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Complex rifted continental margins explained by dynamical
models of depth-dependent lithospheric extension
Ritske S. Huismans
Department of Earth Science, Bergen University, Bergen N-5007, Norway, and
Department of Oceanography, Dalhousie University, Halifax, Nova Scotia B3H 4J1, Canada
Christopher Beaumont
Department of Oceanography, Dalhousie University, Halifax, Nova Scotia B3H 4J1, Canada
ABSTRACT
Observations from a number of rifted margins (e.g., South Atlantic salt basin, midNorwegian margin, Exmouth Plateau) reveal wide regions of extremely attenuated crust and
depositional environments that indicate depth-dependent lithospheric extension. Although the
one-dimensional thermal-kinematic consequences of depth-dependent extension are understood, no comprehensive process-based explanation for the complex style of these margins
exists. Here, we present self-consistent numerical models of passive-margin formation that
explain the depth-dependent extension, the width of the margin, its characteristic tripartite
nature, and why such margins are prone to deposition of evaporites under appropriate climatic conditions. Some features that are important to reproducing the observed characteristics include decoupling between upper and lower parts of the lithosphere during stretching,
contrasting wide and narrow extensional styles above and below the decoupling level, and
progressive focusing of crustal extension toward the rift axis.
Keywords: rifting, passive margins, depth-dependent extension, dynamic modeling, South Atlantic,
salt basins.
INTRODUCTION
Subsidence of rifted continental margins
is explained by the mechanical and isostatic
response to lithospheric stretching (Bond et al.,
1983; McKenzie, 1978; Royden and Keen,
1980; Steckler and Watts, 1978). Despite the
ability of the uniform extension model (lithospheric stretching is uniform with depth;
McKenzie, 1978; Steckler and Watts, 1978)
to explain the general character of rifted margins, observations from some margins, including central South Atlantic (Contrucci et al.,
2004; Karner et al., 2003; Moulin et al., 2005),
Exmouth Plateau (Karner and Driscoll, 1999),
and the central and North Atlantic (Davis and
Kusznir, 2004; Funck et al., 2004), are not consistent with its predictions. For these distinctive
margins, wide regions of extremely attenuated
crust but relatively thin overlying synrift
sediments, the upper layers of which were
deposited in shallow seas, are more compatible with the depth-dependent extension model
(lithospheric stretching varies with depth;
Davis and Kusznir, 2004; Kusznir and Karner,
2007; Reemst and Cloetingh, 2000; Reston,
2007). Although the kinematics of depthdependent stretching are understood (Royden
and Keen, 1980), the conditions favoring this
style and the consequences for a complete rift
zone in space and time require investigation.
We show that dynamical models that lead to
depth-dependent extension explain characteristic features of these distinctive margins, and
we develop a template from the model results
that divides the margins into proximal, sag, and
distal (P, S, D) zones.
CENTRAL SOUTH ATLANTIC
CONJUGATE MARGINS
We first describe the properties of a South
Atlantic–type example of these distinctive
complex margins (Fig. 1). Features that require
explanation include: (1) a wide (~300 km)
region of thin continental crust (Contrucci et al.,
2004; Moulin et al., 2005); (2) faulted early synrift sedimentary basins (Moulin et al., 2005);
(3) undeformed late synrift sediments, including salt, deposited under inferred shallow-water
conditions in “sag” basins (Moulin et al., 2005);
and (4) an enigmatic lower-crustal layer proximal to West Africa (Moulin et al., 2005). The
tectono-stratigraphy of the African margin sedimentary basins indicates (Karner et al., 2003;
Moulin et al., 2005; Vagnes et al., 2005): (5) early
synrift Neocomian to Barremian (144–127 Ma)
strata in a half graben (implying crustal extension); (6) a thick and unfaulted late synrift
middle Barremian to middle Aptian section
(127–117 Ma), indicating shallow-water conditions; (7) late synrift, middle to late Aptian
(117–112 Ma), shallow-water evaporites; and
(8) postrift sediments. The middle to late synrift section (points 6 and 7) is also called the
pre-salt and salt sag basin (Karner et al., 2003).
The lack of identifiable faults in this section is
puzzling given that it formed during rifting and
prior to continental breakup, the latter of which
is inferred to have occurred ca. 112–110 Ma
(Karner et al., 2003; Moulin et al., 2005).
These observations are generally incompatible with uniform extension but are in better
agreement with simple one-dimensional (1-D)
depth-dependent extension kinematics (e.g.,
Fig. DR3 in the GSA Data Repository1), which
explain some of the distinctive features (the combination of points 1, 5, and 7 listed previously);
however, this type of model does not explain
the conditions favoring depth-dependent lithospheric extension, and the consequences for the
distribution and timing of extension across the
complete rift zone. We therefore turn to dynamical thermal-mechanical models (Fig. 2) to investigate depth-dependent extension and to seek an
explanation for the other characteristics.
NUMERICAL MODELS ILLUSTRATING
FORMATION OF RIFTED MARGINS
WITH DEPTH-DEPENDENT EXTENSION
Figure 2 shows two numerical rift models
that build on our previous thermal-mechanical
results (Huismans and Beaumont, 2002, 2003,
2005). An arbitrary Eulerian-Lagrangian finite1
GSA Data Repository item 2008039, supplemental information on South Atlantic passive margin
structure, model setup, and Models 1 and 2, is available online at www.geosociety.org/pubs/ft2008.htm,
or on request from [email protected] or Documents Secretary, GSA, P.O. Box 9140, Boulder, CO
80301, USA.
© 2008 The Geological Society of America. For permission to copy, contact Copyright Permissions, GSA, or [email protected].
GEOLOGY,
February
2008
Geology,
February
2008;
v. 36; no. 2; p. 163–166; doi: 10.1130/G24231A.1; 3 figures; Data Repository item 2008039.
163
B
C
Stratigraphy
Postrift
S America
112
117
121
7
0
6
10
6
o
S
127
30
-20
0
20
5
5
144
0
E
Salt
300
7
Pre-Salt Sag Basin
3
2
5
Crust
600
(m/m.y.)
Brazilian Camamu Margin
Early synrift
West African north Angolan Margin
Salt
Late synrift and postrift
OCB
10
Postrift
8
Late synrift
6
1
Crust
E
8
Salt
10
5
4 7,2-7,4 km/s
20
Crust
20
Mantle
Mantle
30
0
0
100
200
Distance (km)
100
0
Depth (km)
Depth (km)
0
D
6
Time (m.y.)
20
-40
7
Late synrift
Salt
Africa
Early synrift
A South Atlantic Salt Province
-60
Deposition Rate
30
Figure 1. Approximately conjugate crustal cross sections and sedimentation history from
central South Atlantic rifted margins, modified after Contrucci et al. (2004), Henry et al.
(2005), Moulin et al. (2005), Rosendahl et al. (2005), and Vagnes et al. (2005). Numbers 1–8
indicate characteristic features discussed in text. A: Location of South Atlantic Salt Province
on Brazilian and West African passive margins. B–C: Stratigraphy and deposition rate versus
time showing low rates during early synrift stage from 144 to 127 Ma (5), high rates during
late synrift stage (127–122 Ma) (6), and low rates during deposition of salt at end of synrift
interval (7). D: Detail of tectono-stratigraphy (Vagnes et al., 2005) showing fault-bounded
early synrift sediments (2) overlain by undeformed late synrift clastics (6) and evaporites (7).
E: Approximately conjugate crustal cross sections from Brazil and West African Angola
margin between latitudes 2°S and 8°S showing wide, strongly thinned basement (1), thin
early synrift deposits (5), thick undeformed late synrift deposits (6), and intermediate-velocity
body (4) of unknown origin at base of crust. Camamu margin section is closest published
conjugate and is used here as a representative conjugate for Angolan margin (see also GSA
Data Repository [see text footnote 1]). OCB—ocean continent boundary.
element method (Fullsack, 1995) is used to
model extension of a layered lithosphere with
frictional-plastic and thermally activated powerlaw viscous rheologies (for model properties,
see GSA Data Repository). The key to the
behavior of these models is the weak crustal
rheology (reduced wet quartz viscosity, WQ/10),
which decouples deformation in the upper
crust from that in the mantle lithosphere over
a broad region. These simple models (see also
GSA Data Repository) represent many we have
investigated and differ only in that model 1 has a
strong (dry olivine) mantle lithosphere rheology,
whereas model 2 (wet olivine) is weaker.
Both models exhibit a two-phase synrift history of depth-dependent extension. Phase 1 consists of limited early stretching that is distributed
across a wide region in the crust and is matched
by concomitant necking of a narrower region of
underlying mantle lithosphere (Fig. 2A). This
phase continues until mantle lithosphere ruptures beneath moderately extended crust, leaving the base of the middle crust in contact with
upwelled sublithospheric mantle (Fig. 2B). In
phase 2, crustal extension becomes focused in
the distal margin, leading to crustal necking and
lithospheric breakup (Fig. 2C). Crustal stretching ceases in more proximal parts of the margin
because the weak lower crust has been attenuated by shear, and the remaining crust couples to
the underlying strong mantle, thereby limiting or
eliminating differential stretching between crust
164
and mantle (Fig. 2C). In phase 2, ruptured mantle lithosphere is advected laterally (Fig. 2C).
Model 2 is similar to model 1 with the addition
of a small-scale convective mantle instability.
This instability leads to both enhanced removal
of the lower mantle lithosphere and upwelling
of hot sublithospheric mantle, thereby widening the region in which the mantle lithosphere is
partially or completely removed (Fig. 2D). Convective removal also means that geometrically
inferred crust and mantle lithosphere extension
will no longer balance. These particular models
are nearly symmetric, and rupture of the continental crust occurs after ~250 km of extension.
Sensitivity tests show that lithospheric heterogeneities and/or strain softening will likely lead
to asymmetric margins.
Distributed depth-dependent extension has
divided the margins into three zones, proximal
(P), sag (S), and distal (D) (Figs. 2 and 3), each
of which has characteristic patterns of subsidence, sedimentation, and sediment deformation. These patterns can be understood from the
predicted distributions of the crust and mantle
lithosphere attenuation factors, γc (x) and γm (x)
(caption, Fig. 2), and the way the distribution
of strain evolves and is partitioned between rift
phases 1 and 2 (Figs. 2 and 3).
In zone P, proximal to the continent, the crust
undergoes phase 1 extensional attenuation, but
not in phase 2 because the locus of stretching has
migrated into the distal margin, zone D. Although
the mantle lithosphere translates with respect to
the crust, it is not attenuated; therefore, there is
little or no syn- or postrift thermal subsidence.
Zone P generally correlates with the onshore margin rift basins. Zone D experiences both phase 1
and 2 crustal extension. The mantle lithosphere is
necked and excised during phase 1.
Zone S is transitional between zones P and
D. It experiences phase 1 crustal attenuation.
Extension may also be heterogeneous, leading to locally deeper grabens (Figs. 2A and 2C
and inset). Phase 1 and 2 attenuation of mantle
lithosphere increases toward zone D, the locus of
necking and convective flow (Fig. 2). We define
synrift phase 2 as starting after crustal extension
has migrated into the distal margin, zone D, but
we recognize that this migration will be progressive and diachronous. Cooling of upwelled mantle and thermal subsidence dominate in phase 2
(Fig. 2), but continued phase 2 removal of underlying mantle lithosphere can contribute to transient uplift. The amount of thermal subsidence
during the postrift interval correlates with γm (x).
IMPLICATIONS FOR THE ASSOCIATED
SEDIMENTARY BASINS
The implications for the development of
sedimentary basins can be seen from Figure 2
and from the zonal template (Fig. 3) This template summarizes model predictions to be tested
against observations for the case where sedimentation progrades symmetrically across the
rift, thereby filling basins in P and S during both
synrift phases but leaving D underfilled, as in
the dynamical models (Fig. 2). Sediments, once
deposited, will experience and record extension
depending on whether or not stretching of the
underlying crust continues. An important test of
the model is that the template (Fig. 3) should
correctly predict the corresponding distribution
of faulted and unfaulted sedimentary basins.
In zone P (Fig. 3), extension and subsidence
only occur during phase 1. Phase 1 basins will
therefore be cut by normal faults. Any phase
2 sediments that fill residual space will not be
faulted. Given the limited extension and subsidence in zone P, shallow-water sedimentation will dominate. In zone S (Fig. 3), phase 1
sediments will also record crustal extension
by normal faulting. Sediment will accumulate
in phase 2 if accommodation space remains
from phase 1 and/or as a consequence of thermal subsidence. When thermal subsidence is
significant, relatively thick unfaulted phase 2
sedimentary basins will develop overlying the
phase 1 faulted section. We interpret these to
be the pre-salt “sag” basins (zone “S”). Figure 1D (points 5, 6, and 7) shows a typical distal zone S basin with these characteristics, and
a late synrift evaporite, which is also compatible with shallow-water sedimentation during
phase 2 thermal subsidence. The thick postrift
GEOLOGY, February 2008
Model 1
1.0
A
Synrift Phase 1
γm
γc
z (km)
Upper
Lower
Crust
z
Mantle
Lithosphere
900
1100
Phase 1
0
130
1100
Tectonic subsidence
Sediments are faulted
Synrift Phase 2
γc ≤ 0.4
γm ~ 0.4
γc ≤ 0.2
γm ~ 0.2
γc ≤ 0.1
γm = 0
“Sag” basin
Crust
1.0
0.0
8
0
0
500
t = 13 m.y., Δx = –130 km
400
γm
500
γc
600
700
800
600
700
800
Convective
instability
Tectonic subsidence
Sediments faulted
Incremental Strain at Surface
γc → 1.0
γm = 1.0
900
1100
Thermal
subsidence
–50
z (km)
Moho
mantle lithosphere
1350
1300
400
B
γc ≤ 0.3
γm ~ 1.0
900
900
–100
P (Proximal)
Tapering
lower crust owing to shear
Upwelling
lithospheric
mantle
x
–50
S (“Sag”)
Sedimentary basin
Crust
Incremental Strain at Surface
0
0
End Phase 1
γc ≤ 0.4
γc ≤ 0.2 γc ≤ 0.1
γm → 0.4-0.7 γm ~ 0.2 γm = 0
“Sag” basin owing to No subsidence
thermal subsidence
Postrift - Thermal subsidence
1100
–100
No tectonic subsidence
Sediments are not faulted
900
00
13
900
Major subsidence
Major
Minor No subsidence
1350
400
1.0
C
1300
500
600
700
t = 26 m.y., Δx = 260 km
800
γm
γc
0.0
8
0
Incremental Strain at Surface
P
0
S
S
D
P
z (km)
00
91
100
–50
0
130
900
900
–100
1100
End Phase 2
400
1300
500
600
D
1350
Model 2
400
0
700
800
1350
t = 18 m.y, Δx = 180 km
S
500
600
D
S
P
700
800
900
1100
–50
00
13
900
900
1100
1100
400
0
End
1300 Phase 2
13
0
–100
section (Fig. 1D) that overlies this synrift basin
is consistent with a zone S basin with substantial thermal subsidence, which began in synrift
phase 2 and continued, and resulted from significant thinning of underlying mantle lithosphere. Such thinning can cause decompression
melting of upwelled sublithospheric mantle,
consistent with an explanation for the enigmatic
layer (Fig. 1E) as underplated melt. In regard
to our South Atlantic example, the model provides a potential explanation for the faulting in
the basal synrift sediments of rift basins such as
the onshore Inner Kwanza, Congo, and Gabon
basins (Hudec and Jackson, 2004; Karner
et al., 1997). These basins are either classified
as P or S depending on the respective evidence
for/against phase 2 thermal subsidence.
GEOLOGY, February 2008
D (Distal)
t = 7 m.y., Δx = –70 km
0.0
8
z (km)
Figure 2. Numerical
thermo-mechanical models
of lithospheric extension.
A–C: Model 1. D: Model
2. Panels show deformed
Lagrangian mesh, velocity, and temperature.
Crust and mantle attenuation factors are defined
by γc (x) = 1 – 1/δ(x) and
γm (x) = 1 – 1/β(x), where
δ(x) and β(x) are crustal
and mantle lithosphere
thinning factors h0c /hc (x)
and h0m /hm (x), respectively. Increment in strain
at surface is with respect
to previous panel and indicates where sediments
will be faulted during this
interval. Total extension
velocity is V = 1 cm/yr.
Sediments prograde symmetrically onto model surface at constant velocity.
Materials deform plastically or viscously (details
are available in the GSA
Data Repository [see
text footnote 1]). A–B:
Model 1: phase 1, wide
zone of crustal extension, matched by narrow
zone of mantle necking
(model design in Fig. DR4
[see text footnote 1]). C:
Model 1: phase 2, crustal
extension focuses in distal margins and rift axis,
mantle lithosphere is
translated laterally, and
sediments prograde over
nonextending proximal
parts of rift zone. Incremental strain at surface
shows that sediment faulting in phase 2 is concentrated in zone D toward
rift center. D: Model 2:
phase 2, convective
removal of mantle lithosphere. P—proximal, S—
sag, D—distal.
50
13
500
600
700
800
Figure 3. Template summarizing main characteristics of dynamical models classified
according to zones proximal (P), sag (S),
and distal (D). γc and γm are crustal and lithospheric mantle attenuation factors defined in
Figure 2 caption. Phase 1 crustal extension
is distributed, producing limited attenuation
and subsidence; basins are faulted. Mantle
lithosphere extends by focused necking
and ruptures under D with some attenuation
under S. Phase 2 crustal extension migrates
to the rift axis, D. Additional faulting is confined to basins in D. Mantle lithosphere is
advected laterally. Unfaulted “sag” basins
develop where there is cooling and thermal
subsidence in zone S; transient uplift in S
may occur if mantle lithosphere is further
attenuated. Postrift thermal subsidence correlates with γm (x) and is confined to D and
S. Incremental strain at the surface accumulated during each of the stages shown in
A, B, and C indicates early deformed synrift
sediments and late synrift sediments are
undeformed.
x (km)
In zone D (Fig. 3), phase 1 basins may
develop depending on the competition between
subsidence caused by crustal extension and
uplift owing to removal of mantle lithosphere
during necking. During phase 2, there will be
substantial subsidence from the crustal thinning during extension that is now focused in this
zone, combined with thermal subsidence. Both
phase 1 and 2 sediments will be faulted except
where heterogeneous extension leaves pockets
of unextended crust. There will be major postrift
thermal subsidence.
CONCLUSIONS
We have shown how dynamical models
provide a general two-phase explanation for
the distribution of depth-dependent lithospheric
extension, subsidence, and sedimentation across
an entire rift. We have also shown that the tripartite proximal (P), sag (S), and distal (D) zonal
template, derived from the models, compares
favorably with our South Atlantic example, a
case where synrift sediments mostly fill the P
and S basins. The model may apply to other
margins, for example, the Exmouth Plateau
(Karner and Driscoll, 1999) and certain central
and North Atlantic passive margins (Davis and
Kusznir, 2004; Funck et al., 2004), including
those with less synrift sedimentation (Fig. 3).
In the absence of sediments, significant synrift
water depths can develop in the P and S basins
as indicated by the attenuation factors (Fig. 3).
This behavior requires weak crust at depth
that decouples crust and mantle lithosphere
165
during extension. Breakup of continental crust
in the models shown occurs after ~250 km of
extension, leading to a zone of extension that
spans a total width of up to 600 km. The span
includes the initial width of extending zone plus
total amount of extension at breakup. Lowerstrength crust/mantle coupling results in greater
penetration of the initial decoupling and extension into the continental crust and thus a wider
initial rift zone. This rifting style leads to: (1)
respectively wide and narrow distributions of
extension in the crust and mantle lithosphere; (2)
diachronous crustal deformation in which extension migrates to the distal margin by phase 2;
(3) phase 1 excision of mantle lithosphere from
beneath the crust at the rift axis; (4) no surface
flexural flank uplifts; (5) phase 1, proximal (P),
sag (S), and distal (D) zone sedimentary basins
that are relatively shallow and faulted; (6) phase
2 sag and distal sedimentary basins that contain,
respectively, shallow-water unfaulted, and deepwater faulted sediments; and (7) shallow marine
environments (5 and 6, this list), ideal for the
accumulation of evaporite sequences. In particular, the proposed thermal contraction mechanism
for phase 2 subsidence and coeval unfaulted sedimentation in the “sag” basins, zone S, is consistent with continued postrift thermal subsidence
of the same basins. For these basins, thermal
subsidence starts in synrift phase 2 and continues
into postrift time. Dynamical modeling is a key
to understanding the conditions required for this
complex style of rifting, to the calculation of
the time-space distribution of extension, and to
the predicted consequences for the subsidence,
sedimentation, and deformation of the syn- and
postrift sedimentary basins. It is significant that
such complex rifted margins can be attributed to
the evolving extension of a simple lithospheric
system. The only requirement is that crust and
mantle lithosphere decouple during extension.
ACKNOWLEDGMENTS
C. Beaumont was funded by the Canada Research
Chair in Geodynamics. Support was also provided
by an Atlantic Innovation Fund contract and an IBM
Shared University Research Grant. Numerical calculations used software developed by Philippe Fullsack.
We thank Nicky White, Nick Kusznir, and Garry
Karner for the constructive and thorough reviews.
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Manuscript received 28 June 2007
Revised manuscript received 9 October 2007
Manuscript accepted 9 October 2007
Printed in USA
GEOLOGY, February 2008