12.1 Rift Basins
12.1.5 Older Rift Basins
A large number of greatly varying rift basins was
formed during the Triassic and early Jurassic as a
result of the breakup of the Pangea supercontinent
prior to the formation of the Atlantic and Indian
oceans. These rift basins are mostly buried below
younger, post-rift sediments of passive continental
margins (cf. Sects. 12.1.6 and 12.2), but some are
also exhumed as a result of subsequent uplift.
In N orth Africa and Arabia, for example, older rift
structures along the southem rnargin of the Tethys
ocean were folded and inverted during Senonian time
(Guiraud and Bosworth 1997).
Many ancient rift zones which evölved into passive margins were later incorporated into orogenic
belts by convergent plate movements. Widely extended intracontinental basins frequently started as
rift basins which are now deeply buried as, e.g., in
the Paris basin or below the North Sea (cf. Sects.
12.4 and 14.3.3).
Good examples of ancient rift basins have been described
from the Alps (Winterer and Bosellini 1981; Lein 1985;
Eberli 1987; Lemoine and Trümpy 1987; Loup 1993). Rifting in the Alps occurred in the early and middle Jurassic
and pre-dated the opening of the Ligurian Tethys ocean by
spreading and transform movements. Tbe rift basins are
bordered by normal faults with throws up to a few kilometers. However, these are often difficult to identify because
of subsequent intensive deformation. Tbe best evidence for
former rift basins is therefore their basin fill, which is frequently characterized by megabreccias, olistostromes, and
turbidite fans of limited extent. In addition, such basin fills
display marked lateral changes both in facies and in thicknesses.
The High Atlas: an inverted rift zone. An outstanding example of an ancient, relatively short-lived
failed rift zone is the High Atlas in northwestem Africa. It was affected by strike-slip (transtension, cf.
Fig. 12.2c) and inverted by post-rift transpression
and uplift to an intracratonic mountain belt bordered
by elevated plateaus, the Saharan craton to the south,
and Gran Meseta micropiate to the north (Fig.
12.6a,d). The pre- and syn-rift evolution and sedimentary fill of this structure, which was inundated
early by the sea, is summarized in Figure 12.6b,c. It
shows a highly variable depositional history controlled by tectonism, eustatic sea-level fluctuations,
and climatic changes documented in continental and
shallow-marine sediments.
Tbe development of this structure may have resulted from
northward directed crustal detachment (Fig. 12.6b) associated with the breakup of Pangea during the Triassic after a
period of erosion and peneplanation. Rifting was accompanied by strike-slip motion; the rift zone extended from the
Triassic proto-Atlantic to the Tethys ocean over a distance
of more than 1000 km (see, e.g., Beauchamp 1988;
Jacobshagen 1988; Froitzheim et al. 1988; Gomez et al.
553
1996; Beauchamp et al. 1996). Tbe sedimentary fill of this
structure has been described, e.g., by Wurster and Stets
(1982), Warme (1988), Trappe (1991) and others.
During the first phase ofrifting (from the late Triassic to
early Jurassic), the rift graben of the High Atlas accumulated continental red beds and evaporites, and was affected
by syn-rift basaltic volcanism. Due to continuing rifting,
subsidence, and eustatic high sea-level stands, the sea entered the rift zone several times both from the west (protoAtlantic) and the east (Tethys) during the Jurassic, Cretaceous and Tertiary. In the central High Atlas, the widening
and deepening rift basin was subdivided by a central deep
carbonate platform and bordered by narrow carbonate
shelves (Fig. 12.6c). Tbe northern and southern troughs
were partially filled with turbidites and other rnass flow
deposits.
Beginning with the Toarcian high sea level, amiddie
Jurassic thick rnarly sequence was deposited in the central
to eastern regions of the High Atlas. Several shallowingupward cycles and unconformities indicate continued differential subsidence and tilting of blocks within the rift
zone. Periods of rapid subsidence and basin deepening altemated with phases in which the basin was filled up and
the depositional environment changed from marine to continental. During the Paleogene, biogenic carbonates and
reworked phosphorites accumulated in a marine, bay-like
ramp setting in the western regions.
Subsequent transpression caused uplift (basin inversion)
and steep-angle overthrusting and gravitational gliding of
rift sediments, locally including their basement, onto
molasse-like Tertiary sediments of the northem and southem Subatlas zones (Fig. 12.6d;
12.1.6 Rift Basins along Continental Margins
North AHantic
In recent years a great number of buried rift basins
have been detected by geophysical surveys and drilling below the present continental shelves, for exampIe in the northem Atlantic (Fig. 12.7, cf. Fig.
12.2d). These basins result from the breakup of
Pangea and the stretching and thinning of a 500 to
1000 km wide zone of continental crost during the
Mesozoic prior to continental breakup and sea-floor
spreading. Basins of this type are found below the
shelves of all passive continental margins. Some of
these basins are exposed on land where they commonly display simple structures and relatively short
time periods of basin filling, as described in Sects.
11.3 (Fig. 11.12) and in the facies models for halfgrabens (Fig. 12.4). In the following we distinguish
between rift basins related to initial rifting and 10cated distal to the younger ocean (cratonward) and
younger rift basins below the continental shelves 10cated more oceanward. Basins of the first group are
emerged on land, those of the second group are
mostly drowned below the sea.
Emerged early rift basins. The late Triassic-early
Jurassic rift basins along the Atlantic coast of North
12.1.5 Older Rift Basins
A large number of greatly varying rift basins was
formed during the Triassic and early Jurassic as a
result of the breakup of the Pangea supercontinent
prior to the formation of the Atlantic and Indian
oceans. These rift basins are mostly buried below
younger, post-rift sediments of passive continental
margins (cf. Sects. 12.1.6 and 12.2), but some are
also exhumed as a result of subsequent uplift.
In N orth Africa and Arabia, for example, older rift
structures along the southem rnargin of the Tethys
ocean were folded and inverted during Senonian time
(Guiraud and Bosworth 1997).
Many ancient rift zones which evölved into passive margins were later incorporated into orogenic
belts by convergent plate movements. Widely extended intracontinental basins frequently started as
rift basins which are now deeply buried as, e.g., in
the Paris basin or below the North Sea (cf. Sects.
12.4 and 14.3.3).
Good examples of ancient rift basins have been described
from the Alps (Winterer and Bosellini 1981; Lein 1985;
Eberli 1987; Lemoine and Trümpy 1987; Loup 1993). Rifting in the Alps occurred in the early and middle Jurassic
and pre-dated the opening of the Ligurian Tethys ocean by
spreading and transform movements. Tbe rift basins are
bordered by normal faults with throws up to a few kilometers. However, these are often difficult to identify because
of subsequent intensive deformation. Tbe best evidence for
former rift basins is therefore their basin fill, which is frequently characterized by megabreccias, olistostromes, and
turbidite fans of limited extent. In addition, such basin fills
display marked lateral changes both in facies and in thicknesses.
The High Atlas: an inverted rift zone. An outstanding example of an ancient, relatively short-lived
failed rift zone is the High Atlas in northwestem Africa. It was affected by strike-slip (transtension, cf.
Fig. 12.2c) and inverted by post-rift transpression
and uplift to an intracratonic mountain belt bordered
by elevated plateaus, the Saharan craton to the south,
and Gran Meseta micropiate to the north (Fig.
12.6a,d). The pre- and syn-rift evolution and sedimentary fill of this structure, which was inundated
early by the sea, is summarized in Figure 12.6b,c. It
shows a highly variable depositional history controlled by tectonism, eustatic sea-level fluctuations,
and climatic changes documented in continental and
shallow-marine sediments.
Tbe development of this structure may have resulted from
northward directed crustal detachment (Fig. 12.6b) associated with the breakup of Pangea during the Triassic after a
period of erosion and peneplanation. Rifting was accompanied by strike-slip motion; the rift zone extended from the
Triassic proto-Atlantic to the Tethys ocean over a distance
of more than 1000 km (see, e.g., Beauchamp 1988;
Jacobshagen 1988; Froitzheim et al. 1988; Gomez et al.
553
1996; Beauchamp et al. 1996). Tbe sedimentary fill of this
structure has been described, e.g., by Wurster and Stets
(1982), Warme (1988), Trappe (1991) and others.
During the first phase ofrifting (from the late Triassic to
early Jurassic), the rift graben of the High Atlas accumulated continental red beds and evaporites, and was affected
by syn-rift basaltic volcanism. Due to continuing rifting,
subsidence, and eustatic high sea-level stands, the sea entered the rift zone several times both from the west (protoAtlantic) and the east (Tethys) during the Jurassic, Cretaceous and Tertiary. In the central High Atlas, the widening
and deepening rift basin was subdivided by a central deep
carbonate platform and bordered by narrow carbonate
shelves (Fig. 12.6c). Tbe northern and southern troughs
were partially filled with turbidites and other rnass flow
deposits.
Beginning with the Toarcian high sea level, amiddie
Jurassic thick rnarly sequence was deposited in the central
to eastern regions of the High Atlas. Several shallowingupward cycles and unconformities indicate continued differential subsidence and tilting of blocks within the rift
zone. Periods of rapid subsidence and basin deepening altemated with phases in which the basin was filled up and
the depositional environment changed from marine to continental. During the Paleogene, biogenic carbonates and
reworked phosphorites accumulated in a marine, bay-like
ramp setting in the western regions.
Subsequent transpression caused uplift (basin inversion)
and steep-angle overthrusting and gravitational gliding of
rift sediments, locally including their basement, onto
molasse-like Tertiary sediments of the northem and southem Subatlas zones (Fig. 12.6d;
12.1.6 Rift Basins along Continental Margins
North AHantic
In recent years a great number of buried rift basins
have been detected by geophysical surveys and drilling below the present continental shelves, for exampIe in the northem Atlantic (Fig. 12.7, cf. Fig.
12.2d). These basins result from the breakup of
Pangea and the stretching and thinning of a 500 to
1000 km wide zone of continental crost during the
Mesozoic prior to continental breakup and sea-floor
spreading. Basins of this type are found below the
shelves of all passive continental margins. Some of
these basins are exposed on land where they commonly display simple structures and relatively short
time periods of basin filling, as described in Sects.
11.3 (Fig. 11.12) and in the facies models for halfgrabens (Fig. 12.4). In the following we distinguish
between rift basins related to initial rifting and 10cated distal to the younger ocean (cratonward) and
younger rift basins below the continental shelves 10cated more oceanward. Basins of the first group are
emerged on land, those of the second group are
mostly drowned below the sea.
Emerged early rift basins. The late Triassic-early
Jurassic rift basins along the Atlantic coast of North