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Deepwater fold and thrust belt classification, tectonics, structure and hydrocarbon
prospectivity: A review
C.K. Morley
a,
⁎
, R. King
b
, R. Hillis
c
, M. Tingay
b
, G. Backe
b
a
PTTEP, 27th Floor, ENCO Building, Soi 11, Vibhavadi-Rangsit Road, Chatuchak, Bangkok, Thailand, 10900
b
Centre for Tectonics, Resources and Exploration (TRaX), Australian School of Petroleum, University of Adelaide, SA 5005, Australia
c
Deep Exploration Technologies Cooperative Research Centre, c/o University of Adelaide, SA 5005, Australia
abstractarticle info
Article history:
Received 3 April 2010
Accepted 28 September 2010
Available online 4 November 2010
Keywords:
deepwater fold and thrust belts
hydrocarbons
thrust
fold
accretionary prisms
deltas
salt tectonics
shale tectonics
growth faults
continent collision
gravity-driven deformation
near-field stress
far-field stress
Niger Delta
Borneo
Gulf of Mexico
Caspian Sea
Timor
Trinidad
Barbados
Deepwater fold and thrust belts (DWFTBs) are classified into near-field stress-driven Type 1 systems confined
to the sedimentary section, and Type 2 systems deformed by either far-field stresses alone, or mixed near- and
far-field stresses. DWFTBs can occur at all stages of the Wilson cycle up to early stage continent continent
collision. Type 1 systems have either weak shale or salt detachments, they occur predominantly on passive
margins but can also be found in convergent-related areas such as the Mediterranean and N. Borneo.
Examples include the Niger and Nile deltas, the west coast of Africa, and the Gulf of Mexico. Type 2 systems are
subdivided on a tectonic setting basis into continent convergence zones and active margin DWFTBs. Continent
convergence zones cover DWFTBs developed during continent–arc or continent–continent collision, and
those in a deepwater intracontinental setting (e.g. W. Sulawesi, Makassar Straits). Active margins include
accretionary prisms and transform margins. The greatest variability in DWFTB structural style occurs between
salt and shale detachments, and not between tectonic settings. Changes in fold amplitude and wavelength
appear to be more related to thickness of the sedimentary section than to DWFTB type. In comparison with
shale, salt detachment DWFTBS display a lower critical wedge taper, more detachment folds, long and
episodic duration of deformation and more variation in vergence. Structures unique to salt include canopies
and nappes. Accretionary prisms also standout from other DWFTBs due to their relatively long, continuous
duration, rapid offshore propagation of the thrust front, and large amount of shortening. In terms of
petroleum systems, many similar issues affect all DWFTBs, these include: the oceanward decrease in heat
flow, offshore increase in age of mature source rock, and causes of trap failure (e.g. leaky oblique and frontal
thrust faults, breach of top seal by fluid pipes). One major difference between Type 1 and Type 2 systems is
reservoir rock. High quality, continent-derived, quartz-rich sandstones are generally prevalent in Type 1
systems. More diagenetically reactive minerals derived from igneous and ophiolitic sources are commonly
present in Type 2 systems, or many are simply poor in well-developed turbidite sandstone units. However,
some Type 2 systems, particularly those adjacent to active orogenic belts are partially sourced by high quality
continent-derived sandstones (e.g. NW Borneo, S. Caspian Sea, Columbus Basin). In some cases very high rates
of deposition in accretionary prisms adjacent to orogenic belts, coupled with uplift due to collision, results in
accretionary prism related fold belts that pass laterally from sub-aerial to deepwater conditions (e.g. S.
Caspian Sea, Indo-Burma Ranges). The six major hydrocarbon producing regions of DWFTBs worldwide (Gulf
of Mexico, Niger Delta, NW Borneo, Brazil, West Africa, S. Caspian Sea) stand out as differing from most other
DWFTBs in certain fundamental ways, particularly the very large volume of sediment deposited in the basins,
and/or the great thickness and extent of salt or overpressured shale sdetachments.
© 2010 Elsevier B.V. All rights reserved.
Contents
1. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42
2. Outline of the DWFTB classification . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42
2.1. Introduction. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42
2.2. Stress-system terminology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43
Earth-Science Reviews 104 (2011) 41–91
⁎ Corresponding author. Tel.: +6625374264.
E-mail address: chrissmorley@gmail.com (C.K. Morley).
0012-8252/$ – see front matter © 2010 Elsevier B.V. All rights reserved.
doi:10.1016/j.earscirev.2010.09.010
Contents lists available at ScienceDirect
Earth-Science Reviews
j o ur n a l h o m e pa g e: w w w. el s evi e r. c o m/ l oc at e /e a rs c i r e v
Author's personal copy
2.3. Are near-field stress systems confined to passive margins? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44
2.4. Tectonic setting: active margins and convergent zones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44
2.5. Classification scheme. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 45
3. The characteristics of salt and shale detachments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46
3.1. Gravity sliding and gravity spreading . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46
3.2. Salt detachment zones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46
3.3. Shale detachments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 49
3.4. Do thick mobile shale zones exist? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51
4. Structural and petroleum system characteristics of the different types of DWFTB . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52
4.1. Near-field stress-driven linked systems . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52
4.1.1. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52
4.1.2. Type 1a, DWFTBs associated with regional offshore-dipping detachments along overpressured shales . . . . . . . . . . . . . 54
4.1.3. Type 1a, DWFTBs associated with a widespread hinterland-dipping basal detachment zone (large deltas) . . . . . . . . . . . 56
4.1.4. Type 1b, DWFTBs associated with an oceanward-dipping salt detachment zone . . . . . . . . . . . . . . . . . . . . . . . . 58
4.2. Type 2, continental convergent tectonic setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62
4.2.1. Type 2a convergent zones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 64
4.2.2. Type 2bi weakly linked/unlinked DWFTBs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67
4.3. Type 2bii active margins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 71
4.3.1. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 71
4.3.2. Barbados . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 72
4.3.3. The Makran . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73
4.3.4. Andaman subduction zone–Indo-Burma Ranges . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 75
4.3.5. South Caspian Sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 76
5. Synthesis of DWFTB characteristics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78
5.1. Petroleum systems . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78
5.1.1. Source rock types . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78
5.1.2. Temperatures and hydrocarbon maturation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 79
5.1.3. Hydrocarbon migration . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 79
5.1.4. Reservoir quality . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80
5.1.5. Structural traps . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 81
5.1.6. Seal . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83
5.2. Structural development in different tectonic settings . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83
6. Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 85
Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 86
References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 86
1. Introduction
There has been long-standing academic interest in deepwater fold and
thrust belts (DWFTBs), particularly along subduction zones and early-
stage collisional margins. Over the last two decades interest has surged
following advances in deepwater drilling technology by the oil industry.
Deepwater exploration encompasses many different potential traps and
geological settings including DWFTBs, which have featured prominently
because they contain numerous large anticlines with associated hydro-
carbon seeps. At present, six areas are the main focus of deepwater
development and production in DWFTBs, these are: the Gulf of Mexico,
Niger Delta, NW Borneo, the Brazilian Margin, West Africa (Angola,
Congo) and the South Caspian Sea (mostly shallow water, but extending
into deepwater) (Fig. 1). Exploration is being conducted in many more.
The academic and industry activity mentioned above has produced
a wealth of information about DWFTBs that is not only important for
high-grade areas in which very expensive exploration and develop-
ment programs are being conducted, but also for fundamentally
modifying and improving our understanding of how passive margins,
and the early stages of orogenic belts develop. This paper is the first
detailed review of the considerable body of published data on modern
DWFTBs (not ancient, exhumed DWFTBs); it reviews the significant
characteristics of key deepwater fold and thrust belts, and their
petroleum systems. The DWFTBs are reviewed as a group because
despite the very different tectonic processes that can operate in the
hinterlands of the DWFTBs, there are many similar characteristics to
the active DWFTB in all the different tectonic environments. The
examples of DWFTBs in different settings are described in detail in
Section 4 because they either represent well-documented examples,
or they illustrate some of the significant variations found within the
theme of DWFTBs. The broad division of DWFTBs into passive and
active margins, shale vs. salt detachments, and gravity sliding vs.
gravity spreading mechanisms has been proposed previously (e.g.
Morley and Guerin, 1996; Rowan et al., 2004; Krueger and Gilbert,
2009). This paper discusses refinement of this classification, because a
simple grouping into gravity-driven passive margins and lithospheric-
stress driven, tectonically active margins (e.g. Rowan et al., 2004;
Hamilton and De Vera, 2009) does not encompass all settings of
DWFTBs. Instead, a classification system for DWFTBs is proposed
based on the driving mechanism, detachment type and tectonic
setting (Fig. 2).
The justification for the classification scheme used in this paper is
addressed in Section 2. Then the nature of the detachment zones is
discussed in depth, particularly for shale detachments, since their
variations exert a strong control on DEFTB structural style. Currently,
the nature of mobile shale detachments is a subject of controversy,
with some workers doubting whether thick mobile shale masses can
exist at depth in the subsurface. The fourth and largest part of the
paper (Section 4) then describes the tectonic, structural geology and
petroleum system characteristics of the different types of DWFTBs in
the context of the proposed classification scheme. The differences and
similarities in petroleum systems within the different types of DWFTB
are discussed in Section 5.
2. Outline of the DWFTB classification
2.1. Introduction
Previous classification schemes have noted that DWFTBs can either
develop under the influence of gravity where deformation is limited to
42 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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thesedimentary section above a basal detachment (e.g. passive margins),
or in response to stresses that affect much of the crust or the entire
lithosphere, such as those found in accretionary prisms and collision belts
(e.g. Rowan et al., 2004; Hamilton and De Vera, 2009). Consequently
DWFTBs have been subdivided into active and passive margin settings
(Hamilton and De Vera, 2009; Krueger and Gilbert, 2009). Further
subdivisions are made on the basis of whether the basal detachments to
the DWFTBs are developed within salt or shale lithologies.
There are three main problems with the existing classification
schemes: 1) inaccurate terminology, 2) failure to encompass all types
of DWFTBs within the classification, and 3) gravity and lithospheric
stresses driving mechanisms are not exclusively confined to passive
and active margins respectively. These problems, the terminology
used in the classification, and the nature of the basal detachment
zones are discussed in this section in order to justify the new DWFTB
classification presented herein and in Fig. 2.
2.2. Stress-system terminology
Within the classification scheme proposed herein, the separation
into gravity-driven and lithospheric stress driven systems seems a
reasonable, simple subdivision. However, gravity plays an important
role in driving lithospheric plates and can also drive flow of hot,
ductile crust in some regions (e.g. Tibetan Plateau, see review in
Ghosh et al., 2006; SE Asia, Hall, in press). Consequently, use of the
term gravity-driven for just the systems confined to the sedimentary
Fig. 1. Digital elevation map (DEM) of the world, showing locations of deepwater fold and thrust belts, and their classification (see Fig. 2). Note: only some examples of accretionary
prisms are shown. 1 =McKenzie Delta, 2 = Cascadia accretionary prism, 3 = Perdio DWFTB, 4=Mississippi Fan DWFTB, 5=Mexican Ridges DWFTB, 3, 4 and 5 = Gulf of Mexico,
6=Scotia Basin, 7 =Barbados accretionary prism, 8 = Columbus Basin, 9 = Amazon Delta, 10=Para-Maranhao, 11=Sergipe-Alagoas, 12=Espirito Santo Basin, 13 =Campos Basin,
14=Santo Basin, 15 = Pelotas Basin, 16=Straits Gibraltar (Prerif'nappe'), 17 =Essaouira Basin, 18=Agadir and Tarfaya Basins, 19 =Aaiun Basin, 20=MSGBC (Mauritania-
Senegal-Gambia-Bissau-Conarky) Basin, 21 =Greater Niger Delta area (3 squares), 22 = Astrid Thrust Belt, 23 = Lower Congo Basin, 24 = Kwanza Basin, 25=Namibe Basin,
26=Orange Basin, 27 = Majunga Basin, 28 = Rovuma Basin, 29 Kenya, 30 = Somalia, 31 = deep shale detachment Nile Delta, 32 = Messinian salt DWFTB deepwater Niger Delta-
Levant Basin, 33=Cyprus arc accretionary fold belt, 34 = Adana-Cilicia and Iskenderun-Latakia basins, 35 = Black Sea, 36 = South Caspian Sea, 37 = Makran accretionary prism,
38=Krishna-Godavari Basin, 39 = Indo-Burma Ranges, 40=West Luconia Delta, 41 =NW Borneo (Brunei, Sabah), 42 = Sandakan Basin, 43 = Mahakam Delta, Makassar Straits,
44=East Sulawesi-Makassar Straits, 45=Banggai-Sula, 46=Molucca Sea Collision Complex, 47 = Cenderawasih Basin, 48 = Seram, 49 = Timor, 50 = Sumatra-Java accretionary
prism, 51 = Bight Basin, 52 = Hikurang accretionary prism. DEM from http://www.ngdc.noaa.gov/mgg/image/2minrelief.html.
Predominantly/
exclusively
near field stress
Mixed near field
and far field stress
Predominantly/
exclusively far field
stress
Detachment type
Detachment dip
Tectonic setting
Stress type
Potentially any type of
setting with a slope
to deepwater. In
practice predominantly
passive margins
Continental
convergence
zones
Continental
convergence
zones
Accretionary
prisms
Shale
Salt
Type 1 Type 2
Type 2a Type 2b
Type 1a Type 1b
Type 2bi Type 2bii
Shale Shale Shale
Landward = L, Oceanward = O
LLLLO
O
Fig. 2. Classification scheme for modern deepwater fold and thrust belts used in this review. For ancient type 2a and 2bi DWFTBs the basal detachment can be located in shale or in
salt, however no modern examples of salt detached type 2a and 2bi DWFTBs are known.
43C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
Author's personal copy
section could be accused of inaccuracy. Instead, where the potential
energy for driving deformation arises from uplift or sediment loading
and results in gravity-driven deformation confined to deformation
above one or more detachments that reside entirely within the
sedimentary section, we use the term near-field stress systems. Far-
field stress is used for thin-skinned DWFTBs driven by lithospheric
stresses, and/or where gravity drives deformation in parts of the
middle or lower crust that are sufficiently hot and weak.
2.3. Are near-field stress systems confined to passive margins?
The prerequisites for near-field stress deformation are a basin with a
consistent regional slope into a deepwater area, a weak detachment
zone, and a trigger of some kind (usually uplift or delta progradation).
This set of conditions is not exclusive to passive margins, although the
greatest number of near-field stress DWFTBs does occur on passive
margins (e.g. Tari et al., 2003; Rowan et al., 2004). For example the
collisional setting of the Mediterranean produced the Messinian
evaporates (Cita, 1983), which also acted as detachment zones to
DWFTBs, e.g. NE Mediterranean Cyprus Arc (Adana-Cilicia and
Iskenderun-Latakia basins; Bridge et al., 2005), and deepwater Nile
Delta (Loncke et al. 2006). The Sandakan Delta (Fig. 3) developed when
uplift of Borneo triggered sedimentation and forced progradation of the
delta into the southern axis of the developing Celebes Sea spreading
centre (Fig. 3; Balaguru and Hall, 2009). The crust of northern Borneo is
composed of uplifted deepwater sediments, overthrust by ophiolites
thatprobably overlie oceanic or forearc crust (e.g. Hutchison et al. 2000).
Consequently, in terms of both the crustal type, and tectonic setting the
Sandakan Delta did not form on a typical passive margin.
One of the success stories of hydrocarbon exploration in DWFTBs is
the NW Borneo margin where up to 10 km thickness of deltaic
sediments accumulated during the Neogene (e.g. Sandal, 1996; Hall
and Nichols, 2002; Ingram et al., 2004; Morley and Back 2008).
Shortening in the DWFTB was driven both by near field stresses with a
linked growth fault system on the shelf, and by far-field stresses
associated with regional shortening (e.g. Morley et al., 2003; Hesse
et al., 2008; Morley et al., 2008; King et al., 2009). This mixture of
shallow gravity, and deep-seated stresses in an early-stage collision
zone illustrates there is not necessarily a simple distinction between
gravity-dri ven passive margin DWFTBs and t ectonically driven
DWFTBs on active margins (see Section 4.2.1.1. for details). Passive
margins are not synonymous with near-field stress driven deforma-
tion, and the same applies to active and collisional margins with
respect to far-field stresses. It should also be noted there is potential
for far field stresses to cause deformation (e.g. inversion structures)
on Atlantic-type passive margins (e.g. Cloetingh et al., 2008), and to
cause folding of oceanic crust (e.g. the Central Indian Ocean fold belt,
Beekman et al., 1996).
2.4. Tectonic setting: active margins and convergent zones
The term active margin covers either convergent or Pacific-type
margins associated with subduction of oceanic crust, or transform
margins, but not continent–continent, continent–arc collision, or
intracontinental convergence settings (e.g. Condie, 1997). Unfortu-
nately all these settings are lumped together within the term active
margin in the classifications of Hamilton and De Vera (2009) and
Krueger and Gilbert (2009). In this classification far-field stress driven
DWFTBs are subdivided on the basis of tectonic setting into active
margins (i.e. subduction zones and transform margins) and conti-
nental convergent zones.
Continental convergent zones encompass deep water fold and thrust
belts that formed during continent–arc collision or the early-stage of
continent–continent collision, and in regions where a complete Wilson
Cycle has failed to occur, e.g. thinned crust in an intraplate setting
marginal to an orogenic belt (intracontinental convergence zones).
SE Asia displays a wide range of DWFTB settings, including those
related to different stages of subduction and early-stage continent–
continent or continent–arc collision, and obduction (e.g. NW Borneo
and the Banda Arc, Fig. 3; Searle and Stevens, 1984; Hall and Wilson,
2000), as well as intracontinental convergent zones. The Makassar
Straits between Borneo and Sulawesi are an example of an
intra contin ental convergence zone, whe re a deepwater trough
containing a DWFTB on the west side of Sulawesi has developed
over thinned, rifted continental crust (Figs. 4 and 5; Hall et al., 2009).
Intracontinental convergence zones have no place in classification
Fig. 3. Distribution of deepwater fold and thrust belts in SE Asia, oblique DEM image.
44 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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schemes where the subdivision is into active and passive margins
(Rowan et al., 2004; Hamilton and De Vera, 2009; Krueger and Gilbert,
2009).
2.5. Classification scheme
On the basis of the tectonic and structural divisions discussed
above, DWFTBs are be divided into the following categories (Fig. 2):
1) Near-field stress-driven linked extensional-compressional sys-
tems (mostly, but not exclusively found on passive margins),
commonly related to gravitational instability of a seaward
prograding sedimentary wedge.
1a) Predominantly offshore dipping, basal over pressured
shale detachment (e.g. Orange Basin, South Africa; Bight
Basin, S. Australia; Niger Delta, Nigeria – mixed offshore
and landward-dipping detachment).
1b) Predominantly offshore-dipping salt detachment (e.g.
West African margins of Gabon, Congo, Angola; Gulf of
Mexico, Brazil Santos Basin).
2) DWFTBs partially or completely driven by far-field stresses, with
a basal shale detachment that typically dips landwards.
2ai) Linked extensional–compressional systems in zones of
inter- or intra-continent convergence (gravity sliding/
differential loading/far-field lithospheric stresses), e.g.
NW Borneo margin; Columbus Basin, Trinidad.
2aii) Weakly linked or unlinked extensional–compressional
structures in DWFTB systems on early stage zones of
inter- or intra-continent convergence, mostly or entirely
driven by far-field lithospheric stresses, e.g. East Makassar
Straits, Indonesia; Banda Arc, Indonesia.
2b) Weakly linked or unlinked extensional–compressional
structures in DWFTB systems on active margins (sub-
duction and transpressional margins), mostly or entirely
driven by lithospheric stresses, e.g. Barbados accretionary
prism; Makran, Iran and Pakistan, Indo-Burma Ranges,
Bangladesh and Myanmar, South Caspian Sea.
These c a te go ri es and associate d hydroca rbo n provinc e s are
discussed in the following sections.
Luconia
Platform
Rajang/West
Luconian Delta
NW Borneo Trough
Fold and Thrust belt
Sandakan
Delta
Tarakan
Delta
West
Sulawesi
Fold belt
Mahakam
Delta
M
a
kassa
r
S
tr
a
it
s
Borneo
110°
112°
8°
6°
4°
2°
0°
2°
F
200 m
200 m
1000 m
1000 m
100 km
Area containing deepwater
fold and thrust belt
Shale mini-basin province
offshore Sabah
Fold axes
114°
116°
118°
Normal faults
Thrust faults
Seafloor
bathymetry
Fig. 4. Regional map of Borneo showing the location of DWFTBs around the island.
45C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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3. The characteristics of salt and shale detachments
3.1. Gravity sliding and gravity spreading
Gravity gliding (or sliding) and gravity spreading are the two basic
types of near-field stress driven deformation of a sedimentary wedge
that progrades into deeper water. Gravity gliding occurs by the rigid
translation of a rock mass down a slope, while gravity spreading is the
flattening and lateral spreading of a rock mass under its own weight
(Ramberg, 1 981; DeJong and Scholten, 1973; Fig. 6). Gra vity
spreading rarely affects the complete sedimentary section and is
often limited to a thick mobile zone (usually either overpressured,
undercompacted muds or salt) at the base of the gravity-driven
system. In contrast to the original definition, gravity spreading of a
mobile zone in DWFTBs would occur in response loading by overlying
sediments, not just under it's own weight.
Gravity gliding is usually associated with linkage of up-dip
extension with a down-dip contractional toe region via a detachment
zone and covers a wide range of temporal and spatial scales. There are
the geologically instantaneous, shallow slides associated with mass
wasting that can range in size from small landslips covering a few
hundred square meters in area, to giant submarine slides that can
cover over 100 km
2
(e.g. Embly, 1976, Canals et al., 2004; Gee et al.,
2007). Such slides are often triggered by short-term events e.g.
storms, earthquakes and high rainfall. Of greater interest to this
review are the gravity gliding systems on passive margins where the
detachment is buried deeper (~1 km or greater), extends for tens of
kilometers in the transport direction and develops as a result of long-
term geological processes (e.g. high sedimentation rates, uplift of the
adjacent continental area resulting in tilting of the margin). Gravity
gliding can occur along a sharp, narrow detachment surface or along a
zone of ductile deformation (where gravity spreading may also
operate).
Passive margins commonly develop extensive oceanward-inclined
detachment surfaces that facilitate oceanward shallow gravity-driven
translation of sediments overlying the basal detachment (e.g. Cobbold
and Szatmari, 1991; Duval et al., 1992; Maudit et al., 1997; Rowan
et al., 2004). Loading by large deltas like the Niger, which create
depocenters over 10 km thick, can cause a large area of the basal
detachment to rotate during deposition and dip towards the continent
(at least under the DWFTB, diapir and outer growth fault provinces;
Worrall and Snelson, 1989; Morley and Guerin, 1996). More localized
variations in loading can initiate flow in a thick mobile unit from
regions of high pressure to low pressure, without the need for a
surface or basal slope (e.g. Kehle, 1989; Fig. 6B). Gravity spreading
becomes a more important driving mechanism where a thick mobile
unit is present (e.g. Morley and Guerin, 1996; Vendeville, 2005), but
in such circumstances gravity sliding and gravity spreading tend to
operate in tandem (Brun and Merle, 1985).
3.2. Salt detachment zones
DWFTBs are either associated with shale or salt detachment zones.
Important differences in structural style arise from these different
detachment types and consequently classification schemes intro-
duced subdivisions on this basis (Rowan et al., 2004; Krueger and
Gilbert, 2009), which are also used in this classification (Fig. 2).
Although the typical structural components of salt and shale
detachment systems are similar (up-dip growth fault s, diapirs,
compressional toe region), in detail the structural styles, timing of
deformation, location of deformation, and trigger mechanism can be
very different. Fig. 7 is a schematic illustration of two simila r
structure s, one with a mobile shale unit the other with a salt
detachment, to illustrate some of the differences between the types
of mobile unit. Highly overpressured saline fluids can be encountered
immediately below a salt intrusion and in a brine halo surrounding the
salt as warm, overpressured fluids rising from depth dissolved salt and
followed fault and fracture permeability adjacent to salt masses and
either became trapped in sealed compartments or migrated upwards
towards the surface (Warren, 2006). Conversely overpressured shale
2
40
3
4
5
6
7
Cenozoic sedimentary rocks
Sea floor
Makassar Straits
Thinned continental crust
Mantle
Deepwater fold
and thrust belt
0
W
W
E
E
20 km
Seafloor
West Sulawesi
Ophiolite
Offshore Onshore
Inverted rift basins
Broad detachment
Detachment folds
Break thrusts and detachment folds
Fig. 5. Regional cross-section through the Makassar Straits based on Calvert and Hall (2007). The West Sulawesi DWFTB is developed on thinned continental crust of the Makassar
Straits.
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sequences tend to show a downwards increase in overpressure over a
broad interval typical of disequilbrium compaction. Shale diapirs are
highly overpressured and will rise by hydraulically fracturing into the
overlying strata, and stoping blocks of country rock (Morley, 2003a,b).
As discussed by Morley and Guerin (1996) mobile shales tend to
deactivate with time due to dewatering and loss of overpressure, or
may change with time from disequilbrium compaction overpressures
to compacted shales fractured by inflationary overpressured fluids
Conversely the intrinsic rock properties of salt means that it can
deform, flow, migrate through the sedimentary section, and still
maintain its ductile nature, and re-activate even when long time
periods have elapsed between deformation events (see review in
Warren, 2006).
Actively subsiding basins with restricted communications to the
open seas are required for the development of large salt deposits.
Seepage of salt water across a barrier that partly separates the sea
from a restricted evaporite basin can lead to extensive deposition of
evaporites (Warren, 2006). Two of the main occurrences of such
conditions are during the late syn-rift stage (e.g. Central Atlantic), and
early rift to drift transition on passive margins (e.g. South Atlantic
margins of Gabon and Angola and Brazilian Santos and Campos
basins). Hence salt detachments associated with DWFTBs are most
common on passive margins (the collisional setting of the Messinian
evaporite in the Mediterranean is an important exception). The west
African–east South American margin is a classic example of a passive
margin, post-rift salt setting (e.g. Guardado et al., 1989; Teisserenc
and Villemin, 1989; Duval et al., 1992; Spathopoulos, 1996; Tari et al.,
2003. Salt is very weak, and seaward tilting during post-rift thermal
subsidence, regional uplift to form gentle slopes or differential loading
by sediment can trigger deformation. (e.g. Kehle, 1989; Worrall and
Snelson, 1989; Duval et al., 1992; Spathopoulos, 1996, Marton et al.,
2000).
In collision belts salt is typically present within the platform and/
or foreland basin sequences (e.g. Idaho–Wyoming segment of the
Rocky Mountains; Romanian Carpathians; Jura Mountains, Alps;
Zagros Mountains; Siwaliks, Pakistan; Oman Mountains, see review
in Warren, 2006). Salts form, apart from shales or even porous
sandstones, important detachment horizons in the external zones of
fold and thrust belts, but generally salt is not reported as playing an
important role in deformation of the more internal zones. As the
South Atlantic margin demonstrates, salt can extend from onshore
into the deepwater on passive margins, but it is questionable
whether salt can survive long enough to act as a regional
detachment horizon during collision of an accretionary wedge
with a passive margin, or whether it has been mostly extruded
and recycled in the deepwater sequences prior to collision or during
its earliest stages. An example of extensional salt-related raft
tectonics along the Cotiella thrust sheet in the Southern Pyrenees
has been recognized, but this thrust reactivates a salt weld (McClay
et al., 2004). In this context the example of the Moroccan Rif is
discussed below.
The onshore Prerif Thrust Sheet of the Moroccan Rif is an example
of the deformation style associated with a passive margin to
deepwater section that has been extensively affected by salt diapirs
and nappes (Suter, 1980a,b; Wildi, 1983; Chalouan et al., 2008). The
rocks of the Prerif Thrust Sheet lay in a passive margin setting until
Pre-deformation top of mobile detachment zone
Mobile detachment zone (viscous or plastic
material behaviour)
Brittle sedimentary rocks within gravity-driven
deformation belt
Top of sediments prior to
gravity sliding
Compressional
‘toe’ domain
Compressional
‘toe’ domain
Compressional
‘toe’ domain
Extensional domain
Extensional domain
Extensional domain
Diapir
Diapirs
A
B
C
12
Narrow detachment zone
Fig. 6. Schematic illustration of gravity spreading and gravity gliding mechanism. A) Gravity gliding down an inclined slope on a thin detachment, B) Sediment loading causing
gravity spreading employing a thick detachment zone on a flat detachment surface, C) mixed gravity gliding and gravity spreading on a thick detachment zone with both offshore
and onshore inclined detachment surfaces.
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collision of the N African margin with the Intrarif continental
fragment and the Alboran Domain induced shortening during the
Oligocene and Neogene (Michard et al., 2007). This thick, extensive,
poorly organized Prerif Thrust Sheet is several hundred kilometers
long and 30–50 km wide. It is composed of large and small oval to
linear blocks of rock that range in size up to entire hillsides several
kilometers long that are arranged either chaotically, or in a loosely
imbricated patter n within a fine grained matrix of Cenozoic
mudstone. Blocks of Triassic salt up to 100's m across are extensively
present. The unit has been interpreted as an olistostrome or tectonic
melange, with most blocks c onsisting of Mesozoic shelf-
slope
sediments; deepwater, trench associated lithologies are generally
lacking ( Wildi, 1983; Chalouan et al., 2008). Cenozoic turbidites have
been thrust over the Prerif Thrust Sheet (Suter, 1980a,b). The geology
of the Prerift Thrust Sheet is best explained as a highly deformed
halokinetic passive margin. Such a pre-weakened and deformed passive
margin when subject to regional compression expelled the salt and
further broke into highly discontinuous imbricates. The Prerif example
together with the Permo-Triassic Haselgebirge Formation in the
Northern Calcareous Alps (Spötl, 1989; Mandl, 1999) illustrates that
thick depositional salt masses in a passive margin setting may be too
deformed by gravity-driven flow to form a continuous detachment
horizon during latercollision. Sucha margin will form a poorly organized
tectonic unit during continent–continent collision. However, evidence
from other ancient DWFTBs indicates that deeper water flysch deposits
in Type 2 thrust belts can be associated with evaporite detachments,
examples include Albania (Velaj, 2001; Roure et al., 2004)(Permo-
Triassic platform sequence), Carpathians (Burdigalian & Badenian
evaporites in foreland basin sequence) , and the Apennines (Messinian
evaporites in foreland basin sequence) (e.g. Letouzey et al., 1995).
For the collisional and active margin DWFTB settings discussed
here overpressured shale is considered as the primary mobile unit.
Compacted, previously
overpressured shale
Migration and concentration
of overpressured fluids
Sediments with gas effects and
inflationary overpressures , which
reduce reflectivity on seismic data
Undercompacted,
overpressured
shales
Shale intrusions
Weld, where mobile
shale was expelled
Overpressured,
but not mobile shales
Ductile folds in
mobile shale
Fluid flux
Weld, where salt
was expelled
Fault zone or gumbo, commonly overpressured
Zone of overpressure
Zone of overpressure higher beneath
salt canopy
Stoped blocks
within mobile
shales
Poorly imaged bedding
Mud chamber
Mud volcano
A
B
Brine halo from salt dissolutionSalt
Fig. 7. Schematic comparison of structural style and overpressure distribution associated with thick mobile shale and mobile salt detachments. A) Mobile Shale, B) Salt. Partly based
on Warren et al. (in press) and observations from the Niger Delta and Brunei Darussalam (Morley and Guerin, 1996; Van Rensbergen et al., 1999; Morley, 2003a,b).
48 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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3.3. Shale detachments
Shale detachments in thrust belts fall into four main categories: 1)
compacted shales that act as weak, easy slip horizons due to inherent
material weakness (in relation to adjacent units) without large
overpressures, 2) compacted shales with large overpressures, 3) thin
(meters to tends of meters thickness) undercompacted shales with
high overpressures, and 4) thick (100's meters–kilometers thickness)
undercompacted, overpressured shales capable of large-scale viscous
flow.
Undercompacted shale detachment zones require preservation of
porosity during burial prior to DWFTB formation. Such preservation
arises due to disequilibrium compaction, in which the rate of fluid
expulsion (due to compaction) in a formation is slower than the rate
of vertical or tectonic loading, and results in formation fluids
supporting some of the overburden or tectonic stresses and becoming
overpressured as a result (e.g. Osborne and Swarbrick, 1998 ).
Undercompaction of shales overpressured by disequilbrium compac-
tion has been detected from 3D seismic and well data in the
developing decollement of the Nankai Trough (Bangs and Gulick,
2005).
Shale detachments in DWFTBs are w idely associated with high
por e fluid pressures that result from disequilbrium compaction,
inflation or a combination of the two (e.g. Dahlen, 1990; Fish er and
Hounslow, 1990; Moore et al., 1995; Fisher et al., 19 96; Fisher and
Zwart, 1997; Moore and Tobin, 1997; Scre aton et al., 1997; Bilotti
and Shaw, 2005; Morley, 2007a). Indirect evidence of widespread
overpre ssured con ditions, a ssociated with shale detachments, is
pro vided b y the numerous mud volcanoes and fluid escape features
pre sent in most DWFTBs (Fig . 8B). The deep-seat ed nature of th ese
Brittle deformation, dilation, common extension veins, scaly shear fabrics
Protothrusts, cataclasis, brittle-ductile shear
Smectite to illite phase transition
100-150° C
Porosity reduction due to compaction and tectonic deformation
drives fluids from sediments at low tempertaure and pressure
(<5 km, ~10-20% ~60-70%)
(> 80% fluid accreted is released)
Clay mineral dehydration and hydrocarbon
generation.
(3 m m yr H 0 added to fluid system)
3
-1
2
Water saturated incoming
trench sediments
Late Cenozoic accretionary wedge
Imbricated foundation of pre-subduction
Cretaceous-Palaeogene rocks
Mocene-Recent
slope basin fill
Subducting pelagic sequence
Decollement
PACIFIC PLATE
Protothrusts
HIKURANGI TROUGH
Shelf
0
3
6
9
12
km
10 km
VE ~ 2
50°C
100°C
150°C
Base of gas hydrates
Fluid migration pathway
Subduction rate 42 mm/yr
9
9
9
10
11
12
1 = Mud pipe to surface
mud volcano field
2 = BSR, base of gas hydrates
3 = shallow fluid escape pipes
and pockmarks at surface
4 = Crestal normal faults
5 = Saddle area in folds
6 = Deepwater fan
1
2
3
4
5
Simple anticline with
crestal normal faults and
break thrust on forelimb
Mature anticline with landslides and
degradation complex in forelimb
6
7
8
8
7 = Anastamosing channel
system feeding fan
8 = Piggy-back basin in syncline
9 = depression associated with
mud volcano field
10 = channel accommodating
down-slope flow of mud
from mud volcanoes
11 = slump scar from landsides on
forelimb of anticline
12 = Submarine slide MTC
13 = Basal detachment
13
Stratigraphy permeability
Fluid movement along faults
Free gas
Gas hydrate
0
3 km
A
B
Fig. 8. Illustrations of fluid flow in deepwater fold and thrust belts associated with shale detachments. A) Cross section through the Hikurang Subduction zone cross section (Barnes
et al., 2010). B) Schematic 3D section showing typical features associated with growing deepwater folds (based on Morley, 2009a and Barnes et al. 2010).
49C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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features is shown by seismic reflection data that reveals fluid pipes
rising from large thrust faults (e.g. Oregon margin, MacKay et al.,
1992; Niger Delta, Cobbold et al., 2009; Baram Delta Province,
Morley, 2009a;Hikurangsubductionzone,Barnes et al., 2010;
Fig. 8) and sea floor sampli ng of mud volcanoes which shows man y
have high heat flows and the fluids contain thermogenic hydro-
car bons (e.g. Deville et al., 2 003; Dolan et al., 2004; Deville et al.,
2006; Zielinski et al., 2007; Warren et al., in press ; Fig. 9). High pore
fluid pressures are the primary explanation for the very low
detachment strength required to generate the low critical taper of
DWFTB wedges (e.g. Dahlen, 1990; Bilotti and Shaw , 2005; Morley,
2007a).
The presence of low friction coefficient (below 0.25) clay minerals
also plays an important role in shale detachment weakness (e.g. Kopf
and Brown, 2003), and the association of DWFTBs with high pore fluid
pressures is not universal. For example Suppe (2007), Hung et al.
(2007) and Yue (2007) describe the weak Chinshui Shale detachment
in the Chenglungpu Fault, from Taiwan where the surrounding rocks
and the fault zone exhibit only hydrostatic pore fluid pressures.
However, the Chenglungpu Fault links to the basal detachment in
Taiwan, which probably is associated with high pore fluid pressures
(Yue, 2007), and a transient pulse of high overpressure during faulting
cannot be eliminated as a possibility (Hung et al., 2007). The
Chenglungpu and adjacent Sanyi fault zones offers a rare insight
into one type of shale thrust zone, since two wells were drilled by the
Taiwan Chenglungpu-fault Drilling Project, which obtained continu-
ous core across the fault zones (depths between 1013 and 1710 m).
The Chenglungpu Fault was activated during the Chi-Chi earthquake
(Hung et al., 2007). As described by Sone et al. (2007), the main
Chenglungpu Fault zone is about 200 m wide, and displays a
transition from a broad zone of weakly to intensely fractured rock,
three narrow zones (about 3–4 m wide) of breccia, foliated breccia
and fault gouge form the principal displacement zones. Black, fine-
grained, slicken-lined layers less than 10 cm thick within the principal
displacement zones may represent the products of focused, intense
co-seismic shearing. The paucity of veins within the fault zone is
considered to be typical of low-t emperature thru st fault zones
dominated by fracturing (Hung et al., 2007). According to Hung et
al. (2007) the physical properties of the Chelungpu Fault zone with
respect to adjacent host rocks are: relatively low resisitivity (breccia
zone about 40% lower) and density (20–25% reduction), and high
Poisson's ratio (~0.4). The low resistivity is due to infiltration of
drilling mud into highly fractured breccia, while the low density and
high Poission's ratio in the gouge are related to high clay content and
or
fluid. The gouge shows relatively high concentrations illite or
mixed layer smectite/illite compared with other inactive fault zones in
the area, suggesting conversion of smectite to these minerals caused
by (minor) frictional heating during the Chi-Chi earthquake (Hung
et al., 2007).
Shale detachments commonly have high permeability related to
fractures, or conditions favoring disequilbrium compaction, which
enables the presence and flow of overpressured fluids (e.g. Mann and
Kukowski, 1998). Shale detachments contain numerous sources of
permeability heterogeneity related to clay mineralogy, variations in
sedimentary unit type, diagenesis and porosity distribution. For
example radiolarians can cause high initial porosities that abruptly
collapse during burial, while turbidite sands encased in shales can
compartmentalize fluid flow until they cement up (Underwood,
2007). More lithified detachments may generate a permeable dilatant
breccia (Maltman et al., 1993).
Fluid is typically pumped along a detachment immediately after it
fails (Maltman et al., 1997). Shearing of the detachment and loss of
overpressured fluids will tend to reduce porosity in relation to
adjacent sediments beyond the shear zone (Taylor et al. 1990; Brown
et al., 1994). The extensional shear and hydrofracture networks
observed in many accretionary prisms require that pore fluid pressure
is at about lithostatic pressure in order for fractures to remain open
(Berhmann, 1991). These fractures will tend to be low-angled since
the minimum principal is vertical (Brown et al., 1994). A high fluid-
pressure, dilatant zone of high porosity about 14 m thick characterizes
the basal decollement in the northern Barbados Ridge (Shipley et al.,
1994). An abrupt transition from lateral compression within the
Fig. 9. Results of bottom sampling of the deepwater area of Brunei. A) Heat flow, B) Hydrocarbon content from piston cores. Data from Zielinski et al. (2007). Data points are overlain
on a map of the waterbottom based on 3D seismic data, showing the locations of anticlines (from Morley, 2007a).
50 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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Barbados accretionary wedge, to vertical compaction within the upper
part of the detachment zone has been identified from a magnetic
susceptibility survey (Housen et al., 1996), the transition is marked by
high pore fluid pressures (Shipley et al., 1995), that were developed in
conditions of brittle rather than ductile deformation (Housen, 1997).
The Amazon Fan provides an example of an overpressured, compacted
shale detachment on a passive margin (Cobbold et al., 2004).
The causes of long-term fluid fluxes into the basal detachment
zone vary depending upon tectonic setting and these variations can
influence deformation sty le and history (Table 1). There is a
potentially significant difference between the detachments in parts
of large deltas and accretionary prisms, related to sedimentation rates.
In typical accretionary prism settings deposition rates are low, and
deformation can oc cur over many tens of milli ons of years .
Consequently detachments tend to occur in fractured, overpressured
compacted shales, fed by overpressured fluids originating from the
down going slab. Variations in inflationary overpressure production in
the down going slab can affect the structural style. For example in the
Cascadian accretionary prism changes in vergence are accompanied
by differences in the detachment level from 400 m to 1200 m above
the igneous basement, and oblique and strike-slip faults are present at
the abrupt changes in vergence (MacKay et al., 1992; MacKay, 1995).
In the oceanward-verging zone sediments in the down going slab
reach a maximum temperature range of 120°–145 °C, while in the
landwards-verging zone maximum temperatures reach 135°–160 °C
(Underwood, 2007). The additional 15°–20 °C temperature under the
landwards-verging zone probably caused significant extra compac-
tion and permability loss associated with the smectite–illite transi-
tion, which consequently favoured the development of relatively high
disequilbrium compaction overpressures. As summarized in Table 2,
Moore et al. (2007) provide an excellent summary of the changes in
detachment behaviour found in accretionary prisms with respect to
changing temperature and pressure.
In deltas associated with high sedimentation rates, detachments
are developed within the pre-delta or earliest delta-stage marine
shales. Due to their rapid burial, overpressures are first developed in
them by means of disequilibrium compaction (e.g. Morgan et al.,
1968; Evamy et al., 1978; Knox and Omatsola, 1989; James, 1984;
Osborne and Swarbrick, 1998; see review in Tingay et al., 2009).
Subsequently inflationary overpressures internal to the shale body
may develop (e.g. hydrocarbon cracking, tectonic compaction,
smectite–illite transition (Osborne and Swarbrick, 1998; Tingay et
al., 2009). Such overpressure mechanisms have the potential to
preserve porosity within the shale mass and may promote a more
ductile type of detachment than is generally observed in accretionary
prisms (Ings and Beaumont, 2010
). One possible example comes
from eastern Bangladesh, where a mixed disequilibrium compaction
and inflationary origin for the overpressures across a 1000 m thick
interval has been determined (Zahid and Uddin, 2005; Hossain et al.,
2009). Such interpretations are controversial, and whether broad
zones (i.e. 100's m thick) of ductile, ‘mobile’ overpressured shales
can exist at depth is discussed separately in the following section
(Section 3.4).
3.4. Do thick mobile shale zones exist?
The presence of thick mobile shales in a deltaic setting up to
several kilometers thick, buried between depths of 3 and 10 km, and
composed of undercompacted, overpressured shales that deform in
highly ductile manner has been widely accepted in scientific literature
(e.g. Evamy et al. 1978; Morley and Guerin, 1996; Sandal, 1996;
McClay et al., 1998; Ajakaiye and Bally, 2002). However, the existence
of thick, deeply buried mobile shales are, at present, a controversial
aspect of delta tectonics due to improved seismic resolution, and new
observations about the extent of chemical diagenesis in shales. For
example Ruarri et al. (2009) concluded that shale movement is not
akin to salt, and that: 1) deeply sourced shale diapirs are highly
unlikely, 2) increasing seismic resolution will eventually demonstrate
thick mobile shale masses do not exist, and 3) lateral flowage of shale
is highly unlikely. Investigation of shale diagenesis from Upper
Cretaceous mudstones, offshore Norway by Thyberg et al. (in press)
demonstrated the smectite–illite transition releases fine-grained
Table 1
Summary of the key characteristics of shale detachments in DWFTBs. 1=Morley and Guerin (1996),2=Cobbold et al. (2009),3=Colten-Bradley (1987)), 4=MacKay et al. (1992),
5=MacKay (1995),6=Cobbold et al. (2004),7=Underwood (2007).
Type 1a Types 2a, 2bi Type 2bii
Origin of detachment weakness Overpressure, weak, poorly lithified mudstones,
or compacted shale with weak clay minerals
Overpressure, poorly lithified mudstones,
or compacted shale with weak clay minerals
Compacted shale with weak clay
minerals and overpressure (3–7)
Origin of overpressure Disequilbrium compaction, inflationary
overpressures due to diagenesis during
burial (1, 2)
Inflationary due to burial diagenesis during
burial and underthrusting of sediments
beneath basal detachments
Inflationary, mostly from fluids
derived from sediments subducted
in the lower plate (7)
Table 2
Summary of the observed cements, veins and relevant structural fabrics of shale detachment zones in accretionary prisms at various temperature and metamorphic facies ranges.
From Moore et al. (2007).
Temperature and pressure range Observations and source localities
b 100 °C Unmetamorphosed No veins and cements through nearly all cored rocks in ODP. Minor carbonate in Barbados drill cores. Extremely rare quartz,
carbonate and sulphides in Nankai Trough. Accreted terrigenous deposits on Barbados Island lack any carbonte/quartz veins or
cements. Faulting and stratal disruption includes cataclasis of sand sized particles with widespread development of scaly fabric
in mudstones
100°–150 °C; 5 km Typical Zeolite Facies Evidence for cementation and veining variable. Locally significant carbonate cementation along fault zones. Quartz cementation/
veins rare and minor. Cataclasis of sand–sized particles widespread. Development of scaly fabric in mudstones. Incipient
pressure solution cleavage.
150°–300 °C; 5–15 km Typical Prehnite
Pumpellyite Facies
Quartz- and carbonate-veining common. Cataclasis in sands and veined materials temporally alternates with pressure solution.
Cleavages common, especial ly visible in more coherent units. Quartz veins/cements common in the 200°–300 °C ranges. Rare
pseudotachylyte. Veining and cementation facilitate change from velocity-strengthening to a velocity weakening, earthquake-
prone rheology.
150°–350 °C; 15–27 km Classic HP-LT
(e.g. Franciscan Blueschists of eastern belt)
Often coherent layering with multiple fold generations, local stratal disruption predating metamorphism. Pressure solution
common. Progressive replacing of brittle fabrics with more ductile fabrics
N 300 °C; 8–27 km Prehnite–Actinolite,
and Blueschist–Greenschist Facies
Widespread foliation development, pressure solution, microscopically apparent recrystallisation during development of
metamorphic assemblages. Many terranes surprisingly coherent, with multiple phases of folding. Brittle deformation during
peak metamorphism not obvious, but a previous history is suggested by the stratal disruption in some areas.
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quartz that helps stiffen mudstones. 1–3 μm spherical, discrete grains
or short chains and clusters of quartz form a pervasive network that
developed when burial reached depths around 2500 m or 80°– 85 °C.
Helset et al. (2002) have proposed that while overpressure retards
mechanical compaction, it does not retard chemical compaction, and
consequently diagenetic processes may control the timing and
magnitude of overpressure rather than disequilibrium compaction.
Mud volcanoes apparently provide evidence for deep, under-
compacted, overpressured shales. However, if they originated as
diatremes this may not be the case (Brown, 1990). Diatreme-type
mud pipes develop as a consequence of deep gas generated
overpressures that rise upward in a hydraulically fracturing column
and mix with higher undercompacted muds and fluids to generate
mudflows at a shallower level (Brown, 1990). However, the existence
of diatreme-type pipes does not mean that all mud pipes can be
explained in this way (Brown, 1990). In Brunei for example Middle
Miocene shoreface clastics were intruded in two phases (pre-folding
and post-folding) by overpressured mudstones derived from the
underlying Setap Shale Formation (Morley et al., 1998; Morley,
2003b). The Setap Shale Formation forms the basal ‘mobile’ shale zone
for the Baram Delta Province and in the subsurface overpressures
increase downwards towards the Setap Shale Formation as expected
for a classic disequilbrium compaction detachment zone. Furthermore
the porosity-effective stress relationship follows the classic disequil-
brium comp action model within the Setap Formation, not the
inflationary overpressure trend expected for a chemical compaction
model (Tingay et al., 2009).
The issue of whether thick mobile shales are present at depth
arises in regions of shale mini-basins and diapirs, and large growth
faults; the Niger Delta is particularly important to the debate. The
presence of downbuilding synformal depocentres, and strong evi-
dence for pinching and swelling of the inferred mobile shale zone
from seismic reflection data led Morley and Guerin (1996) to
conclude that parts of the Niger Delta (particularly the outer shelf
and slope) were underlain by a thick mobil e shale unit, an
interpretation reconfirmed by Ajakaiye and Bally (2002). Numerical
modeling of shale-detachment structures like the Baram Deltaic
province and the Niger Delta indicates that a thick mobile shale
section is required to generate the counter-regional fault province
that lies in the outer shelf region (a thin detachment does not produce
counter-regional faults), and that a thin shale detachment is present
in the outer fold and thrust belt (Ings and Beaumont, 2010). A study
by ExxonMobile of the inner fold and thrust belt and diapir belt of the
Niger Delta utilized seismic reflection geometries, velocities and
gravity data to conclude that the seismically transparent ‘mobile’
shale zone was a thick, low density, low velocity, overpressured,
ductile zone (Wiener et al., 2009). This zone corresponds with the
region of thick mobile shales required by the numerical modeling of
Ings and Beaumont (2010).
A distinction needs to be made between whether thick shale
sequences, capable of viscous flow exist, and whether shale diapirs
similar to large salt-like diapirs can form (e.g. Van Rensbergen et al.,
1999; Morley, 2003a). Possibly there are no shale equivalents to
classic salt diapirs, instead shale ‘diapirs’ are either narrow pipe like
intrusions, or the larger ‘diapirs’ are actually shale-cored anticlinal
folds, related to buckling an d compression, not active diapiric
processes (e.g. Duerto and McClay, 2002). This folding origin for
diapir mini-basin provinces may be applicable to both the Niger Delta
(Wiener et al., 2009), and NW Sabah (King et al., 2009).
In summary there remains a strong case for the presence of thick
mobile shales in parts of large deltas, although undoubtedly thin
detachment zones are present in many deltas and DWFTBs. Thick,
ductile mobile shales may have been too widely applied in the past,
but equally it is premature to completely eliminate them from all
interpretations (see the Timor example, Section 4.3.1). Ductile mobile
shale detachments may still be present even if evidence for shale
diapirs diminishes, and diapirs are re-interpreted as shale-cored
anticlines. Present understanding of chemical compaction assumes
there is no overpressure control on retardation
. Conversely, the
evidence in favor of mobile shales suggests such diagenesis
retardation mechanisms exist. For example, there is some evidence
for overpressures causing either a change in the type of diagenesis
(Xie et al. 2003), or retardation of diagenesis in sandstones (e.g.
Osborne and Swarbrick, 1999) and chalk (Hardman, 1982), particu-
larly if hydrocarbons are part of the overpressured fluids.
4. Structural and petroleum system characteristics of the different
types of DWFTB
4.1. Near-field stress-driven linked systems
4.1.1. Introduction
Near-field stress-driven DWFTB have been reviewed by Rowan et
al. (2004) in considerable detail, although, as noted in Section 2.3 we
have not limited this category exclusively to passive margins. Tari et
al. (2003) proposed sub-divisions of salt-detachment passive margins
types based on observations along the West African margin. They
noted that the extensive post-rift Aptian-age salt found offshore
Cameroon, Equatorial Guinea, Gabon, Congo and Angola was
relatively thick and widespread, and comprised a system of up-dip
extension, and down-dip contraction where salt nappes, canopies and
toe-thrusts and folds accommodate most of the shortening. In
contrast the Late Triassic and Early Jurassic syn-rift salt in northern
West Africa (Morocco, Mauritania, Senegal, The Gambia and Guinea–
Bissau) was isolated by syn-rift highs and did not necessarily link with
up dip extension. Tari et al. (2003) identified the rate of sediment
loading and resulting slope steepness, and efficiency (i.e. thickness of
salt, salt detachment continuity) of the salt detachment as the first
order controls on the observed variations in salt deformation style
along the West African margin.
The first order control scheme of Tari et al. (2003) was tailored
particularly for sal t, and the differences in continuity of salt
detachment between salt deposited in a syn-rift and post-rift setting.
Such distinctions are not exactly applicable to shale detachments,
since continuous, thick shale horizons, with the potential to become
overpressured are present much more frequently than thick salt
layers, whilst overpressures in shales are transitory and bleed off over
time (unlike salts that are always mechanically weak at typical
subsurface temperatures and pressures). Nonetheless the general
approach of the Tari et al. (2003) classification is also appropriate for
shale detachments (Fig. 10).
There does not appear to be any clear-cut variation in basic
characteristics, such as fold wavelength, between the different
detachment lith ologies (Fig. 11). The clearest control on fold
wavelength, amplitude and spacing is related to the maximum
stratigraphic thickness of units overlying large-scale shale detach-
ments, which varies from about 1 km to ten or more kilometers
(Fig. 12). The use of stratigraphic thickness is a crude parameter, but is
the most straightforward one that can be applied to all DWFTBs, and
has been identified as an important parameter in numerical modeling
of folding vs. faulting behaviour above detachments (Simpson, 2009).
Ideally the thickness of the dominant competent member should be
used. But in many DWFTBs the generally fine grained, poorly lithified
nature of the section suggests that the presence of strong mechanical
contrasts within the stratigraphy are unlikely, or will be only locally
present. Elastic strength will gradually increase with depth due to
compaction and diagenesis, which is why stratigraphic thickness
shows a correlation with fold wavelength. The presence of controlling
competent layers will be more important on passive margins affected
by salt, where the age range of units involved in the DWFTB is greater,
and highe r degrees o f lithification and compaction will have
developed more mechanical contrasts within the stratigraphy. Larger
52 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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folds, and longer duration of activity tend to be associated with
stratigraphically thicker, more long-lived deltaic systems. The
thickness a nd viscosity of the detachment also have a first-order
control on structural style, (Simpson, 2009). Fold trains, or fold
trains with break thrusts, and thrusted diapirs tend to be associated
with thick mobile shale units (e.g. Me xican Ridges, Inner Fold Belt,
Niger Delta Fig. 10B). Conversely imbricate thrusts, pop- up and
triangle zones and associated folds tend to be found along discrete
bas al thrust or thin detach ment zones (e.g. Corr edor et al., 2005;
Cobbold et al., 2009; Fig. 10A, C, D). This application of the Tari et al.
(2003) scheme is fur ther modified to capture one important aspect
of shale detachments: that the detachment does n ot ne cessarily
only dip oceanward. In the case of the Niger Delta the DWFTB is
underlain by a landward dipping detachment (Morley an d Gueri n,
1996; Bilott i and Shaw, 2005). Fig. 13 shows the range of near field
stress-related shale detachments based on stratigraphic thickness,
mobile shale thickness and detachment d ip. Five examples of shale
detachments are plotted; their wide separation illustrates the
considerable variation in basic parameter s withi n the Type 1a
DWFTBs.
Offshore dipping detachments tend to be driven by steepening of
the slope due to hinterland uplift, and hence several episodes of
deformation may affect a margin over tens of millions of years
(Hudec and Jackson, 2004; Rowan et al., 2004). Major deltas are a
more continuous source of near-field stress due to differential
loading and gravity slidin g, particularly where sediment thickness
is great enough to cause a landward dip to the basal detachmen t.
Loading by maj or delta s ultimately imposes offshore progradation
and offshore propagation of folds and thrust s. The two outstanding
examples of this are the Niger Delta and northern Gulf of M exico.
Convers ely, the offsh ore dipping salt detachment systems along th e
SW African margin show a propensity fo r lat e-state l andward
pro pagation of deformation (Hudec and Jackson, 2004; Jackson et
al., 2008; as described below).
Sediment thickness
Mobile shale thickness
Narrow fault/detachment zone
Thick mobile shale
Thin overburden
Thick overburden
0 10’s m 100’s m > 1km
Atlantic
Top Pliocene
Top Miocene
Top Cretaceous
Base Rift
Top Cenomanian
Depth (km)
SW NE
0
40 km
10 km
Two Way Time (s)
W
E
Top Oligocene
Top Mesozoic
Faja de Oro Fault
Top Pliocene
MEXICAN RIDGES
Top Miocene
Cretaceous
Eocene
Seafloor
Annapurna Field
2
4
6
TWT
(s.)
SW
NE
25 km
Extensional Domain
Transitional Domain
Contractional Domain
Break-up
Unconformity
Seabed
M. Aptian
Maast.
u/c
A
B
C
D
0
5
10
0
1
2
3
4
5
6
7
8
Fig. 10. Range of structural styles associated with shale detachments. A = Amazon Fan, (based on seismic section in Cobbold et al., 2004), B = Mexican Ridges (based on seismic
section in Ambrose et al. (2005), C = Krishna-Godavari Basin, India, (redrawn from Cairn Energy online seismic image), D=Orange Basin, De Vera et al. (2010),E=first order control
on structural style scheme for shale detachments based on the scheme proposed for salt tectonics by Tari et al. (2003).
53C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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4.1.2. Type 1a, DWFTBs associated with regional offshore-dipping
detachments along overpressured shales
Examples: Bight Basin S. Australia (Totterdell and Krassay, 2003);
Orange Basin S. Africa (de Vera et al., 2010); West Luconia Delta, W.
Borneo; Sandakan Basin, N. Borneo (Samsu et al., 2000); the NW flank
of the Niger Delta, W. Africa, (Morley and Guerin, 1996). Krishna–
Godavari Basin (Annapurna Field, Cairn Energy); Mexican Ridges, Gulf
of Mexico (Ambrose et al., 2005), Rovuma basin N. Mozambique,
(http://www.hgs.org/en/cev/930/).
Despite displaying regional offshore dips, large tracts of passive
margins do not exhibit well-developed gravity tectonics related
structural provinces with linked, large-scale extensional and contrac-
tional deformation provinces. However, extensive regions of nested,
small-scale normal faults are commonly present and indicative of
widespread deformation related to mass movement, diagenetic
processes and dewatering (see for example Cartwright and Huuse,
2005). An offshore-dipping slope is generally insufficient to induce
widespread, large-scale gravity sliding in shale-prone passive margin
sequences; instead such systems occur where high sedimentation
rates associated with deltas are present. Type 1a systems tend to
display relatively simple geometries with a system of predominantly
oceanward-dipping normal (regional) faults, that pass down dip into
a zone of mostly landwards-dipping imbricate thrust faults and
oceanward-verging folds (Fig. 10). In the transition zone between the
extensional and compressional provinces “shale diapirs” or mud pipes
can be developed.
A belt of regular, landward-dipping imbricate thrusts dominates the
DWFTBs of the offshore Bight Basin, S. Australia and the Orange Basin,
South Africa. The width of the toe thrusts and folds in the Bight Basin is
20–30 km, the thickness of section is up to 3 km, with ~3–4 km spacing
between imbricate thrusts (Totterdell and Krassay, 2003). In the Orange
Basin, toe folding and thrusting is spectacular, and occupies a zone up to
60 km wide in the transport direction, and lies within a section ~3 km
thick (Fig. 10D). Spacing between the imbricate thrusts is 1–5 km. As is
commonly the case in thrust-associated deepwater folds, the back-limb
dip of the anticline is gentler than the dip of the imbricate thrust (de
Vera et al., 2010). Hence the geometry does not conform to the classic
(no shear) fault bend fold or fault propagation fold model (e.g. Suppe,
1983; Suppe and Medwedeff, 1990), but instead appears to be more
appropriately described by a shear fault-bend fold model (Suppe et al.,
2004), or a break thrust model (e.g. Eisenstadt and De Paor, 1987;
Morley, 1994, 2009b).
The up-dip extensional province in the Orange Basin is strongly
segmented by lateral faults, parallel or sub-parallel to the transport
direction (de Vera et al., 2010). This segmentation appears to
continue, although not necessarily along the same structures, down-
dip into the fold and thrust belt where it causes abrupt along-strike
changes in the location of folds and thrusts (de Vera et al., 2010 ). Such
10
20
30
0
Perdido
Nile Delta (Messinian)
Gabon (Astrid)
Angola
Kwanza (Congo)
5.7
5.6
5.4
Scotia Basin
Amazon Fan
Cyprus
Mexican Ridges
Niger Delta
Zagros
16.7
Makran
4.2
Indo-Burma Ranges
Seram
NW Borneo
South Caspian Sea
Makassar Straits
15
10
12
3
7.9
4
Orange Basin
Fold wavelength (km)
Type 1 Type 2
Salt detachmentShale detachment Shale detachment
4
Mean value of fold population wavelength (km)
values only given where 15 or more folds were measured
vertical bar = range of values
Fig. 11. Fold wavelengths for a range of DWFTB settings. The Zagros Mountains although not a DWFTB are illustrated for comparison with a classic orogenic fold and thrust belt.
54 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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segmentation is likely to exist in other, similar DWFTBs. The origin of
oblique and lateral faults is highly variable and for both salt and shale
detachments and includes: inheritance from underlying structures,
inheritance of lateral fault s from up-slope, lateral changes in
detachment zone thickness or material properties, lateral changes in
thickness or lithology in the detachment upper plate, the effects of
variable strain around an object in the lower plate, and conjugate
strike-slip faults in type-1 fold-related fracture-type orientations,
(Stearns and Friedman, 1972), e.g. Lewis et al. (1988), Jackson et al.
(2008), Morley (2009b), Clark and Cartwright (2009).
The Mexican Ridges DWFTB in the western Gulf of Mexico is
developed above an Oligocene shale detachment (Fig. 10B; Buffler,
1983, Feng et al., 1994, Wawrzynice et al., 2003, Ambrose et al., 2005).
The ridges are predominantly offshore verging folds, with some
break-thrusts in their forelimbs. The folding style contrasts strongly
with the imbricate thrust style of many other Type 1ai DWFTBs
(Fig. 10A, C, D). Probably this is due to a relatively thick zone of mobile
shales forming the Oligocene detachment zone.
It is expected that extension should approximat ely balance
contraction in the toe area in linked systems. However, quantifying
the contraction can be difficult since poorly lithi fied, water-rich
sediments have been deformed and consequently simple line-length
or area balancing of folds and thrusts will not necessarily account for
all the deformation. Structurally induced compaction and dewatering
may also play a significant role (e.g. Morgan and Karig, 1995; Henry et
al., 2003; de Vera et al., 2010; Butler and Paton, 2010). In the Orange
Basin, the difference between extension (24 km) and shortening
(16 km) measured from large-scale structures visible on seismic lines,
suggests sub-seismic deformation processes may account for about
24% shortening (de Vera et al., 2010).
Deformation along a shale detachment will tend to cause
dew atering and loss of the initial weak characteristics of the shale
layer particularly if disequilibrium compa ction is the main mech-
anism for generating overpressures. Consequently new deltaic
cyc les on the margin, or new epi sodes of uplift may not cause
reactivation of the initial detachment, the margin may cease to be
prone to large-scale gravity sliding (as seen i n the different
deformation styles of th e Albian-C enomanian White Pointer, and
Santonian–Masst rictian Hammerhead deltas of the Bight Basin,
Totterdell and Krassay, 2003). Alter natively several stacked detach-
ments of different age may be present where highly overpressured
shales are present, e. g. the Orange and Krishna–Godavari ba sins
(Fig. 10C and D).
The Annapurna Field, with mean resources of 76 MMBOE, was
discovered by Cairn Energy in the Krishna–Godavari Basin, offshore
eastern India. The basin contains one of the few fields found so far in
relatively small Type 1 shale detachment provinces. The field is
located in a broad anticline set up between the compressional toe
thrust and the extensional faults up dip (Fig. 10C). This is an atypical
structure compared with the detachment, fault bend and fault
propagation folds found in the larger delta DWFTBs of the Niger
Delta and Gulf of Mexico.
Stratigraphic thickness (km)
30 km0
0
5 10 15 20 25
10
5
1
1 = Orange Basin
2 = Mexican Ridges
3 = Outer fold and thrust
belt Niger Delta
4 = Indo-Burma Ranges
5 = South Caspian Sea
6 = NW Borneo
7 = Seram
8 = Makassar Straits
2
3
4
5
6
7
= range of values
= mean value
Fold wavelength
Fig. 12. Wavelength vs. stratigraphic thickness of deformed series for eight shale detachment-associated DWFTBs.
Detachment dip (oceanward)
Detachment thickness
Thickness of stratigraphy (km)
0
5
10
10’s m
100’s m
1000’s m
-2°
0°
10°
1
5
3
2
1 = Main Niger Delta
2 = NW margin of Niger Delta
3 = Amazon Fan
4 = Mexican Ridges
5 = Orange Basin
4
Fig. 13. Triaxial scheme illustrating the wide range of basic parameters (thickness of
stratigraphy, detachment thickness, detachment dip direction) associated with Type 1a
shale detachment DWFTBs.
55C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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4.1.3. Type 1a, DWFTBs associated with a widespread hinterland-dipping
basal detachment zone (large deltas)
The Niger Delta displays the classic, well-developed zones of a
large delta province with an onshore to shelfal zone of growth fault-
controlled depocentres, a well developed belt of “shale diapirs”
around the outer shelf-slope area, and a fold and thrust belt in the
slope area (e.g. Evamy et al., 1978; Doust and Omatsola, 1989; Morley
and Guerin, 1996; Haack et al., 2000; Ajakaiye and Bally, 2002;
Fig. 14). The Cenozoic stratigraphy of the delta is divided into three
main formations (e.g. Evamy et al., 1978), the Benin Formation
(alluvial), Agbada Formation (deltaic-inner shelf mixed sands and
shales), and the Akata Formation (outer shelf, slope, and bathyl
deepwater shales with deepwater sands). For the deepwater area the
Akata Formation contains source, seal and reservoir rocks.
The delta province extends about 300 km from the onshore
extensional province to the thrust front of the DWFTB, and covers a
distance of about 500 km along strike (Fig. 14). The DWFTB forms two
lobes, each about 200 km along strike, with a central area that is
largely unaffected by thrusting and folding that coincides with the
Charcot Fracture Zone (Fig. 14). The DWFTB extends about 80–100 km
in the transport direction, and has undergone maximum shortening of
about 17–25 km (Corredor et al., 2005). Some folds display pop-up
and triangle-zone geometries, or even a series of back-thrusts (e.g.
Bilotti and Shaw, 2005), but, in general, most folds and thrusts verge
offshore. The Niger DWFTB forms a low critical-taper wedge with a
1°–2° surface slope (α) and 1°–1.5° basal detachment dip (β) with an
overall wedge taper of (α + β) 2.5°± 0.4°, which is explained by high
pore fluid pressure ratios (λ =~0.9) along the basal detachment
(Bilotti and Shaw, 2005). Typical for a critical wedge model, the
sequence of thrusting tends to young offshore, but there are instances
of out-of-sequence thrusting, synchronous deformation or break-back
sequences on a number of thrusts and folds (Morley, 2003a; Corredor
et al. 2005). Spatial and temporal changes in the strength of the Akata
Formation, in particular due to variations in pore fluid pressure and
sand content, are likely to be responsible for many of the changes in
structural style, sequence of deformation and location of detachment
levels observed in the DWFTB (Corredor et al., 2005; Briggs et al.,
2006). Basement topography influences the DWFTB structure in
places too, for example in the location of lateral ramps, and the
presence of basement highs in front of well-developed imbricate
structures (e.g. Morley and Guerin, 1996; Morley 2003a; Corredor et
al., 2005).
Corredor et al. (2005) divided the DWFTB into an inner fold and
thrust belt characterized by oceanward verging, commonly imbricat-
ed, thrusts and folds, including detachment folds, a transition zone
which is largely undeformed except for large, broad folds, and an
outer fold and thrust belt with basinward- and hinterland-verging,
conjugate thrust faults and associated folds (Fig. 15). Many of the
structures are active and their geometry is reflected in the
bathymetry. The outer fold and thrust belt has been the main focus
of deepwater exploration, in particular the southern lobe. For example
the Agbami Field (1998; Fig. 14) is located in the southern lobe, and is
the largest of the deepwater discoveries in Nigeria, with estimated
recoverable volumes of 900 MMBBO, and production rates of
250,000 BBO/day (Maksoud, 2008).
The low-angle of the basal detachment and fore-and back-thrust
fold style in the outer fold and thrust belt (Fig. 15) implies a low
resistance to basal slip and high pore fluid pressures for the sharp
detachment zone that lies within the Akata Formation (Cobbold et al.,
2009). Relatively slow seismic velocities within the Akata Formation
support this inference of high pore fluid pressures, and a pressure
ramp near the top of the formation observed in wells ( Morgan, 2003;
Cobbold et al., 2009). Frost (1996) and Cobbold et al. (2009) have
argued for that hydrocarbon generation triggered by burial within the
thrust belt is responsible for the large-volume overpressures. The
cycle of burial by thrusting and sedimentation that lead to
Edge of thrust
and fold belt
Transition zone
Transition zone
0 50km
4° 5° 6° 7° 8°
3°
4°
5°
3°
4°
5°
Zone dominated by counter
regional growth faults
Zone dominated by
regional growth faults
Diapir province
Inner fold and thrust
belt, including thrusted
diapris
Thrust fault
Normal fault
Diapir ridge axis
Fig. 16
Fig. 17
Oblique thrust
at basement fault
trend
Outer thrust
and fold belt
Outer thrust
and fold belt
Diapir belt at
thrust front
Folds
Ch
a
rcot Ridge
Fig. 15
Akpo Field
Agbami Field
Bonga Field
Erha Field
Fig. 14. Regional map of the main structural provinces of the Niger Delta (modified from Morley and Guerin, 1996 and Cobbold et al. (2009).
56 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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overpressure generation probably caused a feedback mechanism that
aided oceanward propagation of the thrust front (Cobbold et al., 2009).
Evidence that a significant portion of the up-dip extension is
accommodated by shortening landwards of the present DWFTB
includes: 1) the relatively young age of most of the folds (Late
Miocene–Holocene) compared with the Eocene–Holocene age of the
delta. 2) Isolated examples on the shelf of individual growth fault
depocentres tha t have self contained extensional–contractional
systems with growth faults passing down-dip into either “shale
diapirs” or an individual toe thrust (Fig. 16), and 3) The extensive
“diapir” belt that lies between the DWFTB and the extensional
province has taken up some horizontal extensional displacement by
vertical motions and fluid expulsion (Morley, 2003a). Progradation of
the delta over older, more landwards fold and thrust belts possibly
caused them to be lost into the deeper, poorly imaged section.
However, it is also likely that the structural style has changed with
time as the sand-prone delta prograded over thick mobile over-
pressured shales that thin oceanward. Consequently, earlier up-dip
extension was mostly accommodated by movement of thick mobile
shales to form the “shale diapir” belt (Fig. 16B, C; see discussion in
Section 3.4, and Ings and Beaumont, 2010).
Once the delta began to prograde over the “diapir” belt, the DWFTB
started to develop oceanward of the “diapirs” with a relatively thin
basal detachment zone in most places (Morley, 2003a). In Fig. 16D the
latest stage sedimentary sequence (mid Pliocene–Holocene) displays
a broadly synformal geometry in the domain of the “diapir” province,
and is not strongly fault controlled. This geometry suggests a broad
subsidence of the basal mobile shale layer to accommodate the
sedimentation, possibly largely achieved by fluid loss. The geometry
implies broad differential loading, rather than extensional faulting
may have driven the later stage of deformation in the outer fold and
thrust belt in the southern lobe.
Improvements in seismic data quality commonly show that chaotic
regions interpreted as diapirs on older vintages of seismic data shrink in
2 km
1 s. TWT
Akata Fm.
Agbada Fm.
SSW NNE
Ocean floor
Fig. 15. Outer fold and thrust belt of the Niger Delta, drawn from 2D seismic line in Cobbold et al. (2009). See Fig. 14 for location.
Mobile shales
limit of strong overpressured,
mobile shale
?
Mobile shales, or partially
mobile shales, or just
poorly imaged stratified shales?
10 km
A
B
C
Migration of
overpressured fluids
Mobile shale
Upwards moving front of overpressure
Seafloor
Seafloor
Seafloor
S N
Outer fold and
thrust belt
Inner fold and
thrust belt
Late stage synformal geometry
to Mid Pliocene-Holocene section
Late stage synformal geometry
to Mid Pliocene-Holocene section
Mid Pliocene-Holocene
Palaeogene-Near top Miocene
Near top Miocene-Mid Pliocene
Fig. 16. A) Cross-section based on 2D seismic reflection data (Ajakaiye and Bally, 2002) across the southern DWFTB lobe of the Niger Delta (see Fig. 14 for location). B) Conceptual
evolutionary diagram of southern Niger DWFTB: A) Early stage of development with movements on growth faults passing down-dip into shale ridges, B) seaward propagation of
growth faulting and development of second, thrusted shale ridge. C) Cessation of growth faulting, collapse and movement on mobile shale due to differential loading in a saucer-
shaped basin (only minor fault control). During this differential loading stage the outer fold and thrust belt developed.
57C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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size, or are replaced by layered sequences on newer data (e.g. Van
Rensbergen et al., 1999; Van Rensbergen and Morley, 2003). The DWFTB
of the Niger Delta contains folds and imbricates that in places clearly
only affect strongly reflective sequences, but in other places are weakly
reflective or chaotic (Ajakaiye and Bally, 2002). Some recent publica-
tions (e.g. Corredor et al., 2005; Kostenko et al., 2008) have emphasized
the regular thrust belt geometries that clearly occur in many parts of the
DWFTB. However, in parts of the delta, particularly around the outer
diapir province and the eastern margin of the DWFTB, there is a different
structural style,where mobile, overpressured shale diapirs seem to form
the core to imbricated structures (Fig. 16; Morley, 2003a). Ajakaiye and
Bally (2002) show the potential for mobile shale involvement within the
fold and thrust belt, in particular their Fig. 1 profile D3 suggests mobile
shale with a diapiric geometry maybe involved out to the thrust front in
places. The eastern termination of the southern lobe is a narrow feature,
with a very different style from the relatively widely spaced imbricates
of the central southern lobe (e.g. Corredor et al., 2005; Cobbold et al.,
2009). Imbricated slivers are stacked much more vertically and a narrow
zone of mobile shale is possibly involved in the imbricate stack (Fig. 17).
In a different deepwater setting Barber et al. (1986) have shown
from outcrop that large chaotic masses of squeezed diapirs form in
overpressured DWFTBs, and may be an important mechanism of
mélange formation. It is reasonable to suspect that such processes
could also operate within the DWFTB of the Niger Delta. At the very
least there is a cycle where very large volumes of fluid and mud are
pumped from the deeper parts of the delta to the surface along
vertical pipes to form mud volcano fields, and then the mud becomes
re-deposited within the delta sediments and buried (mud pipe-
volcano systems may extrude volumes of material in the order of 1–
11 km
2
over their lifetime, Guliev, 1992; Graue, 2000). This cycle,
coupled with significant shortening by lateral, compression-related
compaction cast doubts on the utility of regional balanced cross-
section solutions for entire deltaic provinces that treat deformation
only in terms of linked fault systems, narrow detachment faults,
simple fault-related folds and assume constant area or line length (e.g.
Butler and Paton, 2010).
The Niger Delta is larger, and contains more structural variety than
the shale detachment systems described above and summarized in
Fig. 10. In particular, the presence of large “diapirs” and an external
thrust belt developed above a hinterland-dipping basal detachment
(e.g. Evamy et al., 1978; Morley and Guerin, 1996). This sets the Niger
Delta apart from other near-field stress driven Cenozoic deltas. One
other exception may be the McKenzie Delta, but its structure is not
well described in the literature and so the Niger Delta is discussed
here as a lone example.
4.1.4. Type 1b, DWFTBs associated with an oceanward-dipping salt
detachment zone
We estimate that thick salt units are found along about 13,000 km
length of the world's passive margins. The most important areas are
the Atlantic margins of South America (Campos, Espirito Santo,
10 km
TWT (s.)
Mobile shales (Akata Formation)
Upper pre-kinematic sequence
Upper syn-kinematic sequence
WE
Lower pre-kinematic sequence
Lower syn-kinematic sequence
Middle syn-kinematic sequence
0
5
8
Fig. 17. Stacked imbricates and mobile shale eastern Niger Delta, based on Ajakaiye and
Bally, 2002, their Fig. 1 profile D3; see Fig. 14 for location).
Angola Salt
Nappe
Thickened salt palteau
Diapir domain
Monocline domain
Atlantic Hinge Zone
Raft Domain Mock Turtle Domain Updip Wedge
Eastern RimInner Salt Basin
Flamingo Platform
Outer Salt Basin
Probable Gabon-Angola horst and graben system
Distal Ramp
Congo Craton
50 km
Depth (km)
Coast line
WSW ENE
Outer Kwanza Basin (passive margin)
Inner Kwanza Basin (interior basin)
C
Abyssal
Plain
Two-way time (s.)
Allochthonous salt fringe
Basement ramp
Distal wall
Astrid thrust belt
Sea floor
Inverted extensional
domain
Contractional domain
Extensional domain
Truncated shelf
S
N
20 km
Aptian salt
Aptian salt
SE
NW
3.0
4.0
5.0
6.0
7.0
8.0
9.0
TWT (s)
10 km
Complex fold-belt
salt-cored anticlines, forced folds above basement blocks
Thin-skinned contraction
Thin-skinned extension
Albian-
Cenomanian
Upper and Middle Jurassic
0
2
4
6
8
0
1
2
3
4
5
6
7
A
B
Fig. 18. Cross-sections across Type 1b (salt detachment) margins. A) Regional cross-section across the Kwanza Basin, Angola, redrawn from Hudec and Jackson (2004). B) Cross-
section across the Sable sub-basin, offshore Nova Scotia based on seismic reflection data, (Deptuck et al., 2009). C) Regional cross-section from the inner continental shelf to the
deepwater Astrid thrust belt, Gabon, redrawn from Jackson et al. (2008).
58 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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Sergipe and Santos basins), West Africa and Canada (Scotia Basin), the
Gulf of Mexico, and the Red Sea. Unlike the Type 1a margins, which
are associated with a shale detachment, gravity-driven deformation
above a salt detachment can occur outside of areas with large deltaic
input onto the margin. Commonly, uplift of the continental margin
triggers salt movement. For example, the Angolan margin underwent
three phases of uplift (early Albian, Campanian and Miocene) that
triggered reactivation of the same Aptian salt layer (e.g. Hudec and
Jackson, 2004). Structures associated with mobile salt can be very
complex, and vary depending upon the thickness of the salt layer,
geometry of pre-existing structures, and uplift history and sediment
loading history of the margin (Fig. 18).
Five basins are described below to illustrate the primary variations
in margin setting, salt thickness, salt timing (syn-rift or post-rift
deposition), and triggering mechanism (thermal subsidence, hinter-
land uplift, deltaic loading).
4.1.4.1. Kwanza Basin, Angola. The Kwanza Basin is an outstanding
example of a long-lived thick salt mass on a passive margin. The
following description is based on Hudec and Jackson (2004). The
thickness of the deformed section is highly variable (Fig. 18A). Up to
4 km of salt, capped by a thin veneer of sediment, was present in the
western part of the basin at the time of initial deformation. Passing
eastwards, up to 3–4 km thickness of sediment overlay thinner, more
isolated salt bodies. The isolated character was due either to salt
movement, or salt depositional thickness not exceeding 50–150 m.
The Kwanza Basin is divided into two sub-basins, the Inner Kwanza
Basin (predominantly located onshore at present-day) and the Outer
Kwanza Basin (located on the shelf to deepwater). A broad high
separated the Inner Kwanza Basin from the Outer Kwanza basin for
most of the Cretaceous–Cenozoic, so there was no linked detachment
system between the two basins for most of their history. This
description focuses on the Outer Kwanza Basin, which had a DWFTB
linked up-dip to an extensional system throughout its history.
The presence of the thick salt caused the structural styles in the
DWFTB to be much more varied than the type 1a DWFTBs, and include
a large-scale salt nappe, thickened salt plateau, and folded and
thrusted (squeezed) diapirs (Figs. 18 and 19). Above the salt nappe,
are relatively short wavelength (2–3 km), broad, open folds affected
by both normal faults and thrusts, or with no faults at all. There is no
strong vergence to the structures.
According to Hudec and Jackson (2004), three phases of regional
uplift are each associated with up slope extension and down slope
contraction, but the location and structural style of the deepwater
zone of contraction differs for each phase. The Albian phase occurred
over a wide area and involved short-wavelength buckle folds (up to
~5 km) and compressional diapirs. Most shortening occurred up-dip
of the basinward salt pinch-out. The Late Cretaceous–Middle Miocene
contraction was focused on the basinward limit of the salt by advance
of the Angola salt nappe. There appears to be little contraction within
synclinal mini-basins created during Albian buckling. About 20 km
basinward translation affected the Outer Kwanza Basin during the
Oligocene–Holocene. Miocene deformation is marked by acceleration
in nappe translation, and the contractional zone broadened to a 150-
km wide zone from the leading edge of the allochthonous salt to the
base of the continental slope. Hudec and Jackson (2004) estimate that
relatively minor uplift and tilting of the continental shelf resulted in
an order of magnitude increase in translation rate (from 300 to
3200 m/my).
Increased sedimentation rates from 8 to N 130 m/my in the
abyssal-plain region occurred during the Late Miocene. One effect
was to lower the slope on the continental rise and thus decrease the
potential energy driving the Angola salt nappe. A second effect was
aggradation of sediment oceanward of the distal salt margin. The roof
of the nappe became effectively buttressed, and its advance hindered.
Consequently oceanward-propagating deformation was limited and
instead there was a landward shift in shortening (Hudec and Jackson,
2004). The landward propagation caused older salt structures to be
reactivated and laterally shortened. Some reactivated structures
display a strong influence of sediment package geometry developed
under extensional or diapiric conditions.
4.1.4.2. Lower Congo Basin, Astrid thrust belt. In many parts of the
Congo passive margin Basin relatively thick salt has led to the
development of allochthonous salt canopies and structures akin to the
Kwanza Basin (Marton et al., 2000; Tari et al., 2003). However, the
Astrid Thrust Belt of the Lower Congo shelf displays deformation
associated with a much thinner salt layer, and relatively thicker
sedimentary section (Congo Fan) than the Kwanza Basin described
above (Fig. 18C; Basin; Jackson et al., 2008).
The following account is based on Jackson et al. (2008). Gravity
driven shortening during the Aptian caused the development of
gentle, west-trending salt-cored anticlines. The main thrust belt
formed during the Late Cretaceous, when a landward propagating
thrust belt developed in response to epirogenic uplift. The Astrid
thrust belt is associated with a detachment that dips oceanward
between 1° and 3° and extends over 130 km in the transport direction.
Fig. 19. Conceptual cross-sections showing the range of salt involving compressional
structures ev ide nt in t he Astrid, DWFTB, Ga bon , (re drawn from Jackson et al., 2 008).
A) Thrusted low-amplitude salt anticline, B) sm all bu ried diapir with an overhanging
bulb, and C) precursor diapir with an allochth onous shee t and pinched-off feede r. The
growth stages are described in the right margin.
59C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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The importance of pre-existing structures is emphasized in this thrust
belt because the salt layer was relatively thin at the time of
contraction, so pre-existing salt diapirs and gentle anticlines influ-
enced the contractional structures (Fig. 19), resulting in a regular
spacing and strike to the thrusts, and curvature of thrust tips to link
with pre-existing diapirs. Late expulsion of salt from buried anticlines
caused folding of overlying thrusts in some places. There is a strong
seaward vergence due to the commencement of thrusting after
welding had occurred along the basal salt. Like the Kwanza Basin, the
landward edge of the thrust belt propagated landwards, probably due
to buttressing down dip by the overburden. During thrusting, early
passive diapirs were squeezed and extruded to form small allochtho-
nous sheets (Fig. 19 C). The propagation of shortening up-dip towards
the shelf into previously extensional domains resulted in inversion of
older normal faults.
4.1.4.3. Scotia Margin. The Scotia Margin is noteworthy in its different
behavior compared with the West African examples described above.
First, the salt is of Triassic age, deposited during the syn-rift stage
(Wade and MacLean, 1990), hence there is a strong influence of the
rift tilted fault block geometry on the location of folds, diapirs and
faults (Fig. 18B). Secondly, deformation was initially driven by delta
loading during the Middle and Upper Jurassic syn-rift stage (Ziegler,
1989). Most of the extension in some places has been accommodated
by extrusion of a salt nappe at the sea floor, with little development of
a contractional zone (Ings and Shimeld, 2006). However, an up dip
extensional system is balanced by a broad zone of down-slope
contraction in the Sable sub-basin, in a region between a major salt
canopy system to the east, and isolated salt diapirs to the west
(Fig. 18B, Deptuck et al., 2009).
4.1.4.4. Northern Gulf of Mexico. The Gulf of Mexico is a very large,
complex and very well documented near-field stress-driven system.
In this kind of review it is impossible to do justice to the depth of
information and understanding that exists for the Gulf, instead some
highly generalized comments must suffice.
The Gulf is an unusually long-lived gravity driven system, in terms
of size, longevity, mixture of salt and shale mobile units and resulting
structural complexity, it can be regarded as the most complex end-
member of the near-field stress systems. The uniqueness of the Gulf
lies in its tectonic setting: the basin opened by continental extension
from ~160 to 150 Ma, followed by a short period of seafloor spreading
from 150 to 140 Ma and then became inactive (see review in Bird et
al., 2005). Seawater influxes during the syn-rift stage permitted
deposition of the massive Middle Jurassic Louann Salt (e.g. Peel et al.,
1995). The Northern Gulf of Mexico contains a great thickness of
section that excee ds 10 km in places and both salt and shale
detachments are present. Underlying the system is the oceanward-
dipping the Louann Salt basal detachment (Fig. 20).
Depositional loading during the Cenozoic triggered the main
episodes of gravity deformation i n the Gulf. Changing pa tterns of
deposition in the Gulf have been identified in a num ber of studies
as reviewed by Galloway et al. (2000), and show that the southern
Rocky Mountains were the main sedim ent sourc e area during the
earlyPalaeogene, and the Houston embayment was the main
depocentre (Fig. 21). Uplift in northern Mexico of the Sierra
Madre and Trans-Pecos in Texas increased sedimentation in the Rio
Grande embaym ent sli ghtl y later (upper Wilc ox). The majo r
depocentre shifted between the Houston and Rio Grande embay-
ments until the Middle Miocene, after which time drainage sh ifted
eastwards to the Mississippi Embaym ent as a result of uplift of t he
Western Interior.
The deepwater area of the Northern Gulf of Mexico is critical to
sustained large-scale hydrocarbon production from the region.
DWFTB plays are only one of three main exploration plays in the
deepwater province, the other two being the Flex Trend and Mini-
Basin trends (Cossey, 2004). Deepwater production as a percentage of
total Gulf of Mexico production has risen from ~6% Oil (21 MMSTBO
and 0.8% gas in 1985 to 73% oil (308 MMSTBO) and 43% gas (~1 TCFG)
in 2008. According to Nixon et al. (2009), over 141 producing projects
existed at the end of 2008 in the deepwater Gulf of Mexico.
There are three main DWFTBs in the Northern Gulf of Mexico, the
Perdido, Mississippi Fan and Port Isabel fold and thrust belts (Fig. 21).
Reviews of the deepwater exploration in these trends are provided by
Meyer et al. (2005, 2007). A cluster of deepwater hydrocarbon
development projects (BAHA, 2002; Trident, 2001; Great White, 2002;
Silvertip, Tobago) exist in the Perdido Fold Belt, but more numerous
development projects are associated with the Mississippi Fan Fold Belt.
Major discoveries in the Mississippi Fan Fold Belt include Neptune
(1995), Mad Dog (1999), Tahiti (2002), Knotty Head (2005), Genghis
Khan (2005) and Big Foot (2005). Thunder Horse (1999) is the largest of
the DWFTB fields, with an estimated mean size of 1 billion barrels of oil
equivalent (BBOE) (Cossey, 2004). InMarch 2009, the field was ramping
up to produce 300,000 BOE per day from only seven wells. It is the
largest producing deepwater asset in the world (http://www.bp.com/
genericarticle.do?categoryId=9004519&contentId=7009088, http://
www.businessweek.com/magazine/content/09_37/b4146000578301.
htm). Conversely, the Port Isabel Fold Belt has no discoveries (Cossey,
2004).
The northwestern Gulf of Mexico contains the deepwater Perdido
fold and thrust belt. This belt is, however, just the external portion of a
more extensive belt of contraction that includes the Port Isabel fold
Fig. 20. Cross-section across the northwestern Gulf of Mexico, redrawn from Peel et al. (1995) and Trudgill et al. (1999). Gravity-driven deformation occurs along a complex system
of salt and shale detachments.
60 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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belt and the salt canopy area (Fig. 20; Peel et al., 1995; Trudgill et al.,
1999). Up-dip extension of 60+ km suggests the down-dip fold and
thrust belts, and the salt canopy must accommodate a similar amount
of contraction, of which the Perdido fold belt only accommodates 5–
10 km (Peel et al., 1995; Trudgill et al., 1999). The Perdido fold belt
mainly underwent deformation during the lower Oligocene, but
minor deformation continued until the Middle Miocene (Trudgill et
al., 1999). In map view, the folds range from linear, to more dome
shaped, fold length appears to be limited in places by underlying NW–
SE trending syn-rift basement structures (Trudgill et al., 1999). The
fold belt underwent a late oceanwards tilting as a result of a wedge of
salt being injected into, and inflating, the basal detachment (Fig. 22;
Trudgill et al., 1999). A complex system of extensional detachments
exists up-dip, with detachments forming at different levels partly in
response to large salt sheets migrating up through different
stratigraphic levels over time (e.g. Diegel et al., 1995).
The switch in depositional systems with time res ulted in
abandonment of the Perdido fold belt, and formation of the
Mississippi Fan fold belt during the Miocene–Pliocene (Wu et al.,
1990; Weimer and Buffler, 1992). The two fold belts are separated by
a wide region without demonstrable contractional deformation
(Diegel et al., 1995). The Mississippi Fan Fold Belt is characterized
by broad, upright folds to asymmetric, offshore and landward verging
folds, cut by thrusts with wavelengths of about 8–12 km. They are
salt-cored detachment anticlines, with welds at the base of the
synclines (Rowan et al., 2004; Grando and McClay, 2004; Fig. 23).
Precursor, small wavelength salt pillows formed in places during the
Late Jurassic–Cretaceous in response to early contractional deforma-
tion, and later helped control the Cenozoic age folding (Grando and
McClay, 2004). The fold belt is partially overlain by the Sigsbee Salt
Nappe, whose emplacement is responsible for the withdrawal of salt
along the basal detachment of the Mississippi Fan Fold Belt. This
situation contrasts with the thick salt detachment of the Perdido Fold
Belt (Fig. 22; Trudgill et al., 1999).
4.1.4.5. Nile Delta. The eastern Mediterranean in the Levant region of
the African Plate underwent riftin g duri ng the early Mesozoic
(Gardosh and Druckman, 2006). Extension probably ceased prior to
the creation of oceanic crust, but was accompanied by extensive
volcanism. Subsequently, during the late Cretaceous to Miocene
several phases of inversion affected the older normal faults giving rise
to the Syrian Arc fold belt. The Messinian evaporites that form the
upper detachment in the deepwater Niger Delta arose as a result of
isolation of the Mediterranean basin during plate convergence
between African and Europe (see review by Jolivet et al., 2006). At
present the African Plate is being subducted northwards along the
Cyprus Arc or undergoing sinistral motion along that trend (Cavazza
et al., 2004; Robertson and Mountrakis, 2006).
The Nile Delta is a large delta that was activated during the Late
Miocene (Tortonian) in conjunction with uplift of the Red Sea-Gulf of
Suez rift shoulders (Guiraud et al., 2001) and is associated with
deformation along two main offshore-dipping detachments. The
deeper detachment occurs in shales, while the upper one is in
Messinian evaporites (Aal et al., 2000; Dolson et al., 2000; Tingay et
Salt Dome-Mini Basin
Province
Mini B
as
in Province
Sa
lt Dome
Wilcox F
au
l
t
S
ystem
s
Upper Eoce
ne
F
au
lt
S
ystem
s
0
300 km
-200 m
-3000 m
-200 m
-3
000 m
Sig
sb
ee
E
scarpment
Mississippi Fan
Fold Belt
Perdido
Fold Belt
Port Isabel
Fold Belt
Mexican
Ridges
Fold Belt
MEXICO
USA
Houston
New Orleans
95°W 90°W
30°N
25°N
US Waters
Mexican Waters
Little Tertiary
extension or
compression
Plio-Pleistocene Detachment Province
Oligo-Miocene Deta
chment Province
Vick
sbu
rg Det
a
chment
Tabular Salt Mini Basin Province
Local collision of salt-
withdrawal basins
Fig. 21. Map of the Gulf of Mexico showing main structural units and water depths, based on maps in Trudgill et al. (1999) and Ambrose et al. (2005).
61C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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al., 2010,in press). This stacking sequence is due to the location of the
delta on a passive margin situated on a northward subducting lower
plate. This convergent setting resulted in the deposition of evaporites
during a late stage of passive margin development as the basin began
to close, as opposed to the classic passive margin sequences seen on
the Atlantic margins, where halites were deposited during basin
opening. Up-dip extensional faults and down-dip squeezed diapirs
and folds are associated with the Messinian evaporite detachment
(Loncke et al., 2006). These authors show the whole systems is about
200–250 km long in the transport direction, with the compressional
zone being about 50–100 km wide. The structures display a wide
range of orientations particularly in the extensional province, which
reflect the interplay between pre-Messinian topography and the
distribution of Messinian evaporites, varying slope gradients at the
base of the salt layer, and major features such as the Eratosthenes
Seamount in the northeastern part of the deepwater delta (Loncke et
al. (2006).
4.2. Type 2, continental convergent tectonic setting
Continental convergent zones contain DWFTBs affected by both
near- and far-field stress systems (type 2a), and type 2b DWFTBs, which
are entirely or predominantly affected by far-field stress systems. The
island of Borneo is a remarkable area for deepwater exploration,
encompassing five different deltaic provinces with deepwater potential
or proven fields that formed in a variety of Type 2 DWFTB settings. These
provinces are the West Luconia, Sandakan, Tarakan and Mahakam
deltas, and the NW Borneo Trough (Brunei–Sabah) (Fig. 4). In addition,
the west Sulawesi Fold Belt lies opposite the Mahakam Delta on the west
Sulawesi margin in the Makassar Straits. Of these provinces the NW
Borneo Trough, Mahakam and Tarakan deltas all have proven
deepwater hydrocarbon fields. The proliferation of deepwater prospects
in the area is due to the uplift of Borneo during the Miocene, which
resulted in rapid erosion of the uplifting island and rapid deposition
along the fringes (e.g. Hall and Nichols, 2002).
Sea floor
4.0
5.0
6.0
7.0
8.0
5 km
Late Miocene Unconformity
Top Cretaceous
Lower Cenozoic Wilcox
NW SE
Louann Salt
Salt Canopy
TWT (s.)
Fig. 23. Section across the Mississippi Fan Fold Belt, based on seismic line in Meyer et al. (2005).
Fig. 22. Cross-section illustrating the development of the Perdido fold belt, Gulf of Mexico, by regional gravity spreading linked to up-dip extension. A) Paleogene initial loading and
extension caused primary extrusion of salt and the development of a regional allochthonous salt canopy. B) Closure of the feeder salt diapirs transferred stress to the toe of the
system, causing development of the Perdido fold belt. C) Further salt movement resulted in inflation of a salt wedge along the Perdido fold belt detachment. Complex deformation in
the Paleogene salt canopy is not represented. Redrawn from Trudgill et al. (1999).
62 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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Complex, early-stage continent–arc collision is ongoing in the
Banda Arc area, east of Borneo, where the Australian Plate is colliding
with the arc system at the eastern end of the Sumatra–Java subduction
zone (Fig. 24). This area has reached the stage where the deepwater
thrust and fold belts are starting to be uplifted and eroded, and the
deepwater troughs are being reduced in extent and depth due to
continued convergence and crustal thickening. Smaller, DWFTBs exist
in the region too, and appear to be associated with strike-slip
deformation arising from collision (e.g. the Banggai-Sulu and
Cenderwasih DWFTBs, Ferdian et al., 2010; Sapiie et al., 2010;
Fig. 3). The small size of the Cenderwasih DWFTB is particularly
remarkable, it has a (NE–SW) strike-length of only about 100 km, a
width of 30 –40 km and lies in water depths up to 2000 m. The DWFTB
is developed in a zone of intracontinental convergence where a
Neogene foredeep basin overlying an older Permo-Triassic syn-rift
and Mesozoic post-rift sequence was folded and thrusted during the
Late Miocene–Holocene (Sapiie et al., 2010).
In this section, the DWFTBs associated with Borneo and Sulawesi
are discussed; these regions may, however, differ from classic collision
zones. These DWFTBs lie within a region of hot, weak continental
crust accreted during the Late Palaeozoic–Cenozoic that is unlikely to
deform in the typical style of cold, older continental crust (Hall and
Morley, 2004). Hall (in press) observes that, for Borneo and Sulawesi,
there was initial collision and shortening, but that later uplift and
erosion of the island interior was not accompanied by sufficient
shortening to drive the observed uplift. Consequently Hall (in press)
links sediment accumulation in basins fringing the island to uplift of
its interior in response to crustal-scale gravity-driven processes. There
appears to be a feedback mechanism between sediment loading in the
fringing basins, and uplift of the island interior possibly involving the
flow of hot, ductile crust from beneath the basins towards the
uplifting island interior, which acts as the sediment source areas.
Deformation in the DWFTBs can be caused both by near-field gravity
processes limited to the sedimentary basin triggered by sediment
Water depths between
sea level and -500 m
Water depths greater
than -4000 m
17
100
77
32
77
Wetar Thrust
Timor Trough
Tanimbar Trough
Seram Tr
ough
126
Aru
Islands
Tanimba
Haimahera
Misool
Ambon
Seram
Timor
Bird’s Head
Sulaweisi
Molucca
Sea
Pacific Ocean
Timor
Sea
Arafura
Sea
Australia
126°E122°E
130°E
134°E
0°
4°
8°
12°
Irian Jaya
Thrust front of deepwater
thrust wedge (early collisional thrust
wedge or accretionary prism)
Arc volcanics
GPS defined plate convergence
rates in cm/yr
-2000 m sea
floor contour
-4000 m sea
floor contour
Banggai-Sula
Platform
Cenderawasih
Bay
Banda Arc
Fig. 24. Regional tectonic map of eastern Indonesia, Banda Arc area (partly based on Hall and Wilson, 2000; Sapin et al., 2009 and Ferdian et al., 2010).
63C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
Author's personal copy
loading of an overpressured mobile shale (Setap Shale), and by deeper
crustal and lithospheric far-field stresses. This crustal-scale coupled
sediment loading-uplift model is similar in principal to the model for
post-rift subsidence proposed by Morley and Westaway (2006) for
another hot, weak part of the Sundaland crust in the Gulf of Thailand.
The model advanced by Hall (in press) needs further evaluation but
represents an interesting possibility. Because this model requires
collisional process to initiate the orogenic zones and subsequent flow
of ductile crust, these processes can be regarded as deep seated in
origin. Consequently these regions remain grouped with convergent
zones, and far-field stresses. Nevertheless, the model contains distinct
differences with classic orogenic belts, particularly as the later stage
development does not require plate boundary forces to drive the
deformation.
The quality of clastic reservoir rock is strongly related to the tectonic
setting and is generally a much higher risk for active margins and
collision zones than it is for passive margins (e.g. Dickinson and Suczek,
1979). For this reason the regionalsetting for sedimentation is discussed
for Type 2 systems, while it was ignored for Type 1 systems.
4.2.1. Type 2a convergent zones
4.2.1.1. Baram Delta Province and Sabah Margin, NW Borneo. The Sabah
margin, NW Borneo is marked by a NW dipping slope that passes into
a deepwater area more than 3000 m deep that marks the NW Borneo
trough. There is little earthquake or volcanic activity, so despite the
trench-like morphology, the present day setting is not one of active
subduction. The presence of a train of deepwater folds along the NW
dipping slope was established from 2D data in the 1980's (e.g. James,
1984; Hinz et al., 1989). The fold belt is about 300 km long and 50–
80 km wide and runs from the Baram Delta Province of the Brunei and
northern Sarawak offshore area in the south, along the Sabah margin
towards Palawan (Fig. 25). The NW Borneo shelf lies to the SE of the
DWFTB and can be divided into two provinces: 1) the Baram Delta
Province in the SW (locations 2 and 3, Fig. 25), and 2) the Sabah Basin
in the NE (locations 4 and 5, Fig. 25). The two basins are transitional to
each other and share the same characteristics, but to differing degrees.
The Baram Delta Province is dominated by a thick shelfal deposits
affected by classic gravity-driven structures (overpressured shales,
growth faults, mud pipes). Conversely the shelfal area of the Sabah
Basin, while preserving some thick depocentres, growth faults and
overpressured shales, is widely deformed by thick-skinned thrusts
and strike-slip faults that have caused multiple unconformities,
inversion structures, folding, uplift and erosion (e.g. Levell, 1987;
Tan and Lamy, 1990; Madon et al., 1999).
During the early Cenozoic, the Proto-South China Sea oceanic crust
was subducted to the SSE below NW Borneo, and the Cenozoic
Crocker Formation developed in an accretionary prism setting (e.g.
James, 1984; Levell, 1987; Hutchison, 1996; Sandal, 1996; Hall, 2002).
Jamming of the subduction zone by the entry of thinned Eurasian
continental crust (Dangerous Grounds) during the early Miocene
terminated subduction and caused uplift and erosion of the area that
1
2
4
5
+
+
+
Sarawak
Sarawak
Brunei
Sabah
NW Borneo Trough
200 m
1000 m
2000
m
2500 m
Offshore
Brunei
Darussalam
South China
Seas
σ1 orientations
from borehole breakouts
A quality
B quality
C quality
D quality
Zone of uplift
including Kinabalu
culmination
Melinau Fm
(Eocene-Oligocene)
113 ° 114 ° 115° 116°
Upper Cretaceous-Eocene
tectonized ophiolite with
remobilized serpentinite
Paleogene Rajang
Group flysch (on oceanic crust?)
Neogene
igneous rocks
Undifferentiated Paleogene Miri Zone,
(on continent basement?)
Belait, Liang and Setap
Shale Formations
(Miocene-Holocene)
Meligan Sandstone
Formation (Oligocene-
Early Miocene)
0
40 km
Normal
fault
Thrust
fault
Syncline
Bathymetry
Anticline
Brunei
B
n
un
B
1 = Deepwater fold and thrust belt
2 = Extensional growth faults in
Baram Delta Province
3 = Inverted zone of Baram
Delta Province
4 = Mobile shale deformation province,
Sabah Basin
5 = Zone of strong inversion,
Sabah Basin
Mt Kinabalu
granite
3
+
7°
6°
5°
4°
Fig. 25. Regional tectonic map of NW Borneo (modified from Sandal, 1996; Morley et al., 2008; King et al., in press).
64 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
Author's personal copy
comprises present-day Borneo (e.g. Levell, 1987; Hutchison et al.,
2000). Middle Miocene–Late Miocene shortening has been largely
confined to the onshore area, with episodic inversion of offshore
structures occurring along the NW Borneo shelf of Brunei and Sabah
(Levell, 1987; Hutchison et al., 2000; Morley et al., 2003). The
progressive uplift and erosion of NW Borneo from the latest Early
Miocene onwards resulted in the rapid deposition of thick deltaic
sequences up to 10 km thick around the margin of Borneo, notably the
West Luconia, Baram, Sandakan, Tarakan and Mahakam deltas (e.g.
Hall and Nichols, 2002; Fig. 4).
The most prominent manifestation of Miocene orogenesis in Sabah is
Mount Kinabalu, which is formed from a Late Miocene granodiorite
sheet (Cottam et al., 2008, 2010). The summit rises up to 4100 m, and
was glaciated during the Pliocene. The granodiorite was crystallized
between 7.9 and 8.2 Ma and rapidly exhumed by ~4–8 km (Hutchison
et al., 2000; Cottam et al., 2008). The granite is interpreted as a product
of collision-related thickening and its exhumation could indicate deep
lithospheric processes, such as loss of a lithospheric mantle root (Cottam
et al., 2008; Hall, 2009). These authors suggest uplift would have had a
profound effect on the Sabah margin and induced gravity-driven folding
and thrusting in the DWFTB. Alternatively the granite could be a product
of slab breakoff (e.g. Von Blanckenburg and Davies, 1995), which would
also explain the late-stage uplift of Borneo (e.g. Hutchison et al., 2000;
Morley and Back, 2008).
The Baram Delta Province built out onto the NW Borneo margin
from the latest early Miocene to the present day (e.g. Sandal, 1996;
Hall and Nichols, 2002; Morley and Back, 2008). The primary source of
the deltaic sediments is the hilly spine of NW Borneo. The region is
largely composed of the Sapulut (Late Cretaceous–early Cenozoic),
Trusmadi and Crocker (Cenozoic) Formations, which are mostly
deepwater sediments, probably of forearc origin (Hutchison, 1996;
van Hattu m et al., 2003). The Crocker Formation sands were
predominantly sourced from Cretaceous granites located in southern
Borneo and the Sunda Shelf, although a minor ophiolite-derived
component is also present (van Hattum et al., 2006). The Sapulut and
Trusmadi Formations are compositionally more mature (quartzose
recycled) (van Hattum et al., 2003). It is important to the hydrocarbon
potential of the region that these formations provided quartz-rich
Champion Delta
Province
Baram Delta
Province
Deepwater fold and thrust belt
6 km+
depocentre
Frigate
Fault
Belait
syncline
SABAH
SARAWAK
BRUNEI
Mt. Mulu
Oligocene-Miocene
Setap Shale Fmn.
and Temburong Fmn.
Melinau Limestone
(late Eocene-
Early Miocene).
Eocene Mulu and
Kelalan Formations
Early Miocene
Meligan Fmn.
N
0 25 km
Approximate
position of deep
regional unconformity
Labuan
Island
Belait
anticline
Zone 1
Zone 2
Zone 2
Zone 3
Zone 4
Seafloor high
XI
IX
Shelf edge
Fig. 26. Map of the structural zones of Brunei, partly based on Sandal (1996), Morley et al. (2003).
65C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
Author's personal copy
(albeit predominantly fine-grained) sands to the Baram Deltaic
Province (Sandal, 1996).
The NW Borneo margin can be interpreted both in terms of a
tectonically active margin and gravity tectonics (e.g. Franke et al.,
2008; Morley et al., 2008), particularly in the Baram Delta Province
(Brunei, SW Sabah, and the northeastern Sarawak). The onshore
Baram Delta Province is characterized by basement-involved defor-
mation that affects the sediments of the early (Middle Miocene) stage
of deltaic deposition. Seismic reflection data from onshore and in the
near offshore area shows well developed folds related to inverted
normal faults (e.g. Levell, 1987; Sandal, 1996; Morley et al., 2003,
2008; Fig. 26). These folds account for up to 3 km shortening as
calculated by Hesse et al. (2008). This proximal zone of inversion is
not seen in the shallow gravity-driven systems, and reflects the
impact of deep-seated stress affecting NW Borneo (Morley et al.,
2003; Tingay et al., 2005).
The structural style of the NW Borneo DWFTB is predominantly
landward dipping imbricate thrusts, with hangingwall folds (Fig. 27).
The imbricate thrusts lie within a critical taper wedge with a surface
slope that ranges between 1° and 2.5° and a basal detachment that
dips 2° to 5° (Morley, 2007a). The wedge thickens from about 3 km
near the thrust front to 10 km near the shelf-slope break. The most
external folds exhibit classic imbricate fault splays off the basal
detachment, however, the fault geometry at depth becomes unclear
passing into the thicker part of the wedge the fault geometry (Cullen,
2010). The folds have a strong expression at the sea floor and
influence gravity-driven sediment pathways down the slope (McGilv-
ery and Cook, 2003; Morley, 2009a; Hesse et al., 2010). The results of
NW
SE
5 km
TWT (s.)
Possible mobile shale zone
(poorly imaged region on
seismic data)
Lower pre-kinematic sequence
(Miocene?)
Upper pre-kinematic sequence
(Miocene?)
Early syn-kinematic sequence,
Late Miocene (?)
Mid syn-kinematic sequence,
Pliocene (?)
Late syn-kinematic sequence,
Pliocene - Holocene
Region of shale diapirs and mini basins,
diapirs squeezed by later compression (?)
Deepwater fold and thrust belt
Lower Plate of Dangerous Grounds
continental crust
Mantle
Kinabalu
Culmination
Mt. Kinabalu
Offshore Onshore
Suture zone and
detachment zone
km
Pliocene-Holocene sediments
Miocene sediments above the
Deep Regional Unconformity
West Crocker Formation
East Crocker/Rajang Formation
LC
UC
South China Sea
Eurasian Plate
Forearc
Crust
Telupid
High
Sabah
Deepwater
Fold Belt
Shelf
Neogene Basin
Ductile wedge of Eurasian Plate
crustal rocks
Oceanic slab broken off?
NW SE
Kinabalu Granite, and migmaties
Miocene melange
+
+
+
+
+
+
+
+
+
+
+
++
+
+
+
+
+
+
+
+
+
+
B
500 ms
10 km
NW SE
10 km
0
1 km
Flow of lower crust?
0
10
20
30
40
50
0
8
A
B
C
D
NW SE
Fig. 27. Structural cross-sections across NW Borneo. A) Regional crustal scale transect through Sabah partly based on Hall and Wilson (2000), B) Detailed cross-sction through Sabah
DWFTB, based on 2D BGR seismic line illustrated in Hesse et al. (2008); see inset in transect. C) Cross-section across Brunei DWFTB based on 3D seismic line illustrated in McGilvery
and Cook (2003).
66 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
Author's personal copy
piston coring in areas of mud volcanoes indicate the larger ones are
expelling hot fluids, which cont ain ther mogenic hydrocarbons
(Zielinski et al., 2007; Warren et al., in press; Fig. 9). These data
support seismic reflection surveys that indicates the mud volcanoes
are sourced from deep, overpressured fluids, and travel along thrust
faults before breaking vertically to the surface in the cores of
anticlines (Fig. 8; Morley, 2009a,b ). The Sabah margin has proven to
be a successful DWFTB play including the following discoveries: Kikeh
(~1300 m, 400–700 million barrels of oil reserves), Gumusut-Kakap
(~1300 m water depth, 2.2 trillion cubic feet (TCF) of gas reserve),
Limbayong (0.8 TCF gas reserve, 1000 m water depth), Malikai
(~565 m water depth) Kebabangan and Kamunsu fields (e.g. Ingram
et al., 2004; Milton, 2006).
The Baram Delta province shows well-developed extensional
growth faults in many places. However, the zone of currently active
extension is narrow (20–30 km), lacks well developed shale “diapirs”,
and lies between a region of folds associated with inverted growth
faults in the near shore and onshore area, and the deepwater fold and
thrust belt (Fig. 26). This arrangement of inverted zone-extensional
zone-DWFTB is also seen in the modern stress pattern, where the
onshore and inner shelf display maximum horizontal stress (Shmax)
orientations approximately sub-perpendicular to the coastline. The
maximum horizontal stress directions for the outer shelf and upper
slope (zone of active extensional faulting) rotate to lie sub-parallel to
the coastline, then in the DWFTB area, Shmax orientations revert to
sub-perpendicular to the coast (e.g. Tingay et al., 2005; King et al.,
2009; Fig. 25).
A newly identified mobile shale-associated region (Shale Prov-
ince) in northern offshore Sabah displays approximately N–S Shmax
orientations (King et al., in press; Figs. 25 and 27). This province is a
region of Late Miocene–Pliocene mobile shale deformation charac-
terized by “diapirs” and synclinal mini-basins, that were later
squeezed and folded (Fig. 27B). King et al. (in press) suggest the
clockwise rotation of stresses within this province with respect to
wells in adjacent areas can be explained by contrasting geomecha-
nical properties of a soft, thickened mobile shale and surrounding
stiffer sediments of the DWFTB.
An effect of the exhumation of Mt. Kinabalu would have been the
latest Miocene and Pliocene transport of large amounts of sediment
onto the Sabah margin. Figure 25 shows that the Shale Province wraps
around the northern end of the Kinabalu Culmination and deposition
within the province coincides with the exhumation of Mt. Kinabalu.
Hence differential loading of the DWFTB may have been caused by
rapid deposition in the Shale Province as sediment supply increased
due to uplift of Mt Kinabalu.
Total shortening across the NW Borneo DWFTB is ~8–13 km. The
extent of the DWFTB and amount of shortening does not appear to be
strongly related to the distribution of extension on the shelf (Hesse
et al., 2008; Morley et al., 2008). Line-length shortening (which may
considerably underestimate shortening by compaction) is generally
equal to, or greater than extension on the shelf (maximum ~6 km
extension opposite 6 km contraction for the late Pliocene-Recent;
Hesse et al. 2008; King et al., 2010; Cullen, 2010). To the NE from
offshore Brunei to offshore Sabah the discrepancy between the
amounts of shortening in the DWFTB compared with extension on the
shelf increases (Hesse et al., 2008). The discrepancies between
shortening and extension have been interpreted as demonstrating a
tectonic component to past and ongoing deformation in the DWFTB
(e.g. Hesse et al., 2008; King et al., 2010).
4.2.1.2. Trinidad (deepwater Columbus Basin). The DWFTB in Trinidad
is not well described in the literature, but because examples of mixed
near- and far-field stress systems are few, as are examples of DWFTBs
associated with strike-slip margins, the deepwater Columbus Basin is
briefly discussed in this section to help fill this gap. The paper by
Garciacaro et al., (in press) provides an excellent overview of the
deepwater setting of this basin.
The Trinidad–Barbados region formed as a result of subduction of
the North and South American plates beneath the Caribbean Plate
since the middle Eocene (Pindell and Kennan, 2007). Eastwards
movement of the Caribbean Plate with respect to South American
presently occurs at a rate of ~20 mm/yr along much of their plate
boundary (Perez et al., 2001; Weber et al., 2001; Saleh et al., 2004).
The El Pilar Fault and associated strike-slip faults are the focus of
eastwards translation of the Caribbean Plate with respect to South
America (Fig. 28). The strike-slip faults pass into the complex,
extension to transtension-dominated province of the Columbus
Basin east of Trinidad. Evolution of the Caribbean–South America
collision zone of eastern Venezuela–Trinidad–Barbados appears to
have changed from a more transpressional structural style prior to
10–12 Ma to a transcurrent style, which has continued to the present
day ( Pindell and Kennan, 2007).
Trinidad is well known for its offshore hydrocarbon fields, and is
extensively affected by both thrusting and folding, and normal faults
(Wood, 2000; Escalona et al., 2008; Fig. 28). Yet exploration in the
DWFTB has so far not achieved commercial success (Rajnauth et al.,
2004). Overpressured mud pipes and mud volcanoes are widespread
and the predominantly thermogenic gas and clasts found in the fluids
from mud volcanoes indicates a deep, overpressured shale origin,
most likely in Cretaceous and/or Paleogene formations (Deville et al.,
2003). The Columbus Basin displays a number of similar elements to
the Niger Delta. It is in large-parts a gravity-driven system associated
with rapid loading of overpressured marine shales by the Orinoco
delta. It displays large growth fault systems and shale ridges at the toe
of the delta on the slope (Wood, 2000; Wood and Mize-Spansky,
2009; Fig. 28). The details of how the near-field and far-field stresses
interact in Trinidad remain to be assessed. They have the potential to
be complex since both gravity-driven deformation and far-field
stress-related major strike-slip faults and thrusts affect the area
(Truempy et al., 2004; Garciacaro et al., 2011).
4.2.2. Type 2bi weakly linked/unlinked DWFTBs
4.2.2.1. Makassar Straits. The Makassar Straits lie between the islands
of Borneo and Sulawesi. West of Sulawesi an extensive deepwater fold
and thrust belt is present, in an area whose geological setting
has proven controversial. North of the Makassar Straits lies the
Celebes Sea, which is floored by oceanic crust generated during an
Eocene phase of seafloor spreading (see review in Hall, 2002). While
there is general agreement that the southern Makassar Straits are
underlain by continental crust, determining how far oceanic crust
formation propagated southward into the northern Makassar Straits
has proved controversial (e.g. Bergman et al., 1996; Hall, 2002). The
latest review of the data indicates that the northern Makassar Straits
also are underlain by highly thinned continental crust (Hall et al.,
2009; Fig. 5).
The Cenozoic tectonic history of Sulawesi is also controversial. A
very brief outline of the tectonics of this highly complex area is given
as background to the setting of the Makassar Straits. Lying on the
eastern margin of the continental core of SE Asia, the island has been
subject to a succession of orogenic events which include the accretion
of continental fragments derived from Australia, subduction, intense
volcanic activity, rifting and ophiolite obduction during the Oligo-
cene–Miocene (see reviews in Bergman et al., 1996; Hall, 2002).
Central and eastern Sulawesi is home to the world's third largest
ophiolite (East Sulawesi Ophiolite; Fig. 5), which is a 15 km thick slab
of oceanic plateau crust of possible Cretaceous origin (Kadarusman
et al., 2004). According to Parkinson (1998) obduction occurred
during the Late Oligocen e. However, there are oth er ophiolite
fragments present on the island with different geochemistry, hence
the Sulawesi ophiolites are likely to be composite and were emplaced
67C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
Author's personal copy
during more than one tectonic event (Hall, 2002). Obduction of the
East Sulawesi Ophiolite is commonly assumed to have been followed
by Late Oligocene–Early Miocene continent–continent collision of an
Australian craton-derived block (eastern Sulawesi) with western
Sulawesi, which resulted in fold and thrust belt development and
extensive lithospheric melting-related magmatism (see reviews in
Bergman et a l., 199 6, and Hall, 2002). Marine sedimentation
continued uninterrupted by contractional deformation in western
Sulawesi and the Makassar Straits throughout the Miocene (Hall,
2002). Deformation of the DWFTB in the Makassar Straits dates from
the Early Pliocene (Calvert and Hall, 2007).
The DWFTB is part of a dominantly east-vergent system of folds
and thrusts that affect western onshore Sulawesi and the eastern
Makassar Straits. This DWFTB is a system of detached folds and thrusts
(Fig. 5; Bergman et al., 1996; Calvert and Hall, 2007). However,
beyond the thin-skinned thrust front, and beneath the thin-skinned
folds, seismic reflection data shows evidence for inversion of older rift
basins, and some seismic lines indicate that the DWFTB may be linked
A
A
A’
A’
N
St Lucia
GRENADA
BASIN
Grenadines
Grenada
Tobago
Trinidad
COLUMBUS
BASIN
BARBADOS RIDGE
BARBADOS BASIN
BARBADOS CREST
Barbados
TOBAGO BASIN
0 50 100 km
Mud volcanoes
Normal faults Strike-slip faults
Thrust faults
ORINOCO
DELTA
60° 00’ 59° 00’ 58° 00’61° 00’62° 00’
13°00’
12°00’
11°00’
10°00’
El Pilar Fault
20 km
Northern basin
Darien Ridge
Fold-thrust belt
Columbus Foreland Basin
Quaternary Neogene Palaeogene Cretaceous
0
2
4
6
8
10
12
Two-way travel time (s.)
Seafloor
+
+ +
NW SE
?
?
A
B
Fig. 28. A) Regional tectonic map of the Trinidad, Columbus Basin, Barbados area (compiled from Deville et al., 2006). B) Line drawing of interpretation from seismic reflection data,
across the western part of the Barbados accretionary prism, modified from Garciacaro et al., (2011), see A) for location.
68 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
Author's personal copy
with, or affected by movement on basement-involved fault systems.
The timing of deformation, and deep-seated nature of the stresses
affecting the DWFTB deformation in the Makassar Straits is similar to
DWFTB deformation in NW Borneo, and in also the Mahakam Delta
area.
Most of the thrusting and folding in the Mahakam Delta is located
onshore or on the shelf, not in deepwater (Bergman et al., 1996;
Ferguson and McClay, 1997), but is a manifestation of the regional far-
field stresses that affect both the onshore areas and the DWFTBs. The
well established NW–SE present-day maximum horizontal stress
direction in the Mahakam delta area fits with a heavy influence from
far-field stresses (Tingay et al., 2010, in press) The West Sulawesi
DWFTB in the Makassar Straits is not associated with a significant
delta. The western margin of Sulawesi plunges abruptly into
deepwater, with virtually no shelf present, unlike the well-developed
shelf of NW Borneo (Fig. 29). Hence, there is no ambiguity regarding
the nature of the driving stress (near- or far-field) in the Makassar
Straits as discussed above for NW Borneo (Baram Delta Province and
Sabah Margin, NW Borneo). The absence of a delta and landward-
dipping basal detachment indicates there is no component of shallow
gravity driven deformation.
The absence of a shelf today offshore western Sulawesi is also in
strong contrast with the older Cenozoic history of the basin. Following
Eocene rifting, the Makassar Straits underwent post rift subsidence
(with extensional activ ity on some faults continui ng into the
Oligocene; Calvert and Hall, 2007, Hall et al., 2009). The centre of
the Makassar Straits is characterized by deepwater sedimentation
during the Neogene. However, Calvert and Hall (2007) describe
extensive shallow marine conditions as being present throughout the
Early Miocene onshore in NW Sulawesi, and persisting in places until
the Middle or Late Miocene. This former shelf area has been uplifted
by folding and thrusting, which was sufficiently young and rapid that
a new, broad shelf has yet to form.
The uplifted, predominantly ophiolitic and volcanic-sourced fold
and thrust belt to the east sourced Neogene sediments within the
DWFTB. Several exploration wells have been drilled in the shallow
water and onshore equivalent section of the DWFTB. The potential
reservoir rocks are fine-grained sandstones, with a large basic igneous
component. Diagenetic alteration of the igneous-derived clasts to
clays and the fine-grained nature of the deposits resulted in a low-
permeability, low-effective porosity reservoir. Hence no good quality
reservoir rock has been established to date for the DWFTB system. The
apparent lack of reservoir is the most significant risk to the presence
of an economic petroleum system associated with this DWFTB.
Another DWFTB complex lies north of the Banggai-Sula micro-
continent, between Sulawesi and the Banda Sea(Fig. 3). This
microcontinent probably collided with the eastern ophiolite and
northern volcanic arms of Sulawesi in the Early Miocene (Ferdian et
al., 2010). The area also underwent Early Pliocene and south-directed,
E–W trending dextral transpressional deformation that formed a
DWFTB as the Molucca Sea Collision Complex (Fig. 3) impinged on the
Sula Platform (Ferdian et al., 2010).
4.2.2.2. Banda Arc. The Banda Arc lies in eastern Indonesia and Timor-
Leste (Fig. 24). The region provides an example of the early stages of
continental margin-arc collision (Karig et al., 1987; see review in
Hall
and Wilson, 2000; Standley and Harris, 2009). The Banda Arc has a
pro longed and complex hi story of Cenozoic deformation, with
Standley and Harris (2009) reporting 5–6 deformation phases of
alternating extension and compression, that are associated with
predominantly top to the SE collisional deformation and top to the
south to SE extension. Deformation prior to about 25 Ma was largely
related to subduction of the Indian Ocean lithosphere at the Sunda
(Java) and Celebes Trenches. Subsequently, events testifying to
continued Australia–SE Asia convergence are the collision of the
Sula Spur with the North Sulawesi volcanic arc, ophiolite obduction in
Fig. 29. Comparison of topography across the Baram and Mahakam deltas, and west Sulawesi.
+
+
+
+
+
+
+
+
+
+
+
+
+
+
+
+
+
+
+
+
+
+
TIMOR TROUGH AUSTRALIAN PLATE
TIMOR
EURASIAN PLATE WETAR STRAIT
INNER BANDA ARC ONNER BANDA ARCBANDA ALLOCHTHON
Forearc crust
Permo-Triassic distal
Australian margin sediments
Post-rift slope and
rise deposits
Syn-orogenic deposits
Post-rift shelf deposits
Asian mantle
WETAR THRUST
WETAR
SUTURE
Australian
oceanic lithosphere
Cambrian-Carboniferous
pre-rift sediments
Australian mantle
Australian crystalline basement
Permo-Triassic
pre-rift sediments
0
50
100
Kilometres
NNW SSE
Fig. 30. Regional crustal-scale transect, illustrating the early-stage collision of the Australian Plate with the Timor island arc. Redrawn from Hall and Wilson (2000).
69C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
Author's personal copy
South Sulawesi, and broad counter-clockwise rotation of the Borneo–
Java region of Sundaland (Hall, in press). At about 15 Ma, the oceanic
Banda back-arc basin began to form as a result of subduction rollback
(see review by Hall, in press). Thinned crust of the Australian
continental margin was subducted beneath the Banda Arc, whilst the
attached oceanic lithosphere was broken off during the late Neogene
(Hall and Wilson, 2000; Fig. 30). The subduction channel may have
become clogged by Australian margin crust only about 3 my ago
(Standley and Harris, 2009). The deepwater troughs fringing the
Banda Arc such as the Timor Trough have evolved from a subduction-
trench setting to a collisional flexural (deepwater) foreland basin
(Carter et al., 1976; Audley-Charles, 1986; Ziegler et al., 1998).
Notably the greatest water depths are found not in the sites of the
former accretionary prisms, but in the back arc location (Fig. 24).
The upper plate of the collision zone includes the islands of Timor,
Seram, Tanimbar and Kei. Hydrocarbon seeps on a number of islands,
including Timor, have been observed, but only Seram is at present an
established hydrocarbon province (Charlton, 2004). The broad zones
of deformation comprise both thin-skinned thrusting, and thick-
skinned deformation of the basement, including inversion of older
normal faults; the degree of overthrusting is considerable in places.
For example, Permo-Triassic sedimentary rocks of the distal Austra-
lian margin observed on Timor have been thrusted onto Mesozoic–
Cenozoic sedimentary sequences of the Timor Trough flexural forearc
foreland basin. In turn the Australian margin sediments that crop out
in Timor have been overridden to the SSE by the Banda Arc forearc.
Oceanic lithosphere in west Seram was obducted over an Australian-
derived microplate of continental crust during the Miocene (Linthout
et al., 1996).
Extensional basins ranging in age from Permian to Paleogene may
contain excellent Mesozoic source rocks (Charlton, 2004). Hydrocar-
bon prospects are available in shallow marine to onshore environ-
ments in these inverted rift basins, as demonstrated by the Oseil Field
in Seram, where the Upper Triassic–Lower Jurassic Manusela
Limestone reservoir is sealed by the external, thin-skinned thrust
and mélange belt (Nilandaroe et al., 2002). In such a setting, large and
relatively simple structures may occur beneath the structurally
complex,
thin-skinned fold and thrust belt (Charlton, 2004; Pairault
et al., 2003; Sapin et al., 2009; Fig. 31) The most external and youngest
structures of the Timor, Tanimbar and Seram fold-and-thrust belt
(Fig. 24) are relatively simple hangingwall anticlines (Fig. 31), located
in water depths between 1500 and 2000+ m. However, in places,
such as the Onin and Kumawa domes east of Seram (Fig. 32) and the
Kei Islands compressional reactivation of deeper seated Mesozoic
extensional fault systems caused late stage uplift of the deepwater
fold and thrust belt to shallow waters, or subaerially, inducing its
erosion (Fig. 31).
Extensive diapirism, forming diapir fields up to 100 km long and
30 km wide, is observed in Australian passive margin sediments
before they enter the Timor accretionary complex (Breen et al., 1986;
Masson et al., 1991). These diapirs were interpreted by Barber and
Brown (1988) as resulting from overpressured fluids generated
initially by disequilbrium compaction, due to the overburden of the
accretionary complex, that were transferred along Jurassic breccias
and sandstones marking the break-up unconformity, to mobilize well-
consolidated Lower Cretaceous clays. The Lower Cretaceous clays
formed diapirs that were incorporated into the accretionary complex
during subduction. Outcrops in Timor also indicate that diapiric
1) Tortonian
2) Messinian
3) Late Pliocene
4) Present
South-western Bird’s Head Microplate Margin ??
Seram fold and thrust
belt front
Misool-Onin-Kumawa Ridge
Ductile deformation ??
High Bouger anomaly
on 250 km high-pass filter
Permian-Paleocene shales
New Guinea Limestone
Klasafet Formation
Steenkook Formation
Brittle/Ductile crust interface
Active fault
Inactive fault
10 km
SW NE
Fig. 31. Conceptual evolutionary diagram of the Seram fold and thrust belt (redrawn from Sapin et al., 2009). For location see Fig. 32.
70 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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mélange is a very important component of the Neogene accretionary
prism (Barber et al. (1986). This mélange is derived from the Lower
Cretaceous clays discussed above, and in the north from Permo-
Triassic clays. On Timor, very large areas are occupied by the Bobonaro
Scaly Clay that is associated with intrusive features, including forcing
of the clay matrix along internal joints and fractures and the break-up
of numerous blocks under high overpressures (Barber et al., 1986).
This evidence needs to be considered as a counter point to the
arguments that large, mobile shale masses do not exist. While regular
folds and thrusts can occupy parts of accretionary prisms, large areas
may also be broken down into more chaotic units by transfer of huge
volumes of overpressured fluids to the distal edge of the accretionary
complex.
4.3. Type 2bii active margins
4.3.1. Introduction
Active margins extend for over 30,000 km around the world, and
their associated accretionary prisms represent a vast accumulation of
sediment. Yet, despite their great volume, their contribution to
conventional world petroleum reserves is small, particularly com-
pared with the prolific DWFTBs of salt-associated passive margins
whose total length is about one third of the active margins. There is
one exception to this statement; it is the South Caspian Sea, which is a
highly unusual example of an active margin. Accretionary prisms form
a wedge-shaped profile, where new material is accreted to the thin
end of the wedge. The oldest and thickest part of the accretionary
prism can be thickened internally and increased in volume by
underplating of material. Consequently the wedge becomes uplifted
and exposed, to commonly form a chain of islands. These islands, such
as the Andamans, Nicobar, Barbados, Timor, Ramree (Burma), or
onshore ranges (e.g. the Makran) may display encouraging signs of
hydrocarbon generation with oil and gas seeps from deepwater
sediments. The seeps are often from mud volcanoes. The sediments of
the islands tend to have undergone multiple episodes of folding and
thrusting, shale diapirism, and sometimes a phase of extensional
tectonics (Pindell and Kennan, 2007; Maurin and Rangin, 2009). This
complex deformation can be understood in the context of critical
taper theory, where out-of-sequence deformation is necessary to keep
thickening the internal part of the wedge so oceanward propagation
of the thrust front can continue (Platt, 1986). Conversely, thickening
of the wedge also promotes gravitational collapse and episodes of
extensional faulting (Platt, 1986). Such a complex deformation history
means that trapping geometries, if they ever exist at all, will tend to be
small. Typically, the older, deeper parts of the accretionary prism have
been exhumed in the islands, hence compaction and diagenesis have
considerably reduced reservoir porosity and permeability. With the
advent of deepwater exploration, there is the possibility that the
Fig. 32. Tectonic map of the Seram collision zone (based on Sapin et al., 2009), showing location of Fig. 31.
71C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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simpler structures further offshore will avoid the disadvantages of the
proximal part of the wedge, and permit exploration for younger, less
deeply buried turbidite reservoirs in larger, simpler structures.
Out-of-sequence faulting is also thought to signify the onset of
significant lithification (i.e. increasing elastic strength) within the
accretionary prism, which permits slip to occur along discrete large-
scale blocks. The modern activity of out-of-sequence thrusts suggests
the seawardmost out-of-sequence thrust approximately defines the
seaward edge of the elastically rigid upper plate (Moore et al., 2007).
The following discussion focuses on present day active margin
DWFTBs and their hydrocarbon accumulations. Therefore, rather than
reviewing the vast range of known active margin settings, we center
on the more atypical settings with the highest hydrocarbon potential.
This means classic example of accretion ary wedges, which are
associated with the subduction of large oceanic plates from areas
such as Japan, Sumatra–Java, and Chile are not discussed herein, since
they have a relatively minor conventional hydrocarbon potential
(although the widespread occurrence of gas hydrates in accretionary
prisms, represents a potentially vast un conventional resource).
Improved conventional hydrocarbon potential occurs where the
more simply deformed external zones of DWFTBs lie in water depths
of 3000 m or less, and where large fluvio-deltaic system(s) may have
provided quartz-rich turbiditic sands to the evolving DWFTB. Areas
that match these criteria are the narrow, complex accretionary prisms
of the Caribbean, (e.g. Barbados, offshore Cuba), and those on the
trend of major continent–continent collision zones (e.g. The Makran,
South Caspian Sea and the Indo-Burma Ranges). Such settings are
invariably transitional to continent–continent collision zones: the
Indo-Burma Ranges grade from an active margin setting northwards
to an early-stage collisional zone. The Barbados accretionary prism
passes laterally into the transpressional setting of Trinidad to the
south and Cuba to the north. While the Makran passes into continent–
continent collision zones to the west (Zagros) and east (Himalayas),
and is associated with the remnant of the Neotethys subduction zone
(e.g. Stampfli and Borel, 2004; Hafkenscheid et al., 2006). Develop-
ment of major Cenozoic orogenic belts (Andes, Himalayas) was
accompanied by the development of large sediment-laden river
systems that debouched i nto the following deepwater ocean ic
domains: Trinidad–Barbados–Orinoco River; Makra n–Indus and
other rivers; Indo-Burma Ranges–Ganges–Bengal Fan, and precursor
systems. A schematic illustration of the general relationships between
large deltas prograding into an active trench system and providing an
evolving accretionary prism with sand-prone sediments is shown in
Figure 33. The deep-water Makran fold and thrust belt trends normal
to the drainage systems of its mountainous hinterland, while the
Indo-Burma Ranges and Trinidad–Barbados accretionary prisms are
charged along strike by fluvial-deltaic systems which prograde from
the related onshore foreland basins (Figs. 28 and 37). The ultimate
end-member to these settings is the South Caspian Sea, which has
undergone in excess of 10 km subsidence since the Pliocene due to the
influx of sediments from adjacent orogenic belts, particularly the
Greater and Lesser Caucasus (Allen et al., 2002; Brunet et al., 2003).
Accretionary prisms display a great range in width depending
upon their stage of development and sediment supply, which is
derived both from fluvio-deltaic sediments supplied to the upper
plate, and sediments scraped off the lower plate. For example the
frontal 50 km of the Apennines displays an average cross-sectional
area of 500 km
2
, while the Northern Barbados Ridge is only 100 km
2
(Bigi et al., 2003). This variation in sediment volume resulted in the
Apennines displaying a deeper detachment and higher elevations
than the Barbados Ridge (Bigi et al., 2003).
4.3.2. Barbados
The Columbus Basin of Trinid ad (Fig. 28) has been briefly
described in Section 4.2.1.2. There is a northwards transition from
the Columbus Basin, to the accretionary prism setting of Barbados.
Elements of the Trinidad geological history are applicable passing into
the deepwater Barbados accretionary prism. For example the basal
overpressured mobile shales, and the hydrocarbon source rocks in
both cases appear to be upper Cretaceous marine shales with La Luna
Formation-affinity (e.g. Hill and Schenk, 2005; Deville et al., 2006).
The Woodbourne Field, onshore Barbados has pro duced upper
Cretaceous-sourced oil from Cenozoic reservoirs since the 1970's,
and has provided encouragement for the new offshore exploration in
deeper water (Dolan et al., 2004; Hill and Schenk, 2005).
The Barbados Ridge Complex narrows from 250 km in the south to
100 km in the north, and its thickness varies from 200 m to 7000 m from
north to south (Moore et al., 1990). This thickness change reflects the
Onshore/nearshore
Shelf
Slope/deep marine
Trench
Continental
orogen or
strike-slip belt
Major fluvial
system and delta
Ophiolite and volcanic-arc
terrane
Anticline
Sediment transport
from arc terrane
Sediment transport from
continental area
Fig. 33. Schematic illustration of an accretionary prism with a dual sedimentary source due to the proximity of the DWFTB to a continental orogenic belt. The situation can occur
when a late-stage accretionary prims lies oblique to a continent-continent collision zone. For example the N– S River Ganges traverses E–W trending Himalayan structures, and
enters the Bay of Bengal and feeds a vast volume of sediment into the N–S trending foredeep basin to the Indo-Burma Ranges. Smaller drainage systems enter the Bay of Begal from
the Indo-Burma Ranges to the east.
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importance of the Orinoco Delta, located to the south of the ridge, as the
Barbados Ridge Complex's primary sediment source (Belderson et al.,
1984; Faugères et al., 1993; Dolan et al., 2004; Fig. 28). Active thrusting
and folding occurs up to 30–100 km from the toe of the prism (Mascle et
al., 1990; Huyghe et al., 1999, 2004, Fig. 28B). According to Mascle et al.
(1990) the pris m is domi nat ed by a serie s of nort hwe st- dip pin g imb ri cat e
thrusts, spaced between 2 and 12 km (average spacing ~7 km).
Deformation becomes more intense to the NW, with faults being more
closely spaced, and shortening increasing from less than 1 km on the
external imbricates to up to 5 km on thrusts further to the NW. Fold axial
trace length in the south of the prism ranges between 10 and 20 km long,
and increases to 30–40 km further north. The decrease in sediment
thickness from south to north affects fold and fault style within the prism:
relatively low-angle thrusts (~20° dip) deform the thin sequences, while
faults with higher dip (up to 40°), andbetterdevelopedpiggy-backbasins
characterize the thicker sequences (Mascle et al., 1990).
The basal detachment of the accretionary prism is strongly
overpressured, and has flowing fluids with pressures N 90% of
lithostatic (Housen et al., 1996). Chlorine stable isotope ratios
measured from mud volcanoes at the outer edge of the Barbados
accretionary prism provide evidence for a fluid reservoir at the base of
the decollement containing seawater and water released during gas
hydrate destabilization (Godon et al., 2004).
Hydrocarbon exploration in the offshore area of Barbados began in
1996 with Conoco (now ConocoPhillips; Dolan et al., 2004). The
exploration concept required the presence of oil-prone source rocks
equivalent to the La Luna Formation facies, Neogene reservoir rocks
(turbidites) derived from a proto-Orinoco sediment source, and
accretionary prism folds and thrusts providing the main traps (Dolan
et al., 2004). The exploration effort resulted in the collection of over
36,000 km of 2D seismic data, 3D seismic data, plus an extensive
program to identify hydrocarbon seeps by the collection of high-
resolution multi-beam swath bathymetry, piston coring and dredge
sampling. Upper Cretaceous and Cenozoic source rocks were
identified (Dolan et al., 2004).
Seismic r eflection data revea led that multiple episodes o f
compression, followed by extension and collapse, had affected the
accretionary prism. Mud pipe intrusions and mud volcanoes occur
extensively (Fig. 28). Conjugate (NW–SE and NE–SW trending) strike-
slip fault systems consistent with E–W compression are also present.
Many folds trend NE–SW and are bounded by NW–SE transpressional
faults. According to Dolan et al. (2004) seal integrity is regarded as a
major challenge due to pervasive shale diapirism and young faulting.
Migration of overpressured fluids is commonly linked with fault
systems. The first offshore exploration well (Sandy Lane-1) estab-
lished the presence of thick, well-developed, sand-rich turbidite lobes,
with biogenic and thermogenic gas saturations up to 15% recorded in
some sands. However, the structure is thought to have been breached
by Late Pleistocene strike-slip fault activity. NW–SE strike-slip fault
trends appear to be the most significant hydrocarbon migration
pathways, and the main cause of trap breach (Dolan et al. (2004).
4.3.3. The Makran
The Makran DWFTB is an accretionary prism about 900 km long,
which lies in relatively modest water depths of up to 3000 m and
grades to the NW into the frontal structures of the Zagros fold and
thrust belt. This accretionary prism developed on a fragment of
Cretaceous Semail Oceanic crust (Stampfli and Borel, 2004) that
survived in the gap between the Arabia–Eurasia collision zone of the
Zagros Mountains to the west, and the India–Eurasia collision zone of
the Himalayas to the east. The Makran accretionary prism developed
throughout the Cenozoic as the Arabian Plate was subducted under
the Eurasian Plate (Harms et al., 1985), with subduction probably
commencing during the Paleocene–Eocene (Platt et al., 1988; Byrne et
al., 1992). The proximity to the India–Eurasian collision zone resulted
in a large sediment supply to the Makran accretionary prism (Fruehn
et al., 1997). Large turbidite systems were deposited during the late
Paleogene to Miocene, incorporated into the accretionary prism, and
subsequently uplifted, eroded and recycled into the younger, ocean-
ward propagating parts of the accretionary prism (Harms et al., 1985;
Platt et al., 1985). In this fashion the accretionary complex has
propagated seawards at an estimated rate of ~10 cm yr
− 1
(White,
1982; Platt et al., 1985). Plate convergence at present is about 2.7–
4.2 cm yr
− 1
(DeMets et al., 1990; Vernant et al., 2004).
Large fluvio-deltaic systems have fed onto the continental margin,
so that offshore growth faults are imaged on seismic reflection data
(Grando and McClay, 2007; Fig. 34). These growth faults appear to be
linked to the toe thrusts. The range of active structures in the Makran
accretionary prism comprises near-shore growth faults, upper slope
to mid-slope shale diapirs or mud pipes and an outer-slope imbricate
zone (Grando and McClay, 2007; Fig. 34), These characteristics are, to
a degree, akin to those of a large delta-driven Type 1ai system.
Nevertheless, and as evident from Fig. 34, the amount of shortening
achieved in the distal imbricate zone exceeds by far the amount of
extension along the proximal normal fault systems. Correspondingly
the Makran system qualifies as a Type 2bii DWFTB.
In the older, uplifted, onshore part of the Makran system satellite
images show a large-scale ex ample of a Miocene growth fault
controlled depocentre that has been subsequently folded (Fig. 35).
In this regard the Makran resembles the Baram Delta Province. The
growth fault depocentre in Figure 35 is of a similar size to major
growth fault depocentres in the Baram and Niger Deltas.
Development of major normal faults in accretionary prisms has
been explained by Platt (1986) and Grando and McClay (2007) in
terms of a critical wedge model (Fig. 36A). Thickening of the back end
of the wedge (for example by underplating) can trigger extensional
collapse of the wedge either by differential loading or gravity sliding.
This model implies extension occurs regardless of the nature of
sedimentation affecting the margin. However, the growth fault
example from the Makran (Fig. 35) indicates a large delta was
present during the Miocen e, and dif fer ent ial loading probably
triggered the development of the large depocentre. Like in NW
Borneo, the normal faults could have linked with compression down-
N S
Chaotic zone
on seismic
Imbricate zone
Deformation Front
TWT (s.)
10 km
Himalayan Turbidites Makran Sands Growth I Growth II Growth III
0
2
4
6
8
Fig. 34. Cross-section through the Makran DWFTB (Grando and McClay, 2007). The chaotic zones on the seismic data could be indicative of shale diapirs, narrower shale pipes, or gas
chimneys.
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0
20
40
0
20
40
km
50 km
Slope too high: extension
Shale diapirs or pipes
Slope too low: contraction
low rate of frontal accretion
large-scale
underplating
km
High sediment supply - low slope
Shale diapir or pipe Slope too low: contraction
Deformation in accretionary prism
driven by both near-field stress
and far-field stress
large-scale
underplating
Inverted/folded growth fault depocentres
Active growth fault depocentre
A
B
Folded inactive extensional detachment
Fig. 36. Models for the development of extensional faults in an accretionary prism setting. A) Extensional collapse in response to overthickening of the inner part of a critical taper
wedge, (Platt, 1986). B) Extension due to differential loading of the inner wedge by a large delta. Older extensional fault depocentres have become inverted.
Fig. 35. Satellite image of a large extensional growth-fault bounded depocentre in the southern Makran, Pakistan. This depocentre has been folded and is now involved in a syncline.
74 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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dip (Fig. 36B), and contributed to shortening within the wedge, then
was subsequently inverted. Such features are seen in both the
Pakistan (Fig. 35) and Iranian Makran (Hosseini-Barzi and Talbot,
2003).
To date, only four wells have been drilled in the offshore Makran,
some of which had gas shows, but apparently established no
commercial accumulations. Numerous gas seeps associated with
mud diapirs are located near the coast (as reviewed in Khan et al.,
2007). Temperatures within the accretionary prism are relatively low,
with geothermal gradients averaging about 2 °C/100 m, and average
heat flow is 47 mWm
− 2
, which suggests maturation of hydrocarbons
between depths of 3000 and 5000 m (Ahmad, 1969; Kaul et al., 2000;
Khan et al., 2007). Several potential source horizons are indicated
including: the Upper Miocene Parkini Shales, kerogen type II–III, TOC
0.5–1% (Khan et al., 2007); and the Middle Miocene (abyssal-slope)
Panjgur Shales, type II–III, TOC 0.7–5.6%. The offshore Parkini turbidite
sandstones of Miocene–Pliocene age appear to offer reservoir
potential, particularly the Panjgur Sandstones (Platt et al., 1985;
Khan et al., 2007). However, there are indications of ophiolitic and
volcanic contribut ions f rom the sediment source area in the
quartzolithic Panjgur Sandstones, which gives rise to a complex
diagenetic history with secondary porosity creation and destruction
and formation of a range of cements (chlorite, ferroan calcite, ankerite
and dolomite; Garzanti et al., 1996; Grigsby et al., 2008).
4.3.4. Andaman subduction zone–Indo-Burma Ranges
The N–S trending Indo-Burma Ranges lie in western Myanmar and
eastern Bangladesh (Fig. 37). Onshore they are composed primarily of
Palaeogene deepwater sediments, and Cretaceous mélanges, contain-
ing blocks of gabbro, pillow basalt, serpentinite, banded chert,
limestone and schist, interpreted as having developed as result of N
to NE directed subduction during the Mesozoic–Cenozoic (Mitchell,
1993). The western, external parts of the Indo-Burma Ranges consists
of f olded Miocene–Pliocene fluvio-deltaic sediments, essentially
sourced by the river G anges, which wer e f olded during the
Pliocene–Holocene (Sikder and Alam, 2003; Maurin and Rangin,
2009).
The Indo-Burma Ranges initially developed as a result o f
subduction of the Neo-Tethys Ocean, and then, from the Senonian
onward, of the Spongtang Ocean beneath the active southern margin
of Eurasia (Stampfli and Borel, 2004; Hafkenscheid et al., 2006; Hall et
al., 2009). The Indo-Burma Ranges are, in part, an old accretionary
prism complex that developed during the Late Cretaceous and
Paleogene and show a complex history of deformation, ophiolite
emplacement, uplift and erosion (see: Mitchell, 1993; Acharyya,
2007). The coupling of India with NW Myanmar, and the collision of
India with Eurasia terminated the early accretionary history during
the late Paleogene (Searle and Morley, in press). However, earthquake
activity from the Burma seismic zone indicates that subduction
processes are still active beneath western Myanmar and Bangladesh
(Stork et al., 2008). This section focuses on the post India–Eurasia
coupling history of the Indo-Burma Ranges.
India has undergone highly oblique convergence with continental
core of SE Asia during the Cenozoic. GPS measurements indicate India
is presently moving about 35 mm/yr
-1
northwards with respect to
Sundaland. This motion is accommodated by distributed deformation
on numerous faults across the Burma microplate,of which the Sagaing
Fault, with a dextral displacement of b 18–20 mm/yr
− 1
, is the largest
(Fig. 37; Vigny et al., 2003).
The outer Indo-Burmese wedge forms the western margin of the
Burma microplate and lies predominantly on oceanic crust (Curiale et
al., 2002). It underwent folding from the late Miocene–Recent
(Maurin and Rangin, 2009). The wedge is characterised by a broad
region of folding in the north, covering the onshore and offshore areas
of Bangladesh and Myanmar. The fold and thrust belt narrows to the
south and becomes increasingly focussed offshore. Folding has
occurred in the north above an effi cient Oligocene deep marine,
overpressured detachment and covers a broad region up to 150 km
wide (Sikder and Alam, 2003; Maurin and Rangin, 2009). Five
hundred kilometres further south, the young fold and thrust belt is
Ganges R.
E
O
M
P
10 km
WellA4
1
2
3
4
5
6
7
0
TWT (s)
P = top Pliocene
M = top Miocene
O = top Oligocene
E = top Eocene
NW SE
Gravity slide
Mud diapir
1 km
0
1
2
3
4
5
Mud diapir?
Wedge
NE
SW
A
B
B
C
Main Boundary Thrust
Shillong Plateau
Sagaing Fault
Ya n g o n
Andaman Basin
Kabaw Fault
Kaladan Fault
95°E
90°E
15°N
20°N
25°N
Indo-Burma
Fold and Thrust Belt
Myanmar Central
Basins and Pegu
Yoma
India/Eurasia
Boundary
Andaman Trench
Coco Island
Shan
Plateau
Burma
Microplate
9
12
18
A
Section B
Section C
C
Fault
Fold axis
Oceanic Crust
Continental
Fig. 37. A) Tectonic map of the Indo-Burma Ranges showing location of structural cross-
sections given in B and C. (compiled from Maurin and Rangin (2009) and Nielsen et al.
(2004)). Arrows (A,B,C) and numbers indicate the direction and relative convergence
rate of India with respect to the upper plate margin (all rates in mm/yr), which
progressive rotate from 035°/18 mm/yr (south M yanmar) to E–W/9 mm/yr
(Bangladesh fold and thrust belt), from Nielsen et al. (2004).
75C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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just 20 km wide and is composed of just 2–3 folds along a narrow
slope, that are bounded by a major transpressional fault on the
western margin (Fig. 37A,B). Young deformation also affects the shelf
area, with the development of normal faults, and mud diapirs;
subsequently the entire region underwent extensive i nversion
(Nielsen et al., 2004; Maurin and Rangin, 2009).
The change in structural styles from south to north indicates that
the highly distributed oblique deformation of the Burma microplate
becomes increasingly focussed towards the south on the oblique-
dextral Andaman Subduction Zone. In the southern region (near the
north Andaman Islands), the 3.5 cm/year motion of India relative to
Sundaland is taken up, in equal amounts, by the Andaman Subduction
zone and Sagaing–West Andaman Fault system. However, this motion
is fully partitioned across the whole Burma microplate in the north
(Nielsen et al., 2004). Deepwater troughs are present both east and
west of the Coco Islands. Considerable folding and thrusting has
inverted deepwater basins on the east side of the islands. The
occurrence of a deepwater basin on the east side is probably due to the
presence of the Andaman Sea Spreading Centre to the south that was a
source of recent subsidence in the region, coupled with the sediment-
starved nature of the area.
The Bay of Bengal is well known for the massive accumulation of
Cenozoic deltaic sediment associated with the Bengal Fan, which
exceeds 20 km near the Bangladesh coastline (e.g. Curray and Moore,
1971; Metevier et al., 1999; Curiale et al., 2002; Clift, 2006). Much of
the northern fold belt area in Bangladesh accumulated sediments in a
deep marine environment during the Early Miocene (Gani and Alam,
2003). By the Late Miocene, the area was undergoing deposition by
braided fluvial systems, and folding developed in a continental
environment (Sikder and Alam, 2003). Much of the Outer Burma
wed ge al ong the prese nt coast of Myanmar lay in deepwater
conditions during the Pliocene. The deepwater gas discoveries in
NW Myanmar are stratigraphically trapped in Pliocene turbidites and
indicate that a viable deepwater petroleum system exists in the area.
However, the timing of folding in this area is young, during the last
2 Ma (Maurin and Rangin, 2009). A combination of uplift due to
folding, and infilling of accommodation space by rapid north to south
progradation of the Bengal Fan, together with lesser input from the
Indo-Burma Ranges to the east (Allen et al., 2008) has resulted in the
northern part of the Outer Burma Wedge rapidly developing into a
shallow marine environment in the relatively recent geological past.
At present, the northern 550 km of the Outer Burma Wedge lies either
in shallow water or onshore. This rapid progradation means that, in
places, plays for deepwater Late Miocene or Pliocene turbidites folded
by accretionary prism-related deformation can be found in a shallow
marine setting. Yet these plays have not met with economic success
offshore so far. One problem may be that folding is too young, with
respect to hydrocarbon migration.
4.3.5. South Caspian Sea
The South Caspian Sea is a prolific oil province with reserves likely
to be in excess of 30 billion barrels of oil equivalent (BBOE). The South
Caspian Basin extends 680 km N–S and 500 km E–W, with maximum
water depth of ~1000 m (Fig. 38). The South Caspian Sea is thought to
be underlain by about 10 km thick oceanic crust, which is covered by
up to 26–28 km of Mesozoic and Cenozoic sediments (Berberian,
1983; Jackson et al., 2002; Knapp et al., 2004; Fig. 39A). This crust
originated during the Mid-Late Jurassic opening of the Izmir–Ankara–
South Caspian back-arc basin (Golonka, 2004; Stampfli and Borel,
2004).
The South Caspian Basin is bounded to the north by the Apsheron
ridge that links the Great Caucasus and Great Balkhan –Kopet-Dagh,
-400 m
Kopet Dagh
Alborz Mtns.
South Caspian Sea Basin
North Caspian Sea Basin
Ap
s
heron
Ridge
Gre
a
ter C
a
u
casus
Kura Basin
Talysh
100 km
48°
52°
40°
37°
56°
-400 m
-400 m water depth
large thrust or strike-slip fault
fold
Fig. 38. Regional tectonic map of the Southern Caspian Sea and surrounding areas. (modified after Allen et al., 2002).
76 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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which evolved by Late Eocene and younger inversion of rifted basins.
The South Caspian Basin is bounded by the Alborz fold and thrust
belt to the south, which remained active during Neogene times
(Allen et al., 2002; Jackson et al., 2002; Knapp et al., 2004). During
the complex late Mesozoic and Cenozoic history the area closed up
as the Arabian Plate and Eurasia converged and intervening oceans
(Neotethys, Sistan oceans) were subducted northwards and lost (e.g.
Golonka, 2004). In the south Caspian area the main shortening and
basin inversion stage occurr ed duri ng the Late Eoc ene–Early
Oligocene as the NeoTethys oceanic domains were closed ( Kopf
et al., 2003). The Late Miocene–recent phase of compression, uplift
and magmatism coincides with the final stages of Arabia –Eurasia
collision (Nikishin et al., 2001). Mountains adjacent to the South
Caspian Sea have been uplifting and providing a sediment source
since the Paleogene. However, rapid Plio-Pleistocene uplift, which
occurred in the Lesser and Greater Caucasus Mountains, the Kopet
Dagh Mountains and Alborz Mountains triggered a very high influx
of sediment into the South Caspian Sea (Jackson et al., 2002; Brunet
et al., 2003).
Allen et al. (2002) investigated subsidence in the South Caspian
Basin. They noted that the ~2.4 km of tectonic subsidence since 5.5 Ma
calculated from backstripping was at a very rapid rate in comparision
with typical foreland basins, and more than half the sediment
thickness of the basin was accumulated during the Pliocene–
Holocene. To explain such anomalous subsidence, Allen et al. (2002)
concluded that “subduction” of the South Caspian basement (either
thick oceanic crust or thinned, high-velocity continental crust) began
ca. 5.5 Ma, as a result of regional reorganization of deformation in the
Arabian–Eurasian collision zone. Present day convergence rates to the
NW are 8–10 mm/yr (Jackson et al., 2002). Knapp et al. (2004) have
shown evidence for an oceanic slab extending over 100 km northward
beneath the Eurasian continental crust (schematically shown in
Fig. 39A). This slab shows seismicity at depths of 30–80 km and is
described as incipient subduction (Jackson et al., 2002; Knapp et al.,
2004).
The close proximity of the South Caspian Basin to continent–
continent collision zones has resulted in the following late-stage
characteristics of the area undergoing ‘incipient subduction’: 1)
subduction cannot proceed much further from its present state, the
remnant of oceanic crust is surrounded by continental crust that is
shortening in collision zones and will prevent large-scale subduction,
although underthrusting will continue, 2) The swamping of the basin
by continent-derived sediments deposited in a fluvio-deltaic setting,
and 3) the lateral transition from deepwater folds to shallow or
subaerial folds, similar to the Indo-Burma Ranges.
According to Devlin et al. (1999) and Brunet et al. (2003) the South
Caspian Basin is thought to comprise up to 5 km of Mesozoic
sediments and volcanics that overlie crys talline basement. The
Mesozoic units are overlain by marine Paleocene and Eocene marls,
clays, bituminous shales and sandstones of up to 1700 m thickness. A
tectonic re-arrangement in the basin configuration occurred during
the early Oligocene, in which some basins were eliminated, and
deposition in the Caspian Sea area occurred in a low-relief setting. The
Oligocene–Early Miocene clay-rich Maykop Series is over 3 km thick
and is an excellent hydrocarbon source rock (Katz et al., 2000). The
Maykop Series are overlain by the up to 10 km of the late Miocene–
Holocene Productive Series. Rapid deposition resulted in widespread,
high magnitude overpressures, and the extensive development of
mud pipes, diapirs and volcanoes that are mostly sourced from the
Maykop Series.
Type 1 near field stress-driven, mobile shale-floored growth faults,
shale diapirs and fold and thrust belts are present in the eastern part
of the South Caspian (Turkmenistan), and at least, in part, developed
earlier than the Pliocene–Holocene folds that dominate the area
Mesozoic sediments
Caspian Sea
South Caspian North Caspian
Mesozoic-Eocene
North
Caspian
Crust
Oceanic (?) crust
Maykop Series
Productive Series
Productive Series
South Caspian Sea
Depth (km)
20 km
Oceanic (?) crust
Maykop Fm.
Folded and thrusted complex
Continental
crust
Mud volcano
South Caspian
Sea
Line of section
Apsheron
Ridge
South North
A
B
5
0
10
15
20
Fig. 39. A) Crustal-scale model for the tectonic setting of the Southern Caspian Sea and Apsheron Ridge. B) Structural cross-section based on 2D deep reflection seismic data,
illustrating structural style of the boundary zone between Absheron Ridge and Southern Caspian Sea (after Stewart and Davies, 2006).
77C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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(Devlin et al., 1999; Lawrence and Babaev, 2000). The spectacular
folds of the South Caspian Sea vary considerably in map view
geometry, ranging from long (50–150 km) linear, to strongly curved
(NW–SE to NE–SW strike directions), and dome and basin forms.
Folds in the South Caspian Sea display a similar structural style
whether they are in deep or shallow waters. Unlike most DWFTBs,
where there is a distinct relationship between water depth and the
location of folds, someof the larger fold trends on the western side of the
South Caspian Basin, which lie predominantly in Azerbaijan, have a
strike length of 150 km, trend NW–SE and pass from being present on
land in the NW to deepwater conditions at the SE part of the trend
(Fig. 38). Despite the unusual fold-water depth relationship the
province has much in common with DWFTBs because it is affected by
shallow gravity driven deformation associated with deltaic deposition,
displays growth faults and down-dip folds and thrusts, and has
extensive, overpressured mobile shales (Devlin et al., 1999; Lawrence
and Babaev, 2000). Three related factors contribute to the shallow into
deep water lateral transition along the fault axes: 1) rapid sedimenta-
tion rates that caused marked infilling of the basin during the Pliocene
from fluvio-deltaic systems arranged radially around the basin, 2) the
inland sea setting where a remnant of oceanic crust has been preserved
surrounded by continental crust, and 3) the presence of active mountain
belts (Alborz, Kopet Dagh and Caucus mountains) that both act as
extensive sediment sources and are associated with Cenozoic regional
folding, thrusting and uplift adjacent to, and within the basin.
The Mesozoic–Eocene section appears to be deformed by thrust
faults and possibly strike-slip or oblique slip faults (schematically
illustrated in Fig. 39). The basement-involved deformation is related
to right-lateral oblique thrusting and inversion of the Caucasus
Trough, which contains very thick sedimentary sequences. Folding in
the South Caspian Basin displays such a variety of orientations and
strong vertical disharmony for several key reasons including: 1)
lateral variations in sediment thickness, 2) tectonic position, 3)
basement type and structure, and 4) location of overpressured shales,
particularly the Maykop Series.
Fowler et al. (2000) provide images of a fold in water depths of
300–600 m, across the Shah Deniz structure (70 km SE of Baku). The
fold is larger than the other deepwater folds described in this paper. It
is a broad, upright fold in a sedimentary section at least 12 km thick
and, consequently, has a large wavelength of 20 km. A few low-
displacement thrusts are present at depth. Although the overall
structure is simple, in detail the presence of extensive mud volcanoes
in the upper 3–5 km of the fold create considerable local complexity.
Large mud pipes (3–5 km diameter) are present, and they feed
smaller mud pipes (100's m diameter) and are linked with complex
associated faulting (such as ring faults, Stewart and Davies, 2006).
Furthermore, thrusts, normal faults and strike-slip faults located at
depth make deformation in the lower part of the fold more complex.
The Shah Deniz field is BP's largest field worldwide, it covers a
surface area of 250 km
2
, ranges in water depth from 50 m to 600 m, and
has recoverable reserves estimated at 22 TCF gas and 750 MMBO
condensate (BP in Azerbaijan, Sustainability Report, 2005, www.bp.
com/caspian). Large reserves have been also been established in the
vicinity of the Apsheron Ridge area, most notably the Azeri–Gunashli–
Chirag field (432.4 km
2
,5–6 billion barrels recoverable; Abrams and
Narmanov, 1997). Yet there remain considerable risks exploring the
large anticlines, and a number have proven to be expensive failures.
Coping with overpressure distribution is one problem, reservoir
distribution/quality and top-seal breach by mud pipes are other factors.
5. Synthesis of DWFTB characteristics
5.1. Petroleum systems
The main aim of this section is to identify broad similarities and
differences in the main elements of the petroleum systems that
characterize the different categories of DWFTB. The most widespread
exploration and development activities have been on salt-associated
passive margin plays in the South Atlantic (e.g. Gulf of Mexico, Brazil,
Angola, Congo, Gabon, Mauritania). There is considerable detailed
published information for these plays (see Weimer et al., 2006 for a
review), which contrasts strongly with the paucity of data for other
plays. Consequently the comparisons made below are based on
uneven data quality.
5.1.1. Source rock types
Large deltas can be up to 10 km thick beneath the shelf, hence
associated source rocks generally have to be present within the deltaic
sediments. Even if potential source rocks are present in pre-deltaic
sediments, the thickness of the delta wedge, and generally late
development of the DWFTB precludes section below the basal mobile
salt or shale detachment rocks from being mature at the time of
structural development, except for the outer part of the wedge. For
example in general the Paleogene Agbada–Akata Formations are
thought to be the main source rocks for the Niger Delta including the
deepwater (e.g. Evamy et al., 1978; Bustin, 1988). Yet the observation
from the Gulf of Mexico that the deepwater area is sourced by older
source rocks than on the shelf led Haack et al. (2000) to propose
hypothetically that the deepwater plays of the Niger Delta might be
sourced in part by Late Cretaceous source rocks. Recent geochemical
analysis of deepwater Niger Delta oils by Samuel et al. (2009) suggests
that this prediction is probably correct.
Source rocks in the deepwater Gulf of Mexico source rocks are
primarily located within the Tithonian sequence, but there are
secondary source rocks of Oxfordian and Mid-Cretaceous age (Cole
et al., 2001). Additionally, the source rocks in the western Gulf are
from carbonate-marl sequences, while source rocks are predomi-
nantly carbonate-marl to clastic sequences in the eastern Gulf (Cole et
al., 2001). These variations in source rock type are reflected in the oil
chemistry: the deepwater oils are moderately sour (0.5–20 wt.%
sulfur) and predominantly 25–35 API, while ultra-deepwater oils and
shelf oils are sweet, low sulfur and high gravity. In contrast to the pre-
deltaic source rocks of the Gulf of Mexico, the source rocks for
deepwater hydrocarbons associated with the Baram and Mahakam
deltas of Borneo appear to have originated from within the deltaic
depocentre as a result of re-working of Neogene shallow marine-
estuarine organic material (resins, waxy cuticulae of t ropical
vegetation, mangrove coals) during lowstands into the deep marine
environment (e.g. Saller et al., 2006; Warren et al., in press).
For smaller deltas, and many salt detachment-dominated passive
margins, the potential age-range of source rocks is wide since the
effects of burial discussed above are not as strong. The South Atlantic
margin is a very well documented example, where source rocks are
closely related to the different stages of passive margin development,
with early continental rift-related lacustrine source rocks, early-drift
stage, evaporite associated source rocks, drift-stage Cretaceous
marine source rocks, and more gas-prone, source rocks deposited by
Tertiary deltas (Katz and Mello, 2000; Schiefelbein et al., 2000). The
passive margin setting also means that some source rock intervals
may thicken passing offshore. For example Cameron et al. (1998)
attributed a broader suit of Cretaceous and Tertiary oil-prone, algal-
rich source rocks to expansion offshore of the Atlantic Hinge offshore
Angola. The main liability in the deepwater area is that in places the
sedimentary overburden of potential source rocks is not sufficiently
thick for their maturation. In some areas of salt tectonics (e.g. West
African margin) hydrocarbon kitchen can be restricted to individual
mini-basins between folds or diapirs, causing such areas to have
laterally highly variable source potential. This variation has significant
consequences for the hydrocarbon charge to adjacent structures.
The crust underlying the DWFTB in collision zones tends to be
thinned continental crust that underthrusts the DWFTB and is buried
during deformation. Consequently, mature source rocks may be present
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from within the thrust belt wedge, and also below (particularly from
syn-rift, post-rift passive margin and foreland basin sequences, e.g.
Ziegler and Roure, 1999). In the collision zones of the Bandar Arc area of
eastern Indonesia, maturationof Paleogenesyn-rift sedimentsis thought
to be important in explaining the occurrence of oil seeps (Charlton,
2009). In the Timor Sea, the NW Australian passive margin series with
their Mesozoic petroleum systems are presently overridden by the
Timor DWFTB (e.g. O'Brien et al., 1999; Charlton, 2002). According to
Baillie (2009), many deepwater areas of Indonesia that were once
thought non-prospective (e.g. Bone Bay, Cendarawasih Bay, Sumatran
forearc), are now being re-examined since new seismic, multibeam and
core data has identified extensive areas with thermogenic hydrocarbon
seeps that in places can be linked with significant thicknesses of
sedimentary section. Many of these areas are associated with DWFTBs.
In accretionary prisms, the subducting lower plate consists of
oceanic crust and lithospheric mantle. Hence, the main potential for
source rocks resides in the upper plate within the thickened
accretionary prism (e.g. Makran: Khan et al., 2007; Indo-Burma Ranges:
Maurin and Rangin, 2009; Sumatra: Baillie, 2009). For example, prism
three distinct groups of oils are recognized in the Barbados accretionary
that are related to generation at multiple levels within the prism from
Late Cretaceous marine source rocks deposited under dysoxic condi-
tions (Hill and Schenk, 2005). Source rocks of this age are widespread
along the southern Caribbean area west of the trench.
5.1.2. Temperatures and hydrocarbon maturation
Several decades of exploration have established that the conditions
for hydrocarbon maturation in deepwater environments is wide-
spread, and in general, there is nothing specific about the deepwater
environment that precludes source rock maturation. However, lower
geothermal gradients may require a different, generally older, source
rock in the deepwater compared with the shelf. Furthermore the
uncertainty associ ated with s ource rocks re aching maturation
increases as exploration extends into ultra deep areas.
Published temperature data for DWFTBs are scarce, but some
generalizations can be made. There is generally a decrease in heat flow
from continental areas to oceanic crust as radiogenic material within
the crust is reduced, such as the Amorican Margin, where typical
continental values of heat flow (N 90 mW/m
2
) are higher than oceanic
values (b 60 mW/m
2
)(Louden and Mareschal, 1996). Geothermal
gradients in passive margins reflect more than just the radiogenic heat
production. The presence of major faults, thick salt and young
sediment depocentres can cause significant lateral variations (e.g.
Forrest et al., 2005; Nagihara and Opre Jones, 2005). Nevertheless
diminishing heat flow or geothermal gradients passing offshore are
the general rule. Maturation can be a significant risk for the outer
parts of DWFTBs on passive margins near the continent–ocean
transition, as the dual decrease in heat flow and sediment thickness
reduce the source rock maturation potential. However, in some areas
this change can be advantageous, for example the hotter shelf area in
the Gulf of Mexico is more gas-prone, while the cooler deepwater area
is more oil-prone.
The Niger Delta slope has an average heat flow of 58 mW/m
2
, with
little decrease in heat flow between the upper and lower slope
(Brooks et al., 1999). Nevertheless the outer fold and thrust belt is
characterized by four-way dip-closed anticlines that are either under
filled by thermogeni c hydrocarbons or filled by biogenic gas,
suggesting that in the outer-most structures, the volume of thermo-
genically generated hydrocarbons is relatively low (Kostenko et al.,
2008).
Convergent setting DWFTBs are underlain by thinned continental
crust. Hence, any offshore increase in depth to the oil and gas
windows is likely to be due to the increased water depth rather than
to a decrease in crustal radiogenic heat production. Heat flow data for
the Brunei deepwater area suggests mean heat flow decreases down
slope (upper slope heat flow ~84.0 ± 66.5 mW/m
2
; lower slope heat
flow ~59.0 ± 22.6 mW/m
2
), but the variations in heat flow value are
highly complex (Zielinski et al., 2007). These authors suggested that
high standard deviation in heat flow values reflect hydrothermal
circulation, which is similar to the heat and fluid flow regimes of
accretionary complexes. They also note that the heat flow values for
the Baram Delta Province are significantly higher than the generally
low (~40–60 mW/m
2
) mean heat flows associated with accretionary
prisms (e.g. Ferguson et al., 1993). Fluid flow along thrust faults
systems and particularly the basal detachment is a very important
mechanism in explaining heat flow, and hydrocarbon migration in
both convergent DWFTBs and accretionary prism systems (e.g. Fisher
and Hounslow, 1990; Zielinski et al., 2007; Barnes et al., 2010). In
accretionary prisms, heat flow values can increase considerably as the
wed ge thins due to fluid flow and expulsion along the basal
detachment (Ashi and Taira, 1993).
5.1.3. Hydrocarbon migration
In deepwater areas, the pattern of hydrocarbon migration can be
assessed by a variety of methods. Seismic reflection data permits
mapping of direct indicators of fluid migration such as mud pipes, gas
chimneys and various other fluid escape features. These features
create sub-vertical zones of low-reflectivity on seismic sections, pock
marks on time slices, affect seismic attributes (direct hydrocarbon
indicators, such as bright spots and flat spots) and caus e the
preferential accumulation of gas hydrates (marked by bottom
simulating reflections) in the crests of anticlines; (e.g. Fowler et al.,
2000; McConnell and Kendall, 2003; Deville et al., 2003, 2006; Gay et
al., 2006; Stewart and Davies, 2006; Morley, 2009a; Barnes et al.,
2010). Large pipes, 100's m wide, appear to originate deep (kms) in
the seismic data, while smaller pipes 10's m across can arise from
depths in the order of 100's m (e.g. Morley, 2009a). Overpressured
fluids may travel through permeable (silty or sandy) beds deep in the
section before rising vertically (e.g. Gay et al., 2006). These permeable
beds would filter mud from the fluids leaving them carrying little clay
fraction (although they may subsequently incorporate more mud),
unlike the pipes sourcing mud volcanoes. Side scan sonar or multi-
beam surveys coupled with targeted piston coring programs, or
simple grid piston coring programs have also provided important
regional coverage that commonly show preferential hydrocarbon
seepage along specific types of structure, such as young faults, mud
pipes, salt diapirs and gas chimneys (e.g. Kaluza and Doyle, 1996; Cole
et al., 2001; Dolan et al., 2004; Zielinski et al., 2007; McConnell et al.,
2008; Warren et al., in press).
There are clear differences between the different types of DWFTBs
with respect to migration and trap timing, in particular between salt
and shale detachment margins, and between margins with landward
and oceanward-dipping detachments. Types 2a and 2b DWFTB with
shale detachments are also all associated with a landward-dipping
detachment and a critical-taper wedge. As described above in the
discussion of heat flow, fluid flow within the critical-taper wedges
tends to be along the thrust faults. Consequently, models for the
migration of hydrocarbons in convergent DWFTBs and accretionary
prisms closely resemble each other (Figs. 8B and 40B; Deville et al.,
2003; Morley, 2009a; Cobbold et al., 2009; Warren et al., in press;
Barnes et al., 2010
). Hydrocarbons and overpressured fluids are
generated within the wedge, and when the overpressure reaches a
critical magnitude fluids move to areas of lower pressure along zones
of weakness. The wedge geometry forces the fluids to migrate
oceanward, and the most pervasive zones of weakness tend to be
thrust faults. Hence fluids tend to migrate laterally and upwards along
thrust fault surfaces until they reach a critical depth where hydraulic
fracturing permits a vertical column of overpressured fluid to force its
way to the surface (Fig. 8). Typically fluid pipes develop in the crests of
anticlines and create mud volcanoes at the sea floor. Fluid pipes and
migration of gas up carrier beds along the backlimb of anticlines
appears to focus gas hydrates (as seen by well-defined bottom
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simulating reflections) in the crests of anticlines. The fluid migration
events can be pulsed, and the geochemistry of oils from fields in the
Niger Delta and seep studies in Barbados shows variations that are
consistent with differing degrees of maturity or biodegradation
associated with multiple charging events (e.g. Katz, 2003; Hill and
Schenk, 2005).
There is reason for concern that insufficient charge may get to the
most distal folds in DWFTBs like Brunei and Niger (Morgan, 2003,
Kostenko et al., 2008). The seep map for offshore Brunei shows only
weak oil seeps in the lower slope area (Fig. 9). A detailed investigation
of the causes of limited column heights in the outer DWFTB of the
Niger Delta concluded that insufficient hydrocarbon charge, and not
trap integrity, was the main problem (Kostenko et al., 2008). This
conclusion suggests lateral migration can be of limited efficiency and/
or that the source rocks are too lean to generate sufficient
hydrocarbons to fill the available trap capacity.
Passive margin salt provinces are characterized by a very different
migration scenario in comparison with the vertical to ocean-wards
migration dominated critical taper wedge settings described above.
The overa ll oceanwards dip of the margin means hydrocarbon
migration is generally either vertical or landwards. As salt is an
effective migration barrier, source rocks must either overlie the salt
unit, or depositional gaps in the salt, large faults, or areas of total salt
withdrawal must be present to permit vertical migration of hydro-
carbons generated by sub-salt source rocks into post-salt structural
traps (e.g. Jackson et al., 1994). Cretaceous and Cenozoic post-salt
source rocks underlie effective petroleum systems in large parts of the
South Atlantic margins (Cameron and White, 1999). However, the
charge of post-salt prospects by hydrocarbons generated by pre-salt
syn-rift source rocks is important in many areas. Regions of thick salt,
such as the salt plateau area of the Outer Kwanza Basin (Fig. 18), have
major problems with hydrocarbon charge since migration from sub-
salt sources is blocked. The post-salt sediments are generally too thin
to be mature, and lateral migration pathways would have to be very
tortuous. The prolonged history of movement of salt on passive
margins (over 100 My in the case of West Africa), the ability of salt to
migrate to different stratigraphic levels with time, and the changing
thickness of sediment overburden with time, mean that hydrocarbon
migration patterns on such margins can be complex, and will have
changed considerably spatially and temporally.
In type 1a DWFTBs ass ociated with relatively small deltas,
prograding sedimentary prisms are commonly not thick enough for
the generation of hydrocarbons within them. Consequently, either the
DWFTB needs to be relatively old, contain source rocks, and be buried
to depths suitable for hydrocarbon generation, or a sub-DWFTB source
is required, with vertical migration pathways to the traps. The timing
of sub-DWFTB source rocks is critical. Post-kinematic burial of the
DWFTB may be necessary in order to mature the source rocks within
or beneath them. Alternatively, structural thickening in the DWFTB
maybe sufficient to bury deeper source rocks to maturity in the largest
deltas. Establishing a vertical migration pathway can be problematic,
and acts contrary to the dip of the basal detachment and dip of the
margin (Fig. 40A). Landward-dip ping detachments offer greater
opportunity for the hydrocarbon charge of a DWTFB from source
rocks located both above and below the detachment ( Fig. 40B). For
example, the Mexican Ridges play in the Laguna Madre–Tuxpan area
(Neogene folds and reservoirs) requires an underlying Upper Jurassic
source (Ambrose et al., 2005). Deep-seated basement faults provide
vertical migration pathways in the southern part of the Laguna
Madre–Tuxpan area, but further north, the offshore-dipping low-
angle detachment system above the source rock is likely to form an
effective barrier to hydrocarbon migration into the DWFTB (Ambrose
et al., 2005).
5.1.4. Reservoir quality
The first-order concern pertaining to reservoir quality in DWFTBs
is related to the sediment source, which in turn is related to the
tectonic setting (Dickinson and Suczek, 1979). In general, the main
DWFTB reservoirs are turbidite sandstones reworked from clastic
shelf deposits that were transported to the deepwater areas during
lowstand in sea level events or tectonic uplift events (Ofurhie et al.,
2002; Fainstein, 2003; Ingram et al., 2004; Saller et al., 2008). A large
number of Type 1 DWFTBs are associated with passive margins, such
as those of eastern North and South America and Africa. These
margins are fed by large river systems, which drain the continental
interior. While the geology of the continental interiors is complex and
highly variable, generally a considerable component of the sediments
are derived from acidic basement, or Paleozoic–Cenozoic clastics and
carbonates. The long fluvial transport paths tend to eliminate the
more reactive (generally more basic-ultrabasic igneous or metamor-
phic) minerals, and consequently in general passive margin deepwa-
ter turbidites are predominantly composed of high reservoir quality
quartzose, or quartzo-feldspathic sands, characteri zed by high
porosities and permeability (e.g. Dickinson and Suczek, 1979;
Marsaglia et al., 1996; Fontanelli et al., 2009).
Type 2 DWFTBs are associated with tectonically active regions
where the contribution of more highly diagenetically reactive
minerals to the sand composition can be much higher, and quartz
content is commonly less than 50% (e.g. Dickinson and Suczek, 1979;
Dickinson, 1982). The uplifted hinterlands of many, if not all, the Type
2 DWFTB provinces, both convergent zones and accretionary prisms,
include associated ophiolites and basic-intermediate volcanic pro-
vinces. C onseq uentl y, the lik elih ood of good quality r eserv oir
development is generally low. The poor quality, volcano-clastic
Neogene sandstones of Western Sulawesi (Bergman et al., 1996) are
the highest risk factor concerning exploration in the convergent zone
DWFTB. Conversely, the deltas of the NW and eastern Borneo margin
have high quality, quartz-rich deepwater sandstones (e.g. Ingram et
al., 2004; Saller et al., 2008). As discussed in Section 4.2.1.1, the reason
for this high reservoir quality is that the older accretionary prism
sediments were uplifted to form the main sediment source area, and
were themselves predominantly supplied by axially transported
sands (axially along the trench) from an uplifted continental area to
the west and southwest. Volcanic and lithic-rich sediments were
A
B
Growth fault
Depocentre
Deepwater fold and thrust belt
Onset of oil window
Immature source rocks
Mature source rocks
Migration pathway
Weak migration anticline
likely to be underfilled
Fig. 40. Model illustrating the effect of the dip of the basal detachment surface on
hydrocarbon migration pathways. A) Seaward dipping detachment, B) Landward
dipping detachment.
80 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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transported we stwards from a s eparate s ource area into the
accretionary prism, but were considerably diluted by sediments
derived directly or indirectly from the continental source area.
The hypothesis that, under some circumstances, accretionary
prisms can involve high-quality reservoir rocks is discussed in the
introduction to Section 4.3 Type 2bii Accretionary Prisms. The
justification for describing the Barbados, Makran and Indo-Burma
accretionary prisms is their potential for economically viable reservoir
rocks. In these areas, major fluvio-deltaic systems, with source areas
in continent–continent collision belts, have supplied the respective
accretionary prism with sediments and thus offer an alternative to the
typical accretionary prism sedimentary source terranes.
Even when good quality reservoir is present, turbidite sandstone
reservoirs are commonly strongly compartmentalized both by
structural features (faults, and possibly shale intrusions; Morley,
2003b), diagen esis, and sa nd body geomet ry (particularly for
channelised, and channel–levee complex sandstones, e.g. Weimer et
al., 2006; Saller et al., 2008). A number of deepwater projects (perhaps
most notably Thunder Horse and Mad Dog in the Gulf of Mexico) have
not produced at the peak rates anticipated, and production has
declined much more rapidly than predicted, both as a consequence of
the reservoir complexity, strength of the water drive, and technology
problems arising from remote, high pressure, high temperature
conditions (Cohen, 2007; http://www.theoildrum.com/node/6415).
Given the enormous investment required, the unanticipated problems
with understanding even the largest deepwater reservoirs, presents a
most significant challenge to the industry.
It was noted previously that the long, slow development of salt-
associated DWFTBs impacted the development of migration pathways.
This slow development also influences the reservoir associations of
salt-related DWFTBs compared with all types of shale-related DWFTBs.
Salt on passive or rifted margins was deposited in a restricted marine
environment, consequently the overlying sediments tend to range
from marginal to shallow marine to deepwater depositional environ-
ments (e.g. Guardado et al., 1989; Spathopoulos, 1996; Katz and Mello,
2000; Hudec and J acks on , 2004). The transition to deepwat er
deposition can occur over tens of millions of years permitting
reservoirs to develop in a range of environments, including shelfal
carbonates as well as clastics. Consequently the range of reservoir
types is much greater in salt DWFTBs than other settings. In most
shale-detachment deepwater folds the reservoir sands were deposited
prior to folding. Syn-kinematic turbidites tend to onlap the fold crests,
and occur in shallow piggyback or mini-basins only, and hence they
are not penetrated by crestal wells (e.g. Ingram et al., 2004; Morley and
Leong, 2008; Morley, 2009a). There are a greater variety of trap types
in salt related DWFTBs, for example syn-kinematic turbidite sands can
contain economic hydrocarbon accumulations in mini-basins in up-
dip pinch-out traps (e.g. Flemings and Lupa, 2004).
A final feature of salt-related DWFTBs is that fold growth begins
early in the post-salt depositional history, is slow, and the fold
wavelength tends to grow with time (Rowan et al., 2004). Since the
folds form positive features at the sea floor the growth of salt related
folds can affect sand distribution and transport pathways over a much
longer time-period (sometimes episodically over a 100 my time span)
than shale-detachment DWFTBs (generally up to ~10 my).
5.1.5. Structural traps
There are more similarities than differences in structural style
between the different types of DWFTB, except for the salt detach-
ments. The main characteristics have been listed in Table 3 and are
based on the descriptions of the DWFTBs in the previous sections. In
DWFTBs, four-way dip-closed anticlines or three-way dip closures
with a lateral thrust seal are the most likely structures to contain
economically viable hydrocarbons. Other types of prospective closure
can include oblique or lateral fault seals, and footwall prospects.
Four-way dip-closed anticlines typically arise from detachment
fold s with no associate d thru sts in either limb (
Figs. 22, 23),
detachment folds with break thrusts, or fault propagation folds
(Fig. 8B; Fig. 41). Salt diapirism may significantly modify the basic fold
geometry (Fig. 19). While detachment folds occur above both salt and
Table 3
Summary of key characteristics of DWFTBs. Fold and thrust style, I =imbricate thrusts, FPF=fault propagation fold, FBF=fault bend fold, DF (TF) = detachment fold associated with
thrust fault, PU = pop-up, BT = break thrusts (associated with folding), CF = crestal normal faults. Overpressure, BDC = burial disequilbrium compaction, C = chemical,
I=inflationary, S=slab-derived fluids.
Type 1a Type 1b Type 1 Type 2a Type 2bi Type 2bii
Shale detachment Salt detachment Very large deltas
Stress Near Field Near Field Near Field Near and far field Predominantly
far field
Predominantly far field
Presence of large delta Yes Generally no Yes Yes No Occasionally
Detached style Yes Yes Yes Yes Yes Yes
Detachment dip direction Oceanward Oceanward Landward Landward Landward Landward
Basement-involved deformation
below DWFTB
No No No No known example,
though it is possible
Yes No (?)
Mobile unit deformation style Absent or simple
shale diapirs,
mud pipes
Salt nappes,
canopies, diapirs,
squeezed anticlines
Salt nappes, canopies,
diapirs, squeezed
anticlines, complex and
simple shale diapirs
Mostly mud pipes,
complex, occasional
large mobile shale
masses (Sabah)
Mostly mud
pipes
Mostly mud pipes, but
stratified accretionary
complexes may pass
laterally into zones of
chaotic deformation (1)
Inversion of DWFTB-related
growth faults
No Yes, in deepwater Uncertain,
probably no (2)
Deepwater no,
nearshore yes
No Deepwater no, nearshore yes
Propagation of DWFTB Oceanward Landward propagation
can be important (3)
Oceanward, lateral
migration of DWFTB(4)
Oceanward Oceanward Oceanward
Duration 1's my–10's my Episodic up to
~100my duration
b 40 my? b 15 my? (5) b 15 my? 10's my–100 my+
Critical taper 3°b 0–3° 2°–3° 3°–8° 3°–11° 4°–11°
Fold and thrust style I, FPF, FBF, DF(TF) DF, I, FPF, PU, CF DF, I, PU, CF, FBF, FPF DF (TF), I, FPF, FBF DF (TF), I,
FPF, FBF
DF (TF), I, FPF, FBF.
Fold wavelength* 1–12 km 2–15 km 2–15 km 5–15 1–30 km 2–30
Dominant vergence Offshore Mixed Offshore Offshore Offshore Offshore
Overpressure BDC, C BDC, C BDC, I , C BDC, I, C BDC, I, C BDC, I, C, S
Shortening amount (km) b 20? b 100(7)
b 30 (8) b 30 (?) b 30? 10's–100's
Other features Inflation of basal
salt layer (9)
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shale detachments, they are most f requently found above salt
detachments (Figs. 22 and 23). The Perdido Fold Belt, Gulf of Mexico
stands out as a particularly well-developed example (e.g. Peel et al.,
1995; Trudgill et al., 1999). In a deepwater setting where risk
mitigation is vital, simple, large four-way dip closures are the most
prospective structures. As discussed by Trudgill et al. (1999) and
Kostenko et al. (2008), thrusts are sometimes over-emphasized in
seismic interpretations, so that many folds once interpreted as being
affected by thrusts in their limbs are now often interpreted as
detachment folds without thrusted limbs.
Most deepwater folds above a shale detachment are associated
with thrusts and commonly a component of lateral thrust fault closure
is required to render such prospects viable. The risks associated with
lateral seal are discussed below in the section on seal (Section 5.1.6).
Trap size is a fundamental parameter that may systematically vary
between the different types of DWFTB. Hamilton and De Vera (2009)
presented data for 23 DWFTBs that showed a considerable variation in
along-strike fold length and fold wavelength between Type 1a
(passive margin shale detachment), Type 1b (passive margin salt
detachment) and Type 2b (active margin shale detachment) DWFTBs.
These authors concluded that the average fold strike-length of salt-
detachment systems is half that of shale-detachment systems, whilst
average fold amplitudes and wavelengths were similar. For detach-
ment and buckle folds, the wavelength is strongly controlled by the
thickness (and rheological characteristics) of the controlling compe-
tent stratigraphic member. Hence stratigraphic thickness of the
section overlying the detachment is likely to have the most significant
impact on fold wavelength. Provinces that tend to have shallow
depths to detachment (e.g. Type 1ai and 1aii DWFTB) will generally
have shorter wavelengths than provinces prone to have deeper
depths to detachment (e.g. Type 1b and 2). The examples of the
wavelength range for large folds, as given in Fig. 11, indicate that the
structural setting and detachment lithology are largely immaterial.
The two largest wavelength provinces (South Caspian Sea, Zagros
Mountains) are both associated with 10 km +thickness of sedimen-
Classic fault bend fold model
Pure-shear fault bend fold
Simple-shear fault bend fold
Simple-shear
detachment fold
A
B
C
D
E
F
Footwall syncline
Break thrust
With time a more discrete
detachment may develop
Thickening by a range of
small-scale structural processes
within core of fold
Fig. 41. Forward models from starting points A) and B), of C) simple-shear fault-bend folds, E) pure-shear fault bend folds, E) simple-shear detachment fold, and F) classic fault bend
folds. Simple-shear folds do not require slip along a discrete detachment; a detachment zone undergoes an externally imposed bedding-parallel simple shear. In pure-shear fault-
bend folds, slip occurs along a discrete detachment, and a zone in the hangingwall undergoes shortening and thickening above the fault ramp. In the detachment-folding example,
heterogeneous shear occurred, with a basal detachment zone (like the simple-shear fault bend fold), but also layer thickening (like the pure-shear model). The formation of a break
thrust after initial folding results in folding and thickening in both the hangingwall and footwall areas of the break thrust. A–D and F based on Suppe et al. (2004), and Corredor et al.
(2005), E) from Morley, 2009b.
100
200
150
50
0
Indo-Burma Ranges
Seram
NW Borneo
South Caspian Sea
40 km
17 km
44 km
44 km
60 km
73 km
82 km
Perdido
First order fold length (km)
Nile Delta (Messinian)
Zagros
Gabon (Astrid)
Makassar Straits
Fig. 42. First-order fold length in DWFTBs. Solid lines are lengths of folds with one main
crest. Dashed lines are folds with multiple crests assumed to be formed by along-strike
linkage.
82 C.K. Morley et al. / Earth-Science Reviews 104 (2011) 41–91
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tary rock overlying the detachment. Despite not being a DWFTB, the
Zagros Mountains are shown in Fig. 42 since they are the best
representative of a thick sedimentary sequence developed above a
salt detachment. However, they also contain thick, Mesozoic and
Cenozoic carbonates that form a competent beam, which differentiate
them from most DWFTBs.
Figure 42 shows the length of anticlines for a number of thrust and
fold belts. The dashed line shows the length of the longest anticlines in
the respective belt. Such anticlines are generally characterized by 3–5
axial culminations and represent composite structures. The solid black
lines represent the length range for a single periclinal fold or fold
segment. Hamilton and De Vera (2009) noted that salt DWFTBs (Type
1b) were associated with relatively short anticlines, yet the great
length of the Zagros anticlines indicates that there is nothing intrinsic
to salt detachments that causes short fold lengths. While fold
wavelength and along-strike length are related to fold growth, other
factors can affect the lateral continuity of folds. These include amongst
others oblique and lateral ramps in the basal thrust, relief at the
detachment level due to underlying structures (e.g. syn-rift fault
systems) salt diapirs and welds, and lateral discontinuities in the basal
detachment horizon (salt, or shales). The low fold length in Type 1b
DWFTBs suggests that in passive margins salt detachment has to
contend with more along-strike obstacles than shale detachments.
These obstacles are primarily related to the structural setting under
which the salt-bearing sequence was deposited (early post-rift, or
syn-rift).
While the majority of compressional structures above shale
detachments are folds and related thrusts, simple fold or fold and
thrust belts above a salt detachment are more unusual. As the Gulf of
Mexico demonstrates, much more shortening in the outer wedge is
accommodated by diapirism, salt nappe and canopy emplacement
than by thrusting and folding. In this respect large classic salt
detachment fold-thrust provinces such as the Zagros do not resemble
DWFTBs.
5.1.6. Seal
Perhaps the greatest single risk inherent to deepwater structural
prospects are uncertainties about the seal integrity. Four factors
influencing seal integrity are considered here: top seal quality, oblique
and lateral faults, thrust fault seal and hydraulic fractures.
In Type 1a DWFTBs, prospective structures are typically associated
with shale or mudstone top seals, and if faults contribute towards
closure, lateral seals are provided either by reservoir/clay juxtaposi-
tion or by membrane seals (clay smears). As discussed above, large
thrust faults are commonly linked with a basal shale thrust
characterized by very high overpressures. These faults provide
conduits for fluids to rise and form feeder pipes for mud volcanoes
(e.g. Deville et al., 2006; Morley, 2009a; Barnes et al., 2010; Fig. 8).
Hence faults are a major risk as lateral seals. The young development
of many fold and thrust belts means that fault reactivation is a
significant liability for many structures. Older faults in the DWFTB that
have been inactive for some time have more potential to act as lateral
seals.
Lateral and oblique faults appear to be commonly associated with
hydrocarbon seeps, as evidenced by deepwater coring programs.
Hence, if such faults form a component of a structural closure, or
control the maximum hydrocarbon column height, their reactivation
risk is likely to be more significant than that of the frontal thrust. For
example, oblique faults are associated with the majority of hydrocar-
bon seeps in the Barbados Accretionary Prism (Dolan et al., 2004). The
occurrence of oblique and lateral faults is highly variable. In some
settings they can be closely spaced (e.g. accretionary prisms, Dolan et
al., 2004; Barnes et al., 2010, Type 1a slides, De Vera et al., 2010, and
thin sedimentary sequences overlying salt detachments, Clark and
Cartwright, 2009). Local topographic features, such as seamounts and
transform faults, that may lie below the basal detachment can also
cause the development of lateral and oblique faults (e.g. Niger delta
Cobbold et al., 2009; Morley, 2009b).
In deepwater environments, shale top sea ls are capable of
holding back large hydrocarbon columns (Kos tenko et al., 2008).
However, the top seal can be breached by fl
uid plumes and mud
volcano feeder pipes, which are commonly found at the crests of
anticlines (e.g. Deville et al., 2006; Mo rley, 2009a ; Warren et al., in
pre ss, Fig. 8B). Crestal n ormal faults resulting from tangential
stresses or from gravity-related stresses (e.g. Morley, 2007b)can
also breach top seal s.
According to Ingram et al. (2004), discoveries in the deepwater of
Sabah demonstrate that sufficient recent oil and gas charge was
available to fill young structures. These discoveries were made in
inactive, buried structures, the trap and seal integrity of which was
apparently maintained. In the same area there are other, more
actively growing folds that have not yet been tested, which may have
a greater trap integrity risk due to high fluid pressures and top seal
hydrofracturing (Ingram et al., 2004). Similarly the Bobo discovery in
the Niger Delta is on an inactive, buried anticline, while an adjacent,
recently active structure shows fluid venting to the surface (Cobbold
et al., 2009).
5.2. Structural development in different tectonic settings
The different tectonic settings described in this paper might be
expected to produce very di fferent DWFTBs. Howe ver, broad
comparisons suggest more similarities than differences (Table 3),
particularly between DWFTBs with a shale detachment irrespective of
their tectonic framework. The biggest differences occur not between
tectonic settings, but between salt and shale detachments. The key
differences include:
1) Deformation above salt detachments occurs as a result of
continental margin post-rift subsidence and loading by sediments
(deltas in particular) combined with uplift of the hinterland.
Deformation of a passive margin sedimentary prism above a shale
detachments results either from far-field stresses or sediment
loading; margin uplift is less significant as a driving mechanism.
2) Salt material properties, combined with a thick sediment overbur-
den, permit the development of salt diapirs, salt nappes and salt
canopies. These features can interfere with the development of
regular folds and thrusts in DWFTBs. Thrusts and imbricate wedges
can form in response to the activation of a basal shear along the salt
allochthon (Hudec and Jackson, 2009). Salt nappe development may
control the amount of salt available for fold detachments (Perdido vs.
Mississippi Fan fold belts). Detachment-type fold belts with little
associated thrusting are a feature of salt detachments.
3) Mobile shale provinces produce more extensive tracts of folds and
thrusts than salt detachments, and cannot develop extensive
nappes and canopies. Their equivalents are probably recycled
mudstones (originating from mud volcanoes) in the syn-kinematic
series. Even large shale cored diapirs may be rare features. With
increasing seismic resolution, most “shale diapirs” are revealed as
steep-flanked, tight anticlines that in some cases are pierced by
pipe-like features feeding mud volcanoes (Van Rensbergen et al.,
1999). Shale detachment provinces are associated with extensive
piercement features due to overpressured fluids (mud pipes, mud
chambers, mud volcanoes, gas chimneys). Changes in deformation
style in shale detachment provinces are, in part, related to the
nature of the detachment, whether it forms a thin, discrete zone of
sliding, or a thick mobile shale zone. Generally, shale detachment
folds are associated with thrusts (usually fault bend fold, fault
propagation fold, or break-thrust styles).
4) Vergence of structures above mobile shale detachments tends to
be seaward, although local variations can occur. The vergence of
structures above highly overpressured shale detachments is
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sometimes landwards. The vergence of structures above salt
detachments is considerably more variable.
5) In salt detachments the sequence of deformation can be ocean-
ward, but along the African margin buttressing of the salt province
in the deepwater area resulted in landward propagation of the
compressional belt (Hudec and Jackson, 2004; Jackson et al., 2008).
In a shale detachment setting, deformation tends to propagate
oceanward, although synchronous and out-of-sequence deforma-
tion within the wedge is common (e.g. Morley, 1988,2003a; Moore
et al., 2007; Cubas et al., 2008).
6) Shale detachments form DWFTBs that can be treated as critical
taper wedges. Salt detachments do not conform to a clear wedge
geometry (e.g. Davis and Engelder, 1985; Dahlen, 1990; Bilotti and
Shaw, 2005; Cubas et al., 2008; Fig. 18).
7) Overpressures are generally trapped below salt units. The onset of
high overpressures is abrupt (sometimes over a thickness of
centimeters) at the base of the salt unit (e.g. Mostofi and Gansser,
1957). In shale detachments overpressures progressively increase
within shales approaching the detachment. Highest overpressures
may be attained within the detachment, not below it.
8) On passive margins, shale detachment activity is related to the
location and duration of delta progradation. This can be short
(millions of years) for small deltas, may last for tens of millions of
years for large deltas. Upon termination of delta progradation, the
loss of overpressure along the detachment results in the
termination of DWFTB deformation. However, salt detachments
are able to keep reactivating simply due to material properties and
hence, the duration of activity (albeit episodic) can exceed 100 My
even in areas where n o major delta provides a continuous
sediment supply to the margin.
Differences in deformation style between the various tectonic
settings of shale detachment include the following:
1) DWFTBs in a Type 1 setting (both salt and shale detachments) are
characterized by a landward (shelfal) zone of extension (margin
sub-parallel SHmax direction) and an oceanward (slope) zone of
compression (margin sub-p erpendicular SHmax direction;
Fig. 43). DWFTBs in a Type 2a setting have an onshore-inner
shelf zone of compression/inversion, a narrow shelfal zone of
active extension and an oceanward (slope) zone of compression,
reflecting mixed near-field and far-field stresses affecting the
margin. Types 2aii and 2b DWFTBs are generally characterized by
compression throughout the wedge, because there is no near-field
effect from deltas.
Extensional fault Thrust
AB C
Maximum and
minimum horizontal
stress orientation
(extensional stress)
Maximum and
minimum horizontal
stress orientation
(compressional stress)
Near-field stress
driven systems
Mixed near- and far-
field stress driven systems
Far-field stress driven
systems
I
I
I
II
II
II
III
III
IIIIV IV
Z
X
Y
Fig. 43. Diagram illustrating idealized cross-sections and stress orientations in three main types of DWFTBs. A=near field stress systems, AI) gravity sliding on oceanward-dipping
system, AII) mixed oceanward and landward-dipping system, AIII) salt-detachment with diapirism and salt nappe formation (salt in black), AIV) map view of stress orientations.
B=mixed near-and far-field stresses, BI) continental thrust belt, (or possibly accretionary prism) linking with near-field stress system towards foreland, BII) collapse of thickened
continental crust driven by flow of crust at base towards hinterland (Z), and ‘thick-skinned’ far-field gravity collapse of crust (Y) linking with ‘thin-skinned’ near-field gravity driven
deformation within the sedimentary section (X), BIII) map view of stress orientations. C) Far-field stress diven systems, CI) accretionary prims, CII) continent–arc, or continent–
ophiolite collision, CIII) intracontinental convergent zone (inversion and thin-skinned thrusting of rift-post-rift basin section on margin of collision zone, CIV) map view of stress
orientations. Grey = sedimentary section involved in DWFTB.
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2) Far-field stresses tend to produce both thick-skinned and thin-
skinned deformation; in particular normal fault-bounded basins
are prone to inversion in Type 2, but not Type 1 settings.
3) The period of DWFTB development in Type 2a and 2bi settings can
be relatively short (around 10–15 My) since continent collision is
acting to shorten and uplift the DWFTB. The life span of individual
DWFTBs on Type 1 margins also tends to be less than 10 My either
because the activity of small deltas only lasts a few million years, or
because older DWFTBs become overstepped and buried during
progradation of large deltas (which develop over tens of millions
of years, and the present day deformation belt (such as in the Niger
Delta) may only represent the last 1 0 My of de velopment.
Accretionary prisms can contain a record of deformation lasting
many tens of millions of years.
Similarities between some of the different provinces are also
important.
1) It is striking that large growth-fault bounded depocentres can be
present in all settings in shale-detachment DWFTBs: e.g. Makran
(Type 2bii), NW Borneo (Type 2a), not just in Type 1 (Niger Delta).
However, the occurrences of large growth fault depocentres in
Type 2 settings appear to be confined to where large deltas have
interacte d w ith far-field driven systems. Growth faults in
accretionary prisms have been previously explained as a product
of collapse of over thickened critical wedges (Platt, 1986). Such a
mechanism may be one part of the explanation, but in the case of
the Makran (Fig. 36), to fill a large growth faults depocentre, with a
high proportion of sand-prone, shallow marine sediments also
implies an active deltaic system was present. Hence, at least for
certain stages in its development, the Makran may have developed
as a linked near-field and far-field driven system, similar to NW
Borneo.
2) Some shale detachments in Type 1 and 2 settings are associated
with DWFTB structures that show either mixed or landward
vergence (e.g. northern Indo-Burma Ranges; external fold and
thrust belt, Niger Delta; Cascadia accretionary prism). These
detachments are inferred to be super-weak due to very high
pore fluid pressures (Byrne and Hibbard, 1987; Underwood, 2002:
Steckler et al., 2008; Cobbold et al., 2009). The occurrence of very
similar fold geometries in the Indo-Burma Ranges and the external
Niger Delta, despite their different tectonic settings, is striking.
6. Conclusions
The classification of DWFTBs presented here is, in many aspects,
similar to those of Rowan et al. (2004), Krueger and Gilbert (2009)
and Hamilton and De Vera (2009). It identifies the differences in
deformation styles, duration of activity, and amount of shortening, the
megatectonic setting and whether the basal detachment zone is
located in weak shales or evaporites. Previous classifications assumed
that the tectonic setting could be equated to the driving mechanism of
DWFTBs (i.e. passive margins = near field stress driven systems,
collisional margins= far-field stress driven systems). This paper has
shown that such distinctions are not so clear-cut, and consequently,
has selected driving mechanism, not tectonic location as the primary
subdivision of the proposed classification. Of particular note is that
some DWFTBs in collision belts (e.g. NW Borneo), transpressional
zones (e.g. Columbus Basin, Trinidad) and possibly some accretionary
prisms (e.g. the Makran) are driven by mixed near-field and far-field
stresses, and not just by far-field stresses as assumed in previous
classifi
cations.
The effect of the megatectonic setting of a DWFTBs is particularly
marked in the domains of deformation duration, cause of deformation
driving stresses, amount of shortening and dimensions of the DWFTB.
Therefore, the megatectonic setting of DWFTBs is used as the first
subdivision criterion of the proposed class i fication. This review
highlights that DWFTBs can develop during all the stages of the
Wilson Cycle, from the passive margin stage, through the ocean
subduction stage to early-stage of continent–continent collision, and
in remnant post-collision basins such as the Mediterranean and Black
Sea. The second subdivision of the proposed classification distin-
guishes between salt and shale detachments. However, for Type 2
settings, there are no examples of modern DWFTBs associated with
evaporite detachments. Consequently,Type 2 detachments are only
shown with a shale detachment in Figure 2. The sub-division of near-
field stress systems into salt and shale detachments is very important
because they mark fundamental differences in structural style and, to
a lesser degree, differences in duration of tectonic activity (ability of
the detachment zone to reactivate) and propagation direction of
deformation (Table 3). Shale detachments are by no means homog-
enous, and further subdivision could consider on how different shale
detachment type affects the structural style (e.g. thin, well compacted
and overpressured; narrow weak minerals and hydrostatic pressure;
narrow, undercompacted and overpressured; broad, undercompacted
and overpressured).
The largest hydrocarbon provinces in each setting (Type 1a, Niger
Delta; Type 1b; Gulf of Mexico; Type 2a, NW Borneo; Type 2bii,
Caspian Sea) are not entirely representative of the more frequently
occurring members of that category. The Gulf of Mexico with its short
period of sea floor spreading, and sheltered location away from the
main Atlantic passive margin where sediments were funneled in from
Northern America is an atypical example of passive margins. The
Niger Delta with its broad zone of mobile shale activity (controver-
sially called diapirs), squeezed folds, and thick mobile shales is not
actually typical for most shale-detachment deltas. In NW Borneo, the
coincidence of a system of deltas with good-quality reservoir sands,
building out onto a convergent margin with a DWFTB driven by mixed
near-field and far-field stresses is atypical. In the Caspian Sea, the
Absheron Ridge is associated with a most unusual early-stage
accretio nary prism that de velo ped during a phase of l imit ed
subduction of oceanic crust spanning perhaps as little as 10 My. The
volume of continent-derived sediment that was deposited in the
South Caspian Sea is highly anomalous for accretionary prisms.
For Type 1 DWFTBs, the volume of accumulated sediments is
crucial for their hydrocarbon potential, particularly in terms of source
rock development and whether these have been buried sufficiently to
attain to maturity to expel hydrocarb ons during and after the
formation of structural traps. In this respect the hydrocarbon potential
of smaller deltas, or passive margins without significant deltas may be
limited. Even in DWFTBs with the greatest exploration success (Gulf
of Mexico, Niger Delta), the distribution of commercial hydrocarbon
accumulations is very uneven, raising questions about source rock
development and maturity in deepwater environments. The issue of
source rock maturity goes beyond just sediment thickness and also
includes the general decrease in heat flow passing oceanwards across
passive margins. Smaller Type 1a systems are also prone to develop
shorter-wavelength folds, or folds in the hangingwalls of imbricate
thrusts, as compared to large deltas. Such characteristics produce
relatively small trap volumes or hydrocarbon column heights, and
render prospects marginal to sub-economic in deepwater environ-
ments. Typical salt detachment passive margins tend to have poorly
developed DWFTBs; by far the best-developed ones are associated
with the very large deltaic province of the Gulf of Mexico.
Type 2a and 2bi DWFTBs such as NW Borneo and Timor involve
continental crust entering the subduction zone. Consequently the
geothermal gradient passing offshore is not subject to the same
oceanward decrease as on passive margins. The presence and
maturity of source rocks are still important issues, but may not be
as critical as on passive margins. Reservoir presence and quality is a
major concern for both Type 2a and 2b DWFTBs, and is a major
liability for some Type 2bi DWFTBs (e.g. West Sulawesi) and most
Type 2bii DWFTBs. However, reservoir quality need to be addressed
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on a case-by-case basis, since there are examples, such as NW Borneo,
Trinidad, South Caspian Sea and the Indo-Burma Ranges where good-
quality, continent-derived sands have entered the systems. Proximity
to a large continental area is a prerequisite. Even with good quality
reservoirs, compartmentalization can create significant problems for
economic exploitation of the reserves. As a general characteristic, the
number and size of structural traps is a highly favorable aspect of Type
2 DWFTBs, although in some areas relatively small wavelength folds
can develop, or the intensity of deformation may be too great. A
significant risk, particularly in Type 2b settings, is the presence of
oblique faults (commonly strike-slip faults) that appear to be non-
sealing, leak fluids and breach potential hydrocarbon traps.
Acknowledgements
Reviewers Peter Ziegler and Chris Talbot provided very useful,
detailed and encouraging reviews that must have been very time
consuming. These reviews hel ped considerably to improve the
manuscript and their contribution is gratefully acknowledged. John
Warren is thanked for an earlier review of the manuscript. Professors
Ahmed Chalouan and André Michard are gratefully acknowledged for
promptly providing comments about the tectonic setting of the Prerif
zone, and Tony Barber is thanked for his comments about the Timor
Sea. The contributions of Ros King, Richard Hillis, Mark Tingay and
Guillaume Backe to this article form TRaX record 109.
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