(4) Precipitation of cement (e.g. quartz cement) will
increase the strength of the grain framework and
prevent further mechanical compaction. The sandstone is then overconsolidated – not due to previously higher stress, but due to cementation. This is
sometimes referred to as pseudo-overconsolidation.
Further compaction will then mostly be controlled
by the rate of dissolution and precipitation.
In the following the diagenetic processes typical of
different burial depths will be discussed.
4.3
Early Diagenesis
As soon as sediments are deposited, early diagenetic
reactions start to modify the primary sediment composition. At very shallow burial depth (<1–10 m),
sediments have the maximum potential to react with
the atmosphere or water, both by fluid flow and diffusion. Transport of dissolved solids by diffusion and
fluid flow (advection) is most efficient near the surface;
in the case of diffusion within about 1 m of the seabed.
The potential for sediments to change their bulk composition after deposition is therefore much higher at
shallow depth than at greater burial. Near the surface
on land, and also in the uppermost few centimetres of
the seabed, the conditions may be oxidising, while at
greater depth in the basin they are always reducing.
Precipitation of minerals due to porewater concentration by evaporation can only occur on land or at
shallow depth within enclosed basins (see Chap. 6).
On land, sediments are exposed to air and fresh
(meteoric) water. Weathering is partly due to reactions
with oxygen in the atmosphere and partly due to dissolution of minerals in freshwater, which is initially
undersaturated with respect to all the minerals present.
These are soil-forming processes which can be considered to be examples of early diagenesis.
In desert environments groundwater and occasional
rainwater may become concentrated through evaporation, causing precipitation of carbonates and also
silicates. Coatings of red or yellow iron oxides and
clays frequently form on desert sand and this may subsequently retard or prevent quartz cementation at greater
depth.
In the sea, the water above the seabed is normally
oxidising. Only where there is poor water circulation
(poor ventilation) is the lower part of the water column
likely to be reducing, though the phenomenom is more
widespread in lakes and inland seas like the Black Sea.
However, even below well-oxygenated water, oxidising
conditions extend in most cases for only a few
centimetres into the sediments, since oxygen is quickly
consumed by the oxidation (decay) of organic matter in
the sediment. This is for the most part facilitated biologically by bacteria. Accumulating sediments normally contain sufficient organic matter to serve as
reducing agents in the porewater. This organic matter
is comprised of both the remains of bottom fauna and of
pelagic organisms, including algae, accumulating on
the seafloor, and also often includes terrestrial plant
debris transported into the basin.
4.4
Redox-Driven Processes on the
Seafloor
Across the redox boundary there is a high gradient in
the concentration of oxygen and sulphate, and of ions
that have very different solubilities in oxygenated and
reduced water. The redox boundary is usually just
1–20 cm below the seafloor and represents equilibrium
between the supply of oxygen by diffusion, and its
consumption by the (mostly biological) oxidation of
organic matter. The oxygen content in the porewater
thus decreases rapidly below the water/sediment interface, providing a concentration gradient for the downward diffusion of oxygen into the uppermost sediments.
The rate of downward diffusion of oxygen is a
function of the concentration gradient of oxygen in
the porewater and the diffusion coefficient in the
sediments. The diffusion coefficient in coarse-grained
sand is higher than in mud and therefore sand tends to
have a deeper redox boundary than mud.
Oxygen can also be consumed in the sediments by
the oxidation of elements like iron and manganese in
minerals, but this is rare in marine environments and
more common in continental deposits. In most marine
environments there is enough organic matter to serve
as reducing agents and therefore little oxidation of iron
in minerals takes place, which explains why marine
sediments do not normally acquire a red colour. A
notable exception is red oxidised mud which may
form in marine environments characterised by slow
sedimentation rates and low organic productivity.
These muds are not very common but occur in some
deep-water facies and also in shallower water
environments with low sedimentation rates.
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K. Bjørlykke and J. Jahren
increase the strength of the grain framework and
prevent further mechanical compaction. The sandstone is then overconsolidated – not due to previously higher stress, but due to cementation. This is
sometimes referred to as pseudo-overconsolidation.
Further compaction will then mostly be controlled
by the rate of dissolution and precipitation.
In the following the diagenetic processes typical of
different burial depths will be discussed.
4.3
Early Diagenesis
As soon as sediments are deposited, early diagenetic
reactions start to modify the primary sediment composition. At very shallow burial depth (<1–10 m),
sediments have the maximum potential to react with
the atmosphere or water, both by fluid flow and diffusion. Transport of dissolved solids by diffusion and
fluid flow (advection) is most efficient near the surface;
in the case of diffusion within about 1 m of the seabed.
The potential for sediments to change their bulk composition after deposition is therefore much higher at
shallow depth than at greater burial. Near the surface
on land, and also in the uppermost few centimetres of
the seabed, the conditions may be oxidising, while at
greater depth in the basin they are always reducing.
Precipitation of minerals due to porewater concentration by evaporation can only occur on land or at
shallow depth within enclosed basins (see Chap. 6).
On land, sediments are exposed to air and fresh
(meteoric) water. Weathering is partly due to reactions
with oxygen in the atmosphere and partly due to dissolution of minerals in freshwater, which is initially
undersaturated with respect to all the minerals present.
These are soil-forming processes which can be considered to be examples of early diagenesis.
In desert environments groundwater and occasional
rainwater may become concentrated through evaporation, causing precipitation of carbonates and also
silicates. Coatings of red or yellow iron oxides and
clays frequently form on desert sand and this may subsequently retard or prevent quartz cementation at greater
depth.
In the sea, the water above the seabed is normally
oxidising. Only where there is poor water circulation
(poor ventilation) is the lower part of the water column
likely to be reducing, though the phenomenom is more
widespread in lakes and inland seas like the Black Sea.
However, even below well-oxygenated water, oxidising
conditions extend in most cases for only a few
centimetres into the sediments, since oxygen is quickly
consumed by the oxidation (decay) of organic matter in
the sediment. This is for the most part facilitated biologically by bacteria. Accumulating sediments normally contain sufficient organic matter to serve as
reducing agents in the porewater. This organic matter
is comprised of both the remains of bottom fauna and of
pelagic organisms, including algae, accumulating on
the seafloor, and also often includes terrestrial plant
debris transported into the basin.
4.4
Redox-Driven Processes on the
Seafloor
Across the redox boundary there is a high gradient in
the concentration of oxygen and sulphate, and of ions
that have very different solubilities in oxygenated and
reduced water. The redox boundary is usually just
1–20 cm below the seafloor and represents equilibrium
between the supply of oxygen by diffusion, and its
consumption by the (mostly biological) oxidation of
organic matter. The oxygen content in the porewater
thus decreases rapidly below the water/sediment interface, providing a concentration gradient for the downward diffusion of oxygen into the uppermost sediments.
The rate of downward diffusion of oxygen is a
function of the concentration gradient of oxygen in
the porewater and the diffusion coefficient in the
sediments. The diffusion coefficient in coarse-grained
sand is higher than in mud and therefore sand tends to
have a deeper redox boundary than mud.
Oxygen can also be consumed in the sediments by
the oxidation of elements like iron and manganese in
minerals, but this is rare in marine environments and
more common in continental deposits. In most marine
environments there is enough organic matter to serve
as reducing agents and therefore little oxidation of iron
in minerals takes place, which explains why marine
sediments do not normally acquire a red colour. A
notable exception is red oxidised mud which may
form in marine environments characterised by slow
sedimentation rates and low organic productivity.
These muds are not very common but occur in some
deep-water facies and also in shallower water
environments with low sedimentation rates.
122
K. Bjørlykke and J. Jahren