8.4 Subsidenee Related to Teetonie Loading, Subduetion, ete.
411
the South Australian continental margin, corrected
for sediment loading. The good fit of many synthetic
curves with the real curves is taken as evidence that
the model assumptions are adequate.
The simulated isochrones for post-rift development (Fig. 8.17a and c) display decreasing sediment
accumulation with time. This results from the assumption that the average water depths in the cross
section have remained the same all the time. In reality, the sedimentation rates varied considerably and
therefore brought about changes in the water depths
as well as in the location of the shelf edge which often does not coincide with major structural boundaries of the continental margin.
The subsidence his tory can be subdivided into two
phases, a rifting stage with rapid tectonic subsidence
lasting around 25 to 30 Ma, and a post-rift margin
development (drift phase, about 100 Ma) with distinctly slower rates of subsidence as a consequence
of cooling and contraction of the lithosphere. Minor
modifications of these models are discussed in Sect.
8.2; sea level-induced unconformities in coastal areas
are described in Sect. 7.2.
The sedimentary sequences of subsiding passive
continental margins are discussed in Sects. 12.2 and
12.3.
8.3.3
Summary (Simulated and Observed Subsidence)
The various "forward models" of rift basins and
passive continental margins can adequately predict the overall thermo-tectonic and total subsidence history of these basins. They therefore
lead to a better understanding of the processes
controlling basin formation.
Crustal stretching during rifting leads to rapid, ±
linear initial subsidence which is followed by
long-Iasting, exponentially decaying thermal
subsidence.
The subsidence rates on passive margins tend to
increase toward the outer margin due to crustal
thinning and adjustment to thecooling oceanic
crust.
8.4 Subsidence Related to
Tectonic Loading, Subduction, and
Strike-Slip Motion
8.4.1 Foreland Basins
General Aspects and Problems
Foreland basins develop on continental crust loaded
by the vertical weight of an approaching overthrust
belt as well as by the lateral force exerted by plate
compression (cf. Sect. 12.6). The subsidence history
of these basins is controlled mainly byflexure of the
underlying lithosphere but also strongly affected by a
stepwise overthrust process below or above sea level,
irregular unloading due to sub aerial erosion of thrust
sheets, and achanging geothermal gradient. For all
these reasons, modeling the development of foreland
basins is a difficult task and therefore not further
- Real subsidence curves largely confirm these
statements, but indicate a number of complications mentioned in Sect. 8.2.5.
- Rifting and crustal extension mayaiso occur in
several distinct phases and thus cause subsidence curves with repeatedly steepening gradients.
- Sediment-filled basins, in particular, cause flexural bending of the adjacent continental crust
beyond a hinge zone. Passive margin sediments
on transitional crust may be accompanied by a
thick and wide sediment wedge.
treated here. The following remarks may indicate
some of the problems encountered in this modeling
(cf. Sects. 11.3.5 and 12.6).
Although the quantification of overthrust loads and their
temporal change appear to be difficult, several authors have
made an attempt to model such a situation for different
time steps (e.g., Watts et al. 1982; Quinlan and Beaumont
1984; Flemings and Jordan 1989; Stockmal et al. 1992;
Slingerland et al. 1994). There are, however, examples
(e.g., the Apennines and Carpathians) where the overthrust
load of the mountain belt may be insufficient to create the
subsidence and basin fill of the existing foreland basins. In
these cases, a deep basin already existed on top of the foreland plate prior to overthrusting. Thus the approaching
load did not produce a marked topographie high and its
effect on the underlying plate was limited. In many cases
the buildup of collision-related compressional stress seems
to play a major role in crustal deformation (Ziegler et al.
1998).
Another serious problem, as already mentioned in Sect.
8.3 and summarized by Allen et al. (1986), is the rheologi-
411
the South Australian continental margin, corrected
for sediment loading. The good fit of many synthetic
curves with the real curves is taken as evidence that
the model assumptions are adequate.
The simulated isochrones for post-rift development (Fig. 8.17a and c) display decreasing sediment
accumulation with time. This results from the assumption that the average water depths in the cross
section have remained the same all the time. In reality, the sedimentation rates varied considerably and
therefore brought about changes in the water depths
as well as in the location of the shelf edge which often does not coincide with major structural boundaries of the continental margin.
The subsidence his tory can be subdivided into two
phases, a rifting stage with rapid tectonic subsidence
lasting around 25 to 30 Ma, and a post-rift margin
development (drift phase, about 100 Ma) with distinctly slower rates of subsidence as a consequence
of cooling and contraction of the lithosphere. Minor
modifications of these models are discussed in Sect.
8.2; sea level-induced unconformities in coastal areas
are described in Sect. 7.2.
The sedimentary sequences of subsiding passive
continental margins are discussed in Sects. 12.2 and
12.3.
8.3.3
Summary (Simulated and Observed Subsidence)
The various "forward models" of rift basins and
passive continental margins can adequately predict the overall thermo-tectonic and total subsidence history of these basins. They therefore
lead to a better understanding of the processes
controlling basin formation.
Crustal stretching during rifting leads to rapid, ±
linear initial subsidence which is followed by
long-Iasting, exponentially decaying thermal
subsidence.
The subsidence rates on passive margins tend to
increase toward the outer margin due to crustal
thinning and adjustment to thecooling oceanic
crust.
8.4 Subsidence Related to
Tectonic Loading, Subduction, and
Strike-Slip Motion
8.4.1 Foreland Basins
General Aspects and Problems
Foreland basins develop on continental crust loaded
by the vertical weight of an approaching overthrust
belt as well as by the lateral force exerted by plate
compression (cf. Sect. 12.6). The subsidence history
of these basins is controlled mainly byflexure of the
underlying lithosphere but also strongly affected by a
stepwise overthrust process below or above sea level,
irregular unloading due to sub aerial erosion of thrust
sheets, and achanging geothermal gradient. For all
these reasons, modeling the development of foreland
basins is a difficult task and therefore not further
- Real subsidence curves largely confirm these
statements, but indicate a number of complications mentioned in Sect. 8.2.5.
- Rifting and crustal extension mayaiso occur in
several distinct phases and thus cause subsidence curves with repeatedly steepening gradients.
- Sediment-filled basins, in particular, cause flexural bending of the adjacent continental crust
beyond a hinge zone. Passive margin sediments
on transitional crust may be accompanied by a
thick and wide sediment wedge.
treated here. The following remarks may indicate
some of the problems encountered in this modeling
(cf. Sects. 11.3.5 and 12.6).
Although the quantification of overthrust loads and their
temporal change appear to be difficult, several authors have
made an attempt to model such a situation for different
time steps (e.g., Watts et al. 1982; Quinlan and Beaumont
1984; Flemings and Jordan 1989; Stockmal et al. 1992;
Slingerland et al. 1994). There are, however, examples
(e.g., the Apennines and Carpathians) where the overthrust
load of the mountain belt may be insufficient to create the
subsidence and basin fill of the existing foreland basins. In
these cases, a deep basin already existed on top of the foreland plate prior to overthrusting. Thus the approaching
load did not produce a marked topographie high and its
effect on the underlying plate was limited. In many cases
the buildup of collision-related compressional stress seems
to play a major role in crustal deformation (Ziegler et al.
1998).
Another serious problem, as already mentioned in Sect.
8.3 and summarized by Allen et al. (1986), is the rheologi-