Rift Basin Architecture & Evolution
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3.3. Tectonics of Rifting and Drifting: Pangea
Breakup
3.3.1. Rift Basin Architecture and Evolution
Roy W. Schlische & Martha Oliver Withjack
Department of Geological Sciences, Rutgers University, Piscataway, NJ 08854-8066 U.S.A.
Rift basins have been increasingly the focus of research in tectonics,
structural geology, and basin analysis. The reasons for this interest
include: (1) Rift basins are found on all passive (Atlantic-type)
continental margins and provide a record of the early stages of
(super)continental breakup. (2) The architecture of these basins and the
basin fill are strongly influenced by the displacement geometry on the
bounding normal fault systems (e.g., Gibson et al., 1989). Thus, aspects
of the evolution of these fault systems, including their nucleation,
propagation and linkage, can be extracted from the sedimentary record.
(3) Many modern and ancient extensional basins contain lacustrine
deposits (e.g., Katz, 1990) that are sensitive recorders of climate.
Milankovitch cycles (e.g., Olsen and Kent, 1999) recorded in these strata
provide a quantitative test of the predictions of basin-filling models
(e.g., Schlische and Olsen, 1990) that can, in turn, be used to infer
aspects of crustal rheology during rifting (e.g., Contreras et al., 1997).
(4) Many of the major petroleum provinces of the world are associated
with rift basins (e.g., the North Sea basins, the Jeanne d'Arc basin, the
Brazilian rift basins).
This section provides a brief overview of the rift basins related to
Pangean breakup, especially those along the central Atlantic margin
(e.g., Olsen, 1997). In particular, we examine (1) the structural
architecture of rift basins; (2) the interplay of tectonics, sediment
supply, and climate in controlling the large-scale stratigraphy of rift
basins; (3) how the sedimentary fill can be subdivided into
tectonostratigraphic packages that record continental rifting, initiation of
seafloor spreading, basin inversion, and drifting; and (4) how coring can
be used to answer fundamental questions related to these topics.
Structural Architecture
A typical rift basin is a fault-bounded feature known as a half graben
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(Fig. 3.3.1.1a). In a cross section oriented perpendicular to the boundary
fault (transverse section), the half graben has a triangular geometry
(Fig. 3.3.1.1b). The three sides of the triangle are the border fault, the
rift-onset unconformity between prerift and synrift rocks, and the
postrift unconformity between synrift and postrift rocks (or, for modern
rifts, the present-day depositional surface). Within the triangular wedge
of synrift units, stratal boundaries rotate from being subparallel to the
rift-onset unconformity to being subparallel to the postrift unconformity.
This fanning geometry, along with thickening of synrift units toward the
boundary fault, are produced by syndepositional faulting. Core from the
Newark basin confirms the thickening relationships (see Section 3.3.2).
Synrift strata commonly onlap prerift rocks. In a cross section oriented
parallel to the boundary fault (longitudinal section), the basin has a
synclinal geometry (Fig. 3.3.1.1c), although more complicated
geometries are associated with segmented boundary fault systems (e.g.,
Schlische, 1993; Schlische and Anders, 1996; Morley, 1999).
Figure 3.3.1.1. Geometry of a simple half graben. (a) Map-view
geometry. (b) Geometry along a cross section oriented perpendicular
to the boundary fault, showing wedge-shaped basin in which synrift
strata exhibit a fanning geometry, thicken toward the boundary fault,
and onlap prerift rocks. (c) Geometry along a cross section oriented
parallel to the boundary fault, showing syncline-shaped basin in which
synrift strata thin away from the center of the basin and onlap prerift
rocks.
The half-graben geometry described above is directly controlled by the
deformation (displacement) field surrounding the boundary fault system
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(Gibson et al., 1989; Schlische, 1991, 1995; Schlische and Anders, 1996;
Contreras et al. 1997). In a gross sense, displacement is greatest at the
center of the fault and decreases to zero at the fault tips (Fig. 3.3.1.2a);
this produces the syncline-shaped basin in longitudinal section. In
traverse section, the displacement of an initially horizontal surface that
intersects the fault is greatest at the fault itself and decreases with
distance away from the fault. This produces footwall uplift and
hanging-wall subsidence, the latter of which creates the sedimentary
basin (Fig. 3.3.1.2b). However, this geometry is affected by fault
propagation and forced folding (e.g., Withjack et al., 1990; Gawthorpe et
al., 1997). As displacement accumulates on the boundary fault, the basin
deepens through time. Because the width of the hanging-wall deflection
increases with increasing fault displacement (Barnett et al., 1987), the
basin widens through time. Because the length of the fault increases
with increasing displacement (e.g., Cowie, 1998), the basin lengthens
through time. The growth of the basin through time produces
progressive onlap of synrift strata on prerift rocks (Fig. 3.3.1.3).
Figure 3.3.1.2. Fault-displacement geometry controls the first-order geometry
of a half graben. (a) Perspective diagram before (left) and after faulting
showing how normal faulting uplifts the footwall block and produces subsidence
in the hanging-wall block. The yellow dashed line shows the outer limit of
hanging-wall subsidence and marks the edge of the basin. Displacement is a
maximum at the center of the fault (only the right half of the fault is shown)
and decreases toward the fault tip. (b) Traverse section before faulting (left)
and after faulting and sedimentation showing footwall uplift and hanging-wall
subsidence. The latter produces a wedge-shaped basin (half graben).
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Figure 3.3.1.3.
Simple filling model
for a growing
half-graben basin
shown in map view
(stages 1-4),
longitudinal cross
section (stages 1-5),
and transverse cross
section (stages 1-4).
Dashed line
represents lake level.
The relationship
between capacity and
sediment supply
determines whether
sedimentation is
fluvial or lacustrine.
For lacustrine
sedimentation, the
relationship between
water volume and
excess capacity
determines the lake
depth. Modified from
Schlische and Anders
(1996).
The simple structural architecture described above may be complicated
by basin inversion, in which a contractional phase follows the
extensional phase (e.g., Buchanan and Buchanan, 1995). Typical
inversion structures include normal faults reactivated as reverse faults,
newly formed reverse and thrust faults, and folds (Fig. 3.3.1.4, 3.3.1.5).
Basin inversion occurs in a variety of tectonic environments (e.g.,
Buchanan and Buchanan, 1995), including several passive margins
related to the breakup of Pangea (e.g., Doré and Lundin, 1996; Vagnes
et al., 1998; Withjack et al., 1995, 1998; Hill et al., 1995; Withjack &
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Eisenstadt, 1999). The causes of inversion on these passive margins is
not well understood. Section 4.2.1 describes how coring, in combination
with other methods, may help further our understanding of basin
inversion on passive margins.
Figure 3.3.1.4. Examples of positive inversion structures. a) Cross section
across part of Sunda arc. During inversion, normal faults became reverse
faults, producing synclines and anticlines with harpoon geometries (after
Letouzey, 1990). b) Interpreted line drawings (with 3:1 and 1:1 vertical
exaggeration) of AGSO Line 110-12 from Exmouth sub-basin, NW Shelf
Australia (after Withjack & Eisenstadt, 1999). During Miocene inversion,
deep-seated normal faults became reverse faults. In response, gentle
monoclines formed in the shallow, postrift strata.
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Figure 3.3.1.5. Experimental
models of inversion structures.
Cross sections through three
clay models showing
development of inversion
structures (after Eisenstadt and
Withjack, 1995). In each
model, a clay layer (with
colored sub-layers) covered two
overlapping metal plates.
Movement of the lower plate
created extension or
shortening. Thin clay layers are
prerift; thick clay layers are
synrift; top-most layer is
postrift and pre-inversion. Top
section shows model with
extension and no shortening; a
half graben containing very
gently dipping synrift units is
present. The middle section
shows model with extension
followed by minor shortening; a
subtle anticline has formed in
the half graben, and is
associated with minor
steepening of the dip of synrift
layers. Bottom section shows
model with extension followed
by major shortening. The
anticline in the half graben is
more prominent, and is
associated with significant
steepening of the dip of synrift
strata. New reverse faults have
formed in the prerift layers.
Although the inversion is
obvious in this model, erosion of
material down to the level of
the red line would remove the
most obvious evidence of
inversion in the half graben.
Furthermore, the prominent
reverse faults cutting the prerift
units could be interpreted to
indicate prerift contractional
deformation, as is common in
the rift zones related to the
breakup of Pangea.
Stratigraphic Architecture
Numerous non-marine rift basins of varied geography and geologic age
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share a remarkably similar stratigraphic architecture (Lambiase, 1990;
Schlische and Olsen, 1990; Fig. 3.3.1.6). Known as a tripartite
stratigraphy, the section begins with basin-wide fluvial deposits overlain
by a relatively abrupt deepening-upward lacustrine succession overlain
by a gradual shallowing-upward lacustrine and fluvial succession. The
key to understanding the significance of this tripartite stratigraphy rests
in the relationships among basin capacity and sediment and water
supply (Schlische and Olsen, 1990; Carroll and Bohacs, 1999). Tectonics
creates accommodation space or basin capacity. Sediment supply
determines how much of that basin capacity is filled and whether or not
lake systems are possible (Figure 3.3.1.7). In general, fluvial deposition
results when sediment supply exceeds capacity, and lacustrine
deposition results when capacity exceeds sediment supply.
Figure 3.3.1.6. Stratigraphic
architecture of
Triassic-Jurassic rift basins of
eastern North America. For
tectonostratigraphic (TS)
package III, nearly all basins
exhibit all or part of a
tripartite stratigraphy: 1,
basal fluvial deposits; 2,
"deeper-water" lacustrine
deposits; 3, "shallow-water"
lacustrine and fluvial
deposits. The southern basins
do not contain TS-IV. TS-I is
only recognized in the Fundy
basin and may or may not be
a synrift deposit. Where TS-II
is recognized, a significant
unconformity (in terms of
missing time) commonly
separates it from TS-III.
Modified from Olsen (1997),
Olsen et al. (2000), and
Schlische (2000).
Figure 3.3.1.7. [BELOW] Relationships among basin capacity, sediment supply,
and volume of water determine the large-scale depositional environments of
terrestrial rift basins. In example 1, basin-wide fluvial sedimentation is predicted.
In example 2, shallow-water lacustrine sedimentation is predicted. For the basin
capacity and available sediment supply shown in this example, no very deep lakes
are possible because the excess capacity of the basin (and thus lake depth) is
limited. Thus, under these conditions, climate is a relatively unimportant control
on lake depth. In example 3, deep-water lacustrine sedimentation is predicted.
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The relationships shown in Figure 3.3.1.7 allow us to interpret the
large-scale stratigraphic transitions observed in many non-marine rift
basins. The fluvial-lacustrine transition may result from an increase in
basin capacity and/or a decrease in sediment supply. The shallow-water
lacustrine to deep-water lacustrine transition may result from an
increase in basin capacity, a decrease in sediment-supply, and/or
increase in the available volume of water. The deep-water lacustrine to
shallow-water lacustrine transition may result from a decrease or an
increase in basin capacity (depending on the geometry of the basin's
excess capacity), an increase in the sediment supply, and/or decrease in
the available volume of water. How do we go about choosing the more
likely interpretation? Interestingly, all of the major stratigraphic
transitions can be explained by an increase in basin capacity, for which
a simple basin-filling model is shown in Figure 3.3.1.3. Other basin
filling models are described by Lambiase (1990), Smoot (1991), and
Lambiase and Bosworth (1995). As discussed in Section 3.3.3, long cores
from rift basins, combined with basin modeling (e.g., Contreras et al.,
1997) and seismic reflection data (e.g., Morley, 1999), are required to
test the predictions of these basin-filling models.
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Figure 3.3.1.8. Idealized rift basin showing unconformity-bounded
tectonostratigraphic packages. Thin black lines represent stratal truncation
beneath unconformities; red half-arrows represent onlaps. In eastern
North America, TS-I may not be a synrift deposit, and thus the geometry
shown here would be incorrect. TS-II is much more areally restricted and
more wedge-shaped than TS-III. The transition between TS-III and TS-IV
is likely related to an increase in extension rate. An offset coring
technique (vertical orange lines), as used in the Newark basin coring
project, does not sample TS-I and most of TS-II. A deep core (vertical
yellow line) is necessary to recover TS-I and TS-II. Modified from Olsen
(1997).
Tectonostratigraphic Packages and Basin Evolution
Olsen (1997) subdivided the synrift strata of central Atlantic margin rift
basins into four tectonostratigraphic (TS) packages (Fig. 3.3.1.6,
3.3.1.8). An individual TS package consists of all or part of a tripartite
stratigraphic succession, is separated from other packages by
unconformities or correlative conformities, and generally has a different
climatic milieu compared to other TS packages. TS-I is a Permian
deposit that may or may not be synrift, whereas TS-II, TS-III, and TS-IV
are Late Triassic and Early Jurassic synrift deposits (Olsen et al., 2000).
The unconformities between TS-I, TS-II, and TS-III represent significant
geologic time. However, it is not yet clear if these unconformities are
related to regional tectonic changes (e.g., pulsed extension) (Olsen,
1997) or to relatively local processes such as strain localization (a
change from distributed extension on lots of small faults to extension on
a few large ones; e.g., Gupta et al., 1998) (Fig. 3.3.1.9). Given their
geometry and location in the rift basin, TS-I and TS-II can generally only
be sampled through deep coring and not the relatively shallow offset
coring utilized in the Newark basin (Section 3.3.3). The rift-onset
unconformity between prerift rocks and various synrift units should not
be taken as evidence of regional uplift preceding rifting; rather, it more
likely reflects erosion and non-deposition occurring over a
topographically elevated region resulting from the assembly of Pangea.
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Figure 3.3.1.9. Stages in the evolution of a rift basin. (a) Early rifting
associated with several minor, relatively isolated normal faults. (b) Mature
rifting with through-going boundary fault zone, widespread deposition, and
footwall uplift and erosion.
TS-III and TS-IV were deposited in much larger basins or subbasins than
was TS-II, and the unconformity between them is small to non-existent
(Olsen, 1997). TS-IV includes the widespread CAMP basalts that were
erupted in a geologically short interval at ~202 Ma (e.g., Olsen et al.,
1996; Olsen, 1999) (The CAMP basalts comprise a large-igneous
province or L.I.P.; see Section 3.1.3). Significantly, TS-IV is absent in all
of the southern basins of the central Atlantic margin. As discussed more
fully in Withjack et al. (1998), TS-IV was probably never deposited in
this region, indicating that synrift subsidence had ceased prior to TS-IV
time. [A postrift basalt sequence, which may or may not be the same
age as CAMP, is present in the southern region and plausibly can be
connected to a seaward-dipping reflector sequence at the continental
margin (Oh et al., 1995). The temporal and spatial relationships of these
igneous rocks is a critical coring target; see sections 4.2.1 and 4.2.2.]
Also significantly, basin inversion in the southern basins occurred shortly
prior to and during TS-IV time, while inversion in the northern basins
occurred after TS-IV time. (During TS-IV time, the northern basins
underwent accelerated subsidence; see Figure 3.3.2.7). Thus, the end of
rifting, the initation of inversion, and probably the initiation of seafloor
spreading are diachronous along the central Atlantic margin (i.e., during
earliest Jurassic time in the southeastern United States and Early to
Middle Jurassic time in the northeastern United States and Maritime
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Canada) (Withjack et al., 1998). Coring, field analysis, and
seismic-reflection profiles of synrift and immediately overlying postrift
deposits and the structures formed in them, are necessary to clarify the
important events occurring at the rift-drift transition.
The inferred diachronous initiation of seafloor spreading along the
present-day margin of the central North America Ocean is part of larger
trend that reflects the progressive dismemberment of Pangea. As the
North Atlantic Ocean continued to develop, seafloor spreading
propagated northward. For example, seafloor spreading between the
Grand Banks and southwestern Europe began during the Early
Cretaceous (e.g., Srivastava and Tapscott, 1986); seafloor spreading
between Labrador and western Greenland began during the early
Tertiary (anomaly 27N) (e.g., Chalmers, et al., 1993); whereas seafloor
spreading between eastern Greenland and northwestern Europe began
slightly later during the early Tertiary (anomaly 24R) (e.g., Talwani and
Eldholm, 1977; Hinz et al., 1993).
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