continents. On the continents there was active rifting
without seafloor spreading and this increased the geothermal gradient and elevated areas which were previously covered by shallow epicontinental seas, to above
sea level.
The short-term variations in sea level are more
difficult to explain, however. For geological periods
which experienced major continental glaciations
(Quaternary-Upper Miocene, Permian-Carboniferous,
Upper Ordovician and late Precambrian) we can resort
to glacioeustatic changes in sea level (changes in sea
level due to glaciation), which are very rapid in geological terms. Widespread sea level changes in the
Lower Tertiary and Mesozoic must be due to plate
tectonic factors.
Subsidence of the seafloor and continental crust of
various thicknesses is a function of time after the
rifting stage. The spreading ridges with hot basalt
(oceanic crust) often lie at about 2 km water depth or
above sea level (island). Seafloor basalt gradually
subsides to almost 6 km due to cooling over 100 ma
without sediment loading (Fig. 8.7a). Sedimentation
on top of the sea floor basalt will gradually reduce the
average density of the crust down to the compensation
depth (about 100 km) and the water depth will then be
reduced. Great water depth is found on old continental
crust where there is little sediments (Fig. 8.7b).
8.8
Sedimentation and Isostatic
Equilibrium
8.8.1 Why Do Sedimentary Basins
Subside?
Sedimentary basins can be assumed to be in isostatic
equilibrium in relation to the Earth’s crust. However,
the crust has a certain rigidity, and it takes time before
equilibrium is attained after loading. Studies of uplift
curves, for example from Scandinavia, show that during the course of 10,000 years the crust has undergone
major adjustments in the form of uplift to compensate
for the unloading of ice after the last glaciation. This is
geologically a very fast response and we can assume
that major sedimentary basins are in approximate isostatic equilibrium with regard to most geological processes. Nevertheless, even the filling of water
reservoirs for hydroelectric plants causes local
subsidence.
For the most part, then, we can apply the classical
Airy isostasy model to sedimentary basins, which
enables us to draw a number of interesting
conclusions.
Uplift and subsidence are also linked to variations
in geothermal gradients and heat flow in sedimentary
basins.
Movement of rocks in relation to the surface affects
the geothermal gradients. Erosion removes the uppermost, colder strata so that warmer strata come closer to
the surface, and the geothermal gradient increases. As
a result of subsidence of a sediment basin and sedimentation, heat flow upward will be partly offset by
rock subsidence, giving lower geothermal gradients.
Sedimentary basins with high rates of sedimentation
are therefore often characterised as “cold basins”.
Tectonic elevation and erosion will increase geothermal gradients.
The geothermal gradient in seafloor rocks is consistently greater than it is over the continents, and on the
continents it is highest in areas of volcanic activity.
Areas with a high geothermal gradient due to volcanism will slowly cool down to normal gradients
when volcanic activity ceases. It may take about 100
million years before a normal geothermal gradient is
re-established, due to cooling and contraction in accordance with the crust’s coefficient of expansion and
isostatic subsidence due to increased density. Seafloor
basalt is hot and flows at relatively shallow depths in
the Earth’s crust, and the spreading oceanic ridges are
only about 2.2 km below the surface of the sea. After
about 180 million years of cooling, the water depth at
isostatic equilibrium is about 5.7 km without sediment
loading. With sediment loading the oceanic crust may
subside to about 17 km (Fig. 8.7a). This is the theoretical maximum thickness of a sedimentary sequence
overlying oceanic crust. Sediment basins on the continental crust have a sedimentary thickness which is a
function of the thickness of the crust and the density of
the sediments and the basement. The thinner the continental crust beneath a sedimentary basin, the more
sediments can accumulate while maintaining isostatic
equilibrium (Fig. 8.8). Along continental margins sedimentary basins have formed on thin (stretched) continental and oceanic crust.
The formation of sedimentary basins clearly
requires that the density of the rocks below the basin
is greater than that of the rocks which surround it. The
sediment and water which fill the basin are lighter and
8 Seismic Stratigraphy, Sequence Stratigraphy and Basin Analysis
263
without seafloor spreading and this increased the geothermal gradient and elevated areas which were previously covered by shallow epicontinental seas, to above
sea level.
The short-term variations in sea level are more
difficult to explain, however. For geological periods
which experienced major continental glaciations
(Quaternary-Upper Miocene, Permian-Carboniferous,
Upper Ordovician and late Precambrian) we can resort
to glacioeustatic changes in sea level (changes in sea
level due to glaciation), which are very rapid in geological terms. Widespread sea level changes in the
Lower Tertiary and Mesozoic must be due to plate
tectonic factors.
Subsidence of the seafloor and continental crust of
various thicknesses is a function of time after the
rifting stage. The spreading ridges with hot basalt
(oceanic crust) often lie at about 2 km water depth or
above sea level (island). Seafloor basalt gradually
subsides to almost 6 km due to cooling over 100 ma
without sediment loading (Fig. 8.7a). Sedimentation
on top of the sea floor basalt will gradually reduce the
average density of the crust down to the compensation
depth (about 100 km) and the water depth will then be
reduced. Great water depth is found on old continental
crust where there is little sediments (Fig. 8.7b).
8.8
Sedimentation and Isostatic
Equilibrium
8.8.1 Why Do Sedimentary Basins
Subside?
Sedimentary basins can be assumed to be in isostatic
equilibrium in relation to the Earth’s crust. However,
the crust has a certain rigidity, and it takes time before
equilibrium is attained after loading. Studies of uplift
curves, for example from Scandinavia, show that during the course of 10,000 years the crust has undergone
major adjustments in the form of uplift to compensate
for the unloading of ice after the last glaciation. This is
geologically a very fast response and we can assume
that major sedimentary basins are in approximate isostatic equilibrium with regard to most geological processes. Nevertheless, even the filling of water
reservoirs for hydroelectric plants causes local
subsidence.
For the most part, then, we can apply the classical
Airy isostasy model to sedimentary basins, which
enables us to draw a number of interesting
conclusions.
Uplift and subsidence are also linked to variations
in geothermal gradients and heat flow in sedimentary
basins.
Movement of rocks in relation to the surface affects
the geothermal gradients. Erosion removes the uppermost, colder strata so that warmer strata come closer to
the surface, and the geothermal gradient increases. As
a result of subsidence of a sediment basin and sedimentation, heat flow upward will be partly offset by
rock subsidence, giving lower geothermal gradients.
Sedimentary basins with high rates of sedimentation
are therefore often characterised as “cold basins”.
Tectonic elevation and erosion will increase geothermal gradients.
The geothermal gradient in seafloor rocks is consistently greater than it is over the continents, and on the
continents it is highest in areas of volcanic activity.
Areas with a high geothermal gradient due to volcanism will slowly cool down to normal gradients
when volcanic activity ceases. It may take about 100
million years before a normal geothermal gradient is
re-established, due to cooling and contraction in accordance with the crust’s coefficient of expansion and
isostatic subsidence due to increased density. Seafloor
basalt is hot and flows at relatively shallow depths in
the Earth’s crust, and the spreading oceanic ridges are
only about 2.2 km below the surface of the sea. After
about 180 million years of cooling, the water depth at
isostatic equilibrium is about 5.7 km without sediment
loading. With sediment loading the oceanic crust may
subside to about 17 km (Fig. 8.7a). This is the theoretical maximum thickness of a sedimentary sequence
overlying oceanic crust. Sediment basins on the continental crust have a sedimentary thickness which is a
function of the thickness of the crust and the density of
the sediments and the basement. The thinner the continental crust beneath a sedimentary basin, the more
sediments can accumulate while maintaining isostatic
equilibrium (Fig. 8.8). Along continental margins sedimentary basins have formed on thin (stretched) continental and oceanic crust.
The formation of sedimentary basins clearly
requires that the density of the rocks below the basin
is greater than that of the rocks which surround it. The
sediment and water which fill the basin are lighter and
8 Seismic Stratigraphy, Sequence Stratigraphy and Basin Analysis
263