposition in hydrothermal/

geothermal fluids having low, reduced sulfur activity. The

dilemma, of course, is that hydrothermal ore fluids contain

reduced sulfur and thioarsenite, and thioantimonite species

(i.e., H3(As,Sb)S3�(As,Sb)(HS)3) will also contribute to the

total dissolved As(III) and Sb(III) in the fluids. The thermody-

namic data for these species are sparse and in poor agreement, as

discussed briefly in the succeeding text:

H3AsO3 aqð Þ ¼ H2AsO3� þHþand

H3SbO3 aqð Þ ¼ H2SbO3� þHþ

Figure 9

100

AuOH0

AuOH0

AuHS0

(a)

AuCl2-

AuCl2-

Au(OH)2-

Au(OH) -

Au(HS)2-

80

60

40

20

0

100

80

60

% A

u%

Au


Carbonate and bicarbonate are ubiquitous ligands in natu-

ral hydrothermal fluids and form stable complexes with many

metals (i.e., with hard/intermediate Lewis acids or class A

acceptors), such as Pb2þ, UO22þ, and REE ions, but there are

essentially no systematic studies of metal carbonate complexes

up to elevated temperature and pressure and from which ther-

modynamic data are available. Such complexes will be pH-

sensitive, and their stabilities will decrease dramatically during

boiling or with the precipitation of carbonate minerals.

Various carboxylic acid and other organic species occur in

sedimentary basin brines, the most important of which is

acetate (Shock and Koretsky, 1993), and a number of metal

acetate complexes have been studied at elevated temperatures

up to 300 �C. These include potentiometric and spectroscopic

measurements of the acetate complexes of Fe2þ (Giordano and

Drummond, 1991; Palmer and Drummond, 1988), Co2þ

(Bridger et al., 1981), Cuþ (Liu et al., 2001), Zn2þ and Pb2þ

(Yang et al., 1989), Cd2þ (Benezeth and Palmer, 2000), Csþ

and Sr2þ (Ragnarsdottir et al., 2001), and Nd3þ, Eu3þ and Y3þ

(Tagirov et al., 2007a,b; Wood et al., 2000; Zotov et al.,

2002). Other literature data on metal–organic complexing

relevant to ore-depositing systems are discussed by Seward

and Barnes (1997).

(b)

(c)

AuHS0

AuHS0

AuCl2-

2

Au(HS)2-

Au(HS)2-

40

20

0

100

80

60

40

20

0

0 2 4pH

% A

u

6 8

Figure 10 The distribution of gold complexes in solutions of differentcomposition at 400 �C and 500 bar. (a) H2SþHS�¼0.001 m,Cl�¼0.001 m; (b) H2SþHS�¼0.001 m, Cl�¼0.5 m;(c) H2SþHS�¼0.5 m, Cl�¼0.5 m. Modified from Stefansson A andSeward TM (2004) Gold(I) complexing in aqueous sulfide solutions to500�C at 500 bar. Geochimica et Cosmochimica Acta 68: 4121–4143.

13.2.4.1.5 Complexing with hydrosulfide/sulfide ligandsThe hydrosulfide ligand, HS�, plays a fundamentally impor-

tant role in the transport and deposition chemistry of some

elements in hydrothermal ore solutions. Auþ is the archtype

soft Lewis acid (class B) electron acceptor and, as such, forms

very stable complexes with HS�. The stoichiometry and stabil-

ity of the AuHS0 and Au(HS)2� species were first reported by

Seward (1973), and there have been numerous studies of the

interaction of Auþ with reduced sulfur ligands since then (e.g.,

Baranova and Zotov, 1998; Benning and Seward, 1996; Dadze

et al., 2000; Gibert et al., 1998; Loucks and Mavrogenes, 1999;

Pokrovski et al., 2009; Renders and Seward, 1989a; Shenberger

and Barnes, 1989; Stefansson and Seward, 2004; Tagirov et al.,

2005, 2006). The study of Stefansson and Seward (2004) pro-

vides thermodynamic data for the formation of the hydrosulfido-

gold(I) complexes, AuHS0 and Au(HS)2�, up to 500 �C and

500 bar, which, together with the data for hydroxide and

chloride complexes (Stefansson and Seward, 2003a,b), enable

the modeling of gold transport and deposition chemistry in

hydrothermal ore fluids over a wide range of temperature and

pressure (Figure 10). As shown in Figure 11, the solubility of

gold in aqueous sulfide solution at elevated temperature is very

sensitive to changes in pH and redox potential. In addition,

any process that causes a loss of reduced sulfur (i.e.,

H2SþHS�) will lead to complex instability and, hence, the

precipitation of gold. Boiling a hydrothermal fluid can be

catastrophic for the stability of the gold(I) hydrosulfide com-

plexes. Partitioning of volatiles into the vapor phase in epither-

mal environments leads to a loss of H2S (i.e., loss of the HS�

ligand), H2 (i.e., a redox change), and CO2 (i.e., an increase in

pH). A change in the redox potential of the residual liquid

phase during boiling and /or mixing with oxygenated meteoric

waters encountered in the upper parts of geothermal/epither-

mal systems removes sulfide sulfur (Seward, 1989). The pre-

cipitation of pyrite and other sulfide minerals also extracts

reduced sulfur from the ore fluid. In supercritical, magmatic

hydrothermal fluids, the condensation of an iron-rich brine

during phase separation allows the less dense, volatile-rich

(H2S-containing) phase to then cool, contract, and ascend

buoyantly in the Earth’s crust without losing most of its dis-

solved sulfide to pyrite precipitation, as discussed by Heinrich

et al. (2004).

Figure 10

-29AuCl2

-

Au(HS)2-

Au(HS)2-

AuHS0

AuHS0

-33

-37

log

f O2

-41

-452 4 6

pH(a)

1 ppb

10 ppb

1 pp

b

1 ppb

1 ppb

10 p

pb

10 p

pb

10 ppb

H2SHS-

SO42-HSO4

-

8 10

100 ppb

-2250 �C Au(HS)2

-

AuHS0

AuOH0

-4

log

mA

u/ m

ol k

g-1

-6

-8

-10

0 2 4 6pH(b)

8 10 12

Figure 11 (a) The solubility of gold (contour lines in ppb) andspeciation at 250 �C and 500 bar as a function of pH and log fO2 in asolution containing 1.0 m NaCl and SS¼0.01 m. The dashed lines defineregions of predominance of the various sulfur species; the regions ofhighest gold solubility are colored yellow (10–100 ppb) and red(>100 ppb). Thermodynamic data for the gold(I) hydrosulfidecomplexes are from Stefansson and Seward (2004). Modified fromWilliams-Jones AE, Bowell RJ, and Migdisov AA (2009) Gold in solution.Elements 5: 281–287. With permission from the MineralogicalAssociation of Canada. (b) The solubility (molal) of gold at 250 �C and500 bar and showing the species contributing to the solubility curve as afunction of pH. Reproduced from Stefansson A and Seward TM (2004)Gold(I) complexing in aqueous sulphide solutions to 500�C at 500 bar.Geochimica et Cosmochimica Acta 68: 4121–4143.


In addition to gold, the other group IB (group 11) elements,

Cu and Ag, also form stable hydrosulfide complexes. Thermody-

namic and local structural data for CuHS0 and Cu(HS)2� have

been studied under hydrothermal conditions by Mountain

and Seward (1999, 2003) and Etschmann et al. (2010), and

Stefansson and Seward (2003a,b,c) have studied the stability of

AgHS0 and Ag(HS)2� up to 400 �C and at 500 bar. At alkaline

conditions, the Cu2S(HS)22� and Ag2S(HS)2

2� stoichiometries

have been identified, but these are not considered to play a

significant role in ore transport. Whether the dominant ore

transport mechanism is chloride or hydrosulfide complexing

depends upon the temperature, pressure, pH, and ligand (Cl�

or H2S/HS�) concentration, as discussed, for example, by

Stefansson and Seward (2003a,b,c). The stoichiometry and sta-

bility of zinc(II) hydrosulfide/sulfide complexes have been stud-

ied by Tagirov et al. (2007a,b) and Tagirov and Seward (2010)

up to 250 �C. At 300 �C and the equilibrium vapor pressure, the

two hydrosulfide species, Zn(HS)20 and Zn(HS)3

�, predominate

in the near-neutral 0.10 m sulfide solutions, but for concentra-

tions of Cl� >0.1 m, the transport of zinc will be dominated by

ZnCln2�n complexes (Ruaya and Seward, 1986). The stability of

the monohydrosulfido-complex of cobalt(II) has recently been

reported by Migdisov et al. (2011a,b), who suggested that the

CoHSþ may play some role in cobalt transport in sulfide-con-

taining fluids at T<200 �C. The relevance of the related mono-

hydrosulfide complexes of iron and nickel to ore transport under

hydrothermal conditions is not known. Mercury(II) forms stable

hydrosulfide/sulfide complexes under hydrothermal conditions

(see Barnes and Seward, 1997, and references therein), and other

studies pertaining to ambient/low-temperature conditions using

solubility, ab initio, and EXAFS methods (Bell et al., 2007;

Lennie et al., 2003; Paquette and Helz, 1997; Tossell, 2001a)

emphasize the stability of the mercury(II) hydrosulfide/sulfide

complexes, but more detailed studies under hydrothermal con-

ditions are required.

13.2.4.1.6 ThioanionsOf particular interest as well are the various thio-substituted

oxyanions of As, Sb, Mo, and W, which have been little studied

at elevated temperatures and pressures. There have been a

number of recent spectroscopic (X-ray absorption and ultravi-

olet) studies of aqueous thioarsenite species (Beak et al., 2008;

Bostick et al., 2005; Zakaznova-Herzog and Seward, 2012, and

references therein) that confirm the simple thioarsenous acid

stoichiometry (i.e., H3AsS3 or As(HS)3), but the system is also

complicated by the formation of intermediate oxythio-species

(i.e., AsOS23� and AsO2S

3�) and their protonated equivalents,

all of whose stabilities are pH-dependent. Thioantimonite and

thioantimonate species have also been reported at elevated

temperatures to 350 �C (Krupp, 1988; Sherman et al., 2000a,b),

but there is a desperate need for further systematic, corrob-

orative studies to confirm stoichiometries and provide high-

quality equilibrium thermodynamic data, especially for

thioantimonite and oxythioantimonite (SbO3�mSm3�) in hy-

drothermal solutions. As noted earlier, the thermodynamics of

molybdic and tungstic acid ionization up to 300 �C (i.e.,

H2(Mo,W)O4) are known, but this is not the case for the

sulfur-substituted analogues. Thiomolybdate and thiotung-

state, ((Mo,W)O4�mSm2�, where m¼1–4), are well known at

ambient temperature (e.g., Erickson and Helz, 2000) and may

be extensively polymerized (e.g., Mo4S156�, Mo4S13

2�, and

Mo2S72�) (Saxena et al., 1968), depending on the moly-

bdenum and/or tungsten activity. Zhang et al. (2012) have

measured the solubility of MoS2 (molybdenite) from 600 to

800 �C at 200 MPa and suggested that an ion-paired dithiomo-

lybdate species, NaHMo2S2, may account for the observed sol-

ubilities. All of these various thioanions will be potentially

important transporting moieties for elements such as As, Sb,

Mo, and W in sulfide-containing hydrothermal fluids in the

Earth’s crust, but there is a dearth of data at elevated tempera-

tures and pressures. In addition, these thioanions may them-

selves act as complexing ligands as in the case of the Cu(I)


thioarsenite complex, CuAsS(SH)(OH) (Clarke and Helz,

2000; Tossell, 2001a,b), but there are no data pertaining to

the stability of such species under hydrothermal conditions.

The transport of gold in ore solutions as a gold(I) thioarsenite

complex has long been hypothesized, but there are no available

data pertaining to the stability and stoichiometry of such spe-

cies at high temperatures and pressures.

It should be noted as well that As(V) and Sb(V) thio-species

have been identified in surface hot springs of active geothermal

systems (Planer-Friedrich and Scheinost, 2011; Planer-Friedrich

et al., 2010). This raises the possibility of thioarsenate and

thioantimonate species facilitating the mobilization of As and

Sb in the upper parts of geothermal/hydrothermal systems in

which the redox potential of the liquid phase has become

elevated due to boiling and mixing with circulating, steam-

heated, oxygenated meteoric waters.

13.2.4.1.7 Complexing with other sulfur-containingligandsThe possible role of intermediate oxidation state sulfur species,

such as polysulfides (SnS2� and Sn

�), thiosulfate (S2O32�), and

sulfite (SO32�), in the hydrothermal transport of gold has been

discussed by Seward (1982), and this topic has recently been

revisited by Pokrovski et al. (2009). The divalent polysulfide

ions (SnS2�) disproportionate to Sn

�, thiosulfate, and monosul-

fide (H2S/HS�) at T>150 �C, as noted by Giggenbach (1974a,

b). The SnS2� ligands form complexes with Auþ up to moderate

temperatures of �150 �C (Berndt et al., 1994; Kakovski and

Tyurin, 1962), and such species may play a role in gold precip-

itation and local remobilization in the upper parts of geo-

thermal systems where fluids of intermediate redox potential

occur. Polysulfide complexes of other elements, for example,

mercury, are known (e.g., Jay et al., 2000; Paquette and Helz,

1997), but there are no studies at hydrothermal conditions from

which thermodynamic data have been obtained.

However, in terms of hydrothermal complexing of gold and

other ‘soft’ metal cations, the S3� ion might be of more impor-

tance. S3� has long been known to occur in ethanol (a low

dielectric protic solvent having some similarities to high-

temperature water with respect to metal complex equilibria)

and salt melts at high temperatures to 600 �C (Giggenbach,

1968, 1971c). This blue polysulfide ion, now considered to be

the S3� species (Seel et al., 1977), was also shown to be

stabilized in high-temperature water by Giggenbach (1971b),

suggesting that it could play a role in metal complex formation

in some hydrothermal ore solutions of intermediate redox po-

tential and at temperatures in excess of 300 �C, as has also been

noted recently by Pokrovski and Dubrovinsky (2011). Tossell

(2012) has also studied the properties of the S3� anion using

quantum mechanical methods and demonstrated that S3� may

form complexes with Cuþ (e.g., Cu(S3)2� and Cu(S3)(H2O)0)

having similar stability to the Cu(I) hydrosulfide moieties.

The intermediate oxidation state oxyanions of sulfur, thio-

sulfate (S2O32�), sulfite (SO3

2�), polythionates (SnO62�), and,

of course, sulfate (SO42�), are known to occur in active geo-

thermal systems, especially in surface hot springs/pools as well

as some well discharges and hot crater lake environments

(Kaasalainen and Stefansson, 2011; Takano et al., 1994;

Webster, 1987; Wilson, 1941; Xu et al., 2000). The sulfur oxya-

nions and divalent polysulfides (i.e., SnS2�) undoubtedly play a

role in metal transport/remobilization in the supergene

environment, and a number of thiosulfate and sulfite minerals

have been identified in oxidized supergene zones of ore deposits

(e.g., sidpietersite, Pb4(S2O3)O2(OH)2, and scotlandite,

PbSO3). Hannebachite, CaSO3H2O, occurs in late-stage,

lower-temperature vesicle fillings of the Eifel volcanics

(Hentschel et al., 1985). In the case of gold(I), for example,

the thiosulfate and sulfite complexes are very stable at 25 �Cwith the equilibrium formation constant, log b2¼26–29.4 (Au

(S2O3)23�) and log b2¼30.1 (Au(SO3)2

3�) (Peshchevitsky

et al., 1970; Pouradier and Gadet, 1969), and the local, near-

linear, structure of the thiosulfate moiety has been studied using

EXAFS and ab initio DFT methods (Bryce et al., 2003).

In addition, the mixed ligand ammine/thiosulfate and sulfite/

thiosulfate gold(I) species have also been studied at 25 �C(Perera and Senanayake, 2004; Perera et al., 2005). However,

the importance of these complexes in hydrothermal ore trans-

port is strongly dependent on the stability of the ligands

themselves, at elevated temperatures and pressures. The poly-

thionates, thiosulfate, and the divalent polysulfides all disp-

roportionate at elevated temperatures (i.e., at T�250 �C), andthe kinetics of these disproportionations are also sensitive to pH

as noted by various workers (Foerster et al., 1923; Giggenbach,

1974a,b; Mizoguchi et al., 1976; Pryor, 1960). Nevertheless,

these intermediate oxidation state, anionic sulfur species may

play some role in the mobilization of some metals such as

gold in the upper, cooler parts of ore-depositing geothermal

systems.

In higher temperature, supercritical hydrothermal systems,

the sulfur redox chemistry is often considered in terms of the

reaction

H2S gð Þ þ 2H2O gð Þ ¼ SO2 gð Þ þ 3H2 gð Þ

In magmatic and volcanic gases, both H2S and SO2 occur,

but the thermodynamics of the reaction favors SO2 at high

temperatures. However, the stability of solvated SO2, that is,

sulfite as SO32� and HSO3

�, at T>300 �C, is poorly known,

although sulfite (SO32� and HSO3

�) has been shown to dispro-

portionate to elemental sulfur and sulfate at temperatures from

150 to 200 �C, sometimes with the formation of transient thio-

sulfate (Foerster et al., 1923; Mizoguchi et al., 1976), depending

on pH. Elemental sulfur hydrolyzes to H2S/HS� and HSO4�/

SO42� in high-temperature water (Ellis and Giggenbach, 1971).

The reactivity of magmatic sulfur dioxide in the presence of

supercritical water with decreasing temperature is given by

4SO2 þ 4H2O ¼ H2Sþ 3HSO4� þ 3Hþ

and this reaction together with the dissociation of HCl0 ion

pairs generates acidity which is then ‘titrated’ by reaction with

silicate minerals (see Giggenbach, 1992).

Sulfate complexes (ion pairs) of the divalent first-row tran-

sition element cations, Fe2þ (Rudolph et al., 1997), Ni2þ

(Madekufamba and Tremaine, 2011), Cu2þ (Mendez de Leo

et al., 2005), and Zn2þ (Rudolph et al., 1999), have been

studied under hydrothermal conditions, but none of these

simple ion pairs (i.e., MSO40) are considered to play a signif-

icant role in ore metal transport in saline hydrothermal fluids

of reduced to intermediate redox potential in the Earth’s crust.


Migdisov and Williams-Jones (2008) have studied the forma-

tion of Nd(III), Sm(III), and Er(III) sulfate complexes up to

250 �C (i.e., (REE)(SO4)þ and (REE)(SO4)2

�) and concluded

that these stable complexes would be the principal REE species

in natural hydrothermal solutions in the absence of elevated

concentrations of other complexing ligands, such as fluoride

and chloride.

13.2.4.1.8 Other complexing ligandsMetal ammine complexes are known at ambient temperature

for many elements, but the stability and stoichiometry of metal

ammino-complexes in hydrothermal solutions are essentially

unstudied, with the exception of the Cu(II) ammine com-

plexes, Cu(NH3)n2þ (1�n�4), in Trevani et al. (2001), who

investigated (visible spectroscopy) their formation up to

150 �C. Amminosilver(I) and amminogold(I) complexes are

known at ambient temperature (Perera and Senanayake, 2004;

Skibsted and Bjerrum, 1974), but there are no studies of the

stability and stoichiometry of these potentially important, ore-

transporting species under hydrothermal conditions.

13.2.4.1.9 Ore fluids with gas-like densityHydrothermal fluids, such as those of magmatic hydrothermal

systems, are commonly supercritical and may vary in density

from low (gas-like) to high (liquid-like), depending on the

level of emplacement of the hydrothermal system. On cooling,

these fluids generally separate into vapor and liquid by con-

densation or boiling, although the gas-like fluid may contract

to a liquid and the liquid-like fluid may expand to a vapor

(Heinrich, 2005). This separation of formerly supercritical hy-

drothermal fluids into vapor and liquid leads to partitioning of

the components of the original fluid between these two phases.

Volatile components, such as acidic gases (e.g., CO2, H2S, and

HCl), partition preferentially into the vapor, whereas compo-

nents with low volatility, such as NaCl and KCl, partition

preferentially into the liquid. This does not mean, however,

that low-volatility components are absent from the vapor. On

the contrary, salts like NaCl (e.g., Armellini and Tester, 1993;

Bischoff et al., 1986) and NH4Cl (Palmer and Simonson,

1993) dissolve in appreciable concentrations in water vapor

(depending on density), although in much lower concentra-

tions than in liquid water. Nonetheless, there are important

differences in the nature of the species dissolved in aqueous

vapor and liquid.

As discussed earlier, ionic species, including simple and

complex ions, are important in aqueous liquids and com-

monly predominate. By contrast, aqueous vapor is dominated

by molecular, uncharged species. For example, the nature of

the predominant reduced sulfur species in aqueous liquids

depends on pH and can be either H2S(aq) or HS�, whereas inthe vapor, the predominant form is molecular H2S. The frac-

tionation of solution components between aqueous liquids

and vapor therefore varies, even at constant temperature and

pressure, and depends strongly on the properties of the aque-

ous liquid, namely, the degree of ionization of the dissolved

components. As a result, the fractionation cannot be described

by a simple partition coefficient. However, it can be predicted

by using Henry’s law to determine the distribution of the

molecular species between gas and liquid, and ionization con-

stants to determine the proportion of the component of inter-

est that is in molecular form in the liquid (e.g., H2S0 relative to

H2S(aq)þHS�). Henry’s law relates the partial pressure of

the gas species over the liquid to its concentration in the

liquid phase and, via the Henry’s law constant, enables the

distribution of the molecular species to be calculated from its

thermodynamic properties. Salting-out effects may give rise

to induced phase separation (e.g., Duan and Li, 2008;

Suleimenov and Krupp, 1994), and this calculation requires

that the liquid should be at the same temperature as the gas.

During the past three decades, a large body of experimental

data has been gathered on the solubility of acid gases in aque-

ous solutions, and these have been used to develop models

that describe the behavior of some of these gases in the two-

phase region (e.g., CO2, CH4, and H2S) (Dubessy et al., 2005;

Majer et al., 2008). Because of nonideal behavior, accurate

prediction of the partitioning of gas species between vapor

and liquid necessitates that these models also employ equa-

tions of state for the gas phase (e.g., Stryjek and Vera, 1986)

and activity models for the liquid solutions (e.g., Helgeson

et al., 1981), in addition to the Henry’s law and ionization

constants referred to earlier.

Deviations from ideal behavior in aqueous vapor are due

largely to the ability of water molecules to form van der Waals

and hydrogen bonds, which, as shown experimentally by Liu

and Cruzan (1996), leads to the formation of weakly bound

gas-phase clusters, even in pure water vapor. In addition,

solutes in water vapor or supercritical, low-density steam do

not exist as ‘anhydrous’ molecular moieties but rather as hy-

drated clusters. For example, calculations by Lemke and

Seward (2008a) showed that even at the conditions of the

Earth’s atmosphere, mixtures of H2O, CO2, and N2O contain

ppm levels (molar) of hydrated CO2(H2O)n, N2O(H2O)n, and

H2O(H2O)n clusters (n�4). Such relatively modest cluster

formation is also typical of mixtures of water vapor with highly

volatile compounds. For example, no significant contribution

of hydrated species to arsenic solubility was detected in the

system As(OH)3–H2O (Pokrovski et al., 2002) and to boron

solubility in the system B(OH)3–H2O (Kukuljan et al., 1999),

although the well-known solubility of both arsenic and boron

in steam suggests the formation of hydrate clusters.

In contrast to volatile compounds, water clusters contribute

significantly to the solubility of weakly volatile species in water

vapor. For example, the solubility of NaCl in water vapor has

been shown experimentally to be many orders of magnitude

higher than that calculated assuming the absence of hydrated

NaCl gas species (e.g., Armellini and Tester, 1993; Bischoff

et al., 1986; Suleimenov et al., 2006). Suleimenov et al.

(2006) showed that the solubility of NaCl in supercritical

‘steam’ at 450 �C and up 350 bar is due primarily to an assem-

blage of the three even-numbered clusters, NaCl(H2O)n, where

n¼4, 6, and 8. Thus, NaCl dissolves in water vapor predomi-

nantly as NaCl(H2O)n clusters, and the contribution of unhy-

drated NaCl gas species to the NaCl solubility is insignificant.

Based on these experimental data, a set of models describing

the formation of these clusters and their stability has been

proposed (Pitzer and Pabalan, 1986; Suleimenov et al.,

2006). Similar behavior has been demonstrated for KCl

(Hovey et al., 1990) and for a set of chlorides and oxides of

economically important metals, that is, Cu, Ag, Au, Sn, and Mo

(Figure 12; Archibald et al., 2001, 2002; Migdisov and

Williams-Jones, 2005; Migdisov et al., 1999; Rempel et al.,

2006; Zezin et al., 2011).

AgCl–HCl–H2O100 ppb

10 ppb

50 ppb

12.0

10.0

8.0

X A

gCl (

10-9

)

pH2O (bar)

6.0

4.0

2.0

0.0

0 50 100

330 �C

360 �C

Ideal solubility

150 200

Figure 12 Experimentally determined concentration of AgCl in vapor asa function of pH2O. Ideal solubility refers to the vapor pressure of AgClover the corresponding solid. Modified from Migdisov AA, Williams-Jones AE, and Wagner T (2009) An experimental study of the solubilityand speciation of rare earth(III) elements in fluoride and chloride bearingaqueous solutions at temperatures up to 300�C. Geochimica etCosmochimica Acta 73: 7087–7109.

400

0.0

0

100

75

50

25

6.0

5.0

4.0

3.0

2.0

1.0

800700600T (�C)

500

400 700650600

AgCl + H2O =

AgCl:(H2O)

Au + HCl + H2O =AuCl:(H2O) + 0.5H2

r = 0.2 g cm-3

r = 0.2 g cm-3

HCl = 0.1 vol%Ni/NiO buffer

T (�C)(a)

(b)

Au

in V

apor

(pp

m)

Au

in V

apor

(pp

m)

500450 550

Figure 13 The predicted concentration of (a) Ag and (b) Au in aqueousfluid of density 0.2 g cm�3 as a function of temperature. The fluid in(a) was saturated with AgCl and in (b) with elemental gold; the fO2 of thelatter system, which contained 0.1 vol% HCl, was buffered by the Ni/NiObuffer. Modified from Migdisov AA and Williams-Jones AE (2013)A predictive model for metal transport of silver chloride by aqueous vaporin ore-forming magmatic-hydrothermal systems. Geochimica etCosmochimica Acta 104: 123–135.


These findings have important implications for theories of

ore formation. It is known, for example, that the hydrothermal

systems responsible for the formation of porphyry deposits are

dominated by vapor, rather than liquid (Henley and McNabb,

1978). There is also compelling evidence that the alteration

associated with the formation of high-sulfidation epithermal

Au–Ag deposits is caused by extremely acidic fluids that prob-

ably can only be produced by the condensation of acidic

magmatic vapors. Consequently, it is attractive to postulate,

as some researchers have, that the ore-forming fluid for these

deposits was a supercritical, low-density fluid that evolved by

condensation and/or contraction to a fluid of liquid-like den-

sity (Williams-Jones and Heinrich, 2005). Calculations of the

solubility of Cu and Au in vapor as CuCl(H2O)n and AuCl

(H2O)n clusters at 360 �C and pressures up to 200 bar reach

levels of 10�2.8 and 10�6.5 m, respectively (Archibald et al.,

2001, 2002). Such concentrations are high enough to form a

hypothetical 50 million tonne porphyry copper deposit with a

grade of 0.5% Cu in 20 500–56 000 years (depending on

composition of the fluid) or a 36 tonne Au deposit in

167 000–238 000 years. However, the experimentally deter-

mined concentrations are considerably lower than those mea-

sured in vapor inclusions, albeit at much higher temperature

and pressure (Ulrich et al., 1999). This apparent disagreement

has recently been resolved as a result of the discovery that

the solubility of metals in chloride-bearing aqueous fluids in-

creases exponentially with the density of the fluid, due to the

formation of increasingly larger clusters of water molecules

around the metal, and that the free energy of cluster formation

varies linearly with reciprocal temperature (Migdisov and

Williams-Jones, 2013). Using these observations, Migdisov

and Williams-Jones (2013) have extrapolated the solubility

data for silver and gold and shown that they reach the ppm

levels measured in vapor inclusions at the near-magmatic con-

ditions of their entrapment (Figure 13).

Cadmium exhibits volatility in high-temperature systems

and occurs in volcanic gases (e.g., Symonds et al., 1987;

Wahrenberger et al., 2002) and in high-temperature fumarole

sublimates (e.g., Chaplygin et al., 2005); however, themolecular

basis for its gas-phase transport and volatility is poorly

known. Recent FT-ICR/MS (Fourier transform ion cyclotron res-

onance mass spectrometry) measurements combined with ab

initio quantum chemical calculations have identified an exten-

sive array of gas-phase cadmium(II) chloride–water clusters

(Lemke and Seward, 2013), such as (CdCl)þ(H2O)n (n¼1–4).

Various dimers, trimers (with respect to cadmium), and higher-

order polynuclear species, such as (Cd4Cl7)þ(H2O) as well as

larger, doubly charged clusters, were also identified mass

spectrometrically.

As noted earlier in the case of volatile Cd chloride–water

clusters, the dominance of molecular solubility in low-density

fluids does not exclude the formation of clusters containing

ionized species in the vapor phase, especially in the super-

critical region. Indeed, proton–water clusters have been

detected over boiling water and in supercritical steam using

infrared spectroscopy (Carlon, 1979), nuclear magnetic reso-

nance (Matubayasi et al., 1997), and high-pressure mass spec-

trometry (Likholyot et al., 2007). Clusters containing ionized

halide ions (Cl�, Br�, and I�) have also been detected experi-

mentally at temperatures from �45 to þ175 �C (Likholyot

et al., 2005). Recent studies based on ab initio calculations

Figure 12

Figure 13


have reported standard thermodynamic properties for such

clusters and predicted them for more complex aggregates,

such as H3Oþ(H2O)m(H2S)n, NH4

þ(H2O)m(H2S)n, and H3Sþ

(H2O)m(H2S)n, where m�6 and n�4 (Lemke and Seward,

2008b).

Although there is compelling evidence for the transport of

metals in water vapor, it seems probable that deposition of

these metals occurs after condensation and/or contraction

of the vapor to liquid. Evidence for this is provided by the

blanket-like distribution of gold in high-sulfidation epithermal

deposits, interpreted to represent a level where the temperature

of the host rocks is below the dew point of the fluid

(Chouinard et al., 2005b). In many cases, the mechanisms of

deposition are likely to be the same as those discussed earlier

for the liquid. However, as the concentrations of the metals

may be relatively low, it is likely that theymay be insufficient to

saturate the ore fluid of interest, particularly in relatively low-

pressure environments such as those of some high-sulfidation

epithermal systems. In these systems, the metals may therefore

concentrate, not by precipitation of an ore mineral of the

metal, but rather by adsorption onto the surface of a gangue

mineral like pyrite or other sulfide minerals (Chouinard et al.,

2005a), as demonstrated by Renders and Seward (1989),

Cardile et al. (1993), and Widler and Seward (2002).

Finally, the intriguing but as yet unstudied role of super-

critical CO2 in the transport of some elements in high-

temperature–high-pressure environments in the Earth’s lower

crust should be mentioned. Numerous authors have reported

the occurrence of fluid inclusions containing pure CO2 and

CO2þCH4þN2 fluids that have been trapped in minerals

formed at granulite and amphibolite facies metamorphic con-

ditions (e.g., Andersen et al., 1997; Cuney et al., 2007;

Munyanyiwa et al., 1993; Touret, 1971). The question arises

as to the role, if any, of supercritical CO2 solvent in element

transport at high temperatures and pressures in the deep crust.

It is well known that molecular (uncharged) metal complexes

and metal chelates have appreciable solubility in supercritical

CO2 (e.g., Ashraf-Khorassani et al., 1997; Haruki et al., 2011).

What might be the solubility of transition metal carbonyl

complexes such as Fe(CO)5 or Ni(CO)4 or even simple ion

pairs (e.g., HCl0 and/or NaCl0) in supercritical CO2 at deep

crustal temperatures and pressures (e.g., 700 to 850 �C and up

to 15 kbar)? Other molecular components such as TiCl4 and

SiCl4 are highly ‘soluble’ in supercritical CO2, and in fact,

the phase equilibria in the binary systems TiCl4–CO2 and

SiCl4–CO2 (Suleimenov et al., 2003; Tolley and Tester, 1989)

indicate extensive miscibility between the two end-member

components. For example, in the binary SiCl4–CO2 system at

T>150 �C and P>125 bar, both components mix/interact

to give a homogeneous fluid phase throughout the entire

composition range.

13.2.5 Epilogue

Despite the extensive archive of information on ore deposits,

which includes knowledge of tectonic settings and regional

permeabilities, mineralogy, trace element and isotope geo-

chemistry, fluid inclusion compositions, temperatures, and

pressures, our understanding of the chemistry of metal

transport and ore mineral deposition (i.e., ore formation) is

still fraught with inadequacy because of the lack of data per-

taining to metal complex stoichiometry and stabilities at ele-

vated temperatures and pressures. At temperatures up to about

300 �C and at moderate pressures not far removed from satu-

rated vapor pressures, we have some considerable knowledge

of complex equilibria for a few metals such as Fe, Cu, Pb, Zn,

Ag, and Au. For most other metals, the data may pertain to

complexing with only one ligand or derive from a single ex-

perimental study at restricted conditions, such as at one ele-

vated temperature and pressure. As summarized in the

discussions earlier, information has also been acquired at a

molecular level from spectroscopic (X-ray absorption, Raman,

and UV–Vis) and quantum chemical computations. While this

information is fundamental and necessary, high-quality ther-

modynamic data with which to model metal transport and

deposition over the range of fluid compositions, temperature,

and pressure occurring in hydrothermal environments

throughout the Earth’s crust are also required. At present,

gold is the only metal for which reliable thermodynamic

data exist over a wide range of temperature and pressure to

supercritical conditions(i.e., up to 600 �C and 1800 bar) for

a number of important complexes (i.e., with HS�, Cl�, andOH�). Thus, Heinrich et al. (2004) were able to model gold

transport and deposition associated with phase separation

and subsequent magmatic vapor contraction from the mag-

matic porphyry to the epithermal environment by combining

the relevant thermodynamic data for the gold(I) complexes

with phase equilibria in the system NaCl–H2O and fluid

inclusion data.

A priority must therefore be to acquire high-quality thermo-

dynamic and molecular data on metal complex equilibria per-

tinent to ore transport and deposition processes over wide

ranges of temperature and pressure. As would be expected, the

ion pairing of metal complexes with the bulk electrolyte salt

cations, such as Naþ and Kþ, at near-magmatic supercritical

conditions will also be of interest, but such species have been

little studied. These species (e.g., NaAu(HS)20 or NaHMoO2S2

0)

may play an important role in enhancing metal solubilities and

partitioning into a magmatic volatile-rich phase as noted, for

example, by Zajacz et al. (2010) and Zhang et al. (2012). This

will be a challenging endeavor at high temperatures >350 �Cand at pressures up to 5000 bar. The lack of such data comprises

the major barrier to the further understanding of the formation

of hydrothermal ore deposits as well as to the molecular intri-

cacies of the chemistry involved.

Acknowledgments

We thank our colleagues, Chris Heinrich and Steve Scott, for

their helpful suggestions on an earlier version of this chapter.

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