12.1 Rift Basins
Rift volcanism. Some of the rift zones and their sediments are strongly influenced by syn-rift volcanism
(Fig. 12.3), others not. Some authors even hold the
opinion that extensive rifting and continental breakup
is initiated by large mantle p1umes (hot spots) as,
e.g., postulated for the central Atlantic (Wilson
1997). When continental lithosphere is thinned by
stretching, the upwelling asthenosphere decompresses and generates partial melt (White 1989). The
hotter the asthenosphere, the more melt is generated
and the more up1ift of the crust can occur. Relatively
small variations in temperature cause major differences in the volume of magma generated. Large volumes of magma are produced if the lithosphere is
underlain by abnormally hot mantle material (forming plumes or "hot spots"). The partial melt is buoyant and rises to be added to the overlying crust or to
extrude at the surface. The phase prior to rifting may
be characterized by widely extended flood basalts.
Later, the rift valley can be fi1led by lava flows and
tephra of variable composition, or volcanic structures
are built up in the graben zone separating it into
subbasins.
Subsidence. The principal mechanism of subsidence in rift zones is not always c1ear. The widely
used stretching models (cf. Sect. 8.3) are not generally applicable. According to Artyushkov and Baer
(1989), there are two main modes of stretching (Fig.
12.1 e): (1) S-type (synthetic) stretching, in which the
fault blocks are not or only slightly tilted toward the
down-thrown side of a boundary fault; and A-type
(antithetic) stretching, in which the blocks are markedly tilted and separated by listric faults. Substantial
subsidence due to significant crustal thinning can be
attained only by A-type processes. If S-type rift basins exhibit great subsidence but limited extension,
the lowermost crust may have been altered or replaced by asthenospheric upwelling and injection of
mantle-derived material. This process is also referred
to as "basification" of the lower crust (e.g., Ziegler
1988). Then, the altered, moderately thinned crust
may break into wide blocks under tensile stresses and
subside isostatically along normal faults.
Fig. 12.2. a Rift zone composed of half-grabens with
alternatin~ polarities. Adjacent half-grabens are separated by mterbasinal ridges oblique to the rift axis
(transfer zones). Model of Lake Tanganyika (cf. Fig.
12.3a; Rosendahl et al. 1986). b Cretaceous branched
rift structure of the Bass Strait region between Austra1ia and Tasmania. (Etheridge et al. 1987). e Model
of Triassic rifting in the High Atlas, Morocco. Combined syndepositional extension and strike-s1ip displacement. Pre-existing N700E faults are reacbvated
by strike-slip motion and produce secondary N300E
normal faults within the graben structure.
(Beauchamp 1988; cf. Fig. 12.6). d Simplified model
547
12.1.3 Sediments of Rift Basins
General Aspeets
The great variety of structural style and depositional
environments mentioned above makes it difficult to
define a specific facies evolution for rift basins.
However, a succession from (1) fluvial deposits, (2)
lake deposits with a tendency towards playa conditions, and finally (3) areturn to fluvial deposits has
been frequently identified in relatively small, simple
rift structures (cf. Sect. 11.3.3). Rift systems along
evolving continental margins, receiving less
terrigenous material, often develop deeper marine
environments with upward deepening and shallowing
sediment successions. Deltaic and shelf facies may
pass into slope and foot-of-slope facies and later return to shallow-marine sediments in response to active and reduced phases of rifting. Varying rates of
subsidence can lead to changes in the sediment accommodation space and thus to prograding or
backstepping facies belts. Rifting below sea level on
thinned transitional crust and at great distances from
the continents may create, at least for some time,
sediment-starved basins with marine sequences of
limited thickness (cf. Fig. 12.14).
Larger, complex continental rift systems often
show less predictable sedimentary successions with
or without significant contributions of volcanic material. Pre-rift topography controls the positions of major rivers, which may or may not enter the basin and
thus govern its hydrological and sedimentological
budget. Some simplified facies models are discussed
in the following section.
Facies Models for Asymmetrie Rift Basins
Some idealized facies models for half-graben basins
exposed to various depositional environments are
shown in Fig. 12.4. All of them indicate a markedly
asymmetric basin fill caused by differing sediment
supply and depositional processes on the two sides of
the basin, particularly if most of the material is supp1ied from the flanks of the half-graben. Sediments
of rifted continental margin between Newfoundland
and Iberian Meseta, Spain (cf. Fig. 12.7a,c) prior to
ocean spreading. e Model of the development of a
secondary detachment in Jeanne d'Arc basin, Grand
Banks off Newfoundland (see d). Superposition of
detachment above and perpendicular to previous
extensional regime generates substantial dip-slip
movement along listric faults and thus strike-slip parallel to boundary fault ofextensional basin. (d,e After Tankard and Welsink 1988). fRift-drift transition
in the Red Sea. Older basins are abandoned and the
evolving new basin is laterally continuous. (After
Bosworth 1994)
Rift volcanism. Some of the rift zones and their sediments are strongly influenced by syn-rift volcanism
(Fig. 12.3), others not. Some authors even hold the
opinion that extensive rifting and continental breakup
is initiated by large mantle p1umes (hot spots) as,
e.g., postulated for the central Atlantic (Wilson
1997). When continental lithosphere is thinned by
stretching, the upwelling asthenosphere decompresses and generates partial melt (White 1989). The
hotter the asthenosphere, the more melt is generated
and the more up1ift of the crust can occur. Relatively
small variations in temperature cause major differences in the volume of magma generated. Large volumes of magma are produced if the lithosphere is
underlain by abnormally hot mantle material (forming plumes or "hot spots"). The partial melt is buoyant and rises to be added to the overlying crust or to
extrude at the surface. The phase prior to rifting may
be characterized by widely extended flood basalts.
Later, the rift valley can be fi1led by lava flows and
tephra of variable composition, or volcanic structures
are built up in the graben zone separating it into
subbasins.
Subsidence. The principal mechanism of subsidence in rift zones is not always c1ear. The widely
used stretching models (cf. Sect. 8.3) are not generally applicable. According to Artyushkov and Baer
(1989), there are two main modes of stretching (Fig.
12.1 e): (1) S-type (synthetic) stretching, in which the
fault blocks are not or only slightly tilted toward the
down-thrown side of a boundary fault; and A-type
(antithetic) stretching, in which the blocks are markedly tilted and separated by listric faults. Substantial
subsidence due to significant crustal thinning can be
attained only by A-type processes. If S-type rift basins exhibit great subsidence but limited extension,
the lowermost crust may have been altered or replaced by asthenospheric upwelling and injection of
mantle-derived material. This process is also referred
to as "basification" of the lower crust (e.g., Ziegler
1988). Then, the altered, moderately thinned crust
may break into wide blocks under tensile stresses and
subside isostatically along normal faults.
Fig. 12.2. a Rift zone composed of half-grabens with
alternatin~ polarities. Adjacent half-grabens are separated by mterbasinal ridges oblique to the rift axis
(transfer zones). Model of Lake Tanganyika (cf. Fig.
12.3a; Rosendahl et al. 1986). b Cretaceous branched
rift structure of the Bass Strait region between Austra1ia and Tasmania. (Etheridge et al. 1987). e Model
of Triassic rifting in the High Atlas, Morocco. Combined syndepositional extension and strike-s1ip displacement. Pre-existing N700E faults are reacbvated
by strike-slip motion and produce secondary N300E
normal faults within the graben structure.
(Beauchamp 1988; cf. Fig. 12.6). d Simplified model
547
12.1.3 Sediments of Rift Basins
General Aspeets
The great variety of structural style and depositional
environments mentioned above makes it difficult to
define a specific facies evolution for rift basins.
However, a succession from (1) fluvial deposits, (2)
lake deposits with a tendency towards playa conditions, and finally (3) areturn to fluvial deposits has
been frequently identified in relatively small, simple
rift structures (cf. Sect. 11.3.3). Rift systems along
evolving continental margins, receiving less
terrigenous material, often develop deeper marine
environments with upward deepening and shallowing
sediment successions. Deltaic and shelf facies may
pass into slope and foot-of-slope facies and later return to shallow-marine sediments in response to active and reduced phases of rifting. Varying rates of
subsidence can lead to changes in the sediment accommodation space and thus to prograding or
backstepping facies belts. Rifting below sea level on
thinned transitional crust and at great distances from
the continents may create, at least for some time,
sediment-starved basins with marine sequences of
limited thickness (cf. Fig. 12.14).
Larger, complex continental rift systems often
show less predictable sedimentary successions with
or without significant contributions of volcanic material. Pre-rift topography controls the positions of major rivers, which may or may not enter the basin and
thus govern its hydrological and sedimentological
budget. Some simplified facies models are discussed
in the following section.
Facies Models for Asymmetrie Rift Basins
Some idealized facies models for half-graben basins
exposed to various depositional environments are
shown in Fig. 12.4. All of them indicate a markedly
asymmetric basin fill caused by differing sediment
supply and depositional processes on the two sides of
the basin, particularly if most of the material is supp1ied from the flanks of the half-graben. Sediments
of rifted continental margin between Newfoundland
and Iberian Meseta, Spain (cf. Fig. 12.7a,c) prior to
ocean spreading. e Model of the development of a
secondary detachment in Jeanne d'Arc basin, Grand
Banks off Newfoundland (see d). Superposition of
detachment above and perpendicular to previous
extensional regime generates substantial dip-slip
movement along listric faults and thus strike-slip parallel to boundary fault ofextensional basin. (d,e After Tankard and Welsink 1988). fRift-drift transition
in the Red Sea. Older basins are abandoned and the
evolving new basin is laterally continuous. (After
Bosworth 1994)