8.3 Rift Basin Models and Observed Subsidence
small amount of uplift, accompanied by a relatively strong
uplift of the graben shoulders. For further information see
Plein (1993).
Jordan rift, Dead Sea. This rift system is still active
and also affected by strike-slip motion. It has experienced mean subsidence rates of at least 200-500
mlMa, possibly more if the youngest evolution of
this structure is considered (cf. Sect. 12.1 and Fig.
11.18).
East African rift, Baikai rift. Other young continental rift zones, such as the western branch of the
East African rift system which started to evolve 7 to
12 Ma ago, exhibit maximum average subsidence
rates of 500 to 1000 mlMa (based on data sumrnarized by Ebinger 1989). These high values are valid
for the master fault in asymmetric half-graben structures, and they differ markedly from basin to basin in
the segmented rift zone.
Recent data from the large Baikai rift system in
Central Asia exhibit two distinct phases of rapid subsidence with exponential decay (Moore Jr. et al.
1997). The rift developed along the transtensional
boundary between the Siberian craton and a mosaic
of microplates to the southeast since the past 30-40
Ma. A shift in the regional stress field in the Pliocene
(pure extension or transtension; Delvaux et al. 1997)
caused the second phase of rapid subsidence. Unloaded basement subsidence in the central basin
reached <:4 km (sediment-loaded approximately 8
km) since the onset of rifting, but much less in the
other parts of the basin. Because sediment supply
was not sufficient to fill the entire rift zone, its central part is still a very deep lake.
Older Continental Rift Zones
Older continental rift zones have frequently evolved
into wider elongate basins and displaya large degree
of total subsidence over the long term, although the
rate of subsidence may be relatively low.
Paris basin. Three wells in the Paris basin have revealed that central parts of the basin, i.e., the former
rift basin, began to subside earlier than locations outside of the rift in the flexurally controlled portion of
the basin (Fig. 8.13b). Areevaluation of the subsidence history of this basin has revealed that the longterm trend is interrupted several times by accelerated
and decelerated subsidence, possibly caused by
phases of compression or transpression (Loup and
Wildi 1994).
North Sea. This basin has a long and complicated
history of subsidence (Ziegler 1988), causing deposition of Permian to Mesozoic sediments up to several
kilometers in thickness (cf. Fig. 14.l1). A general405
ized curve of total subsidence in the sediment-filled
Viking graben, northern North Sea, shows three distinct phases of increased, fault-related subsidence
during the Triassic, late Jurassic, and early Tertiary
(Fig. 8.13b).
This general scheme was interrupted by a phase ofuplift in
the Paleocene followed by accelerated subsidence in the
Eocene (Nadin and Kuznir 1995). These exursions (not
shown in Fig. 8.13b) were probably caused by the activity
of the Ieeland rnagrnatic plume.
Most wells drilled in the central North Sea, representing
a failed rift zone, also indicate accelerated post-Middle
Cretaceous subsidence (around 20 mlMa), the general
trend of which is shown in Fig. 8.13b. Whether this subsidence was entirely due to continental stretching (e.g.,
Sclater and Christie 1980; Sclater and Shorey 1989; Beach
et al. 1987) or caused mainly by mineral phase transition in
the lower ernst (Artyushkov and Baer 1989) still appears to
be an open question (cf. Sects. 8.1.3 and 12.1).
Several phases of subsidence were also found for the
Mesozoic shelves of the Alpine Tethys (e.g Wildiet al.
1989). The Swiss Plateau and Jura Mountains, for example,
experienced major subsidence phases in the Triassie, early
Jurassie, middle Jurassie, and early Cretaceous, but only
part ofthese phases were active at one location.
8.3.2 Passive Margin Models and
Observed Subsidence
General Concepts
Passive continental margins evolve from a rift basin
or aseries of rift basins (cf. Sects. 12.l and 12.2)
into a wider oceanic basin due to the accretion of
new oceanic crust. They commonly develop on both
margins of the growing ocean and therefore include
part of the former rift basin. Continuing tectonic subsidence and sediment accumulation usually cause a
landward and seaward extension of the depositional
area (cf. Fig. 8.16). Under these conditions, flexural
response of the adjacent continental and oceanic
lithosphere becomes a primary factor in passive margin subsidence in addition to local, Airy-type isostatic compensation.
As mentioned in the previous section, some recent
simulations of rift basin development use a two-layer
model of the lithosphere. In addition, many seismic
investigations have revealed that a discontinuity exists between the upper and lower continental crnst,
commonly at a depth of about 10 to 15 km (Fig.
8.15a and b). This discontinuity is explained as the
transition from the brittle to the ductile state of the
crust. Extension of the brittle material leads to the
formation of listric faults and tilted blocks, whereas
the ductile lower crust is deformed by creep (Fig.
8.15b). Although the manner of stretching is different, it is nevertheless a uniform extension of the
crust, i.e., the thicknesses of both layers are reduced
by the same stretching factor (one-Iayer model).
small amount of uplift, accompanied by a relatively strong
uplift of the graben shoulders. For further information see
Plein (1993).
Jordan rift, Dead Sea. This rift system is still active
and also affected by strike-slip motion. It has experienced mean subsidence rates of at least 200-500
mlMa, possibly more if the youngest evolution of
this structure is considered (cf. Sect. 12.1 and Fig.
11.18).
East African rift, Baikai rift. Other young continental rift zones, such as the western branch of the
East African rift system which started to evolve 7 to
12 Ma ago, exhibit maximum average subsidence
rates of 500 to 1000 mlMa (based on data sumrnarized by Ebinger 1989). These high values are valid
for the master fault in asymmetric half-graben structures, and they differ markedly from basin to basin in
the segmented rift zone.
Recent data from the large Baikai rift system in
Central Asia exhibit two distinct phases of rapid subsidence with exponential decay (Moore Jr. et al.
1997). The rift developed along the transtensional
boundary between the Siberian craton and a mosaic
of microplates to the southeast since the past 30-40
Ma. A shift in the regional stress field in the Pliocene
(pure extension or transtension; Delvaux et al. 1997)
caused the second phase of rapid subsidence. Unloaded basement subsidence in the central basin
reached <:4 km (sediment-loaded approximately 8
km) since the onset of rifting, but much less in the
other parts of the basin. Because sediment supply
was not sufficient to fill the entire rift zone, its central part is still a very deep lake.
Older Continental Rift Zones
Older continental rift zones have frequently evolved
into wider elongate basins and displaya large degree
of total subsidence over the long term, although the
rate of subsidence may be relatively low.
Paris basin. Three wells in the Paris basin have revealed that central parts of the basin, i.e., the former
rift basin, began to subside earlier than locations outside of the rift in the flexurally controlled portion of
the basin (Fig. 8.13b). Areevaluation of the subsidence history of this basin has revealed that the longterm trend is interrupted several times by accelerated
and decelerated subsidence, possibly caused by
phases of compression or transpression (Loup and
Wildi 1994).
North Sea. This basin has a long and complicated
history of subsidence (Ziegler 1988), causing deposition of Permian to Mesozoic sediments up to several
kilometers in thickness (cf. Fig. 14.l1). A general405
ized curve of total subsidence in the sediment-filled
Viking graben, northern North Sea, shows three distinct phases of increased, fault-related subsidence
during the Triassic, late Jurassic, and early Tertiary
(Fig. 8.13b).
This general scheme was interrupted by a phase ofuplift in
the Paleocene followed by accelerated subsidence in the
Eocene (Nadin and Kuznir 1995). These exursions (not
shown in Fig. 8.13b) were probably caused by the activity
of the Ieeland rnagrnatic plume.
Most wells drilled in the central North Sea, representing
a failed rift zone, also indicate accelerated post-Middle
Cretaceous subsidence (around 20 mlMa), the general
trend of which is shown in Fig. 8.13b. Whether this subsidence was entirely due to continental stretching (e.g.,
Sclater and Christie 1980; Sclater and Shorey 1989; Beach
et al. 1987) or caused mainly by mineral phase transition in
the lower ernst (Artyushkov and Baer 1989) still appears to
be an open question (cf. Sects. 8.1.3 and 12.1).
Several phases of subsidence were also found for the
Mesozoic shelves of the Alpine Tethys (e.g Wildiet al.
1989). The Swiss Plateau and Jura Mountains, for example,
experienced major subsidence phases in the Triassie, early
Jurassie, middle Jurassie, and early Cretaceous, but only
part ofthese phases were active at one location.
8.3.2 Passive Margin Models and
Observed Subsidence
General Concepts
Passive continental margins evolve from a rift basin
or aseries of rift basins (cf. Sects. 12.l and 12.2)
into a wider oceanic basin due to the accretion of
new oceanic crust. They commonly develop on both
margins of the growing ocean and therefore include
part of the former rift basin. Continuing tectonic subsidence and sediment accumulation usually cause a
landward and seaward extension of the depositional
area (cf. Fig. 8.16). Under these conditions, flexural
response of the adjacent continental and oceanic
lithosphere becomes a primary factor in passive margin subsidence in addition to local, Airy-type isostatic compensation.
As mentioned in the previous section, some recent
simulations of rift basin development use a two-layer
model of the lithosphere. In addition, many seismic
investigations have revealed that a discontinuity exists between the upper and lower continental crnst,
commonly at a depth of about 10 to 15 km (Fig.
8.15a and b). This discontinuity is explained as the
transition from the brittle to the ductile state of the
crust. Extension of the brittle material leads to the
formation of listric faults and tilted blocks, whereas
the ductile lower crust is deformed by creep (Fig.
8.15b). Although the manner of stretching is different, it is nevertheless a uniform extension of the
crust, i.e., the thicknesses of both layers are reduced
by the same stretching factor (one-Iayer model).