derived from an albite-rich gneiss the K-feldspar content is likely to be too low and much of the kaolinite
would then not be illitised.
Similarly, not much illite will be formed in
sandstones with little kaolinite or smectite. Both in
Haltenbanken and the North Sea there are Jurassic
reservoirs where plagioclase is the dominant feldspar
and where the low K-feldspar content is unable to
supply the necessary potassium for illitisation of kaolinite. The low illite content in such reservoirs
preserves better permeability. This is a direct function
of the provenance and could be due to erosion of albite
gneisses rather than granitic gneisses.
The distribution of authigenic illite in sedimentary
basins like the North Sea and Haltenbanken shows that
illite formation is strongly controlled by the present
day burial depth and temperature. The increase in illite
content at about 3.7–4.0 km is usually very sharp,
indicating a temperature-controlled reaction rather
than a high kinetic reaction rate when the association
of kaolinite and K-feldspar becomes thermodynamically unstable. Basin loading from thick Pleistocene
sequences in these areas suggests that the illite formed
recently.
K-Ar dating of illite gives variable ages for the
formation of illite. This is probably because even
very small amounts of detrital (older) mica or feldspar
will produce too-old ages (Hamilton et al. 1989).
4.12 Porewater
In a sedimentary basin the amount of solids dissolved
in porewater is very small compared to the volume of
solids. During burial diagenesis, significant precipitation of authigenic minerals must be accompanied by
dissolution of other minerals or the same mineral, as in
the case of pressure solution. Even if the porewater is
supersaturated with respect to a certain mineral, only
very small amounts can precipitate before the
porewater attains equilibrium. Precipitation of new
minerals requires that other minerals dissolve because
the porewater has very little capacity to store ions.
This has been quantified using chemical modelling
(Giles 1997) from which it can be concluded that
extremely large fluid fluxes are required through the
pores in order to dissolve or precipitate significant
amounts of cements like calcite or quartz.
The porewater reacts with the minerals it is in
contact with, and with increasing temperature the
composition becomes more and more in equilibrium
with these minerals. This is because the kinetics of the
mineral reactions become faster. The porewater in
sedimentary basins consists of solutions buffered by
the minerals present. The pH is controlled partly by the
carbonate reactions (pCO 2 ), and the pH will decrease
from 7.5–8 in the seawater to 4–5 at 3–4 km depth. As
temperature rises above 100
C, silicate reactions
become increasingly important, e.g. between Kfeldspar, kaolinite and illite which will determine the
K
+
/H
+ ratio.
Organic acids are weak acids which can not significantly change the pH in the strongly buffered
porewater. The buffering capacity of organic acids
has been shown to be orders of magnitude lower than
for the carbonate and silicate systems (Hutcheon and
Abercrombie 1990). Organic acids generated in source
rocks like the Kimmeridge Clay Fm in the North Sea
are likely to be neutralised by reactions with calcite
which is commonly present in these source rocks. The
limited effect of organic acids and CO 2 on mineral
dissolution and diagenesis has also been shown experimentally (Barth and Bjørlykke 1993).
When porewater moves it will nearly always cause
some dissolution and precipitation but this is rarely
significant due to low velocities. The volume of
minerals dissolved or precipitated can be calculated
using the following equation:
V c ¼ F t sin α ðdT=dZÞ Δ S=ρ
The volume of precipitated mineral V c is a product
of the fluid flux integrated over time (t), the angle of
fluid flow relative to the isotherms (α), the geothermal
gradients (dT/dZ), the solubility as a function of the
temperature (ΔS) and the density of the mineral (ρ).
The solubility gradient (ΔS) is 1–3 ppm/
C,
depending on the temperature (Wood 1986). This
means that at 100–150
C about 2 ppm of quartz
precipitates for each degree the porewater is cooled.
This gives a solubility gradent of 2.10
–6 /
C. If the
geothermal gradient is 30
C=km ð3 Â 10
À2 C=mÞ,
porewater must move upwards more than 30 m to
reduce the temperature by 1
C and the quartz cement
is distributed through these 30 m of sandstone. Assuming vertical flow (sin α ¼ 1) and geothermal gradients
4 Sandstones and Sandstone Reservoirs
137
would then not be illitised.
Similarly, not much illite will be formed in
sandstones with little kaolinite or smectite. Both in
Haltenbanken and the North Sea there are Jurassic
reservoirs where plagioclase is the dominant feldspar
and where the low K-feldspar content is unable to
supply the necessary potassium for illitisation of kaolinite. The low illite content in such reservoirs
preserves better permeability. This is a direct function
of the provenance and could be due to erosion of albite
gneisses rather than granitic gneisses.
The distribution of authigenic illite in sedimentary
basins like the North Sea and Haltenbanken shows that
illite formation is strongly controlled by the present
day burial depth and temperature. The increase in illite
content at about 3.7–4.0 km is usually very sharp,
indicating a temperature-controlled reaction rather
than a high kinetic reaction rate when the association
of kaolinite and K-feldspar becomes thermodynamically unstable. Basin loading from thick Pleistocene
sequences in these areas suggests that the illite formed
recently.
K-Ar dating of illite gives variable ages for the
formation of illite. This is probably because even
very small amounts of detrital (older) mica or feldspar
will produce too-old ages (Hamilton et al. 1989).
4.12 Porewater
In a sedimentary basin the amount of solids dissolved
in porewater is very small compared to the volume of
solids. During burial diagenesis, significant precipitation of authigenic minerals must be accompanied by
dissolution of other minerals or the same mineral, as in
the case of pressure solution. Even if the porewater is
supersaturated with respect to a certain mineral, only
very small amounts can precipitate before the
porewater attains equilibrium. Precipitation of new
minerals requires that other minerals dissolve because
the porewater has very little capacity to store ions.
This has been quantified using chemical modelling
(Giles 1997) from which it can be concluded that
extremely large fluid fluxes are required through the
pores in order to dissolve or precipitate significant
amounts of cements like calcite or quartz.
The porewater reacts with the minerals it is in
contact with, and with increasing temperature the
composition becomes more and more in equilibrium
with these minerals. This is because the kinetics of the
mineral reactions become faster. The porewater in
sedimentary basins consists of solutions buffered by
the minerals present. The pH is controlled partly by the
carbonate reactions (pCO 2 ), and the pH will decrease
from 7.5–8 in the seawater to 4–5 at 3–4 km depth. As
temperature rises above 100
C, silicate reactions
become increasingly important, e.g. between Kfeldspar, kaolinite and illite which will determine the
K
+
/H
+ ratio.
Organic acids are weak acids which can not significantly change the pH in the strongly buffered
porewater. The buffering capacity of organic acids
has been shown to be orders of magnitude lower than
for the carbonate and silicate systems (Hutcheon and
Abercrombie 1990). Organic acids generated in source
rocks like the Kimmeridge Clay Fm in the North Sea
are likely to be neutralised by reactions with calcite
which is commonly present in these source rocks. The
limited effect of organic acids and CO 2 on mineral
dissolution and diagenesis has also been shown experimentally (Barth and Bjørlykke 1993).
When porewater moves it will nearly always cause
some dissolution and precipitation but this is rarely
significant due to low velocities. The volume of
minerals dissolved or precipitated can be calculated
using the following equation:
V c ¼ F t sin α ðdT=dZÞ Δ S=ρ
The volume of precipitated mineral V c is a product
of the fluid flux integrated over time (t), the angle of
fluid flow relative to the isotherms (α), the geothermal
gradients (dT/dZ), the solubility as a function of the
temperature (ΔS) and the density of the mineral (ρ).
The solubility gradient (ΔS) is 1–3 ppm/
C,
depending on the temperature (Wood 1986). This
means that at 100–150
C about 2 ppm of quartz
precipitates for each degree the porewater is cooled.
This gives a solubility gradent of 2.10
–6 /
C. If the
geothermal gradient is 30
C=km ð3 Â 10
À2 C=mÞ,
porewater must move upwards more than 30 m to
reduce the temperature by 1
C and the quartz cement
is distributed through these 30 m of sandstone. Assuming vertical flow (sin α ¼ 1) and geothermal gradients
4 Sandstones and Sandstone Reservoirs
137