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Hidrotermal Processes and Mineral System - Franco Pirajno

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of a fluid
through the interconnected pore spaces in a rock, without displacement of the
rock particles. Permeability is measured in darcy, which is the rate of flow of a
fluid that passes through a cm2 for the length of 1 cm in a second (permeability
coefficient k= cm/s). Porosity is the sum total of available spaces between rock
particles or grains and is expressed as a percentage (porosity coefficient n) of the
total volume of the rock. Porosity can be primary if it is a relic of deposition,
such that the space between grains was not completely eliminated by compac-
tion or subsequent chemical changes, or secondary when there is partial dis-
solution at grain margins, fracturing or chemical changes. Figure 1.4 illustrates
types of porosity. Permeability, therefore, controls the rate at which fluids flow
through connected pore spaces in a porous medium (Cathles and Adams 2005).
Cathles and Adams (2005) distinguished two types of permeability: intrinsic
and dynamic. Intrinsic permeability is defined by the property of the medium at
standard temperature and pressure; dynamic permeability is that which the
medium might have under non-standard conditions. An example of the former
is provided by sandstone and quartz arenites, which are composed of quartz
and feldspar grains, and their permeability decreases with increasing propor-
tions of feldspars and rock fragments. Another example is clay coatings and oil
1.2 Origin of Water; Sea and Surface Waters 15
staining of the component grains, which inhibit compaction and therefore tend to
preserve permeability. Fluid pressure and temperature control dynamic perme-
ability, as for example thermal contraction at themargins of an igneous intrusion.
This will induce fracturing that will increase permeability. In sedimentary basins
permeability is dynamically controlled where, for example, pore pressures
hydraulically fracture the rocks, or where there is thermal expansion and positive
volume changes through chemical reactions (Cathles and Adams 2005).
Groundwater is known to contain bicarbonates, sulphates, chlorides and
alkali metals, with their amount depending on the composition of the surround-
ing rocks and the length of time that the water has been in contact with them.
Groundwater is the ‘‘meteoric’’ water of economic geologists, and in regions of
high geothermal gradients it will activate as a meteoric hydrothermal system as
it rises along fractures or faults. Where it reaches the surface this water becomes
a hot spring (see Section 1.4.5). Meteoric hydrothermal systems in volcanic
regions are generally heated by subjacent magmas, andmay reach temperatures
above 3508C. During their upward flow they become solutions that deposit
metals and sulphides.Where they discharge at or near the surface these meteoric
waters are known as geothermal fields, such as those of Rotorua in New
Zealand, or the Yellowstone National Park in the USA. Examples of hot
springs and geothermal fields are shown in Fig. 1.7; hot springs are discussed
more fully in Section 1.4.5. Fluids of these geothermal systems usually have near
neutral pH, low S and salinity.
Subsurface flow of groundwater can be very important for the formation of
some ore deposits, such as roll-front U deposits (Ingerbritsen and Sanford
Fig. 1.4 Types of pore spaces and permeability in sedimentary rocks; (A) well sorted sand-
stone with high permeability; (B) poorly sorted sandstone with low permeability; (C) well
sorted sandstone with porous clasts, permeability is very high; (D) well sorted sandstone with
interstitial cementing material, poor permeability; (E) and (F) fractured carbonate rocks,
permeability is high (after Desio 1959)
16 1 Water and Hydrothermal Fluids on Earth
1998). In the case of roll-front U deposits, oxidised groundwater leaches U6+
from the rocks through which it travels and precipitates U4+, where these
waters encounter a reducing medium (redox front).
1.3 Structure and Properties of Water; Hydration and Hydrolysis
The structure of the water molecule is of special importance to its physico-
chemical behaviour, as discussed in Franks (1982) and Neilson and Enderby
(1986). It has been determined that the structure of water-ice is analogous to
that of tridymite, and that the oxygens in ice are tetrahedrally coordinated (the
tetrahedron being the building block of all silicate minerals in the Earth’s crust).
Therefore water appears to have a pseudo-crystalline arrangement similar to
the quartz structure, which possibly serves to explain its higher density relative
to ice (Paton 1978). The oxygen ion in a water molecule is much larger than the
hydrogen ion, with the result that themolecule can be described as a sphere that,
though neutral, has a positive charge on the side of the two hydrogens and a
negative charge on the side of the oxygen. Consequently, the isolated molecule
has a polar character and behaves in solution like a small magnet. This polar
character is the key to the hydration and hydrolysis of silicate minerals, and
thus important in weathering and hydrothermal alteration processes. In the
process of hydration, water molecules are attracted to and, by virtue of their
polar charges, become orientated around other attracting ions, forming hydra-
tion shells (Brimhall and Crerar 1987). Dissolution occurs when successive
layers of water molecules completely surround the ion. Polar water molecules
can enter crystal lattices and orientate themselves against charged mineral
surfaces. Where these molecules come into mutual contact, a lubricated surface
is produced, for example clays which become slippery when wet. Water of
crystallisation is part of the mineral lattice, as in gypsum (CaSO4
.2H2O). On
heating, this water is given off without breaking the lattice, and can be restored
to reform the mineral.
Hydrolysis is the effect of the dissociation of water molecules into H+ and
OH– ions. The process of hydrolysis is responsible for the breakdown of silicate
minerals, and involves the addition of H+ and OH– to bonding sites in the
mineral lattice. Hydrolysis is defined as the reaction between water and the ion
of a weak acid or a weak base (Krauskopf 1979, p. 37). An example of hydro-
lysis is the reaction of fayalite with water at neutral pH:
2FeSiO4 þ 4H2O$ 2Fe2þ þ 4OH� þ 2H2SiO4
Or the hydrolysis reaction of atmospheric SO2 to form sulphurous acid
SO2 þH2O! H2SO3 ! Hþ þHSO�3
1.3 Structure and Properties of Water; Hydration and Hydrolysis 17
Hydrolysis reactions tend to be accelerated under conditions of low pH, as
for example in the vicinity of oxidising sulphide ore bodies. The presence of H+
ions in acid waters promotes the attack on silicate minerals resulting in the
liberation of cations. These cations may remain in the vicinity and become fixed
as stable secondary mineral assemblages, while others go into solution and are
transported elsewhere. The mobility of these cations under different physico-
chemical conditions has important implications for exploration geochemistry
and for the evaluation of gossans (Taylor and Eggleton 2001).
Thus, hydrolysis is dependent on the concentration of H+ ions, and any
process which affects their concentration will also affect the speed and intensity
of the hydrolysis process. The breaking of molecular bonds by water in a silicate
melt, and consequent lowering of viscosity and consolidation temperature, is a
similar process. Hydrolysis of silicate minerals is very important for hydro-
thermal alteration because the hydrogen ions penetrate the silicate lattices
where they compete with cations (K, Ca, Na etc.) to attach to oxygen ions.
The larger concentration of the charge in the hydrogen ions predominates,
resulting in the displacement of the cations which are transferred from the
silicate into the solution, while H+