1]EarthByte Group, School of Geosciences, Madsen Buildg. F09, The University
of Sydney, NSW 2006, Australia 2]Global New Ventures, Statoil ASA, Forus,
Stavanger, Norway 3]formerly with GET GME TSG, Statoil ASA, Oslo, Norway

Christian Heine
<EMAIL_ADDRESS>
# Kinematics of the South Atlantic rift

Christian Heine  Jasper Zoethout  R. Dietmar Müller  [  [  [

###### Abstract

The South Atlantic rift basin evolved as branch of a large Jurassic-Cretaceous
intraplate rift zone between the African and South American plates during the
final breakup of western Gondwana. While the relative motions between South
America and Africa for post-breakup times are well resolved, many issues
pertaining to the fit reconstruction and particular the relation between
kinematics and lithosphere dynamics during pre-breakup remain unclear in
currently published plate models. We have compiled and assimilated data from
these intraplated rifts and constructed a revised plate kinematic model for
the pre-breakup evolution of the South Atlantic. Based on structural
restoration of the conjugate South Atlantic margins and intracontinental rift
basins in Africa and South America, we achieve a tight fit reconstruction
which eliminates the need for previously inferred large intracontinental shear
zones, in particular in Patagonian South America. By quantitatively accounting
for crustal deformation in the Central and West African rift zone, we have
been able to indirectly construct the kinematic history of the pre-breakup
evolution of the conjugate West African-Brazilian margins. Our model suggests
a causal link between changes in extension direction and velocity during
continental extension and the generation of marginal structures such as the
enigmatic pre-salt sag basin and the São Paulo High. We model an initial E–W
directed extension between South America and Africa (fixed in present-day
position) at very low extensional velocities from 140 Ma until late
Hauterivian times ($\approx$126 Ma) when rift activity along in the equatorial
Atlantic domain started to increase significantly. During this initial
$\approx$14 Myr-long stretching episode the pre-salt basin width on the
conjugate Brazilian and West African margins is generated. An intermediate
stage between $\approx$126 Ma and Base Aptian is characterised by strain
localisation, rapid lithospheric weakening in the equatorial Atlantic domain,
resulting in both progressively increasing extensional velocities as well as a
significant rotation of the extension direction to NE–SW. From base Aptian
onwards diachronous lithospheric breakup occurred along the central South
Atlantic rift, first in the Sergipe-Alagoas/Rio Muni margin segment in the
northernmost South Atlantic. Final breakup between South America and Africa
occurred in the conjugate Santos–Benguela margin segment at around 113 Ma and
in the Equatorial Atlantic domain between the Ghanaian Ridge and the Piauí-
Ceará margin at 103 Ma. We conclude that such a multi-velocity, multi-
directional rift history exerts primary control on the evolution of this
conjugate passive margins systems and can explain the first order tectonic
structures along the South Atlantic and possibly other passive margins.

The formation and evolution of rift basins and continental passive margins is
strongly depedendent on lithosphere rheology and strain rates (e.g. Buck et
al., 1999; Bassi, 1995). Strain rates are directly related to the relative
motions between larger, rigid lithospheric plates and thus the rules of plate
tectonics. A consistent, independent kinematic framework for the pre-breakup
deformation history of the South Atlantic rift allows to link changes in
relative plate velocities and direction between the main lithospheric plates
to events recorded at basin scale and might help to shed light on some of the
enigmatic aspects of the conjugate margin formation in the South Atlantic,
such as the pre-salt sag basins along the West African margin (e.g. Huismans
and Beaumont, 2011; Reston, 2010), the extinct Abimael ridge in the southern
Santos basin, and the formation of the São Paulo high. Over the past decades,
our knowledge of passive margin evolution and the sophistication of
lithospheric deformation modelling codes has substantially increased (e.g.
Huismans and Beaumont, 2011; Péron-Pinvidic and Manatschal, 2009; Rüpke et
al., 2008; Crosby et al., 2008; Lavier and Manatschal, 2006), as have the
accuracy and understanding of global and regional relative plate motion models
(e.g. Seton et al., 2012; Müller et al., 2008) for oceanic areas. However, the
connections between these two scales and and the construction of quantified
plate kinematic frameworks for pre-breakup lithospheric extension remains
limited due to the fact that no equivalent of oceanic isochrons and fracture
zones are generated during continental lithospheric extension to provide
spatio-temporal constraints on the progression of extension. Provision of such
kinematic frameworks would vastly help to improve our understanding of the
spatio-temporal dynamics of continental margin formation. The South Atlantic
basin with its conjugate South American and West African margins and
associated Late Jurassic/Early Cretaceous rift structures (Fig. 1) offers an
ideal testing ground to attempt to construct such a framework and link a
kinematic model to observations from marginal and failed rift basins. There
are relatively few major lithospheric plates involved, the motions between
these plates during the early extension phase can modeled using well-
documented intraplate rifts and the conjugate passive margins still exist in
situ. In addition, an extensive body of published literature exists which
documents detailed aspects of the conjugate passive margin architecture.

###  0.1 Aims and rationale

Following Reston (2010), previously published plate kinematic reconstructions
of the South Atlantic rift have only been able to address basic questions
related to the formation of the conjugate South Atlantic margins in pre-
seafloor spreading times and fall short of explaining the impact of plate
kinematic effects on rift margin evolution. In this paper, we present a
completely revised plate kinematic model for the pre-breakup and early
seafloor spreading history of the South Atlantic rift that integrates
observations from all major South American and African rifts. We alleviate
unnecessary complexity from existing plate tectonic models for the region –
such as a number of large, inferred intracontinental shear zones, for which
observational evidence is lacking – while providing a kinematic framework
which can consistently explain relative plate motions between the larger
lithospheric plates and smaller tectonic blocks. The model improves the full
fit reconstruction between South American plates and Africa, especially in the
southernmost South Atlantic and is able to explain the formation of the Salado
and Colorado Basins in Argentina, two Early Cretaceous-aged basins which
strike nearly orthogonal to the main South Atlantic rift, in the context of
larger scale plate motions. This allows us to link changes in plate motions
with well-documented regional tectonic events recorded in basins along the
major rift systems and to explain the formation of the enigmatic pre-salt sag
basins of the central South Atlantic in the context of a multi-direction,
multi-velocity plate kinematic history.

##  1 Plate reconstructions: data, regional elements, and methodology

We build a spatio-temporal kinematic framework of relative motions between
rigid lithospheric plates based on the inventory of continental rifts,
published and unplublished data from the conjugate passive margins and
information from oceanic spreading in the South Atlantic. We utilise the
interactive, open-source plate kinematic modelling software GPlates
(http://www.gplates.org) as an intergration platform and the Generic Mapping
Tools (http://gmt.soest.hawaii.edu; Wessel and Smith, 1998) to generate our
paleo-tectonic reconstructions (See Figs. 13–21 and electronic supplements).
High resolution vectorgraphics of our reconstructions along with the plate
model and geospatial data for use in GPlates are available in the
supplementary data as well as on the internet at
http://datahub.io/en/dataset/southatlanticrift.

###  1.1 Absolute plate motion model and timescale

The plate kinematic model is built on a hybrid absolute plate motion model
with a moving hotspot reference frame back to 100 Ma and an adjusted
paleomagnetic absolute reference frame for times before 100 Ma (Torsvik et
al., 2008). Global plate motion models are based on magnetic polarity
timescales to temporarily constrain plate motions using seafloor spreading
magnetic anomalies (e.g. Seton et al., 2012; Müller et al., 2008). Here, we
use the geomagnetic polarity timescale by Gee and Kent (2007, hereafter called
_GeeK07_) which places the young end of magnetic polarity chron M0 at 120.6 Ma
and the base of M0r at 121 Ma. Agreement across various incarnations of
geological time scales exists that the base of magnetic polarity chron M0r
represents the Barrêmian-Aptian boundary (e.g. Ogg and Hinnov, 2012; He et
al., 2008; Channell et al., 2000, and references therein). However,
considerable debate is ongoing about the absolute age of of the base of M0r
with proposed ages falling in two camps, one around 121 Ma (e.g. He et al.,
2008; Gee and Kent, 2007; Berggren et al., 1995) and a second one around
125-126 Ma (Ogg and Hinnov, 2012; Gradstein et al., 2004). Both choices affect
the absolute ages of the different pre-Aptian stratigraphic stages as
magnetostratigraphy is used to provide a framework for biostratigraphic
events. Our preference of using Gee and Kent (2007)’s 121.0 Ma as base for M0r
is based the arguments put forward by He et al. (2008) and on a review of
global seafloor spreading velocities, in which an older age assigned to the
base of M0r of 125.0 Ma (Gradstein et al., 2004) would result in significantly
higher spreading velocities or larger oceanic crust flux for the pre-M0
oceanic crust (Seton et al., 2009; Gee and Kent, 2007; Cogné and Humler,
2006). Additionally, Stone et al. (2008) report a reversed magnetic polarity
for a Cretaceous-aged dyke from Pony’s Pass quarry on the Falkland Islands
dated to 121.3$\pm$1.2 Ma (Ar-Ar dating). This further supports the notion of
tying the absolute age of M0r old and Base Aptian to closer to 121 Ma than to
125 or 126 Ma as postulated by Ogg (2012); Gradstein et al. (2004).

In pre-seafloor spreading, continental environments, well constrained direct
age control for relative plate motions does not exist. Here, stratigraphic
information from syn-rift sequences and subsidence data needs to be converted
to absolute ages in a consistent framework. As the correlation between
absolute ages and stratigraphic intervals has undergone several major
iterations over the past decades, published ages from South American and
African rift basins require readjustment. To tie stratigraphic ages to the
magnetic polarity timescale predominantly used for global plate kinematic
models, we have converted the estimates given by Gradstein et al. (1994, here
named _Geek07/G94_) and Gradstein et al. (2004, _Geek07/GTS04_) to the GeeK07
polarity chron ages (Fig. 2). We use Geek07/GTS04, which places Base Aptian
(Base M0r old) at 121 Ma (Fig. 2).

For published data we have converted both, stratigraphic and numerical ages to
the new hybrid timescale. A particular example for large differences in
absolute ages are the publications by Genik (1993, 1992) which are based on
the EXXON 1988 timescale (Fig. 2; Haq et al., 1987). Here, the base Cretaceous
is given as $\approx 133$ Ma, whereas the recent GTS timescales (Gradstein et
al., 2004) as well as our hybrid timescale place the base Cretaceous between
144.2 $\pm$ 2.5 and 142.42 Ma, respectively. Issues pertaining to correlate
biostratigraphic zonations across different basins further complicate the
conversion between stratigraphic and absolute ages regionally (e.g. Chaboureau
et al., in press; Poropat and Colin, 2012).

###  1.2 Continental Intraplate Deformation

Deformation between rigid continental plates has to adhere to the basic rules
of plate tectonics and can be expressed as relative rotations between a
conjugate plate pair for any given time (Dunbar and Sawyer, 1987). Depending
on data coverage and quality, a hierarchical plate model can be assembled from
this information (e.g. Ross and Scotese, 1988). In pre-seafloor spreading
environments, the quantification of horizontal deformation on regional scale
is complicated as discrete time markers such as oceanic magnetic anomalies or
clear structural features, such as oceanic fracture zones, are lacking.
However, rift infill, faulting and post rift subsidence allow – within
reasonable bounds – to quantify the amount of horizontal deformation (White,
1989; Gibbs, 1984; Le Pichon and Sibuet, 1981, e.g.). We assume that a
significant amount of horizontal displacement ($>$15 km) is preserved in the
geological record either as foldbelts (positive topographic features) or
fault-bounded sedimentary basins and recognised subsequently. Key elements of
our plate kinematic model are a set of lithospheric blocks which are non-
deforming during the rifting and breakup of western Gondwana (160–85 Ma; Fig.
1). These blocks are delineated using first order structures, such as main
basin-bounding faults, thrust belts or large gradients in reported sediment
thickness, indicative of subsurface faulting, in conjunction with potential
field and published data. The delineation of these blocks is described in
detail in later parts of this manuscript.

We then construct a relative plate motion model based on published data to
constrain the timing, direction, and accumulated strain in intraplate
deformation zones. This information is augmented with interpreted tectonic
lineaments and kinematic indicators (faults, strike slip zones) from various
publications (e.g. Matos, 2000; Genik, 1993; Matos, 1992; Genik, 1992; Exxon
Production Research Company, 1985) as well as potential field data from
publicly available sources (Andersen et al., 2010; Sandwell and Smith, 2009;
Maus et al., 2007) to further refine rigid block and deforming zone outlines.
Stage rotations were derived by identifying the predominant structural grain
and choosing an appropriate rotation pole which allows plate motions to let
relative displacement occur so that the inferred kinematics from strain
markers along the whole deforming zone (e.g. WARS/CARS) are satisfied. The
amount of relative horizontal displacement was then implemented by visual
fitting, using published data and computed extension estimates using total
sediment thickness.

Sediment thickness can serve as proxy for horizontal extension in rift basins
which experienced a single phase of crustal extension. Here, we calculate
total tectonic subsidence (TTS; Sawyer, 1985) by applying an isostatic
correction for varying sediment densities (Sykes, 1996) to the gridded total
sediment thickness (Exxon Production Research Company, 1985). Extension
factors based on total tectonic subsidence, assuming Airy isostasy throughout
the rifting process and not accounting for flexural rigidity, can be computed
along a set of parallel profiles across a rift basin, oriented parallel to the
main extension direction using the empirical relationship of Le Pichon and
Sibuet (1981):

$TTS=7.83(1-\frac{1}{\beta})$ (1) $\gamma=\frac{TTS}{7.83}$ (2)

where $\gamma=1-\frac{1}{\beta}$ and $\beta$ is the stretch factor. It has
been shown that this method allows to compute an upper bound for extension
factors in rift basins, in accordance with results from other methods (Barr,
1985; Heine et al., 2008) Values obtained using this methodology are likely to
overestimate actual extension estimates and main uncertainties in this
approach are input sediment thickness and the estimated sediment compaction
curves along with the assumption that a single rift phase created the basin.
We have applied this methodology for some of the major rift basins
(Muglad/Melut, Doba/Doseo/Bongor, Termit/Ténéré/Grein-Kafra, Gao, Salado,
Colorado; see Tab. Kinematics of the South Atlantic rift and supplementary
material) to constrain the amount of rift-related displacement. We assume the
bulk of the sediment thickness was accumulated during a single rift phase.
However, it is known that both CARS and WARS experienced at least one younger
phase of mild rifting and subsequent reactivation which has affected the total
sediment thickness (e.g. Genik, 1992; Guiraud et al., 2005) and hence will
affect the total amount of extension computed using this method. The extension
estimates extracted from individual rift segments are implemented by
considering the general kinematics of the whole rift system, as individual
rift segments can open obliquely and their margins are not necessarily
oriented orthogonally to the stage pole small circles.

Figure 5 shows estimates for a set of parallel profiles across the Termit
Basin, orthogonal to the main basin axis, ranging between c. 50–100 km. Where
possible, these estimates for extension were verified using published data
(Tab. Kinematics of the South Atlantic rift). Due to the inherent lack of
precise kinematic and temporal markers during continental deformation and
insufficient data we only describe continental deformation by a single stage
rotation (Table 1, Figs. 6 and 7).

###  1.3 Passive margins and oceanic domain

Over the past decade, a wealth of crustal-scale seismic data covering the
conjugate South Atlantic margins, both from industry and academic projects,
have been published (e.g. Blaich et al., 2011; Unternehr et al., 2010;
Greenroyd et al., 2007; Franke et al., 2007, 2006; Contrucci et al., 2004;
Mohriak and Rosendahl, 2003; Cainelli and Mohriak, 1999; Rosendahl and
Groschel-Becker, 1999). We made use of these data to redefine the location of
the continent-ocean boundary in conjunction with proprietary industry long
offset reflection seismic data (such as the ION GXT CongoSPAN lines,
http://www.iongeo.com/Data_Library/Africa/CongoSPAN/) as well as proprietary
and public potential field data and models (e.g. Sandwell and Smith, 2009;
Maus et al., 2007).

As some segments of this conjugate passive margin system show evidence for
hyperextended margins as well as for extensive volcanism and associated
seaward dipping reflectors sequences (SDRs), we introduce the “landward limit
of the oceanic crust” (LaLOC) as boundary which delimits relatively
homogeneous oceanic crust oceanward from either extended continental crust or
exhumed continental lithospheric mantle landward or SDRs where an
interpretation of the Moho and/or the extent of continental crust is not
possible. This definition has proven to be useful in areas where a classic
continent-ocean boundary (COB) cannot easily be defined such as in the distal
parts of the Kwanza basin offshore Angola (Unternehr et al., 2010), the
oceanward boundary of the Santos basin (Zalán et al., 2011) or along the
conjugate magmatic margins of the southern South Atlantic.

Area balancing of the Top Basement and Moho horizons has been used on
published crustal scale passive margin cross sections by Blaich et al. (2011)
to restore the initial pre-deformation stage of the margin, assuming no out-
of-plane motions and constant area. Initial crustal thickness estimates are
based on the CRUST2 crustal thickness model (Laske, 2004). While these
assumptions simplify the actual margin architecture and do not account for
alteration of crustal thickness during extension, Fig. 3 shows that the
differences between a choice of three different limits of the extent of
continental crust (minimum, COB based on Blaich et al., 2011, and LaLOC) only
has relatively limited effects on the width of the restored margin.

Considering the plate-scale approach of this study and inherent uncertainties
in the interpretation of subsalt structures on seismic data, these estimates
provide valid tie points for a fit reconstruction. The resulting fit matches
well with Chang et al. (1992)’s estimates for pre-extensional margin geometry
for the Brazilian margin (compare Figs. 13–21). Area balancing of the
continental basement (Top Basement to Moho) results in stretching estimates
ranging from 2.6-3.3 (Fig. 3). Some of the margin cross sections do not cover
the full margin width from unstretched continental to oceanic crust (e.g.
North Gabon and Orange sections). Here, we allow for a slight overlap in the
fit reconstruction. We have also carried out an extensive regional
interpretation of Moho, Top Basement and Base Salt reflectors on ION GXT
CongoSPAN seismic data to verify results from areal balancing of the published
data.

The conjugate South Atlantic passive margins are predominantely non-volcanic
in the northern and central part and volcanic south of the conjugate Santos/
Benguela segment (Blaich et al., 2011; Moulin et al., 2009, e.g). In the
volcanic margin segments, the delineation of the extent of stretched
continental crust is hampered by thick seaward-dipping reflector sequences
(SDRs) and volcanic build-ups. Previous workers have associated the landward
termination of the SDRs along the South American and southwestern African
continental margins with the “G” magnetic anomaly and a prominent positive
“large magnetic anomaly (LMA)” delineating the boundary between a transitional
crust domain of mixed extended and heavily intruded continental crust (Blaich
et al., 2011; Moulin et al., 2009; Gladczenko et al., 1997; Rabinowitz and
LaBrecque, 1979, e.g.). We follow previous workers in magmatically dominated
margin segments by using the seaward edge of SDRs and transition to normal
oceanic spreading for the location of our LaLOC.

Published stratigraphic data from the margins is integrated to constrain the
onset and dynamics of rifting along with possible extensional phases (e.g.
Karner and Gambôa, 2007). Little publicly available data from distal and
deeper parts of the margins exist which could further constrain the spatio-
temporal patterns of the late synrift subsidence.

To quantify oceanic spreading and relative plate motions between the South
American and Southern African plates before the Cretaceous Normal Polarity
Superchron (CNPS, 83.5–120.6 Ma) we use a pick database compiled by the
EarthByte Group at the University of Sydney, forming the base for the digital
ocean floor age grid (Seton et al., 2012).

We combine these data with the interpretations of Max et al. (1999) and Moulin
et al. (2009) and the WDMAM gridded magnetic data (Maus et al., 2007) to
create a set of isochrons for anomaly chrons M7n young (127.23 Ma), M4 old
(126.57 Ma), M2 old (124.05 Ma), and M0r young (120.6 Ma) using the magnetic
polarity timescale of Gee and Kent (2007). M sequence anomalies from M11 to M8
are only reported for the African side (Rabinowitz and LaBrecque, 1979)
whereas M7 has been identified on both conjugate abyssal plains closed to the
LaLOC (Moulin et al., 2009; Rabinowitz and LaBrecque, 1979). Oceanic spreading
and relative plate velocities during the CNPS are linearily interpolated with
plate motion paths only adjusted to follow prominent fracture zones in the
Equatorial and South Atlantic.

###  1.4 Reconstruction methodology

Deforming tectonic elements and their tectono-stratigraphic evolution from the
South Atlantic, Equatorial Atlantic, and intraplate rift systems in Africa and
South America such as intraplate basins, fault zones or passive margin
segments are synthesised using available published and non-published data to
constructed a hierarchical tectonic model (Fig. 9). Starting with the African
intraplate rifts, we iteratively refine the individual stage poles and fit
reconstructions based on published estimates and our own computations,
ensuring that the implied kinematic histories for a plate pair do not violate
geological and kinematic constraints in adajcent deforming domains. For
example, the choice of a rotation pole and rigid plate geometries which
describe the deformation between our Southern African and NE African plates to
explain the opening of the Muglad Basin is also required to match deformation
along the western end of the Central African Shear Zone in the Doba and Bongor
basins (e.g. Fig. 6).

After restoring the intra-African plate deformation, we iteratively refine the
tight fit reconstruction of South American plate against the NW African and
South African plates by reconstructing our restored profiles (Fig. 3), Chang
et al. (1992)’s restored COB, and key structural elements.

Subsequently, remaining basin elements and rigid blocks of the Patagonian
extensional domain (Salado, North Patagonian Massif, Deseado, Rawson, San
Julian and Falkland blocks) are restored to pre-rift/deformation stage and
integrated to achieve a full fit reconstruction for the latest Jurassic/Early
Cretaceous time in the southern South Atlantic (cf. Fig. 13).

##  2 Tectonic elements: Rigid blocks and deforming domains

Burke and Dewey (1974) pointed out that Africa did not behave as a single
rigid plate during Cretaceous rifting of the South Atlantic. They divided
Africa into two plates separated along the Benue Trough-Termit Graben (WARS in
Fig. 1). Subsequent work identified another rift system trending to the east
from the Benoue area (Genik, 1992; Fairhead, 1988, CARS in Fig. 1).

There is less evidence for Late Jurassic/Early Cretaceous deformation in South
America, although several continent-scale strike slip zones have been
postulated (see Moulin et al., 2009, and references therein). we define four
major plate boundary zones and extensional domains: the Central African
(CARS), West African (WARS), South Atlantic (SARS), and Equatorial Atlantic
Rift Systems (EqRS). We include a “Patagonian extensional domain”, composed of
Late Jurassic/Early Cretaceous aged basins and rigid blocks in southern South
America in our definition of the SARS (Fig. 1). From these four extensional
domains, only the SARS and EqRS transitioned from rifting to breakup, creating
the Equatorial and South Atlantic Ocean basins. In the next sections we review
timing, kinematics, type and amount of deformation for each of these domains.

###  2.1 Africa

The West African and Central African rift systems (Figs. 1 and 4; WARS & CARS)
and associated depocenters document extensional deformation between the
following continental lithospheric sub-plates in Africa, starting in the
Latest Jurassic/Early Cretaceous:

  1. 1.

Northwest Africa (NWA), bound to the East by the WARS/East Niger Rift and
delimited by the Central and Equatorial Atlantic continental margins to the
West and South, respectively (e.g. Guiraud et al., 2005; Burke et al., 2003;
Genik, 1992; Fairhead, 1988).

  2. 2.

Nubian/Northeast Africa (NEA), bound by the WARS to the West, and the
CARS/Central African Shear Zone to the South (Bosworth, 1992; Genik, 1992;
Popoff, 1988; Schull, 1988; Browne et al., 1985; Browne and Fairhead, 1983,
e.g.). To the East, Northeast and North this block is delimited by the East
African rift, the Red Sea and the Mediterranean margin, respectively.

  3. 3.

Southern Africa (SAf) is separated from NEA through the CARS and bound by the
East African Rift system to the East and Southeast. Its southern and western
Margins are defined by the South Atlantic continental passive margins (e.g.
Nürnberg and Müller, 1991; Unternehr et al., 1988).

  4. 4.

The Jos subplate, named after the Jos Plateau in Nigeria, is situated between
NWA, NEA and the Benoue Trough region. We define this plate along its western
margin by a graben system of Early Cretaceous age in the Gao Trough/Graben
area in Mali and the Bida/Nupe basin in NW Nigeria (Guiraud et al., 2005;
Guiraud and Maurin, 1992; Genik, 1992; Adeniyi, 1984; Cratchley et al., 1984;
Wright, 1968). The Benoue Trough and WARS delimit the the Jos subplate to the
South and East. As northern boundary we chose a diffuse zone through the
Iullemmeden/Sokoto Basin and Aïr massive, linking the WARS with the Gao Trough
area (Guiraud et al., 2005; Genik, 1992; Cratchley et al., 1984, e.g.).

  5. 5.

The Adamaoua (Benoue), Bongor and Oban Highlands microplates (Fig. 1) are
situated south of the Benoue Trough and north of the sinistral Borogop fault
zone. This fault zone defines the western end of the CARS as it enters the
Adamaoua region of Cameroon (Genik, 1992; Benkhelil, 1982; Burke and Dewey,
1974). Together with the Benoue Trough in the north, the Atlantic margin in
the west and the Doba, Bongor, Bormu-Massenya basins it encompasses a
relatively small cratonic region in Nigeria/Cameroon which has been termed
“Benoue Subplate” by previous workers (e.g. Moulin et al., 2009; Torsvik et
al., 2009). The Yola rift branch (YB in Fig. 1) of the Benoue Trough as well
as the Mamfe Basin (Mf) indicate significant crustal thinning (Fairhead et
al., 1991; Stuart et al., 1985) justifying a subdivision of this region into
the two blocks.

The CARS and WARS are distinct from earlier Karoo-aged rift systems which
mainly affected the eastern and southern parts Africa (Bumby and Guiraud,
2005; Catuneanu et al., 2005) but have presumably formed along pre-existing
older tectonic lineaments of Panafrican age (Daly et al., 1989).

####  2.1.1 Central African Rift System (CARS)

The eastern part of the CARS, consisting of the Sudanese Melut, Muglad and
Bagarra basins, forms a zone a few hundred kilometers wide with localised NW-
SE striking sedimentary basins which are sharply delimited to the north by the
so-called Central African Shear Zone (Bosworth, 1992; McHargue et al., 1992;
Schull, 1988; Exxon Production Research Company, 1985; Browne and Fairhead,
1983). Subsurface structures indicate NNW/NW-trending main basin-bounding
lineaments and crustal thinning with up to 13 km of Late Jurassic/Early
Cretaceous-Tertiary sediments (Fig. 4; Mohamed et al., 2001; McHargue et al.,
1992; Schull, 1988; Browne et al., 1985; Browne and Fairhead, 1983). Published
values for crustal extension in the Muglad and Melut basins in Sudan range
between 22-48 km (Browne and Fairhead, 1983), 15–27 km (McHargue et al., 1992)
in SW-NE direction and $\beta=1.61$ for an initial crustal thickness of 35 km,
resulting in 56 km of extension (Mohamed et al., 2001) with a first postrift
phase commencing in the Albian ($\approx$110 Ma; McHargue et al., 1992).

Dextral transtensional motions along the western part of the CARS during the
Early Cretaceous created the Salamat, Doseo and Doba basins (Fig. 4; Bosworth,
1992; McHargue et al., 1992; Schull, 1988). These depocenters locally contain
more than 8 km of Early Cretaceous to Tertiary sediments and are bound by
steeply dipping faults, indicative of pull-apart/transtensional kinematics for
the basin opening (Fig. 4; Genik, 1993; Maurin and Guiraud, 1993; Binks and
Fairhead, 1992; Genik, 1992; Guiraud et al., 1992; Exxon Production Research
Company, 1985; Browne and Fairhead, 1983). The structural inventory of these
basins largely follows old Panafrican-aged lineaments and has been reactived
during the Santonian compressive event (Guiraud et al., 2005; Janssen et al.,
1995; Maurin and Guiraud, 1993; Genik, 1992; Daly et al., 1989).

The Borogop Fault (Fig. 1 & 6) defines the western part of the Central African
Shear Zone and enters the cratonic area of the Adamaoua uplift in Cameroon,
where smaller, Early Cretaceous-aged basins such as the Ngaoundere Rift are
located (Fig. 4; Maurin and Guiraud, 1993; Plomerová et al., 1993; Guiraud et
al., 1992; Fairhead and Binks, 1991). The reported total dextral displacement
is estimated to be 40–50 km based on basement outcrops and around 35 km in the
Doseo Basin (Genik, 1992; Daly et al., 1989).

We have implemented a compounded total extension between 20–45 km between 140
and 110 Ma (early Albian) in the Muglad and Melut Basins between, generated
through rotation of the SAf block counterclockwise relative to NEA (Fig. 6,
Tab. Kinematics of the South Atlantic rift). The stage pole is located located
in the Somali Basin, south of the Anza Rift/Lamu Embayment, resulting in
moderate extension ($\approx$10 km) in the northern Anza Rift in Kenya which
is supported by observations faulting of Early Cretaceous age (Morley et al.,
1999; Reeves et al., 1987). The chosen stage rotation results in estimates of
distributed oblique extension/ sinistral transtension along the western end of
the CARS in the Bongor, Doba and Doseo Basins. Based on our choice of rigid
plates, this stage pole accounts for the opening of the Sudanese, Central
African Rifts as well as the Bongor basin.

Transtensional dextral displacement along the Borogop fault zone in our model
amounts to 40-45 km in the Doseo Basin which is in agreement with 25–56 km
extension reported from the Sudanese basins (Mohamed et al., 2001; McHargue et
al., 1992).

Stage poles and associated small circles for the NEA–SAf rotation are oriented
orthogonally to mapped Early Cretaceous extensional fault trends for the Doba
and Doseo Basin (Genik, 1992), and graben-bounding normal faults in the
Sudanese basins (Fig. 6). Other authors have used 70 km of strike
slip/extension for the CASZ and Sudan Basins, respectively (Moulin et al.,
2009), which is about double the amount reported (Genik, 1992; McHargue et
al., 1992). Torsvik et al. (2009) model the CARS but do not specify an exact
amount of displacement between their Southern African and the NE African sub-
plates.

####  2.1.2 West African Rift System (WARS)

The West African/East Niger rift (WARS) extends northward from the eastern
Benoue Trough region through Chad and Niger towards southern Algeria and Lybia
(Fig. 1). The recent Chad basin is underlain by a series of N–S trending rift
basins, encompassing the Termit Trough, N’Dgel Edgi, Tefidet, Ténéré, and
Grein-Kafra Basins containing up to 12 km of Early Cretaceous to Tertiary
sediments (Figs. 4 & 5; Guiraud et al., 2005; Guiraud and Maurin, 1992; Genik,
1992; Exxon Production Research Company, 1985). These basins are extensional,
asymmetric rifts, initiated through block faulting in the Early Cretaceous,
with a dextral strike-slip component reported from the Tefidet region (“Tef”
in Fig. 1; Guiraud and Maurin, 1992; Genik, 1992). The infill is minor
Paleozoic to Jurassic pre-rift, non-marine sediments and a succession of non-
marine to marine Cretacous clastics of up to 6 km thickness with reactivation
of the structures during the Santonian (Bumby and Guiraud, 2005; Guiraud et
al., 2005; Genik, 1992). Towards the north of the WARS, N–S striking fault
zones of the El Biod-Gassi Touil High in the Algerian Sahara and associated
sediments indicate sinistral transpression during the Early Cretaceous
(Guiraud and Maurin, 1992). The main rift development occurred during Genik
(1992)’s Phase 3 from the Early Cretaceous to Top Albian (130–98 Ma) with full
rift development by 108 Ma (Genik, 1992, using the EXXON timescale).

Early Cretacous sedimentation and normal faulting in the Iullemmeden/Sokoto
and Bida Basins in NW Nigeria and the Gao Trough in mali indicates that
lithospheric extension also affected an area NW of the Jos subplate and
further west of the WARS sensu strictu (Guiraud et al., 2005; Obaje et al.,
2004; Genik, 1993, 1992; Guiraud and Maurin, 1992; Adeniyi, 1984; Cratchley et
al., 1984; Petters, 1981; Wright, 1968). Reported sediment thicknesses here
range between 3–3.5 km for the Bida Basin (Obaje et al., 2004, Fig. 4). Our
definition of the WARS hence encompasses this area of diffuse lithospheric
extension.

Palinspastic restoration of 2-D seismic profiles across the Termit basin part
of the WARS yields extension estimates between 40–80 km based (using Moho
depths of 26 km; Genik, 1992). Our computed maximum extension estimates for
the WARS rift basins using total tectonic subsidences results in a maximum
cumulative extension of 90–100 km (Fig. 5, Tab. Kinematics of the South
Atlantic rift). We use 70 km of extension in the Termit Basin region and 60 km
in the Grein-Kafra Basin to accommodate relative motions between the Jos
Subplate and NEA between Base Cretaceous and 110 Ma. Fault and sediment
isopach trends indicate an E–W to slightly oblique rifting, trending NNW-SSE
for the main branch of the WARS. Our stage rotation between NWA and NEA
results in oblique, NNE-SSW directed opening of the WARS (Fig. 7). Other
workers have used 130 km of E–W directed extension in the South (Termit Basin)
and 75 km for northern parts (Grein/Kafra Basins; Torsvik et al., 2009)
between 132–84 Ma or 80 km of SW-NE directed oblique extension (Moulin et al.,
2009). For the Bida (Nupe) Basin/Gao Trough, we estimate that approximately 15
km of extension occurred between the Jos Subplate and NWA, resulting in a
cumulative extension between NWA and NEA of 85–75 km between 143 Ma and 110
Ma.

####  2.1.3 Benoue Trough

The Benoue Trough and associated basins like the Gongola Trough, Bornu and
Yola Basins are located in the convergence of the WARS and CARS in the
junction between the Northwest, Northeast and Southern African plates. The
tectonic position makes the Benoue Trough susceptible to changes in the
regional stress field, reflected by a complex structural inventory (Benkhelil,
1989; Popoff, 1988). Sediment thicknesses reach locally more than 10 km along,
with the oldest outcropping sediments reported as Albian age from anticlines
in the Upper Benoue Trough (Fairhead and Okereke, 1990; Benkhelil, 1989).
Subsidence in the Benoue Trough commences during Late Jurassic to Barrêmian as
documented by the Bima-1 formation in the Upper Benoue Trough (Guiraud et al.,
2005; Guiraud and Maurin, 1992).

The observed sinistral transtension in the Benoue Trough is linked to the
opening of the South Atlantic basin and extension in the WARS and CARS
(Guiraud et al., 2005; Genik, 1993, 1992; Fairhead and Binks, 1991; Fairhead
and Okereke, 1990; Benkhelil, 1989; Popoff, 1988; Benkhelil, 1982; Burke,
1976; Burke and Dewey, 1974). It follows that the onset of rifting and amount
of extension in the Benoue Trough is largely controlled by the relative
motions along the WARS and CARS. maximum crustal extension estimates based in
gravity inversion are 95 km, 65 km, and 55 km in the Benue and Gongola
Troughs, and Yola Rift with about 60 km of sinistral strike slip (Fairhead and
Okereke, 1990; Benkhelil, 1989). Rift activity is reported from the “Aptian
(or earlier)” to the Santonian (Fairhead and Okereke, 1990), synchronous with
the evolution of the CARS and WARS (Guiraud et al., 2005; Genik, 1992).

We here regard the Benoue Trough as product of differential motions between
NWA, NEA and the Adamaoua Microplate which is separated from South Africa by
the Borogop Fault zone. Deformation implemented in our model for the WARS and
CARS result in $\approx$20 km of N–S directed relative extension and about 50
km of sinistral strike-slip in the Benoue Trough between the Early Cretaceous
and early Albian.

The Mamfe Basin at the SW end of the Benoue Trough is separating the Oban
Highlands Block from the Adamaoua Microplate in the east (Fig. 1). It is
conjugate to NE Brazil. 20 km of extension are reported for this region which
we have restored in our plate model.

###  2.2 South America

The present-day South American continent is composed of a set of Archean and
Proterozoic cores which were assembled until the early Paleozoic, with its
southernmost extent defined by the Rio de la Plata craton (Fig. 8; Pángaro and
Ramos, 2012; Almeida et al., 2000). Large parts of South America, in contrast
to Africa, show little evidence for significant and well preserved, large
offset (10’s of km) intraplate crustal deformation during the Late Jurassic to
Mid-Cretaceous. In the region extending from the Guyana shield region in the
north, through Amazonia and São Francisco down to the Rio de la Plata Craton
there are no clearly identifiable sedimentary basins or compressional
structures with significant deformation reported in the literature which
initiated or became reactivated during this time interval.

The Amazon basin and the Transbrasiliano Lineament have been used as the two
major structural elements by various authors to accommodate intraplate
deformation of the main South American plate. Eagles (2007) suggests the
Solimões-Amazon-Marajó basins as location of a temporary, transpressional
plate boundary during South and Equatorial Atlantic rifting, where a southern
South America block is dextrally displaced by $\approx$200 km against a
northern block. The basin is underlain by old lithosphere of the Amazonia
Craton (Li et al., 2008; Almeida et al., 2000) which experienced one main
rifting phase in early Paleozoic times and subsequent, predominantely
Paleozoic sedimentary infill (Fig. 8; da Cruz Cunha et al., 2007; Gonzaga et
al., 2000; Matos and Brown, 1992; Nunn and Aires, 1988). It is covered by a
thin blanket of Mesozoic and Cenozoic sediments which show mild reactivation
with NE-trending reverse faults and minor dextral wrenching along its eastern
margin/Foz do Amazon/Marajó Basin during the Late Jurassic to Early Cretaceous
(da Cruz Cunha et al., 2007; Costa et al., 2001; Gonzaga et al., 2000).
Reactivation affecting the whole Amazon basin is reported only from the
Cenozoic (Azevedo, 1991; Costa et al., 2001). We do not regard this tectonic
element as a temporary plate boundary during formation of the South Atlantic.

The continental-scale Transbrasiliano lineament (TBL; Almeida et al., 2000)
formed during the Pan-African/Brasiliano orogenic cycle. It is a potential
candidate for a major accommodation zone for intraplate deformation in South
America (Feng et al., 2007; Pérez-Gussinyé et al., 2007), however, the amount
of accommodated deformation and the exact timing remain elusive (Almeida et
al., 2000). Some authors suggest strike slip motion along during the opening
of the South Atlantic between 60–100 km along this 3000 km long, continent-
wide shear zone, reaching from NE Brazil down into northern Argentina
(Aslanian and Moulin, 2010; Moulin et al., 2009; Fairhead et al., 2007). Along
undulating lineaments such as the TBL, any strike-slip motion would have
resulted in a succession of restraining and releasing bends (Mann, 2007),
creating either compressional or extensional structures in the geological
record. For comparison, the reported offset along the Borogop Shear zone in
the Central African Rift System ranges around 40 km during the Late
Jurassic–Early Cretaceous and created a series of deep ($>$6 km)
intracontinental basins (Doba, Doseo, Salamat – Fig. 4; Genik, 1992; McHargue
et al., 1992). While we do not refute evidence of reactivation of the TBL
during the opening of the South Atlantic, published geological and geophysical
data do not provide convincing support for the existence of a plate boundary
separating the South American platform along the Transbrasiliano lineament
during the opening of the South Atlantic.

In southern Brazil, previous authors have argued for the Carboniferous Paraná
Basin being the location for a large NW–SE striking intracontinental shear
zone to close the “underfit” problems in the southern part of the South
Atlantic (Moulin et al., 2012, 2009; Torsvik et al., 2009; Eagles, 2007;
Nürnberg and Müller, 1991; Unternehr et al., 1988; Sibuet et al., 1984). This
zone, obscured by one of the largest continental flood basalt provinces in the
world (Peate, 1997; White and McKenzie, 1989), has been characterised as R-R-R
triple junction with 100 km N-S extension (Sibuet et al., 1984), as Paraná-
Coehabamba shear zone with 150 km dextral offset (Unternehr et al., 1988), as
Parana-Chacos Deformation Zone with 60–70 km extension and 20-30 km of strike
slip (Nürnberg and Müller, 1991), as Paraná-Etendeka Fracture Zone – a
transtensional boundary with 175 km lateral offset (Torsvik et al., 2009), or
dextral strike-slip zone with 150 km strike slip and 70 km extension (Moulin
et al., 2009). Peate (1997) ruled out the possibility for a R-R-R triple
junction due to timing of magma emplacement and orientation of the associated
dyke swarm. The minimum amount of deformation proposed by previous authors is
20–30 km strike slip and 60–70 km extension (Nürnberg and Müller, 1991). In
analogy to the well documented CARS and the discussion of the TBL above, such
significant displacement should have resulted in a set of prominent basins
extending well beyond the cover of the Paraná flood basalt province and
manifested itself as major break along the South American continental margin,
similar to the Colorado and Salado basins further south. Although sub-basalt
basin structures have been reported (Eyles and Eyles, 1993; Exxon Production
Research Company, 1985), there is no evidence for a large scale continental
shear zone obscured by the Paraná large igneous province (LIP).

In our model we have subdivided the present-day South American continent in
the following tectonic blocks, partly following previous authors (Fig. 1; e.g.
Moulin et al., 2009; Torsvik et al., 2009; Macdonald et al., 2003; Nürnberg
and Müller, 1991; Unternehr et al., 1988):

  1. 1.

The main South American Platform (SAm), extending from the Guyana Craton in
the North to the Rio de la Plata Craton in the South, with the exception of
the NE Brazilian Borborema Province.

  2. 2.

The NE Brazilian Borborema Province block (BPB).

  3. 3.

The Salado Subplate, located between the Salado and Colorado Basins.

  4. 4.

The _Patagonian extensional domain_ south of the Colorado Basin, composed of
the rigid Pampean Terrane, North Patagonian Massif, Rawson block, San Julian
block, the Deseado block, and Malvinas/Falkland block, and the Maurice Ewing
Bank extended continental crust.

####  2.2.1 Northeast Brazil

The Borborema Province block (BPB; Figs. 1 and 8) is located in NE Brazil and
separated from the South American plate along a N-S trending zone extending
from the Potiguar Basin on the eastern Brazilian Equatorial Atlantic margin
southwards to the Recôncavo-Tucano-Jatobá rift (RTJ, Fig. 8). It is an
exception in the otherwise tectonically stable South American plate and a set
of isolated Early Cretaceous rift basins (e.g. Araripe and Rio do Peixe
basins) as well as abundant evidence of reactivated Proterozoic aged basement
structures and shear zones indicates distributed, but highly localised
lithospheric deformation during the opening of the South Atlantic rift (de
Oliveira and Mohriak, 2003; Matos, 2000; Mohriak et al., 2000; Chang et al.,
1992; Matos, 1992; Milani and Davison, 1988; Castro, 1987). Rifting commenced
in the Latest Jurassic/Berriasian and lasted until the Mid-Barrêmian, well
documented through extensive hydrocarbon exploration (Matos, 2000, 1999;
Szatmari and Milani, 1999; Magnavita et al., 1994; Chang et al., 1992; Milani
and Davison, 1988).

Similar to the smaller tectonic blocks in the Benoue Trough region, we regard
the lithospheric deformation affecting this block caused by the motions of the
larger surrounding tectonic plates. The Borborema province is “crushed” during
the early translation of South America relative to Africa with extension in
the West Congo cratonic lithosphere localised along existing and reactivated
basement structures leading to small, spatially confined basins such as the
Araripe, Rio do Peixe, Iguatu and Lima Campos.

We model the rifting in the Recôncavo-Tucano-Jatobá and Potiguar basins by
allowing for $\approx$40 and 30 km extension, respectively, through relative
motions between South America and the Borborema Province block between 143 Ma
and 124 Ma (Mid-Barrêmian).

####  2.2.2 Southern South America

The WNW-ESE striking Punta del Este Basin, the genetically related Salado
Basin adjacent to the South and the ENE-WSW trending Santa Lucia
Basin/Canelones Graben system, delimit our South American Block towards the
South (Fig. 8; Soto et al., 2011; Jacques, 2003; Kirstein et al., 2000;
Stoakes et al., 1991; Zambrano and Urien, 1970). Well data supports the onset
of syn-rift subsidence around the Latest Jurassic/Early Cretaceous and post-
rift commencing at Base Aptian (Stoakes et al., 1991). The rift-related
structural trend is predominantly parallel to the basin axis, indicating a
NNE-SSW directed extension. Sediment thicknesses reach 6 km with crustal
thicknesses around 20–23 km (Croveto et al., 2007) yielding stretching factors
of around 1.4. We have implemented 40 km of NE-SW transtension between 145 Ma
to Base Aptian for the eastern part of the basin, which is assumed to have
been linked towards the west by a zone of diffuse deformation with the General
Levalle basin. We have split the southern Rio de la Plata craton along the
syntaxis of the Salado/Punta del Este Basin between the South American block
and the Salado Sub-plate.

The rigid Salado block contains the Precambrian core of the Tandilia region
and the Paleozoic Ventania foldbelt (Pángaro and Ramos, 2012; Ramos, 2008) and
is delimited by the Late Jurassic/Early Cretaceous-aged Colorado, Macachín,
Laboulaye/General Levalle and San Luis basins in the South, Southwest and
West, respectively (Fig 8; Pángaro and Ramos, 2012; Franke et al., 2006;
Webster et al., 2004; Urien et al., 1995; Zambrano and Urien, 1970). Basement
trends deduced from seismic and potential field data indicate, similar to the
Salado Basin, E-W trending rift structures, orthogonal to the SARS and point
to to N-S directed extension/transtension (Pángaro and Ramos, 2012; Franke et
al., 2006, J. Autin personal communication, 2012). Pángaro and Ramos (2012)
estimate around 45 km (20%) N-S directed extension for the Colorado basin. Our
model assumes $\approx$50 km of NE-SW directed transtension for the basin from
150 Ma to Base Aptian when relative motions between Patagonian Plates and
South America cease (Somoza and Zaffarana, 2008).

The Colorado Basin marks the transition between the blocks related to the
South American Platform and the Patagonian part of South America (Pángaro and
Ramos, 2012), which we here summarise as Patagonian extensional domain. The
Patagonian lithosphere south of the Colorado Basin is composed of a series of
amalgamated magmatic arcs and terranes with interspersed Mesozoic sedimentary
basins (Ramos, 2008; Macdonald et al., 2003; Ramos, 1988; Forsythe, 1982). For
the purpose of this paper, the North Patagonian Massif, Rawson Block, Deseado
Block and Malvinas/Falkland Island Block are not separately discussed as
deformation in the Colorado and Salado basins largely accounts for clockwise
rotation of the Patagonian South America during the Late Jurassic to Aptian.

####  2.2.3 The Gastre Shear Zone

The Gastre shear zone is used in previous plate tectonic models as major
intracontinental shear zone, separating Patagonian blocks from the main South
American plate (Torsvik et al., 2009; Macdonald et al., 2003). However, no
substantial transtensional or transpressional features along this proposed
faults zone are recognised in this part of Patagonia, nor is the geodynamic
framework of southern South America favouring the proposed kinematics. A
detailed geological study of the Gastre Fault zone area lead von Gosen and
Loske (2004) to conclude that there is no evidence for a late Jurassic–Early
Cretaceous shear zone in the Gastre area. Our model does not utilise a Gastre
Shear Zone to accommodate motions between the South American and Patagonian
blocks.

##  3 Plate reconstructions

The plate kinematic evolution of the South Atlantic rift and associated
intracontinental rifts preserved in the African and South American plates, is
presented as self-consistent kinematic model with a set of finite rotation
poles (Table 1). In the subsequent description of key timeslices we refer to
Southern Africa (SAf) fixed in present day position. We will focus on the
evolution of the conjugate South Atlantic margins. For paleo-tectonic maps in
1 Myr time intervals please refer to the electronic supplements.

###  3.1 Kinematic scenarios

Plate motions are expressed in the form of plate circuits or rotation trees in
which relative rotations compound in a time-dependent, non-commutative way.
The core of our plate tectonic model is the quantified intraplate deformation
which allows us to indirectly model the time-dependent velocities and
extension direction in the evolving South Atlantic rift. Between initiation
and onset of seafloor spreading, the plate circuit for the South America plate
is expressed by relative motions between the African sub-plates (Fig. 9). The
kinematics of rifting are well constrained through structural elements and
sedimentation patterns, however, the timing of extension carries significant
uncertainties due to predominantely continental and lacustrine sediment
infill. Limited direct information from drilling into the deepest parts of
these rifts is publicly available. The regional evolution allows for a
relatively robust dating of the onset of deformation at the Base Cretaceous in
all major rift basins (e.g. Janssen et al., 1995), whereas the onset of post-
rift subsidence in the CARS and WARS, is not as well constrained and further
complicated through which have experienced subsequent phases of significant
reactivation (e.g. Guiraud et al., 2005).

The design of our plate circuit (Fig. 9) implies that the timing of rift and
post-rift phases significantly affect the resulting relative plate motions
between South America and Southern Africa. We have tested five alternative
kinematic scenarios by varying the duration of the syn-rift phase in the
African intracontinental rifts, and along the Equatorial Atlantic margins to
evaluate the temporal sensitivity of our plate model (Tab.2). The onset of
syn-rift was changed between 140 Ma and 135 Ma in the CARS and WARS, along
with the end of the syn rift phase ranging between 110 Ma and 100 Ma. For the
Equatorial Atlantic rift, we evaluated syn-rift phase onset times between 140
Ma and 132 Ma (Top Valanginian). Additionally, we included the plate model of
Nürnberg and Müller (1991) with forced breakup at 112 Ma (“NT91”; Torsvik et
al., 2009) in our comparison. Flowlines for each alternative scenario were
plotted and evaluated against observed lineations and fracture zone patterns
of filtered free-air gravity. The implied extension history for the conjugate
South Atlantic passive margins was used as a primary criterion to eliminate
possible alternative scenarios. In the preferred model PM1, rifting along
SARS, CARS, WARS, and EqRS starts simultaneously around the early Berriasian
(here: 140 Ma). It satisfies geological (timing and sequence of events) and
geophysical observations (alignment of flowlines with lineaments identified in
the gravity data) for all marginal basins. Alternatively tested models
introduced kinematics such as transpression in certain parts of the margins
for which are not supported by the geological record. All tested scenarios
(PM1-PM5), however, show a good general agreement, confirming the robustness
of our methodology of constructing indirect plate motion paths for the
evolution of the SARS through accounting for intraplate deformation.

The largest difference between our model preferred model PM1 and model NT91
are the start of rifting and the implied kinematic history for the initial
phase of extension. Relative motions between SAm and the African plates
commences at 131 Ma in NT91. This results in higher extensional velocites and
strain rates in NT91 for all extensional domains due to a shorter duration
($\Delta t$ = 12 Myrs) between the onset of plate motions and key tiepoint at
Chron M0 (the same across all models).

In the northern Gabon region, the flowlines for model NT91 indicate a rapid
initial NNE-SSW translation of South America relative to Africa by about 100
km for the time from 131–126 Ma (Fig. 10). The resulting transpression along
the northern Gabon/Rio Muni margin during this time interval, is not evidenced
from the geological record (e.g. Brownfield and Charpentier, 2006; Turner et
al., 2003). This initial extension phase is followed by a sudden E-W kink in
plate motions from 126 Ma to 118 Ma before the flowlines turn SW-NE, parallel
to our model(s) for the time of the CNPS. Our modeled extension history for
the Gabon margin implies an initial 40-60 km E–W directed rifting between
South America and Africa until 127 Ma, followed by a 40∘ rotation of the
extension direction and subsequent increase in plate velocities.

Along the Nambian margin, predicted initial extension directions for the
relative motions between South America–Africa and Patagonian terranes–Africa
diverge in the central Orange Basin, conjugate to the Colorado and Salado
Basins (Fig. 11). Northward from here, relative motions for South
America–Africa are directed WNW–ESE, while the Patagonian Terranes–Africa
motions south of the central Orange Basin imply opening of the southernmost
SARS in ENE-WSW direction. This is a result of the relative clockwise rotation
of the Patagonian Terranes away from South America ue to the rifting in the
Salado/Punta del Este and Colorado basins and accordance with structural
observations from the southern Orange Basin (H. Koopmann, personal
communication, 2012). These model predictions are supported by the trend of
the main gravity lineaments in contrast to the plate motions paths of model
NT91 which are not reconcilable with gravity signatures (Fig. 11).

In the Santos basin our preferred model PM1 predicts an initial NW-SE directed
extension in the proximal part, oriented orthogonally to the main gravity
gradients, Moho topography and proximal structural elements (Fig. 12; e.g.
Stanton et al., 2010; Chang, 2004; Meisling et al., 2001). In the western
Santos basin we model the initial extension phase to be focussed between the
São Paulo High/Africa and South America, until the onset of the third
extensional phase around the Barrêmian/Aptian boundary, a time when the
inferred oceanic Abimael spreading becomes extinct (Figs. 9 and 12; Scotchman
et al., 2010) and the São Paulo High is translated from the African to the
South American plate.

###  3.2 Fit reconstruction and the influence of Antarctic plate motions

The fit reconstruction (Fig. 13) is generated by restoring the pre-rift stage
along the intraplate WARS and CARS for the African sub-plates, and in the
Recôncavo-Tucano-Jatobá, Colorado and Salado Basins for the South American
sub-plates using estimates of continental extension (see Sect. 2). We then use
area balanced crustal-scale cross sections along the South Atlantic
continental margins, (Sect. 1.3, Fig. 3) in combination with the restored
margin geometry published by Chang et al. (1992) to construct pre-rift
continental outlines. Margins along the Equatorial Atlantic are generally
narrow (Azevedo, 1991) and associated with complex transform fault tectonics
and few available published crustal scale seismic data, the location of the
LaLOC here is only based on potential field data, with the expection of the
Demerara Rise-French Guiana-Foz do Amazon segment where we have utilised
Greenroyd et al. (2008, 2007). We allow on average 100 km of overlap between
the present-day continental margins in the EqRS (Fig. 14), compared to a few
hundred km in the central part of the SARS. Total stretching estimates for
profiles across the margins for the central and southern South Atlantic
segment range between 2.3–3.8 for the 10 profiles shown in Fig. 3. South
America is subsequently visually fitted against the West African and African
Equatorial Atlantic margins using key tectonic lineaments such as fracture
zone end points and margin offsets (Fig. 13). Transpressional deformation has
affected the conjugate Demerara Rise/Guinea Plateau submerged promontories,
resulting in shortening of both conjugate continental margins during the
opening of the SARS (Basile et al., 2013, 2005; Benkhelil et al., 1995). Our
reconstructions hence show a gap of $\approx$50 km between the Demerara Rise
and the Guinea Plateau at the western EqRS as we have not restored this
shortening (Fig. 13).

The dispersal of Gondwana into a western and eastern part was initated with
continental rifting and breakup along the incipient Somali-Mozambique-Wedell
Sea Rift (e.g. König and Jokat, 2006; Norton and Sclater, 1979). NE-SW
directed displacement of Antarctica as part of eastern Gondwana southwards
relative to South America and Africa creates an extensional stress field which
affects southernmost Africa and the present-day South American continental
promontory comprised of the Ewing Bank and Malvinas/Falkland Island and the
Proto-Weddell sea from the Mid-Jurassic onwards (König and Jokat, 2006;
Macdonald et al., 2003). Our pre-rift reconstruction for the Patagonian
extensional domain takes into account possible earlier phases of extension
which have affected, for example, the San Jorge, Deseado or San Julían Basins
(Jones et al., 2004; Homovc and Constantini, 2001) and hence should represent
a snapshot at 140 Ma, rather than a complete fit reconstruction which restores
all cumulative extension affecting the area in Mesozoic times. Initial rifting
of this region preceeding relative motions between South America and Africa,
is recorded by Oxfordian-aged syn-rift sediments in the Outeniqua Basin in
South Africa as well as subsidence and crustal stretching in the North
Falkland Basin and the Maurice Ewing Bank region (Fig. 8; Jones et al., 2004;
Paton and Underhill, 2004; Macdonald et al., 2003; Bransden et al., 1999). The
overall NE-SW extension causes a clockwise rotation of the Patagonian blocks
away from SAf and SAm commencing at $\approx$150 Ma in our model, with an fit
reconstruction position of the Falkland Islands just south Aghulas
Arch/Outeniqua basin. This early motion results in the onset of crustal
stretching in the North Falkland (150 Ma), Colorado (150 Ma) and Salado (145
Ma) basins which preceeds relative motions between the main African and South
American plates (Fig. 13). The resulting relative motion in the southern part
of the SARS is NE-SW directed transtension (Fig. 11), and E-W extension
between the Malvinas/Ewing Bank promontory and Patagonia. Our fit
reconstruction for the southernmost South Atlantic places the Malvinas/
Falkland Islands immediately south of the Aghulas Arch/Outeniqua Basin. Based
on published extension estimates for relevant basins in southern
Patagonia/offshore Argentina (Jones et al., 2004, e.g.), alternative pre-rift
reconstructions placing the Falkland Islands further east and excessive
microplate rotation (Mitchell et al., 1986) can not be reconciled with the
currently available data.

###  3.3 Phase I: Initial opening – Berriasian to late Hauterivian
(140–126.57 Ma)

Extensional deformation along the WARS, CARS and SARS is documented to start
in the Early Cretaceous (Berriasian) by the formation of intracontinental rift
basins and deposition of lacustrine sediments along the conjugate South
Atlantic margins (Chaboureau et al., in press; Poropat and Colin, 2012; Dupré
et al., 2007; Brownfield and Charpentier, 2006; Bate, 1999; Cainelli and
Mohriak, 1999; Coward et al., 1999; Karner et al., 1997; Chang et al., 1992;
Guiraud et al., 1992). Here, we use 140 Ma as absolute starting age of the
onset of deformation in the South Atlantic, Central and West African rift
systems (lower Berriasian in our Geek07/GTS04 hybrid timescale, Fig. 2). In
the southern part of the SARS, SW-directed extension has already commenced in
the latest Jurassic with syn-rift subsidence in the Colorado, Salado, Orange
and North Falkland Basins (Séranne and Anka, 2005; Jones et al., 2004; Clemson
et al., 1999; Maslanyj et al., 1992; Stoakes et al., 1991). Evidence for
deformation and rifting along the EqRS during the Early Cretaceous is sparse,
however, indications for magmatism and transpressional deformation are found
in basins along the margin (e.g. Marajó Basin, São Luis Rift; Soares Júnior et
al., 2011; Costa et al., 2002; Azevedo, 1991). The Ferrer-Urbanos-Santos Arch
is an E–W trending basement high between the proximal parts of the Barrerinhas
basin and the onshore Parnaiba basin which was generated by transpression
during the Neocomian (Azevedo, 1991).

Extension in all major rift basins occurs at slow rates during the initial
phase, with separation velocities between SAm and SAf around 2 mm a-1 in the
Potiguar/Rio Muni segment and up to c. 20 mm a-1 in the southernmost SARS
segment, closer to the stage pole equator (Fig. 15). Extensional velocities in
the magmatically dominated margin segments of the southern South Atlantic are
well above 10 mm a-1. The predominant extension direction changes from NW-SE
in the northern SARS segment to WNW-ESE in the southern part, and SSW-NNE for
the conjugate Patagonia-SAf segment (Figs. 15 and 13–16. Along the central
SARS segment, this phase corresponds to Karner et al. (1997)’s “Rift Phase 1”
with broadly distributed rifting.

Flowlines derived from the plate model for the early extension phase correlate
well with patterns observed in the free air gravity field along the proximal
parts of the margins, such as in the northern Orange, and inner Santos Basins
(Figs. 11, 12). Along NE–SW trending margin segments, such as Rio Muni/Sergipe
Alagoas and Santos/Benguela, the extension is orthogonal to the margin,
whereas oblique rifting occurs in other segments. In this context we note that
the strike of the Taubaté Basin in SE Brazil (“Tau” in Fig. 8) as well as
observed onshore and offshore extensional structures and proximal Moho uplift
in the Santos Basin (Stanton et al., 2010; Fetter, 2009; Meisling et al.,
2001; Chang et al., 1992) are oriented orthogonal to our modeled initial
opening direction of the South Atlantic rift. The Recôncavo-Tucano-Jatobá
(RTJ) rift opens as we model the BPB attached to the West African Congo Craton
and Adamaoua subplate, with SAm being relatively displaced northwestward.
Here, the Pernambuco shear zone acts as continuation of the Borogop Fault Zone
into NE Brazil (Matos, 1999). Potiguar Basin and Benoue Trough make up one
depositional axis while the RTJ and inner/northern Gabon belong to the central
South Atlantic segment.

Around 138 Ma (Mid-Berriasian), break up and seafloor spreading starts in
isolated compartments between the Rawson Block and the continental margin
south of the Orange Basin. The Tristan da Cuñha hotspot has been located
beneath the Pelotas-Walvis segment of the SARS since 145 Ma (Figs. 13–16, and
paleo-tectonic maps in electronic supplement) and likely caused significant
alteration of the lithosphere during the early phase rifting, with the
eruption of the Paraná-Etendeka Continental Flood basalt province occurring
between 138–129 Ma (Peate, 1997; Stewart et al., 1996; Turner et al., 1994).
While we have not included a temporary plate boundary in the Paraná Basin in
our model, evidence from the Punta Grossa and Paraguay dyke swarms (Peate,
1997; Oliveira, 1989) points to NE-SW directed extension, similar to the
“Colorado Basin-style” lithospheric extension orthogonal to the main SARS
(Turner et al., 1994). Emplacement of massive seaward dipping reflector
sequences dominate the early evolution of the southern SARS segment and are
well documented by seismic and potential field data (Blaich et al., 2011;
Moulin et al., 2009; Bauer et al., 2000; Gladczenko et al., 1997, e.g.). While
the delineation of the extent of continental crust along volcanic margins
remains contentious, we model a near-synchronous transition to seafloor
spreading south of the Walvis Ridge/Florianopolis High around 127 Ma (Figs.
16–17 and supplementary material).

Between 136–132 Ma, the plume center is located beneath the present-day South
American coastline in the northern Pelotas Basin. Our reconstructions indicate
that the westernmost position of a Tristan plume (assumed to be fixed) with a
diameter of 400 km is more than 500 km east of the oldest Paraná flood
basalts, at the northern and western extremities of the outcrop (Turner et
al., 1994). Oceanic magnetic anomalies off the Pelotas/northern Namibe margins
indicate asymmetric spreading with a predomiant accretion of oceanic crust
along the African side (Moulin et al., 2009; Rabinowitz and LaBrecque, 1979),
which in our reconstructions can be explained through plume-ridge interactions
of high magma volume fluxes and relatively low spreading rates (Fig. 15;
Mittelstaedt et al., 2008). The São Paulo – Río de Janeiro coastal dyke
swarms, emplaced between 133–129 Ma, have any extrusive equivalents in the
Paraná LIP (Peate, 1997). The dykes are oriented orthogonal to our modeled
initial extension direction, both along the Brazilian as well as along the
African margin in southern Angola. Along with a predominant NE–SW striking
metamorphic basement grain presenting inherited weaknesses (Almeida et al.,
2000), these dykes are probably related to lithospheric extension along the
conjugate southern Campos/Santos–Benguela margin segment of the SARS.

In the southern segment, the Falkland-Aghulas Fracture zone is established
when the Maurice Ewing block starts moving with the Patagonian plate around
134 Ma. Strain rate estimates from the Orange basin support the onset of
rifting around 150-145 Ma (Jones et al., 2004).

Towards the late Hauterivian (127 Ma), the width of the enigmatic Pre-salt Sag
basin width has been created by slow, relative extension between SAm and SAf,
with extension velocities ranging between 7–9 mm a-1 (Campos/Jatobá, Fig. 15).

###  3.4 Phase II: Equatorial Atlantic rupture (126.57–120.6 Ma)

Following the initial rifting phase which is characterised by low strain
rates, deformation along the EqRS between NWA and northern SAm intensifies due
to strain localisation and lithospheric weakening (Heine and Brune, 2011).
Marginal basins of the EqRS record increasing rates of
subsidence/transpression (Azevedo, 1991) and in the southern South Atlantic
seafloor spreading anomalies M4 and M0 indicate a 3-fold increase in relative
plate velocities between SAm and African Plates (Figs. 17 and 15). Velocity
increases of similar magnitude, albeit with a different timing, are reported
by Torsvik et al. (2009); Nürnberg and Müller (1991). Along the EqRs, Azevedo
(1991) describes a set of transpressional structures in the Barreirinhas Basin
which caused folding of Albian and older strata and reports 50–120 km of
dextral strike slip along the Sobradinho Fault in the proximal Barrerinhas
basin, pointing towards an early dextral displacement. With the changed
kinematics, SAm now rotates clockwise around NWA, causing transtensional
opening of basins in the eastern part of the EqRS and about 20 km of
transpression along the Demerara Rise and Guinea Plateau at the western end of
the EqRS. This is in accordance with observed compressional structures along
the southern margin of the Guinea Plateau and the northeastern margin of the
Demerara Rise (Basile et al., 2013; Benkhelil et al., 1995).

The increase in extensional velocities jointly occurs with a sudden change in
extension direction from NW–SE to more E–W (Fig. 15). Depending on the
distance from the stage pole for this time interval, the directional change is
between 75∘ in the northern SARS (Rio Muni-Gabon/Potiguar-Sergipe Alagoas
segment) and 30∘ in the southern Pelotas/Walvis segment. The directional
change visible in the flowlines agrees very well with a pronounced outer
gravity high along the Namibian margin (Fig. 11) and lineaments in the Santos
Basin (Fig. 12). This change in the plate motions had severe effects on the
patterns and distribution of extension in the SARS. Karner et al. (1997)
report a 100 km westward step of the main axis of lithospheric extension in
the Gabon/Cabinda margin segment during their second rift phase (Hauterivian
to late Barrêmian). Along the Rio Muni/ North Gabon margin a mild inversion in
the pre-breakup sediment is observed (Lawrence et al., 2002) which we relate
to the change in plate kinematics, manifested in this margin segment as change
from orthogonal/slightly oblique extension to transform/strike slip. In the
Santos Basin, the change in extensional direction results in transtensional
motions, most likely responsible for the creation of the Cabo Frio fault zone
and localised thinning in proximal parts of the SW Santos basin, now
manifested as an “Axis of Basement Low” (Modica and Brush, 2004) north of the
São Paulo High. By the time of change in plate motions ($\approx$127 Ma), the
pre-salt sag basin width had been fully generated (Fig. 17). Increasing
extensional velocities resulted in fast, highly asymmetric localisation of
lithosphere deformation in the northern and central SARS segments, most likely
additionally influenced by the inherited basement grain/structures. This has
left large parts of the Phase I rift basin preserved along the Gabon-Kwanza
margin, whereas south of the Benguela/Cabo Frio transform the Phase I rift is
largely preserved on the Brazilian side.

Davison (2007) reports that from about 124 Ma (Early Aptian using Gradstein et
al., 2004), extensive evaporite sequences are deposited in areas north of the
Walvis Ridge, starting with the Paripuiera salt in the northern
Gabon/Jequitinhonha segment and the slightly younger Loeme evaporites in the
Kwanza basin. We note that in our reconstructions, the massive volcanic build
ups of the Walvis Ridge/ Florianopolis Platform form a large barrier toward
the Santos basin in the north. North of the ridge, exension localised and
propagated northwards into the southwestern part of the Santos Basin (the now
aborted Abimael ridge; Scotchman et al., 2010; Mohriak et al., 2010; Gomes et
al., 2009) towards the end of this phase.

The São Paulo High remains part of the African lithosphere, which explains the
“Cabo Frio counterregional fault trend” in the northern Santos Basin (Modica
and Brush, 2004). Prior NW-SE extension, paired with additional thinning and
transform motions along the Cabo Frio fault system (Stanton et al., 2010;
Meisling et al., 2001) could have formed a shallow gateway through the inner
parts of the Santos Basin, allowing the supply of seawater into the isolated
central SARS. Dynamic, plume-induced topography (Rudge et al., 2008) from the
extensive magmatic activity in the Paraná-Etendeka Igneous province might have
caused additional lithospheric buoycancy of the extending lithosphere in the
Santos basin.

Seafloor spreading is fully established in a Proto-South Atlantic ocean basin
(Pelotas-Walvis segment to Falkland/Aghulas fracture zone), while oceanic
circulation in this segment was probably quite restricted as the
Falkland/Malvinas–Maurice Ewing Bank continental promontory had not cleared
the Southern African margin (Fig. 18).

###  3.5 Phase III: Break up – Aptian to late Albian (120.6–100 Ma)

Magnetic anomaly chron M0 (120.6 Ma) is clearly identified in large parts of
the southern South Atlantic and provides, along with established fracture
zones, one major tiepoint for our reconstruction. Further lithospheric weaking
and strain localisation along the EqRS resulted in widespread transtensional
motions and related complex basin formation in the marginal basins (Basile et
al., 2005; Matos, 2000, 1999; Azevedo, 1991) causing a second increase in
extensional velocities between the African and SAm plates and a minor change
in separation direction from 120.6 Ma onwards (Fig. 15). We here linearly
interpolate the plate velocities during the CNPS (120.6–83.5 Ma) only
adjusting plate motion paths along well established fracture zones.

By 120 Ma (early Aptian, Fig. 18) continental breakup has occurred in the
northernmost SARS between the Potiguar/Benin and Pernambuco–Paraiba/Rio Muni
margin segments and between the conjugate northern Gabon/Jequitinhonha–Camamu
margins with incipient breakup in the remaining part of the central SARS
(Espirito Santo/Cabinda-Campos/Kwanza). South of the Cabo Frio/Benguela
transform, changes in plate motions resulted in deformation to jump from the
Avedis and Abimael ridges in the southwestern Santos basin (Scotchman et al.,
2010; Mohriak et al., 2010; Gomes et al., 2009) towards the African side,
rifting the São Paulo High away from the Namibian/Benguela margin. Extensive
evaporite deposition continues in most parts of the deforming central SARS,
peaking in Mid-Aptian times (Karner et al., 2007; Davison, 2007).

By 115 Ma (Middle Aptian, Fig. 19), large parts of the SARS and EqRS have
broken up diachronously and entered post-rift thermal subsidence. Our model
predicts breakup in the Campos/Kwanza segment by 119 Ma, in the westernmost
part of the EqRS at 118 Ma, in line with deep water basin conditions reported
between the Guinea and St. Paul Fracture zones (Jones, 1987). In the EqRS our
model predicts accretion of oceanic lithosphere in the Deep Ghanaian Basin
from 117 Ma onwards, and break up along the southern Campos/Benguela margin
segment in the SARS around 115 Ma. For the Gabon margins, Dupré et al. (2007)
and Karner et al. (1997) assume the onset of rifting at around 118 Ma, and
Late Barrêmian–Early Aptian, respectively, which is in agreement with our
model. Subsidence data from nearly all margin segments in the SARS indicate
cessation of fault activity and a change to post-rift thermal subsidence by
this time, with the exception of the outer Santos Basin, where the final
breakup between SAf and SAm occurs between 113–112 Ma. This timing is in
agreement with the deposition of the youngest evaporites in the outer Santos
Basin around 113 Ma (Davison, 2007). We hypothesise that the early salt
movement in the Gabon, Kwanza, Espirito Santo, Campos and Santos basins and
the observed chaotic salt in the distal part of these basins is related to the
fast localisation of lithospheric deformation, break-up and subsequent rapid
subsidence during the early Aptian, introducing topographic gradients
favouring gravitational sliding and downslope compression in the earliest
postrift (Fort et al., 2004).

Relative plate motions in the CARS and WARS cease around 110 Ma in the early
Albian, with most intracontinental rift basins entering a phase of thermal
subsidence before subsequent minor reactivation occurred in Post-Early
Cretaceous times (Janssen et al., 1995; Genik, 1993; Maurin and Guiraud, 1993;
Genik, 1992; Guiraud and Maurin, 1992; McHargue et al., 1992). The cessation
of rifting in WARS and CARS results in the onset of transpression along the
Côte d’Ivoire-Ghanaian Ridge in the EqRS as the trailing edges of SAm continue
to move westward, while NWA now remains stationary with respect to the other
African plates. The NWA transform margins now provide a backstop to the
westward-directed motions of SAm, causing compression associated with up to 2
km of uplift to occur along the Ghanaian transform margin between the middle
to late Albian (Antobreh et al., 2009; Basile et al., 2005; Pletsch et al.,
2001; Clift et al., 1997). Complete separation between African and South
American continental lithospheres is achieved at 104 Ma (Fig. 21), while the
oceanic spreading ridge clears the Côte d’Ivoire/Ghana Ridge by 100–99 Ma.

We present a new plate kinematic model for the evolution of the South Atlantic
rifts. Our model integrates intraplate deformation from temporary plate
boundary zones along the West African and Central African Rift Systems as well
as from Late Jurassic/Early Cretaceous rift basins in South America to achieve
a tight fit reconstruction between the major plates. Our plate motion
hierarchy describes the motions of South America to Southern Africa through
relative plate motions between African sub-plates, allowing to to model the
time-dependent pre-breakup extension history of the South Atlantic rift
system. Three main phases with distinct velocity and kinematics result through
the extension along the Central African, West African and Equatorial Atlantic
rift systems exerting significant control on the dynamics of continental
lithospheric extension in the evolving South Atlantic rift from the Early
Cretaceous to final separation of Africa and South America in the late Albian
(104 Ma). An intial phase of slow E–W extension in the South Atlantic rift
basin from 140 Ma (Base Cretaceous) to 127 Ma (late Hauterivian) causes
distributed extension in W–E direction. The second phase from 126–121 Ma (late
Hauterivian to base Aptian) is characterised by rapid lithospheric weakening
along the Equatorial Atlantic rift resulting in increased extensional
velocities and a change in extension direction to SW–NE. In the last phase
commences at 120 Ma and culminates in diachronous breakup along the South
Atlantic rift and formation of the South Atlantic ocean basin. It is
characterised by a further increase in plate velocities with a minor change in
extension direction.

We argue that our proposed three-stage kinematic history can account for most
basin forming events in the Early Cretaceous South Atlantic, Equatorial
Atlantic, Central African and West African Rift Systems and provide a robust
quantitative tectonic framework for the formation of the conjugate South
Atlantic margins. In particular, our model addresses the following issues
related to the South Atlantic basin formation:

  * •

We achieve a tight fit reconstruction based on structural restoration of the
conjugate South Atlantic margins and intraplate rifts, without the need for
complex intracontinental shear zones.

  * •

The orientation of the Colorado and Salado Basins, oriented perpendicular to
the main South Atlantic rift are explained through clockwise rotation of the
Patagonian sub-plates. This also implies that the southernmost South Atlantic
rift opened obliquely in a NE–SW direction.

  * •

A rigid South American plate, spanning the Camamu to Pelotas basin margin
segments, requires that rifting started contemporaneously in this segment of
the South Atlantic Rift (Gabon, Kwanza, Benguela and Camamu, Espirito Santo,
Campos and Santos basins), albeit so far stratigraphic data (e.g. Chaboureau
et al., in press) indicates differential onset of the syn rift phase in the
various marginal basins.

  * •

The formation of the enigmatic pre-salt sag basin along the West African
margin is explained through a multi-phase and multi-directional extension
history in which the initial sag basin width was created through slow (7–9 mm
a-1) continental lithospheric extension until 126 Ma (Late Hauterivian).

  * •

Normal seafloor spreading in the southern segment of the South Atlantic rift
commenced at around 127 Ma, after a prolonged phase of volcanism affecting the
southern South Atlantic conjugate margins.

  * •

Strain weakening along the Equatorial Atlantic Rift caused an at least 3-fold
increase of plate velocities between South America and Africa between 126–121
Ma, resulting in rapid localisation of extension along the central South
Atlantic rift.

  * •

Linear interpolation of plate motions between 120.6 Ma and 83.5 Ma yield
break-up ages for the conjugate Brazilian and West African margins which are
corroborated by the onset of post-rift subsidence in those basins.

We conclude that our multi-direction, multi-velocity extension history can
consistently explain the key events related to continental breakup in the
South Atlantic rift realm, including the formation of the enigmatic pre-salt
sag basin.

The plate kinematic framework presented in this paper is robust and
comprehensive, shedding new light on the spatio-temporal evolution of the
evolving South Atlantic rift system. However, problems remain with regard to
the absolute timing of events in the evolution of the South Atlantic rift due
to a lacking coherent stratigraphy for both conjugate margin systems, and
limited access to crustal scale seismic data. Relatively old and limited data
for the African intraplate rifts results in large uncertainties in both
absolute timing as well as structural constraints on the kinematics of rifting
which we have attempted to minimise in this comprehensive approach.

Strengthening the kinematic framework for the South Atlantic rift also offers
an ideal laboratory to investigate the interaction between large scale plate
tectonics and lithosphere dynamics during rifting, and the relationship
between plume activity, rift kinematics and the formation of volcanic/non-
volcanic margins.

Data and high resolution images related to the publication are provided with
the manuscript and can be freely accessed under a Creative Commons license at
the Datahub under the following URL:
http://datahub.io/en/dataset/southatlanticrift.

###### Acknowledgements.

We would like to thank Statoil for permission to publish this work. CH and JZ
are indebted to many colleagues in Statoil’s South Atlantic new venture teams,
specialist groups, and R&D who have contributed in various aspects to an
inital version of the model. In particular we would like to thank Hans
Christian Briseid, Leo Duerto, Rob Hunsdale, Rosie Fletcher, Eric Blanc, Tommy
Mogensen Egebjerg, Erling Vågnes, Jakob Skogseid, and Ana Serrano-Oñate for
discussions and valuable feedback. Feedback from Patrick Unternehr, Ritske
Huismans, Julia Autin and Hannes Koopmann on various aspects of our work are
kindly acknowledged. CH thanks John Cannon and Paul Wessel for numerous fast
software bug fixes. Editor Douwe J. J. v. Hinsbergen, an anonymous reviewer
and Dieter Franke are thanked for constructive reviews and comments which
improved the discussion version of the paper. Dieter Franke also kindly
supplied digitised SDR outlines and structural data for the Argentine margin.
Plate reconstructions were made using the open-source software GPlates
(http://www.gplates.org), all figures were generated using GMT
(http://gmt.soest.hawaii.edu; Wessel and Smith, 1998). Structural restorations
were partly done using Midland Valley’s Move software with an academic
software license to The University of Sydney. CH is funded by ARC Linkage
Project L1759 supported by Shell International E&P and TOTAL. JZ publishes
with permission from Statoil ASA. RDM is supported by ARC Laureate Fellowship
FL0992245.

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