676
For example, during the transformation of skeletal aragonite
(e.g. corals) to calcite, the aragonite may release strontium to
its environment, while the new calcite takes up magnesium
from outside (cf. Fig. 13.16e).
13.3.5 Hydrodynamic Dispersion
Some authors (e.g. Hanor 1994; Morse et al. 1997)
have pointed out that neither molecular diffusion nor
uniform pore water flow can explain a11 phenomena of
diagenesis. A combination ofboth pro ces ses including
local sma11-scale mixing of pore waters of slightly differing composition is accomplished by hydrodynamic
dispersion. This term refers to the different velocities
ofthe fluid flow through a heterogeneous porous medium. The water particles are forced to take paths of
different lengths and they are slowed close to grain
surfaces. This leads to a spatia11y varying pore water
composition with the result that dissolution of a certain
mineral at a specific site may be fo11owed by reprecipitation of the mineral at another site not far away.
Hydrodynamic dispersion is defined by the flow velocity and
an aquifer-specific dispersion coefficient (see textbooks for
hydrogeology, e.g., Freeze and Cherry 1979) and also inc1udes molecular diffusion. Ifthe fluid velocity is very low,
the solute transport is dominated by molecular diffusion (cr.
Fig. l6e). At higher fluid velocities (greater than a few
mmlyear) mechanical dispersion mainly controls solute transfer. An example from the GulfCoast (Fig. 13.18c) indicates
that diffusion transport has been accelerated by hydrodynamic dispersion. In this case, compaction flow is probably
sufficient to account for the additional solute transport. Without the aid of advection, diffusion-controlled highly saline
pore water may subside almost at the same rate as the basin
(B jill"lykke and Gran 1994).
13.3.6 Formation Waters
The processes discussed so far allow a better understanding ofthe great variations in the chemical composition of formation waters. The nature of these fluids is
an important topic in hydrocarbon exploration and the
key for the interpretation of dissolution and precipitation of minerals in the subsurface. Even within the
same stratum, the composition of the formation waters
and thus the diagenetic response may change significantly. One has to distinguish between two end-member types of pore waters:
- Connate water is inc1uded in the pore space of the
sediment at the time of deposition (Latin-derived
"borne with"). Connate marine water contains 32-37
g/I dissolved species of approximately constant composition. Continental connate waters greatly vary in
their composition. In humid climates, connate water or
groundwater has a very low salinity of mostly ::;1 g/l,
the dominant anions being hydrogen carbonate andlor
Chapter 13 Diagenesis
sulfate. In arid regions, the connate water may attain
very high salinites (cf. Sect. 6.4).
- Formation water is buried older pore water which is
not in direct contact with meteoric water. It has reacted
to increased temperatures and pressures and exchanged
components with the solid partic1es ofthe surrounding
sediment matrix. Furthermore, it rnay have taken up
fluids, gases or hydrocarbons from elsewhere, or it
may have exported components to other sites.
In hydrologica11y open diagenetic systems, both connate waters and formation waters can mix with meteoric water, diagenetically altered groundwater, and
evaporitic brines. In this way, pore fluids of greatly
variable composition are generated. Mixed pore fluids
are often undersaturated or oversaturated with respect
to a certain mineral species and therefore lead to dissolution or reprecipitation of this mineral.
Stagnant or very slowly moving formation waters
generally tend to establish a thermodynamic equilibrium with their surrounding solids. Most of their components are buffered by carbonate and silicate minerals
and therefore reach only relatively low cation concentrations (e.g. for Ca and Mg; Fig. 13.18d,e). General
trends are a depletion in Mg, which is consumed by
dolomitization, and enrichments in Ca (released by
dolomitization) and silica (due to a better solubility at
high temperature ).
Marine sediments with connate sea water and devoid of evaporites show moderate changes in their total
dissolved solids, TDS, of their formation waters with
depth, provided the system is largely c10sed (Fig.
13.18a). Sedimentary sequences in contact with halitebearing evaporites develop hypersaline pore fluids (up
to ~300 g/I TDS) with very high chloride contents accompanied either by high Naor Ca values (Fig. 13.18).
Chloride, bromide, and often also sulfate are not buffered by ambient mineral phases. Particularly chloride
can therefore be used as a kind of tracer for the source
of the fluids. The cation concentrations of these
hypersaline pore fluids tend to increase by several orders of magnitude with growing chloride content, but
the absolute values for Mg and K remain much lower
than those for Ca and Na (Fig. 13.18d,e).
Pore fluids cannot migrate from norma11y pressured
into overpressured zones which tend to expe11 pore
water. This is demonstrated by Fig. 13.18b) where
deeply buried mudstones have become overpressured
and therefore maintained low pore water salinities below sand-dominated younger strata with evaporitic
brines.
For example, during the transformation of skeletal aragonite
(e.g. corals) to calcite, the aragonite may release strontium to
its environment, while the new calcite takes up magnesium
from outside (cf. Fig. 13.16e).
13.3.5 Hydrodynamic Dispersion
Some authors (e.g. Hanor 1994; Morse et al. 1997)
have pointed out that neither molecular diffusion nor
uniform pore water flow can explain a11 phenomena of
diagenesis. A combination ofboth pro ces ses including
local sma11-scale mixing of pore waters of slightly differing composition is accomplished by hydrodynamic
dispersion. This term refers to the different velocities
ofthe fluid flow through a heterogeneous porous medium. The water particles are forced to take paths of
different lengths and they are slowed close to grain
surfaces. This leads to a spatia11y varying pore water
composition with the result that dissolution of a certain
mineral at a specific site may be fo11owed by reprecipitation of the mineral at another site not far away.
Hydrodynamic dispersion is defined by the flow velocity and
an aquifer-specific dispersion coefficient (see textbooks for
hydrogeology, e.g., Freeze and Cherry 1979) and also inc1udes molecular diffusion. Ifthe fluid velocity is very low,
the solute transport is dominated by molecular diffusion (cr.
Fig. l6e). At higher fluid velocities (greater than a few
mmlyear) mechanical dispersion mainly controls solute transfer. An example from the GulfCoast (Fig. 13.18c) indicates
that diffusion transport has been accelerated by hydrodynamic dispersion. In this case, compaction flow is probably
sufficient to account for the additional solute transport. Without the aid of advection, diffusion-controlled highly saline
pore water may subside almost at the same rate as the basin
(B jill"lykke and Gran 1994).
13.3.6 Formation Waters
The processes discussed so far allow a better understanding ofthe great variations in the chemical composition of formation waters. The nature of these fluids is
an important topic in hydrocarbon exploration and the
key for the interpretation of dissolution and precipitation of minerals in the subsurface. Even within the
same stratum, the composition of the formation waters
and thus the diagenetic response may change significantly. One has to distinguish between two end-member types of pore waters:
- Connate water is inc1uded in the pore space of the
sediment at the time of deposition (Latin-derived
"borne with"). Connate marine water contains 32-37
g/I dissolved species of approximately constant composition. Continental connate waters greatly vary in
their composition. In humid climates, connate water or
groundwater has a very low salinity of mostly ::;1 g/l,
the dominant anions being hydrogen carbonate andlor
Chapter 13 Diagenesis
sulfate. In arid regions, the connate water may attain
very high salinites (cf. Sect. 6.4).
- Formation water is buried older pore water which is
not in direct contact with meteoric water. It has reacted
to increased temperatures and pressures and exchanged
components with the solid partic1es ofthe surrounding
sediment matrix. Furthermore, it rnay have taken up
fluids, gases or hydrocarbons from elsewhere, or it
may have exported components to other sites.
In hydrologica11y open diagenetic systems, both connate waters and formation waters can mix with meteoric water, diagenetically altered groundwater, and
evaporitic brines. In this way, pore fluids of greatly
variable composition are generated. Mixed pore fluids
are often undersaturated or oversaturated with respect
to a certain mineral species and therefore lead to dissolution or reprecipitation of this mineral.
Stagnant or very slowly moving formation waters
generally tend to establish a thermodynamic equilibrium with their surrounding solids. Most of their components are buffered by carbonate and silicate minerals
and therefore reach only relatively low cation concentrations (e.g. for Ca and Mg; Fig. 13.18d,e). General
trends are a depletion in Mg, which is consumed by
dolomitization, and enrichments in Ca (released by
dolomitization) and silica (due to a better solubility at
high temperature ).
Marine sediments with connate sea water and devoid of evaporites show moderate changes in their total
dissolved solids, TDS, of their formation waters with
depth, provided the system is largely c10sed (Fig.
13.18a). Sedimentary sequences in contact with halitebearing evaporites develop hypersaline pore fluids (up
to ~300 g/I TDS) with very high chloride contents accompanied either by high Naor Ca values (Fig. 13.18).
Chloride, bromide, and often also sulfate are not buffered by ambient mineral phases. Particularly chloride
can therefore be used as a kind of tracer for the source
of the fluids. The cation concentrations of these
hypersaline pore fluids tend to increase by several orders of magnitude with growing chloride content, but
the absolute values for Mg and K remain much lower
than those for Ca and Na (Fig. 13.18d,e).
Pore fluids cannot migrate from norma11y pressured
into overpressured zones which tend to expe11 pore
water. This is demonstrated by Fig. 13.18b) where
deeply buried mudstones have become overpressured
and therefore maintained low pore water salinities below sand-dominated younger strata with evaporitic
brines.