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doi:10.1130/B26398.1
2009;121;55-64 Geological Society of America Bulletin
Xiangyang Xie and Paul L. Heller
Plate tectonics and basin subsidence history
Geological Society of America Bulletin
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© 2009 Geological Society of America
ABSTRACT
Tectonic setting exerts fi rst-order control
on basin formation as refl ected in basin sub-
sidence history. While our approach ignores
the effects of fl exural loading and eustatic
sea-level change, consistency of backstripped
subsidence histories (i.e., with local loading
effects of sediment removed) suggests con-
sistent tectonic driving mechanisms in each
tectonic setting, with the possible exception of
forearc basins.
Based on published subsidence curves and
open-fi le stratigraphic data, we show the sub-
sidence characteristics of passive margins,
strike-slip basins, intracontinental basins,
foreland basins, and forearc basins. Pas-
sive margin subsidence is characterized by
two stages, rapid initial, synrift subsidence
and slow post-rift thermal subsidence, with
increasing subsidence rates toward the adja-
cent ocean basin. Subsidence of intracontinen-
tal basins is similar in magnitude to that seen
in passive margin settings, but the former is
generally slower, longer lived, and lacks initial
subsidence. Long-lived subsidence for many
intracontinental basins is consistent with
cooling following thermal perturbation of
thick lithosphere found beneath old parts of
continents. Basins associated with strike-slip
faults are usually short lived with very rapid
subsidence. Changes in local stress regimes
as strike-slip faults evolve, and migrate over
time, coupled with three-dimensional heat
loss in these small basins likely explain this
subsidence pattern. Foreland basin subsi-
dence rates refl ect the fl exural response to
episodic thrust loading. Resultant subsidence
curves are punctuated by convex-up (acceler-
ating) segments. Forearc basins have the least
consistent subsidence patterns. Subsidence
histories of these basins are complex and may
refl ect multiple driving mechanisms of subsi-
dence in forearc settings.
Second-order deviations in subsidence sug-
gest reactivation or superimposed tectonic
events in many basin settings. The effects of
eustatic sea-level change may also explain
some deviations in curves. For many of these
settings, subsidence histories are suffi ciently
distinctive to be used to help determine tec-
tonic setting of ancient basin deposits.
Keywords: subsidence analysis, passive mar-
gins, intracontinental basins, foreland basins,
strike-slip basins, forearc basins.
INTRODUCTION
Sedimentary basins refl ect prolonged subsi-
dence of Earth’s surface, due to large-scale tec-
tonic processes operating between and within
plates (Kusznir and Ziegler, 1992). To the degree
that tectonic processes are refl ected in subsi-
dence history, basins in similar tectonic settings
should show similar patterns of subsidence.
Subsidence histories are taken to refl ect isostatic
adjustment to lithospheric processes, such as
thermal events, thickness changes, and loading
history. Therefore, subsidence history provides
insight into basin-forming mechanisms. Differ-
ences in subsidence histories between basins
may refl ect how fundamental driving mecha-
nisms vary as well as secondary infl uences, such
as sea-level change and sediment loading. By
comparing subsidence curves between different
basins of similar tectonic setting, it is possible to
determine consistency and/or differences in the
processes that drive subsidence.
Dickinson (1976) and Angevine et al. (1990)
compiled subsidence histories of basins in order
to discriminate between subsidence styles in
various tectonic settings. Since those studies,
more data have become available that may be
used to more fully defi ne subsidence patterns as
a function of tectonic setting. In this paper, we
use some of these data to demonstrate the styles
of subsidence in various plate tectonic settings,
emphasizing those settings where the modes
of subsidence are still poorly understood. Our
objective is to show that these histories can be
used as templates allowing tectonic interpreta-
tions based on subsidence patterns.
SUBSIDENCE ANALYSIS
Subsidence analysis yields a graphic repre-
sentation of the vertical movement of a strati-
graphic horizon, with respect to a datum in a
sedimentary basin. It tracks the subsidence and
uplift history at that location since the horizon
was deposited (van Hinte, 1978). Data needed
to reconstruct subsidence history include strati-
graphic thickness, lithology, estimate of paleo-
water depths, and age control. Subsidence anal-
ysis begins with a plot of sediment accumulation
through time using the present-day thickness
of each dated stratigraphic unit. Second, the
effects of compaction are included based on the
assumption that porosity lost is mostly caused
by mechanical compaction. Third, since sea
level is used as the datum for subsidence analy-
sis, paleobathymetry corrections are needed
to correct the seafl oor position to this datum.
The resulting curve refl ects the total subsidence
history (van Hinte, 1978) including the contri-
bution of tectonic loads, sediment loads, and
sea-level changes. Of these, the local isostatic
effects of sediment loading can be removed by
“backstripping” (Steckler and Watts, 1978).
The resulting subsidence curve, referred to
as “tectonic subsidence,” shows the idealized
subsidence history of a basin that would have
existed if only water, and no sediment, fi lled
the subsiding hole. Tectonic subsidence history
refl ects basin subsidence due to factors other
than sediment deposition and attendant isostatic
adjustment and compaction. More importantly,
it provides a way of normalizing subsidence in
different basins that have undergone very dif-
ferent sedimentation histories.
It is important to be aware of some limitations
to this analysis. These come from the inaccuracy
of data used to reconstruct history and from the
assumptions built into the method. In particu-
lar, age control and water depth often hamper
For permission to copy, contact editing@geosociety.org
© 2008 Geological Society of America
55
Plate tectonics and basin subsidence history
Xiangyang Xie
†*
Paul L. Heller
Department of Geology and Geophysics, University of Wyoming, Laramie, Wyoming 82071, USA
†
E-mail: xiangyang@utig.ig.utexas.edu
*Present address: Institute for Geophysics, Uni-
versity of Texas at Austin, 10100 Burnet Road Build-
ing 196, Austin, Texas 78758, USA.
GSA Bulletin; January/February 2009; v. 121; no. 1/2; p. 55–64; doi: 10.1130/B26398.1; 7 fi gures.
Xie and Heller
56 Geological Society of America Bulletin, January/February 2009
subsidence analysis. Water-depth estimates are
often diffi cult to determine because of a paucity
of unique depth indicators in sedimentary rocks.
However, the impact of poorly constrained water
depth can be reduced by analyzing sections that
are mostly composed of shallow marine deposits,
thus reducing the absolute magnitude of water-
depth uncertainties. In addition, any uncertain-
ties due to water-depth assignments are reduced
by working with relatively thick stratigraphic
successions and relatively shallow water depths.
For our compilation, we have chosen to use a
simplifi ed water-depth scale (Angevine et al.,
1990), assigned based on fossil and/or lithofa-
cies evidence: nonmarine is here considered to
be 50 m above sea level, inner shelf is 50 m, outer
shelf is 150 m, and upper slope is 350 m. Non-
marine water depths are typically from coastal
plain deposits associated with shorelines. For
the most part, errors in water-depth assignment
are relatively small compared to the magnitudes
of the curves. However, signifi cant changes in
water depth occurring within the large uncer-
tainties inherent in bathyal-abyssal water depths
may leave important tectonic signals undetected
(Dickinson et al., 1987) and so are not included
here. Age control is another potential source of
inaccuracy. The number and resolution of age
assignments vary widely for different sedimen-
tary successions. We have focused on sections
that have at least four to fi ve dated stratigraphic
horizons identifi ed as points for subsidence
analysis. For this compilation we accept the
time scales used by the original authors. The
possibility of signifi cant hiatuses can also
lead to error. In order to reduce the impact of
this issue, we limit our study to those sections
where major unconformities, where identi-
fi ed by the original authors, are few. We com-
pensate for changes in stratigraphic thickness
through time due to compaction following the
standard approach outlined by van Hinte (1978)
and Allen and Allen (2005). For simplicity, we
used the exponential porosity versus depth rela-
tionships from Sclater and Christie (1980), and
generalized lithologies to sandstone, shale, or
limestone. Subsidence is calculated with respect
to sea level. Since no universally accepted quan-
tifi ed curve of sea level exists, we have simply
chosen to ignore sea-level changes and assume
that they cause only relatively low-magnitude
variations in our calculated subsidence. Even
the magnitude of long-term sea-level change is
poorly constrained, but is likely small (<200 m)
relative to the magnitude of subsidence seen in
these curves (Haq et al., 1987; Harrison, 1990).
As such, we assume that the sea-level datum
has not changed over time. Our approach to
backstripping utilizes the one-dimensional local
isostatic method of Steckler and Watts (1978).
However, this approach does imply that signifi -
cant variations in sediment loading near the ana-
lyzed site in the basin do not cause subsidence.
For broad basins, where basin thicknesses do
not vary greatly over short distances, this is a
reasonable assumption (Angevine et al., 1990).
This becomes more important for relatively
small basins and/or those formed over relatively
rigid lithosphere. This error is reduced by the
fact that we are looking at comparing the overall
shape of resultant curves in similar tectonic set-
tings and in basins of similar size.
All subsidence curves presented here are tec-
tonic subsidence curves. Subsidence histories
were only collected from sites where the plate
tectonic setting, as reported in the literature, is
well known (Fig. 1). It is possible that future
work will lead to a reinterpretation of tectonic
setting for some of these curves. We document
both the location and the original author used
in our compilation so that these redefi ned cases
can be clearly identifi ed. Subsidence curves are
grouped together by tectonic setting and plot-
ted at the same scale to facilitate comparisons
(Figs. 2–7). For clarity, several simplifying
approaches are used to improve our compari-
son of curves. First, no uncertainties in age and
paleobathymetry are shown on the graphs. Sec-
ond, major unconformities, if present, are shown
by fl at horizontal lines on the subsidence curves,
and small unconformities are not shown.
RESULTS
Our compilation is limited to sources that pro-
vide quantitative data that include stratigraphic
thickness, age assignments, proxy water-depth
data, and where the tectonic setting is well
established. In addition, an individual basin
may undergo different phases of subsi dence as
tectonic settings change through time. We have
tried to isolate those phases, or megasequences
(Allen and Allen, 2005), during which the spe-
cifi c tectonic setting of interest occurs. Fig-
ure 1 locates the sites and tectonic setting for
Intercontinental basins
Foreland basins
Forearc basins
7-7
7-1
7-2
2-4
2-5
2-7
6-1
2-1
2-6
2-8
4-2
3-1
3-2 3-3
3-4 3-5
3-6
3-9
4-1
4-3 4-4
4
6-2
6-4
7-3
7-5
7-6
7-9
7-8
7-4
3-8
3-8
4-5
4-7
4-8
2-3
3-7
6-5
4-6
4
6
62
6-7
6-6
66
6-3
Intercontinental basins
Foreland basins
Forearc basins
2-2
3-1
6
6
6
6
6
6
6
6
6
6
3-8
6-8
Passive margins
Strike-slip basins
6
6
6
6
2-
2
1
Figure 1. Locations of subsidence data by tectonic settings. Numbers refer to specifi c subsidence curves. First value is fi gure number, and
second value is curve number in that fi gure (e.g., 2–3 = Figure 2, curve 3).
Basin subsidence
Geological Society of America Bulletin, January/February 2009 57
data compiled in this study. Figures 2–7 show
the results of subsidence analysis as a function
of plate tectonic settings. On each set of subsi-
dence curves we include, for comparison, a ref-
erence curve that parallels best-fi t thermal subsi-
dence of the seafl oor assuming a semi-infi nite
half-space model from Stein and Stein (1992).
All curves are corrected for compaction and
backstripped assuming local isostasy.
Passive Margins (Fig. 2)
Subsidence following continental rifting and
breakup leads to asymmetric subsidence and
foundering of continental margins (Steckler and
Watts, 1978). As a result, the amount of subsi-
dence increases seaward of the hinge zone. All
subsidence curves show an initial phase of rapid
subsidence followed by a phase in which subsi-
dence rates are reduced (e.g., Watts and Ryan,
1976; Steckler and Watts, 1978) and mimic the
age-depth curve of the seafl oor. Some margins
demonstrate an abrupt change in subsidence
rates between these phases. However, this
abruptness may refl ect a poorly constrained his-
tory of early subsidence in some cases. Initial
subsidence deposits are often coarse nonmarine
deposits that are notoriously diffi cult to bio-
stratigraphically date. In addition, in modern
settings, the initial subsidence deposits are the
deepest and, thus, less frequently penetrated
parts of the sections. As a result there tend to
be few age constraints to delimit the early sub-
sidence history and the transition from rapid to
slower subsidence. Nonetheless, some well-con-
strained curves, such as the one shown from the
Gulf of Lion (Steckler and Watts, 1980), indi-
cate that abrupt changes can be real. Subsidence
in passive-margin settings typically continues
for more than 150 m.y. Maximum subsidence
(Fig. 2) varies up to 4 km, in part depending
on distance seaward of the hinge zone (i.e., the
landward limit of extension).
Passive margin formation and subsidence
mechanisms have been much studied follow-
ing the breakthrough work of Watts and Ryan
(1976) and Steckler and Watts (1978). Rift
basins develop early during continental breakup
followed by passive margin subsidence once
breakup is complete. Not all rifts go to comple-
tion, and many “failed rifts” can be found (cf.
Allen and Allen, 2005, their Fig. 9.11). Theoreti-
cal and analytical studies suggest that tectonic
subsidence can be divided into an initial “synrift”
phase that primarily refl ects isostatic response
to extension and thinning of continental crust,
followed by a “post-rift” phase driven by ther-
mal reequilibration as the lithosphere cools and
thickens back to equilibrium. Synrift stretching
and thinning by factors of less than 2 are com-
mon in rift basins (e.g., Hendrie et al., 1994;
Kusznir et al., 1996a, 1996b; Roberts et al., 1995;
Swift et al., 1987) and variable along individual
passive margins. In general, stretching factors
increase seaward to the point of continental rup-
ture and ocean crust formation. In addition, local
variability in subsidence can refl ect local struc-
ture and thinning as well as superimposed effects
(King and Ellis, 1990; Nadin and Kusznir, 1995).
Various mechanical models have been proposed
to explain details of subsidence curves in this
setting. Such models consider how extensional
strain is partitioned through the lithosphere (e.g.,
pure shear versus simple shear and depth-depen-
dent stretching), character (e.g., symmetric ver-
sus asymmetric and volcanic versus nonvolcanic
6b
6a
8b
3
2
1
300 200 100
Ma
0
0
2
4
kilometers
300 200 100
Ma
0
0
2
4
PASSIVE MARGINS
4 5
600 500 400
Ma
300
Seafloor subsidence
Seafloor subsidence
7
kilometers
8a
Figure 2. Tectonic subsidence curves for passive margin settings.
Locations shown on Figure 1. Solid curves correspond to time scale
at top of graph and dotted lines to time scale at bottom of graph.
Thermal decay curve (dashed) for subsi dence of cooling seafl oor
(Stein and Stein, 1992), minus (i.e., shallowed) 500 m, is shown
for comparison. 1—Paleozoic Miogeocline, southern Canadian
Rocky Mountains (Bond and Kominz, 1984); 2—Moroccan Basin
(Ellouz et al., 2003); 3—Campos Basin (Mohriak et al., 1987);
4—Gippsland Basin (Falvey and Mutter, 1981; P. Yin, 1985, per-
sonal commun.); 5—Gulf of Lion (Benedicto et al., 1996); 6—U.S.
Cordilleran Miogeocline (Bissell, 1974; Armin and Mayer, 1983;
Devlin et al., 1986; Devlin and Bond, 1988); 7—Lusitanian Basin
(Stapel et al., 1996); 8—U.S. Atlantic margin (Steckler and Watts,
1978; Swift et al., 1987).
Xie and Heller
58 Geological Society of America Bulletin, January/February 2009
margins), timing, and rate of heat loss during and
following continental breakup (e.g., Bott, 1980;
Jarvis and McKenzie, 1980; Turcotte, 1980;
Watts, 1981; Wernicke et al., 1982; Cochran,
1983; Nadon and Issler, 1997), as well as super-
imposed tectonic events (e.g., Dore and Stewart,
2002; Nielsen et al., 2002). Variations in these
factors may explain differences in magnitude of
curves shown.
Strike-Slip Basins (Fig. 3)
Basins related to strike-slip faults include a
variety of basin types that result from a com-
bination of transform fault movement that may
include either elements of crustal extension or
shortening (Christie-Blick and Biddle, 1985).
Strike-slip basins show a large variety in basin
size and geometry. However, they are typically
narrow and smaller than those produced by
regional extension or shortening. Different types
of basins can form in strike-slip settings, from
simple pull-apart basins developed along fault
oversteps to more complex basin forms in zones
of transtension and transpression (e.g., May et
al., 1993; Sutherland and Melhuish, 2000). Fault
geometry and related fault-mechanical process
are the critical controls for the development of
strike-slip basins as demonstrated by different
authors (e.g., Crowell, 1974; Mann et al., 1983;
Ingersoll, 1988). Regardless of specifi c basin
shapes and locations, all curves generated for
this setting are characterized by rapid and short-
lived (typically <10 m.y.) subsidence.
Subsidence in this setting is dictated by the
spatial confi guration of the various scales of
strike-slip fault systems within the basin, as well
as the history of displacement and attendant heat
loss (Sawyer et al., 1987; Chen and Nábelek,
1988). Depending on local patterns of defor-
mation, subsidence curves in strike-slip basins
may be episodic and end abruptly (e.g., Crow-
ell, 1974; Mann et al., 1982; May et al., 1993).
Tectonic subsidence typically exceeds 2 km
and reaches 4 km in exceptional cases (e.g., the
Los Angeles basin, Fig. 3, line 5). The magni-
tude and concave-up shape of these curves are
similar to those from passive margin settings
(Fig. 1), although the subsidence rates are much
faster. The fact that most strike-slip basins are
short lived probably refl ects the evolution and
frequent change in position of the master strike-
slip faults (Sylvester, 1988; Cloetingh et al.,
1996; Storti et al., 2003; Allen and Allen, 2005;
Waldron, 2005).
Stretching models have been applied to strike-
slip basins; however, compensation is made for
the small space and time scales associated with
these basins, such as fi nite rifting times, accen-
tuated lateral heat fl ow, and depth-dependent
extension (Cochran, 1983). Heat fl ow increases
due to lithospheric thinning. Theoretical and
fi eld studies suggest that heat is lost rapidly dur-
ing the extension process, in part, by lateral con-
duction (Cochran, 1983; Pitman and Andrews,
1985). The short-lived tails seen at the ends
of most curves may result from cooling of the
small remaining thermal anomaly once the fault
ceased to be active (Pitman and Andrews, 1985).
The result is that there is very little subsidence
that continues once extension stops (Nilsen and
McLaughlin, 1985). In some cases, the absence
of evidence for post-rift thermal subsidence may
also be a result of subsequent deformation of the
basin (Christie-Blick and Biddle, 1985).
Intracontinental Basins (Fig. 4)
Intracontinental, or intracratonic, basins are
large basins formed on an old continental lith-
osphere away from any known active tectonic
margin (Dickinson, 1976). These basins are
typically quite large (>150,000 km
2
in area),
have relatively slow, long-lived subsidence
(typically >200 m.y. in duration), but in most
compiled cases tectonic subsidence is less than
2 km. Cross-sectional geometries of the North
American examples are approximately symmet-
ric (Illinois, Michigan, and Williston Basins);
whereas, others are not.
Subsidence curves of intracontinental basins
in Figure 4 are approximately exponential in
shape, similar to passive margins, but most lack
a rapid initial subsidence phase. Overall, subsi-
dence curves follow the shape and magnitude of
seafl oor subsidence, but with longer decay con-
stants. Such a comparison has led some (Haxby
et al., 1976; Sleep and Sloss, 1980; Cercone,
1984; Nunn and Sleep, 1984; Nunn et al., 1984;
Howell and van der Pluijm, 1999; Kominz et al.,
2001) to suggest a thermal decay origin for at
least some of these basins.
A comparison of intracontinental subsi-
dence curves to simple thermal subsidence
models (Fig. 5) indicates broad consistency.
To model thermal subsidence, we use McKen-
zie’s (1978) simple-stretching model, but only
calculate post-rift subsidence resulting from
lithosphere reequilibration following thinning.
Stretching (thinning) factors in this case only
refl ect thinning of mantle lithosphere due to
thermal perturbation and not necessarily exten-
sion. Stretching factors ranging from 1.1 to 1.5
and equilibrium lithosphere thickness of 125
and 200 km are shown. Notice that thicker lith-
osphere has a longer decay constant to reach
thermal equilibrium. Thermal decay constants
increase as the square of lithosphere thickness.
As a result, it is not too surprising that thick
lithosphere, which tends to exist beneath the
oldest cores of continents (Chapman and Pol-
lack, 1977; Artemieva and Mooney, 2001), has
the longest subsidence histories.
Most models of thermal reequilibration are
similar in approach to that used by Haxby et al.
(1976) for the Michigan Basin. In this model a
large-scale, but undocumented, thermal event
leads to formation of a dense crustal mass that
causes subsequent subsidence as the lithosphere
cools. Flexure broadens the width of basin
defl ection (Nunn and Sleep, 1984).
Notable in the subsidence curves are the
deviations from idealized thermal subsidence
(Fig. 5). These deviations are more pronounced
than those seen in passive margins and suggest
that tectonic reactivation characterizes many
intracontinental basins. The most extreme of
these is the Ordos Basin (Fig. 4, line 6). How-
ever, recent work suggests that this basin may be
the result of constructive interaction of deforma-
tion events around the basin margin and is not
primarily driven by thermal effects (Xie, 2007).
Most of the other examples also show strong
deviations away from simple thermal equilibra-
tion, more than can be reasonably accounted for
by eustatic sea-level changes. Various authors
have suggested interacting tectonic mechanisms
impacting these basins including intraplate
stresses, multiple thermal perturbations, reac-
tivation of inherited structures, far-fi eld effects
of nearby tectonic events, or changes in litho-
sphere rheology (Nunn and Sleep, 1984; Klein
and Hsui, 1987; Bond, 1991; Kaminski and Jau-
part, 2000).
Foreland Basins (Fig. 6)
Foreland basins are asymmetric basins adja-
cent, and parallel, to an attendant contractional
orogenic belt. Foreland basins, or foredeeps, sit
atop a defl ected continental lithosphere of the
underlying plate in both continental collision
zones (peripheral foreland basins of Dickinson
[1976]) and behind volcanic arcs (retroarc fore-
land basin of Dickinson [1976]). Many studies
have demonstrated that, for the most part, these
basins form as a regional isostatic (fl exural)
response to loading by the adjacent orogenic belt
(e.g., Beaumont, 1981; Jordan, 1982; DeCelles
and Giles, 1996).
Foreland basin subsidence curves differ from
thermal subsidence curves seen in most other
basins in that the former are characterized by
their convex-up shape and frequent episodic sub-
sidence events. The convex-up profi le refl ects
accelerating subsidence as the tectonic load
migrates toward the foreland coupled with the
curved fl exural profi le of the basin. As the basin
widens due to migration of the thrust load and
associated sedimentation, the distal parts of the
Basin subsidence
Geological Society of America Bulletin, January/February 2009 59
300 200 100
Ma
0
0
2
4
kilometers
INTRACONTINENTAL BASINS
1
2
3
4
5
6
7
8
500 400
S
e
a
f
l
o
o
r
s
u
b
s
i
d
e
n
c
e
Figure 4. Tectonic subsidence of intracontinental basins. Locations shown in Figure 1. See thermal decay curve
(dashed) for subsidence of cooling seafl oor (Stein and Stein, 1992), minus 1500 m, is shown for comparison. 1—Illi-
nois Basin, Farley well (Bond and Kominz, 1984); 2—Michigan Basin (Bond and Kominz, 1984); 3—Williston
Basin, North Dakota (Bond and Kominz, 1984); 4—Williston Basin, Saskatchewan (Fowler and Nisbet, 1985);
5—Northeast German Basin (Scheck and Bayer, 1999); 6—Southwest Ordos Basin (Xie, 2007); 7—Paris Basin
(Prijac et al., 2000); 8—Parana Basin (Zalan et al., 1990).
STRIKE-SLIP BASINS
100
Ma
0
0
2
4
kilometers
S
e
a
f
l
o
o
r
s
u
b
s
i
d
e
n
c
e
S
e
a
f
l
o
o
r
s
u
b
s
i
d
e
n
c
e
45a
5b
6
Ma
100 0
0
2
4
kilometers
S
e
a
f
l
o
o
r
s
u
b
s
i
d
e
n
c
e
12
3
Ma
100 0
0
2
4
kilometers
7
8
9
Figure 3. Tectonic subsidence curves for strike-slip basins. Locations shown in Figure 1. Thermal decay curve
(dashed) for subsidence of cooling seafl oor (Stein and Stein, 1992), minus 500 m, is shown for comparison. 1—Chuck-
anut Basin (Johnson, 1984, 1985); 2—Ridge Basin (Crowell and Link, 1982; Karner and Dewey, 1986); 3—Death
Valley (Hunt and Mabey, 1966); 4—Salinian block (Graham, 1976); 5—Los Angeles Basin (Rumelhart and Ingersoll,
1997); 6—Gulf of California (Curray and Moore, 1984); 7—Cuyama Basin (Dickinson et al., 1987); 8—Bozhang
Depression (Hu et al., 2001); 9—Salton Trough (Kerr et al., 1979).
Xie and Heller
60 Geological Society of America Bulletin, January/February 2009
basin show time-transgressive subsidence. That
is, while the proximal foreland basin responds
immediately to adjacent thrust loads, the distal
parts of the basin may show later subsidence as
loads migrate basinward over time (Jones et al.,
2004). The result is a time lag as the tectonic
and sediment load propagates across a foreland
basin. In addition, the redistribution of sediment
and subsidence over time in this way leads to
fl attening of the basin geometry and reduction
in size of the attendant forebulge. Jones et al.
(2004) suggest time lags of tectonic signals
on the order of a few million years or less out
across foreland basins. The total duration of oro-
gens, as seen in subsidence histories (Fig. 6), is
typically a few tens of millions of years.
Smaller scale episodes of subsidence super-
imposed on the overall subsidence profi le pri-
marily refl ect intermittent thrust events (e.g.,
Heller et al., 1986), although not every event
signifi cantly changes the confi guration of the
thrust load. Duration and episodicity of subsi-
dence varies from basin to basin, as set by the
pace of growth of the adjacent orogen, and in
different parts of a single basin (e.g., Fig. 6,
lines 8a and 8b), as a function of local loading
history. The maximum magnitude of tectonic
subsidence seen in compiled curves is ~3 km.
Blind thrusts often propagate into the proxi-
mal foreland basin and sedimentation can con-
tinue above these structures. DeCelles and Giles
(1996) refer to this part of the foreland basin
as the “wedge top.” While these basins are not
included here, it is clear that thrust emplacement
will impact subsidence history in these parts of
the proximal foreland (e.g., Vergés et al., 1998).
Other smaller basins may form in concert with
deformation of the adjacent orogen. These
include piggyback and back-bulge basins (Ori
and Friend, 1984; DeCelles and Giles, 1996).
Subsidence history of piggyback basins tends
to be shorter lived and of less magnitude than
their associated foreland basins (e.g., Burbank
et al., 1992; Carrapa et al., 2003), and is not
considered here. Back-bulge basins, if present,
are very subdued features that lie outboard of
foreland basins, beyond the forebulge, and form
as a dampened fl exural response to loading
of an elastic plate. Magnitude of defl ection of
back-bulge basins is very small, typically a few
percent of the depth of the associated foreland
basin, and may be diffi cult to uniquely identify.
Forearc Basins (Fig. 7)
Forearc basins lie between trenches and their
associated, parallel, magmatic arcs (Dickinson,
1995). The sizes and confi gurations of both
modern and ancient forearc basins are highly
variable, but it is clear that typical forearc basins
are narrow and elongate, with thick sediment
packages confi ned to deep structural troughs. We
note that there is often a large range of paleoba-
thymetry in these settings, so that resultant sub-
sidence curves may be less well constrained.
Subsidence curves from forearc basins, as a
group, have the most diverse range of shapes
(Fig. 7). Some show very rapid, short-lived sub-
sidence similar to strike-slip basins. Others have
slower, relatively linear subsidence. Still others
show an abrupt transition from rapid subsidence
to very slow subsidence rates, similar to some
200 300 400
m.y.
0
2
kilometers
1
2
3
4
5
6
8
7
0 100
200 300
m.y.
0
2
kilometers
0 100
β = 1.1
β = 1.2
β = 1.4
β = 1.1
β = 1.2
β = 1.4
β = 1.3
β = 1.3
β = 1.5
β = 1.1
β = 1.1
β = 1.2
β = 1.3
β = 1.3
β = 1.4
β = 1.4
β = 1.5
Figure 5. Comparison of intracontinental basin subsidence curves (numbered heavy lines from Fig. 4) with post-
rift thermal subsidence curves calculated from the McKenzie (1978) stretching model. Time (m.y.) is shown since
basin formation. Thin solid lines assume lithosphere thickness of 125 km; dashed lines assume lithosphere thick-
ness of 200 km. Stretching factors (β) from 1.1 to 1.5 are shown.
Basin subsidence
Geological Society of America Bulletin, January/February 2009 61
curves from passive margins. In addition, some
forearc basins show large uplift events, such as
in the Indonesian forearc basin (Beaudry and
Moore, 1985). Other basins, such as the Chilean
forearc, show signifi cant amounts of rotation
and widening of basin fi lls over time (Coul-
bourn and Moberly, 1977). Most of the curves
show less than 2 km of tectonic subsidence. The
modern Tonga forearc is exceptionally deep
(Fig. 7, line 4).
The range of shapes of subsidence curves in
this setting indicates that a variety of factors may
contribute to forearc basin subsidence. Most
curves are relatively simple in form and imply
a monotonic driving mechanism. The curves
from the Great Valley of California (Fig. 7, line
1), exhibit an abrupt change in subsidence rate
possibly refl ective of a change in driving mecha-
nism. The Great Valley curves also show very
different timings of infl ection points indicating
that the basin is tectonically segmented into dif-
ferentially subsiding zones. Basin segmentation
is seen elsewhere (Izart et al., 1994) and may
be common in these settings. Episodic subsi-
dence and even uplift of some basins is seen in
Figure 7, although it is not clear to what extent
these may refl ect errors in bathymetric assign-
ments. Causes of segmentation include parti-
tioned strain associated with oblique subduc-
tion (Izart et al., 1994), bathymetric changes in
the underlying subducted slab that isostatically
impact the overlying plate (Kobayashi, 1995),
and collision of crustal fragments in the subduc-
tion zone (Clift and MacLeod, 1999).
Possible subsidence mechanisms in forearc
basins include growth, loading, and under-
plating of the accretionary prism, which may
drive tectonic rotation and basin widening in
some settings (Coulbourn and Moberly, 1977).
Basin growth has also been tied to an increase
in width of the arc-trench gap due to fl attening
of the underthrusted plate and resultant migra-
tion of the accretionary wedge and volcanic arc
(Dickinson, 1995). Regional isostatic effects of
changing lithospheric thickness and density due
to age and structure of the underthrusted plate
can account for segmentation, and even uplift,
of forearc basin subsidence (Moxon and Gra-
ham, 1987). Of course, compression associated
with the coupling of the upper and lower plates
across convergent margins suggests that fold-
ing and thrust loading may contribute to subsi-
dence (Fuller et al., 2006). However, extensional
faulting may contribute to subsidence in some
forearc settings (Izart et al., 1994; Unruh et
al., 2007). Thermal subsidence associated with
either cooling of the fl ank of the adjacent arc
massif (Moxon and Graham, 1987) or cooling
of an accreted warm microplate (Angevine et al.,
1990) are possible mechanisms. In fact, thermal
subsidence in forearc settings can be accelerated
due to refrigeration by the underthrusted plate
(Mikhailov et al., 2007). Clift and MacLeod
(1999) discuss the role of tectonic erosion by
the down-going slab as a cause of subsidence
and tilting of the forearc basin. Subsidence of
forearc basins is the least understood and most
poorly constrained of the tectonic settings
explored in this study.
SUMMARY
Tectonic setting exerts primary control on sed-
imentary basin subsidence history. Several basin
settings seem to have distinctive subsidence pat-
terns suggesting a limited range of driving mech-
anisms. As such, calculated subsidence history is
a potential tool for identifying tectonic setting of
ancient basins of unknown origin. Passive mar-
gins show rapid initial synrift subsidence fol-
lowed by prolonged thermal subsidence similar
to that seen for subsiding seafl oor. Strike-slip
basins all demonstrate rapid, albeit short-lived,
subsidence. Foreland basins are characterized by
segmented convex-up subsidence. Intracontinen-
tal basins studied here show long-lived gradual
subsidence. While overall the subsidence pat-
tern of intracontinental basins is consistent with
thermal subsidence of thick lithosphere, most
profi les contain large deviations from predicted
Figure 6. Tectonic subsidence of foreland basins. Locations
shown in Figure 1. Thermal decay curve (dashed) for subsi-
dence of cooling seafl oor (Stein and Stein, 1992), minus 1500 m,
is shown for comparison. 1—Eastern Avalonia, Anglo-Brabant
fold belts (van Grootel et al., 1997); 2—Southern Alberta Basin
(Gillespie and Heller, 1995); 3—San Rafael Swell, Utah (Heller
et al., 1986); 4—Pyrenean foreland basin, Gombrèn (Vergés et
al., 1998); 5—Swiss Molasse basin (Burkhard and Sommaruga,
1998) modifi ed from total subsidence using water:sediment
density contrast); 6—Hoback Basin, Wyoming (Cross, 1986);
7—Green River Basin, Wyoming (Cross, 1986; Heller et al.,
1986); 8—Magallanes Basin (Biddle et al., 1986).
300 200 100 0
0
2
4
kilometers
300 200 100
Ma
Ma
0
0
2
kilometers
FORELAND BASINS
1
3
2
4
5
6
7
8a
8b
4
S
e
a
f
l
o
o
r
s
u
b
s
i
dence
S
e
a
f
l
o
o
r
s
u
b
s
i
dence
Xie and Heller
62 Geological Society of America Bulletin, January/February 2009
thermal curves suggestive of tectonic reactiva-
tion. Forearc basins studied have various subsi-
dence profi les, suggesting there may be a range
of driving mechanisms of subsidence.
Although subsidence analysis can be a use-
ful tool in identifying tectonic setting in ancient
sequences, second-order variations in the sub-
sidence rate provides specifi c information on
important local details of driving forces and
tectonic timing. Finally, caution should be used
given the limitation of data sets available for this
compilation and because basins can span mul-
tiple tectonic settings over time and/or space.
This approach is best used in conjunction with
other structural and basin analysis techniques.
ACKNOWLEDGMENTS
We would like to thank Philip Allen, Barbara
Carrapa, Ken Dueker, William Dickinson, Steve
Graham, Andrew Hynes, Michelle Kominz, Ranie
Lynds, and Osamu Takano for reviews, discussions,
and useful references.
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