SESolid EarthSESolid Earth1869-9529Copernicus PublicationsGöttingen, Germany10.5194/se-9-1239-2018Influence of basement heterogeneity on the architecture of low subsidence
rate Paleozoic intracratonic basins (Reggane, Ahnet, Mouydir and Illizi
basins, Hoggar Massif)Influence of basement heterogeneity on the architecturePerronPaulpaul.perron@u-bourgogne.frpaul.perron@hotmail.frGuiraudMichelVenninEmmanuelleMorettiIsabellePortierÉricLe PourhietLaetitiahttps://orcid.org/0000-0001-9495-4742KonatéMoussaUniversité de Bourgogne Franche-Comté, Centre des Sciences de la Terre, UMR CNRS 6282 Biogéosciences, 6 Bd Gabriel, 21000 Dijon, FranceENGIE SA, 1, place Samuel de Champlain, Faubourg de l'Arche, 92930 Paris La Défense, FranceNeptune Energy International S.A., 9-11 Allée de l'Arche – Tour EGEE – 92400 Courbevoie, FranceSorbonne Université, CNRS-INSU, Institut des Sciences de la Terre Paris, ISTeP UMR 7193, 75005 Paris, FranceDépartement de Géologie, Université Abdou Moumouni de Niamey, BP:10662, Niamey, NigerPaul Perron (paul.perron@u-bourgogne.fr, paul.perron@hotmail.fr)7November2018961239127530May201827June201828September20184October2018This work is licensed under the Creative Commons Attribution 4.0 International License. To view a copy of this licence, visit https://creativecommons.org/licenses/by/4.0/This article is available from https://se.copernicus.org/articles/9/1239/2018/se-9-1239-2018.htmlThe full text article is available as a PDF file from https://se.copernicus.org/articles/9/1239/2018/se-9-1239-2018.pdf
The Paleozoic intracratonic North African Platform is characterized by an
association of arches (ridges, domes, swells, or paleo-highs) and low
subsidence rate syncline basins of different wavelengths (75–620 km). The
Reggane, Ahnet, Mouydir and Illizi basins are successively delimited from
east to west by the Amguid El Biod, Arak-Foum Belrem, and Azzel Matti arches.
Through the analysis of new unpublished geological data (i.e., satellite
images, well logs, seismic lines), the deposits associated with these arches
and syncline basins exhibit thickness variations and facies changes ranging
from continental to marine environments. The arches are characterized by thin
amalgamated deposits with condensed and erosional surfaces, whereas the
syncline basins exhibit thicker and well-preserved successions. In addition,
the vertical facies succession evolves from thin Silurian to Givetian
deposits into thick Upper Devonian sediments. Synsedimentary structures and
major unconformities are related to several tectonic events such as the
Cambrian–Ordovician extension, the Ordovician–Silurian glacial rebound,
the Silurian–Devonian “Caledonian” extension/compression, the late Devonian
extension/compression, and the “Hercynian” compression. Locally, deformation is
characterized by near-vertical planar normal faults responsible for horst and
graben structuring associated with folding during the
Cambrian–Ordovician–Silurian period. These structures may have been
inverted or reactivated during the Devonian (i.e., Caledonian, Mid–Late
Devonian) compression and the Carboniferous (i.e., pre-Hercynian to
Hercynian). Additionally, basement characterization from geological and
geophysics data (aeromagnetic and gravity maps), shows an interesting
age-dependent zonation of the terranes which are bounded by mega-shear zones
within the arches–basins framework. The “old” terranes are situated under
arches while the “young” terranes are located under the basins depocenter.
This structural framework results from the accretion of Archean and
Proterozoic terranes inherited from former orogeny (e.g., Pan-African orogeny
900–520 Ma). Therefore, the sedimentary infilling pattern and the nature of
deformation result from the repeated slow Paleozoic reactivation of
Precambrian terranes bounded by subvertical lithospheric fault systems.
Alternating periods of tectonic quiescence and low-rate subsidence
acceleration associated with extension and local inversion tectonics
correspond to a succession of Paleozoic geodynamic events (i.e., far-field
orogenic belt, glaciation).
Introduction
Paleozoic deposits fill numerous intracratonic basins, which may also be
referred to as “cratonic basins”, “interior cratonic basins”, or
“intracontinental sags”. Intracratonic basins are widespread around the
world (Heine et al., 2008) and exploration for
nonconventional petroleum has revived interest in them. They are located in
“stable” lithospheric areas and share several common features
(Allen and Armitage, 2011). Their geometries are
large circular, elliptical, and/or saucer-shaped to oval.
Their stratigraphy is
filled with continental to shallow-water sediments. Their subsidence rate is
low (5 to 50 m Ma-1)
and long (sometimes more than 540 Myr). Their structural
framework shows the reactivation of structures and emergence of arches also
referred to in the literature as “ridges”, “paleo-highs”, “domes”, and
“swells”. Multiple hypotheses and models have been proposed to explain how
these slowly subsiding, long-lived intracratonic basins formed and evolved
(see Allen and Armitage, 2011 and references therein or Hartley and Allen, 1994). However, their tectonic and
sedimentary architectures are often poorly constrained.
Geological map of the Paleozoic North Saharan Platform (North
Gondwana) georeferenced, compiled and modified from (1) a Paleozoic subcrop
distribution below the Hercynian unconformity geology of the Saharan
Platform (Boote et al., 1998; Galeazzi et
al., 2010), (2) a geological map (1/500000) of the Djado Basin
(Jacquemont et al., 1959), (3) a geological map (1/200000) of
Algeria (Bennacef et al., 1974; Bensalah et al., 1971), (4) a geological map (1/50000) of
Aïr (Joulia, 1963), (5) a geological map (1/2000000) of Niger
(Greigertt and Pougnet, 1965), (6) a geological map (1/5000000)
of the Lower Paleozoic of the central Sahara (Beuf et al.,
1971), (7) a geological map (1/1000000) of Morocco (Hollard
et al., 1985), (8) a geological map of the Djebel Fezzan
(Massa, 1988), and (9) a basement characterization of the different
terranes from geochronological data compilation (see Supplement) and
geological maps (Berger et al., 2014; Bertrand and Caby, 1978; Black et al., 1994; Caby, 2003; Fezaa
et al., 2010; Liégeois et al., 1994, 2003, 2005, 2013). Terrane names and abbreviations:
Tassendjanet (Tas), Tassendjanet nappe (Tas n.), Ahnet (Ah), In Ouzzal
Granulitic Unit (IOGU), Iforas Granulitic Unit (UGI), Kidal (Ki),
Timétrine (Tim), Tilemsi (Til), Tirek (Tir), In Zaouatene (Za), In
Teidini (It), Iskel (Isk), Tefedest (Te), Laouni (La), Azrou-n-Fad (Az),
Egéré-Aleskod (Eg-Al), Serouenout (Se), Tazat (Ta), Issalane (Is),
Assodé (As), Barghot (Ba), Tchilit (Tch), Aouzegueur (Ao), Edembo (Ed),
and Djanet (Dj). Shear zone and lineament names and abbreviations: suture zone East Saharan Craton
(SZ ESC), west Ouzzal shear zone (WOSZ), east Ouzzal shear zone (EOSZ),
Raghane shear zone (RSZ), Tin Amali shear zone (TASZ), 4∘10′
shear zone, 4∘50′ shear zone, and 8∘30′ shear zone.
The main specificities of intracratonic basins are found on the Paleozoic
North Saharan Platform. The sedimentary infilling during ca. 250 Myr is
relatively thin (i.e., around a few hundred to a few thousand meters), of
great lateral extent (i.e., 9 million km2), and is separated by major
regional unconformities (Beuf et al., 1968a, 1971; Carr, 2002; Eschard et al., 2005, 2010; Fabre, 1988,
2005; Fekirine and Abdallah, 1998; Guiraud et al., 2005; Kracha, 2011;
Legrand, 2003a). Depositional environments were mainly continental to
shallow-marine and homogeneous. Very slow and subtle lateral variations
occurred over time (Beuf et al., 1971; Carr, 2002; Fabre, 1988; Guiraud et al., 2005; Legrand, 2003a).
The Paleozoic North Saharan Platform is arranged (Fig. 1) into an
association of long-lived broad synclines (i.e., basins or
subbasins) and anticlines (i.e., arches) of different wavelengths (λ: 75–620 km). Burov and Cloetingh (2009) report deformation
wavelengths of the order of 200–600 km when the whole lithosphere is
involved and of 50–100 km when the crust is decoupled from the lithospheric
mantle. This insight suggests that the inherited basement fabric influences
intracratonic basin architecture at a large scale. In addition, pre-existing
structures, such as shear zones and terrane suture zones, are present
throughout the lithosphere, affecting the geometry and evolution of
upper-crustal structural framework forming during later tectonic events
(Peace et al., 2018; Phillips et al., 2018).
In this study of the Reggane, Ahnet, Mouydir, and Illizi basins, a
multidisciplinary workflow involving various tools (e.g., seismic profiles,
satellite images) and techniques (e.g., photogeology, seismic
interpretation, well correlation, geophysics, geochronology) has enabled us
to (1) make a tectono-sedimentary analysis, (2) determine the spatial
arrangement of depositional environments calibrated by biostratigraphic
zonation, (3) characterize basin geometry, and (4) ascertain the inherited
architecture of the basement and its tectonic evolution. We propose a
conceptual coupled model explaining the architecture of the intracratonic
basins of the North Saharan Platform. This model highlights the role of
basement heritage heterogeneities in an accreted mobile belt and their
influence on the structure and evolution of intracratonic basins. It is a
first step towards a better understanding of the factors and mechanisms that
drive intracratonic basins.
(a) Geological map of the Paleozoic of the Reggane, Ahnet, and
Mouydir basins. For the legend and references see Fig. 1. (b) An E–W cross section of
the Reggane, Ahnet, and Mouydir basins associated with the different
terranes and highlighting the classification of the different structural
units. Localization of the interpreted sections (seismic profiles and
satellite images). W represents well and O represents outcrop. See Fig. 1 for location of
the geological map A and cross section B.
Geological setting: the Paleozoic North Saharan Platform and the Reggane,
Ahnet, Mouydir, and Illizi basins
The Reggane, Ahnet, Mouydir and Illizi basins (Figs. 1 and 2) are located in
southwestern Algeria, north of the Hoggar Massif (Ahaggar). They are
depressions filled by Paleozoic deposits. The basins are bounded to the
south by the Hoggar Massif (Tuareg Shield) and they are separated one
another by the Azzel Matti, the Arak-Foum Belrem, and the Amguid El Biod arches.
Figure 3 synthesizes the lithostratigraphy, the large-scale sequence
stratigraphic framework delimited by six main regional unconformities (A to
F), and the tectonic events proposed in the literature (cf. references under
Fig. 3) affecting the Paleozoic North Saharan Platform.
Paleozoic litho-stratigraphic, sequence stratigraphy, and tectonic
framework of the north peri-Hoggar basins (North African Saharan Platform)
compiled from (1) a chronostratigraphic chart (Ogg et al., 2016),
(2) the Cambrian–Silurian (Askri et al., 1995) and the
Devonian–Carboniferous stratigraphy of the Reggane Basin
(Cózar et al., 2016; Lubeseder, 2005; Lubeseder et al., 2010; Magloire, 1967; Wendt et al.,
2006), (3) the Cambrian–Silurian (Paris, 1990; Wendt et
al., 2006) and the Devonian–Carboniferous stratigraphy of the Ahnet Basin
(Beuf et al., 1971; Conrad, 1973, 1984; Legrand-Blain, 1985; Wendt et al., 2006, 2009a), (4) the
Cambrian–Silurian (Askri et al., 1995; Paris, 1990; Videt et al., 2010) and the Devonian–Carboniferous
stratigraphy of the Mouydir Basin (Askri et al., 1995;
Beuf et al., 1971; Conrad, 1973, 1984; Wendt et al., 2006, 2009a), (5) the
Cambrian–Silurian (Eschard et al., 2005; Fekirine and Abdallah, 1998; Jardiné and Yapaudjian, 1968; Videt et al.,
2010) and the Devonian–Carboniferous stratigraphy of the Illizi Basin
(Eschard et al., 2005; Fekirine and Abdallah, 1998; Jardiné and Yapaudjian, 1968), (6) the
Cambrian–Silurian (Dubois, 1961; Dubois and Mazelet, 1964; Eschard et al., 2005; Henniche, 2002; Videt et
al., 2010) and the Devonian–Carboniferous stratigraphy of the
Tassili-N-Ajjers (Dubois et al., 1967; Eschard et al., 2005; Henniche, 2002; Wendt et al., 2009a), (7) the sequence
stratigraphy of the Saharan Platform
(Carr, 2002; Eschard et al., 2005; Fekirine and Abdallah, 1998), (8) a eustatic and climatic chart
(Haq and Schutter, 2008; Scotese et al., 1999), and (9) tectonic events
(Boudjema, 1987; Coward and Ries, 2003; Craig et al., 2008; Guiraud et al., 2005;
Lüning, 2005). (A) Infra-Tassilian (Pan-African) unconformity, (B) intra-Arenig unconformity, (C) Taconic and glacial unconformity,
(D) isostatic rebound unconformity, (E) Caledonian unconformity, and (F) Hercynian
unconformity.
During the Paleozoic, the Reggane, Ahnet, Mouydir and Illizi basins were
part of a set of the supercontinent Gondwana (Fig. 1). This supercontinent
resulted from the collision of the West African Craton (WAC) and the East
Saharan Craton (ESC), which sandwiched the Tuareg Shield (TS) mobile belt during
the Pan-African orogeny (Craig et al., 2008; Guiraud et
al., 2005; Trompette, 2000). This orogenic cycle followed by the chain's
collapse (ca. 1000–525 Ma) was also marked by phases of oceanization and
continentalization (ca. 900–600 Ma) giving rise to the heterogeneous
terranes in the accreted mobile belt (Trompette, 2000). The Hoggar
Massif is composed of several accreted, sutured, and amalgamated terranes of
various ages and compositions resulting from multiple phases of geodynamic
events (Bertrand and Caby, 1978; Black et al., 1994; Caby, 2003; Liégeois et al., 2003).
Twenty-three well preserved terranes were identified in the Hoggar Massif and
grouped into Archean, Paleoproterozoic, and Mesoproterozoic–Neoproterozoic
juvenile Pan-African terranes (see legend in Fig. 1). In the West African
Craton, the Reguibat Shield is composed of Archean terrains in the west and
of Paleoproterozoic terranes in the east (Peucat
et al., 2003, 2005).
Schematic synthesis of the integrated method of basin analysis in
this study.
Then, there is evidence of a complex and polyphased history throughout the
Paleozoic (Fig. 3), with alternating periods of quiescence and tectonic
activity, individualizing and rejuvenating ancient N–S, NE–SW, or NW–SE
structures in arch and basin configurations
(Badalini et al., 2002; Boote et al., 1998; Boudjema, 1987; Coward and Ries, 2003;
Craig et al., 2008; Guiraud et al., 2005; Logan and Duddy, 1998; Lüning,
2005). The Paleozoic successions of the North Saharan Platform are
predominantly composed of siliciclastic detrital sediments
(Beuf et al., 1971; Eschard et al., 2005). They form
the largest area of detrital sediments ever found on continental crust
(Burke et al., 2003), dipping gently NNW
(Beuf et al., 1971, 1969; Fabre, 1988, 2005; Fröhlich et al., 2010; Gariel et
al., 1968; Le Heron et al., 2009). Carbonate deposits are observed from the
Middle–Late Devonian to the Carboniferous
(Wendt, 1985, 1988, 1995; Wendt et al., 1993, 1997, 2006, 2009a; Wendt and Kaufmann,
1998). From south to north, the facies progressively evolve from continental
fluviatile to shallow marine (i.e., upper to lower shoreface) and then to
offshore facies (Beuf et al., 1971; Carr, 2002; Eschard et al., 2005; Fabre, 1988; Fekirine and
Abdallah, 1998; Legrand, 1967a).
Data and methods
A multidisciplinary approach was used in this study integrating new
data (i.e., satellite images, seismic lines and well-logs data) in particular
from the Reggane, Ahnet, Mouydir, and Illizi basins and the Hoggar Massif
(Fig. 4).
Synthesis of facies associations (AF1 to AF5), depositional
environments, and electrofacies in the Devonian series compiled from
internal (Eschard et al., 1999) and published studies
(Beuf et al., 1971; Biju-Duval et al.,
1968; Wendt et al., 2006).
The Paleozoic series of the Ahnet and Mouydir basins are well-exposed over
an area of approximately 170 000 km2 and are well observed in satellite
images (Google Earth and Landsat from USGS). Furthermore, a significant
geological database (i.e., wells, seismic records, geological reports) has
been compiled in the course of petroleum exploration since the 1950s. The
sedimentological dataset is based on the integration and analysis of cores,
outcrops, well logs, and of lithological and biostratigraphic data. They
were synthesized from internal SONATRACH (Dokka, 1999),
IFP-SONATRACH consortium reports (Eschard et al., 1999), and published articles (Beuf et
al., 1971; Biju-Duval et al., 1968; Wendt et al., 2006). Facies described
from cores and outcrops of these studies were grouped into facies
associations corresponding to the main depositional environments observed on
the Saharan Platform (Table 1). Characteristic gamma-ray (GR) patterns
(electrofacies) are proposed to illustrate the different facies
associations. The gamma-ray peaks are commonly interpreted as the
maximum flooding surfaces (MFS)
(e.g., Catuneanu et al., 2009; Galloway, 1989; Milton et al., 1990; Serra and
Serra, 2003). Time calibration of well logs is based on palynomorphs
(essentially Chitinozoans and spores) and outcrops on conodonts, goniatites,
and brachiopods (Wendt et al., 2006). Palynological data of
wells (W1, W7, W12, W19 and W20) from internal unpublished data
(Abdesselam-Rouighi, 1991; Azzoune, 1999;
Hassan, 1984; Khiar, 1974) are based on biozonations from
Magloire (1967) and Boumendjel et al. (1988). Well W18 is supported by palynological data and biozonations from
Kermandji et al. (2008).
Synsedimentary extensional and compressional markers are characterized in
this structural framework based on the analyses of satellite images (Figs. 5 and 6), seismic profiles (Fig. 7), 21 wells (W1 to W21), and 12 outcrop
cross sections (O1 to O12). Wells and outcrop sections are arranged into
three E–W sections (Figs. 10, 11 and 12) and one N–S section (Fig. 13).
Satellite images (Figs. 5 and 6) and seismic profiles (Fig. 7) are located
at key areas (i.e., near arches) illustrating the relevant structures (Fig. 2). The calibration of the key stratigraphic horizon on seismic profiles
(Fig. 7) was settled by sonic well-log data using Petrel and OpendTect
software. Nine key horizons easily extendable at the regional scale are
identified and essentially correspond to major depositional unconformities:
near top infra-Cambrian, near top Ordovician, near top Silurian, near top
Pragian, near top Givetian, near top mid-Frasnian, near top Famennian, near
base Quaternary and near Hercynian unconformities (Fig. 7). The
stratigraphic layers are identified by the integration of satellite images
(Google Earth and Landsat USGS: https://earthexplorer.usgs.gov/, last access: 29 November 2016), digital
elevation models (DEM), and the 1:200000 geological maps of Algeria
(Bennacef et al., 1974; Bensalah et al., 1971).
Subsidence analysis characterizes the vertical displacements of a given
sedimentary depositional surface by tracking its subsidence and uplift
history (Van Hinte, 1978). The resulting curve details the total
subsidence history for a given stratigraphic column (Allen
and Allen, 2005; Van Hinte, 1978). Backstripping is also used to restore the
initial thicknesses of a sedimentary column (Allen and
Allen, 2005; Angevine et al., 1990). Lithologies and paleobathymetries have
been defined using facies analysis or literature data. Porosity and the
compaction proxy are based on experimental data from (Sclater
and Christie, 1980). In this study, subsidence analyses were performed on
sections using OSXBackstrip software performing 1D Airy backstripping
(following Allen and Allen, 2005; Watts, 2001); available
at: http://www.ux.uis.no/nestor/work/programs.html, last access: 5 January 2017).
(a) Typology of different types of faults (inherited straight
faults vs. sinuous short synlithification propagation faults) in the
Cambrian–Ordovician series of the Djebel Settaf (Arak-Foum Belrem Arch;
interbasin boundary secondary arch between the Ahnet and Mouydir basins).
(b) Structural control of channelized sandstone bodies in Late Ordovician
series of South Adrar Assaouatene, Tassili-N-Ajjers (Tihemboka interbasin
boundary secondary arch between the Illizi and Murzuq basins). The dotted red
line represents Tamadjert Fm. channelized sandstone bodies. The abbreviations used in the figure are as follows: OTh – In Tahouite Fm. (Early
to Late Ordovician, Floian to Katian); OTj – Tamadjert Fm. (Late Ordovician,
Hirnantian); sIm – Imirhou Fm. (Early Silurian); sdAs1 – Asedjrad Fm. 1 (Late Silurian
to Early Devonian); dAs2 – Asedjrad Fm. 2 (Early Devonian, Lochkovian); dSa – Oued
Samene Fm. (Lower Devonian, Pragian). See Fig. 2 for map and cross-section
location.
(a) Normal fault (F2) associated with a footwall anticline and a
hanging wall syncline with divergent onlaps (i.e., wedge-shaped unit DO1) in
the Early to Late Ordovician In Tahouite series (Tassili-N-Ajjers, Tihemboka
interbasin boundary secondary arch between the Illizi and Murzuq basins).
(b) Ancient normal fault (F2) escarpment reactivated and sealed during
Silurian deposition (poly-historic paleo-reliefs) linked to thickness
variation, divergent onlaps (DO2) in the hanging wall synclines, and onlaps
on the fold hinge anticline (Tassili-N-Ajjers, Tihemboka interbasin
boundary secondary arch between the Illizi and Murzuq basins). 1: Early to
Late Ordovician extension, 2: Late Ordovician to Early Silurian extension,
and 3: Middle to Late Silurian sealing (horizontal drape). (c) Normal fault (F5)
associated with forced fold with divergent strata (syncline-shaped hanging
wall syncline and associated wedge-shaped unit DO2) and truncation in the
Silurian–Devonian series of Dejbel Settaf (Arak-Foum Belrem Arch;
interbasin boundary secondary arch between the Mouydir and Ahnet basins).
1: Cambrian–Ordovician extension and 2: Silurian–Devonian extensional
reactivation (Caledonian extension). (d) Blind basement normal fault (F1)
associated with forced fold with in the hanging wall syncline divergent
onlaps of the Lower to Upper Devonian series (wedge-shaped unit DO3) and
intra-Emsian truncation (Arak-Foum Belrem Arch; interbasin boundary
secondary arch between the Mouydir and Ahnet basins). (e) N170∘
normal blind faults F1 and F2 forming a horst and graben system associated with
a forced fold with Lower to Upper Devonian series divergent onlaps
(wedge-shaped unit DO3) and intra-Emsian truncation in the hanging-wall
syncline (in the Mouydir Basin near Arak-Foum Belrem Arch, eastward
interbasin boundary secondary arch). (f) Inherited normal fault (F1)
transported from footwall to hanging wall associated with an inverse fault (F1')
and accommodated by a detachment layer in the Silurian shales series (thickness
variation of Imirhou Fm. between the footwall and hanging wall) and spilled dip
strata markers of overturned folding (Djebel Idjerane, Arak-Foum Belrem
Arch, eastwards interbasin boundary secondary arch). 1: Cambrian–Ordovician extension and 2: Middle to Late Devonian compression.
The abbreviations used in the figure are as follows: OTh – In Tahouite Fm. (Early to Late Ordovician, Floian to Katian); OTj – Tamadjert
Fm (Late Ordovician, Hirnantian); sIm – Imirhou Fm. (Early to
Mid-Silurian); sdAt – Atafaïtafa Fm. (Middle Silurian); dTi – Tifernine Fm. (Middle
Silurian); sdAs1 – Asedjrad Fm. 1 (Late Silurian to Early Devonian); dAs2 – Asedjrad Fm. 2 (Early
Devonian, Lochkovian); dSa: Oued Samene Fm. (Early Devonian, Pragian); diag: Oued
Samene shaly-sandstones Fm. (Early Devonian, Emsian?); d2b – Givetian; d3a – Mehden
Yahia Fm. (Late Devonian, Frasnian); d3b – Mehden Yahia Fm. (Late Devonian,
Famennian); dh – Khenig sandstones (late Famennian to early
Tournaisian); hTn2 – late Tournaisian; hV1 – early Visean. The red line represents unconformity. See Figs. 1,
2, and 5 for map and cross-section location.
The 800 km2 outcrop of basement rocks of the Hoggar Massif provides an
exceptional case study of an exhumed mobile belt composed of accreted
terranes of different ages. To reconstruct the nature of the basement, a
terrane map (Figs. 15 and 16) was put together by integrating geophysical
data (aeromagnetic anomaly map: https://www.geomag.us/, last access: 1 December 2016, Bouguer gravity
anomaly map: http://bgi.omp.obs-mip.fr/, last access: 1 December 2016), satellite images (7ETM+ from
Landsat USGS: https://earthexplorer.usgs.gov/, last access: 29 November 2016) data, geological maps (Berger
et al., 2014; Bertrand and Caby, 1978; Black et al., 1994; Caby, 2003; Fezaa
et al., 2010; Liégeois et al., 1994, 2003, 2005, 2013), and
geochronological data (e.g., U/Pb radiochronology, see Supplement data 1).
Geochronological data from published studies were compiled and georeferenced
(Fig. 1). Thermo-tectonic ages were grouped into eight main thermo-orogenic
events (Fig. 1): The Liberian-Ouzzalian event (Arcehan, > 2500 Ma), the Archean, Eburnean (i.e., Paleoproterozoic, 2500–1600 Ma), the
Kibarian (i.e., Mesosproterozoic, 1600–1100 Ma), the Neoproterozoic
oceanization-rifting (1100–750 Ma), the syn-Pan-African orogeny (i.e.,
Neoproterozoic, 750–541 Ma), the post-Pan-African (i.e., Neoproterozoic,
541–443 Ma), the Caledonian orogeny (i.e., Siluro-Devonian, 443–358 Ma), and
the Hercynian orogeny (i.e., Carbo-Permian, 358–252 Ma).
(a) N–S interpreted seismic profile in the Ahnet Basin near Erg
Tegunentour (near Arak-Foum Belrem Arch, westward interbasin boundary
secondary arch) showing steeply dipping northward basement normal blind
faults associated with forced folding. (b) NW–SE interpreted seismic
profile of near Azzel Matti Arch (interbasin principal arch) showing
steeply dipping southeastward basement normal blind faults associated with
forced folds. The westernmost structures are featured by reverse fault
related propagation fold. (c) W–E interpreted profile of the Ahnet Basin
(Arak-Foum Belrem Arch, westward interbasin boundary secondary arch)
showing horst and graben structures influencing Paleozoic tectonics
associated with forced folds. (d) W–E interpreted seismic profile of Bahar
el Hammar in the Ahnet Basin (Ahnet intra-basin secondary arch) showing
steeply dipping normal faults F1 and F2 forming a positively inverted horst
associated with folding. Multiple activation and inversion of normal faults
are correlated with divergent onlaps (wedge-shaped units): DO0 infra-Cambrian
extension, DO1 Cambrian–Ordovician extension, DO2 Silurian extension with
local Silurian–Devonian positive inversion, and DO3 Frasnian–Famennian
extension–local compression. See Fig. 2 for map and cross-section
location.
Structural framework and tectono-sedimentary structure analyses
The structural architecture of the North Saharan Platform is characterized
by mostly circular to oval shaped basins structured by major faults
frequently associated with broad asymmetrical folds displayed by three main
trends (Fig. 1): (1) near-N–S, varying from 0 to 10
or 160∘ N , (2) from 40 to
60∘ N , and from (3) 100 to 140∘ N (Figs. 1, 3a, and 4). These
fault zones are about 100 km (e.g., faults F1 and F2, Fig. 5) to tens of
kilometers long (e.g., faults F3 to F8, Fig. 5). They correspond to the
mainly N–S Azzel-Matti, Arak-Foum Belrem, Amguid El Biod, and Tihemboka
arches, the NE–SW Bou Bernous, Ahara, and Gargaf arches, and the NW–SE
Saoura and Azzene arches (Fig. 1).
(a) Core description, palynological calibration, and gamma-ray
signatures of well W7 modified from an internal core description report
(Dokka, 1999) and an internal palynological report
(Azzoune, 1999). (b) Devonian sequential stratigraphy of well-log
W7. For the location of well W7 see Fig. 2a.
Synsedimentary extensional markers
Extensional markers are characterized by the settlement of steeply westward or
eastward-dipping basement normal faults associated with colinear
syndepositional folds of several kilometers in length (e.g., Figs. 6a to e and 7a), represented by footwall anticline and hanging wall syncline-shaped
forced folds. They are located in the vicinity of different arches (Fig. 2)
such as the Tihemboka Arch (Figs. 5b and 6a, b), Arak-Foum Belrem Arch
(Figs. 5a, 6c to f and 7a, c), Azzel Matti Arch (Fig. 7b), and Bahar El
Hamar area intra-basin arch (Fig. 7d). These tectonic structures can be
featured by basement blind faults (e.g., fault F1 in Fig. 7a). The
deformation pattern is mainly characterized by brittle faulting in
Cambrian–Ordovician series down to the basement and fault-damping in
Silurian series (e.g., faults F1 to F6 in Fig. 7b). The other terms of the
series (i.e., Silurian to Carboniferous) are usually affected by folding
except (see F1 faults in Figs. 6d, 7b, d, and c) where the brittle
deformation can be propagated to the Upper Devonian (due to reactivation
and/or inversion as suggested in the next paragraph).
In association with the extensional markers, thickness variations and tilted
divergent onlaps of the sedimentary series (i.e., wedge-shaped units,
progressive unconformities) in the hanging wall syncline of the fault
escarpments are observed (Figs. 6 and 7). These are attested using
photogeological analysis of satellite images (Fig. 6) and are marked by a
gentler dip angle of the stratification planes away from the fault plane
(i.e., fault core zone). The markers of syndepositional deformation
structures are visible in the hanging-wall synclines of Precambrian to Upper
Devonian series (Figs. 6 and 7).
The footwall anticline and hanging-wall syncline-shaped forced folds
recognized in this study are very similar to those described in the
literature by Grasemann et al. (2005), Khalil and McClay (2002), Schlische (1995),
Stearns (1978), Withjack et al. (1990, 2002), and Withjack and Callaway (2000). The
wedge-shaped units (DO0 to DO3; Figs. 5, 6, and 7) associated with the
hanging-wall synclines are interpreted as synsedimentary normal
fault-related folding. The whole tectonic framework forms broad extensional
horsts and grabens related to synsedimentary forced folds controlling basin
shape and sedimentation.
Following Khalil and McClay (2002), Lewis et al. (2015), Shaw et al. (2005), and Withjack et al. (1990), we use
the ages of the growth strata (i.e., wedge-shaped units) to determine the
timing of the deformation. The main four wedge-shaped units identified (DO0
to DO3) are indicative of the activation and/or reactivation of the normal
faults (extensional settings) during the Neoproterozoic (DO0),
the Cambrian–Ordovician (DO1), the Early to Mid-Silurian (DO2), and the Middle to Late
Devonian (DO3) times.
In planar view, straight (F1 in Fig. 5a) and sinuous faults (F2, F3, F3',
F4, F4', and F5 in Fig. 5a) can be identified. The sinuous faults are
arranged “en echelon” into several segments with relay ramps. These faults
are ten to several tens of kilometers long with vertical throws of hundreds
of meters that fade rapidly toward the fault tips. The sinuous geometry of
normal undulated faults as well as the rapid lateral variation in fault
throw are controlled by the propagation and the linkage of growing parent
and tip synsedimentary normal faults (Marchal et
al., 2003, 1998). We use the stratigraphic age of impacted layers (here
Tamadjert Fm.) to date (re)activation of the faults.
According to Holbrook and Schumm (1999), river patterns are
extremely sensitive to tectonic structure activity. Here we find that the
synsedimentary activity of the extensional structures is also evidenced by
the influence of the fault scarp on the distribution and orientation of
sinuous channelized sandstone body systems (dotted red lines in Fig. 5b). It
highlights the (re)activation of the faults during the deposition of these
channels, i.e., Late Hirnantian dated by (Girard et
al., 2012).
After the development of the extensional tectonism described previously,
evidence of synsedimentary compressional markers can be identified. These
markers are located and preferentially observable near the Arak-Foum Belrem
Arch (Fig. 6f; F2 in Fig. 7c), the Azzel Matti Arch (2 in Fig. 7v), and the
Bahar El Hamar area intra-basin arch (2 in Fig. 7d). The tectonic structures
take the form of inverse faulting reactivating former basement faults (F1'
in Fig. 6f, F1 in Fig. 7c, F1' in Fig. 7d, F1 in Fig. 7b). The
synsedimentary inverse faulting is demonstrated by the characterization of
asymmetric anticlines and is especially observable in satellite images and
restricted to the fault footwalls (Fig. 5a along F1–F2).
Landsat image analysis combined with the line drawing of certain seismic
lines reveals several thickness variations reflecting divergent onlaps (i.e.,
wedge-shaped units) which are restricted to the hanging-wall asymmetric
anticlines (2 in Figs. 6f, 7b, c and d). The compressional synsedimentary
markers clearly post-date extensional divergent onlaps at hanging-wall
syncline-shaped forced folds (1 in Figs. 7c, c and d). This architecture is
very similar to classical positive inversion structures of former inherited
normal faults (Bellahsen and Daniel, 2005; Bonini et al., 2012; Buchanan and McClay, 1991; Ustaszewski et
al., 2005). Tectonic transport from the paleo-graben hanging wall toward the
paleo-horst footwall (F1, F2–F2', F4–F4' in Fig. 7b; F1–F1' in Fig. 7d) is
evidenced. Further positive tectonic inversion architecture is identified by
tectonic transport from the paleo-horst footwall to the paleo-graben hanging
wall (F1-F1' in Fig. 6f; F1, F5, and F6 in Fig. 7c). This second type of
tectonic inversion is very similar to the transported fault models defined
by Butler (1989) and Madritsch et al. (2008). The
local positive inversions of inherited normal faults occurred during
Silurian–Devonian (F4' Fig. 7b) and Middle to Late Devonian times (Figs. 7b, c and d). A late significant compression event between the end of the
Carboniferous and the Early Mesozoic was responsible for the exhumation and
erosion of the tilted Paleozoic series. This series is related to the
Hercynian angular unconformity surface (Fig. 7b).
Stratigraphy and sedimentology
The whole sedimentary series described in the literature is composed of
fluviatile to braid-deltaic plain Cambrian, not only fluviatile (e.g.,
Brahmaputra River analogue), with a transitional facies from continental to
shallow marine (Beuf et al., 1968a, b, 1971; Eschard et al., 2005, 2010; Sabaou et al., 2009), Upper
Ordovician glaciogenic deposits (Beuf et al., 1968a, b, 1971; Eschard et al., 2005, 2010), argillaceous deep marine Silurian
deposits (Djouder et al., 2018; Eschard et al., 2005, 2010; Legrand, 1986, 2003b; Lüning
et al., 2000), and offshore to embayment Carboniferous deposits
(Wendt et al., 2009a). In this complete sedimentary
succession, we have focused on the Devonian deposits as they are very
sensitive to and representative of basin dynamics. The architecture of the
Devonian deposits allows us to approximate the main forcing factors
controlling the sedimentary infilling of the basin and its synsedimentary
deformation. Eleven facies associations organized into four depositional
environments (Table 1) are defined to reconstruct the architecture and the
lateral and vertical sedimentary evolution of the basins (Figs. 10, 11, 12
and 13).
Facies association, depositional environments, and erosional
unconformities
Based on the compilation and synthesis of internal studies
(Eschard et al., 1999), and published papers on the Saharan
Platform (Beuf et al., 1971; Eschard et
al., 2005, 2010; Henniche, 2002) and on the Ahnet and Mouydir basins
(Biju-Duval et al., 1968; Wendt et al., 2006),
eleven main facies associations (AF1 to AF5) and four depositional
environments are proposed for the Devonian succession (Table 1). They are
associated with their gamma-ray responses (Figs. 8 and 9). They are organized
into two continental fluvial (AF1 to AF2), four transitional coastal plain
(AF3a to AF3d), three shoreface (AF4a to AF4c), and two offshore (AF5a to
AF5b) sedimentary environments.
Continental fluvial environments
This depositional environment features the AF1 (fluvial) and the AF2 (flood
plain) facies association (Table 1). Facies association AF1 is mainly
characterized by a thinning-up sequence with a basal erosional surface and
trough cross-bedded intraformational conglomerates with mud clast lag
deposits, quartz pebbles, and imbricated grains (Table 1). It passes into
medium to coarse trough cross-bedded sandstones, planar cross-bedded
siltstones, and laminated shales. These deposits are associated with rare
bioturbation (except at the surface of the sets),
ironstones, phosphorites, corroded quartz grains, and phosphatized pebbles.
Laterally, facies association AF2 is characterized by horizontally laminated
and very poorly sorted silt to argillaceous fine sandstones. They contain
frequent root traces, plant debris, well-developed paleosols, bioturbation,
nodules, and ferruginous horizons. Current ripples and climbing ripples are
associated in prograding thin sandy layers.
In AF1, the basal erosional reworking and high energy processes are
characteristic of channel-filling of fluvial systems
(Allen, 1983; Owen, 1995). Eschard et al. (1999) identify three fluvial systems (see A, B, and C in Fig. 9) in
the Tassili-N-Ajjers outcrops: braided dominant (AF1a), meandering dominant
(AF1b), and straight dominant (AF1c). They differentiate them by their
different sinuosity, direction of accretion (lateral or frontal), the
presence of mud drapes, bioturbation, and giant epsilon cross-bedding.
Gamma-ray signatures of these facies associations (A, B, and C in Fig. 9)
are cylindrical with an average value of 20 gAPI. The gamma-ray shapes are
largely representative of fluvial environments
(Rider, 1996; Serra and Serra, 2003; Wagoner
et al., 1990). The bottom is sharp with high value peaks and the tops are
frequently fining-up, which may be associated with high values caused by
argillaceous flood plain deposits and roots (Eschard et
al., 1999). AF2 is interpreted as humid floodplain deposits
(Allen, 1983; Owen, 1995) with crevasse splays or
preserved levees of fluvial channels (Eschard et al.,
1999). Gamma-ray curves of AF2 (D, Fig. 9) show a rapid succession of low to
very high peak values, ranging from 50 to 120 gAPI. AF1 and AF2 are typical
of the Pragian “Oued Samene” Formation (Wendt et al., 2006).
In the Illizi Basin, these facies are mainly recorded in the Ajjers
Formation (dated Upper
Cambrian? to Ordovician, see Fabre, 2005; Vecoli, 2000; Vecoli et al., 1995,
1999, 2008; Vecoli and Playford, 1997) and the Lochkovian to Pragian “Barre
Moyenne” and “Barre Supérieure” formations
(Beuf et al., 1971; Eschard et al., 2005).
The main depositional environments (a–l) and their associated
electrofacies (i.e., gamma-ray patterns) modified and compiled from
Eschard et al. (1999).
Transitional coastal plain environments
This depositional environment comprises facies associations AF3a
(delta/estuarine), AF3b (fluvial/tidal distributary channels), AF3c (tidal
sand flat), and AF3d (lagoon/mudflat) (Table 1). AF3a is mainly dominated by
sigmoidal cross-bedded heterolithic rocks with mud drapes. It is also
characterized by fine to coarse, poorly sorted sandstones and siltstones
often structured by combined flow ripples, flaser bedding, wavy bedding, and
some rare planar bedding. Mud clasts, root traces, desiccation cracks, water
escape features, and shale pebbles are common. The presence of epsilon
bedding is attested, which is formed by lateral accretion of a river point
bar (Allen, 1983). The bed surface sets are intensively
bioturbated (Skolithos and Planolites) indicating a shallow marine subtidal setting
(Pemberton and Frey, 1982). Faunas such as brachiopods,
trilobites, tentaculites, and graptolites are present. AF3b exhibits a
fining-up sequence featured by a sharp erosional surface, trough
cross-bedded, very coarse-grained, poorly sorted sandstone at the base and
sigmoidal cross-bedding at the top (Figs. 8 and 9). AF3c is formed by
fine-grained to very coarse-grained sigmoidal cross-bedded heterolithic
sandstones with multidirectional tidal bundles. They are also structured by
lenticular, flaser bedding and occasional current and oscillation ripples
with mud cracks. They reveal intense bioturbation composed of Skolithos (Sk),
Thalassinoides (Th), and Planolites (Pl) ichnofacies indicating a shallow marine subtidal setting
(Frey et al., 1990; Pemberton and Frey, 1982). AF4d is
characterized by horizontally laminated mudstones associated with
varicolored shales and fine-grained sandstones. They exhibit mud cracks,
occasional wave ripples, and rare multidirectional current ripples. These
sedimentary structures are poorly preserved because of intense bioturbation
composed of Skolithos (Sk), Thalassinoides (Th), and Planolites (Pl). The fauna includes ammonoids (rare),
goniatites, calymenids, pelecypod molds, and brachiopod coquinas.
In AF3a, both tidal and fluvial systems in the same facies association can
be interpreted as an estuarine system (Dalrymple et
al., 1992; Dalrymple and Choi, 2007). The gamma-ray signature is
characterized by a convex bell shape with rapidly alternating low to mid
values (30 to 60 gAPI) due to the mud draping of the sets (see E Fig. 9).
These forms of gamma ray are typical of fluvial–tidal influenced
environments with upward-fining parasequences
(Rider, 1996; Serra and Serra, 2003; Wagoner
et al., 1990). AF3a is identified at the top of the Pragian “Oued Samene”
Formation and in Famennian “Khenig” Formation (Wendt et al.,
2006) in the Ahnet and Mouydir basins. In the Illizi Basin, AF3a is mostly
recorded at the top Cambrian of the Ajjers Formation, in the Lochkovian
“Barre Moyenne”, and at the top Pragian of the “Barre Supérieure”
Formation (Beuf et al., 1971; Eschard et al., 2005).
The AF3b association can be characterized by a mixed fluvial and tidal
dynamic based on criteria such as erosional basal contacts, fining-upward
trends, or heterolythic facies (Dalrymple et al.,
1992; Dalrymple and Choi, 2007). They are associated with abundant mud
clasts, mud drapes, and bioturbation indicating tidal influences
(Dalrymple et al., 1992,
2012; Dalrymple and Choi, 2007). The major difference with the estuarine
facies association (AF3a) is the slight lateral extent of the channels which
are only visible in outcrops (Eschard et al., 1999).
The gamma-ray pattern is very similar to the estuarine electrofacies (see F
Fig. 9). AF3c is interpreted as a tidal sand flat laterally present near a
delta (Lessa and Masselink, 1995) and associated with an estuarine
environment (Leuven et al., 2016). The gamma-ray
signature (see G Fig. 9) is distinguishable by its concave funnel shape with
alternating low and mid peaks (25 to 60 gAPI) due to the heterogeneity of
the deposits and rapid variations in the sand/shale ratio. These facies are
observed in the “Talus à Tigillites” Formation of the Illizi Basin
(Eschard et al., 2005). In AF4d, both ichnofacies and facies
are indicative of tidal mudflat/lagoonal depositional environments
(Dalrymple et al., 1992; Dalrymple and Choi,
2007; Frey et al., 1990). The gamma-ray signature has a distinctively high
value (80 to 130 gAPI) and an erratic shape (see H Fig. 9). AF4d is observed
in the “Atafaitafa” Formation and in the Emsian prograding shoreface
sequence of the Illizi Basin (Eschard et al., 2005). It is also
recorded in the Lochkovian “Oued Samene” Formation and the Famennian
“Khenig” sandstones (Wendt et al., 2006).
Shoreface environments
This depositional environment is composed of AF4a (subtidal), AF4b (upper
shoreface), and AF4c (lower shoreface) facies associations (Table 1). AF4a
is characterized by the presence of brachiopods, crinoids, and diversified
bioturbations, by the absence of emersion, and by the greater amplitude of
the sets in a dominant mud lithology (Eschard et al.,
1999). AF4b is heterolithic and composed of fine to medium-grained
sandstones (brownish) interbedded with argillaceous siltstones and
bioclastic carbonated sandstones. Sedimentary structures include oscillation
ripples, swaley cross-bedding, flaser bedding, cross-bedding, convolute
bedding, wavy bedding, and low-angle planar cross-stratification. Sediments
were affected by moderate to highly diversified bioturbation by Skolithos (Sk),
Cruziana, Planolites, (Pl) Chondrites (Ch), Teichichnus (Te), Spirophytons
(Sp) and are composed of ooids, crinoids, bryozoans,
stromatoporoids, tabulate and rugose corals, pelagic styliolinids, neritic
tentaculitids, and brachiopods. AF4c can be distinguished by a low
sand/shale ratio, thick interbeds, abundant hummocky
cross-stratification (HCS), deep groove marks, slumping, and intense bioturbation (Table 1).
AF4a is interpreted as a lagoonal shoreface. The gamma-ray pattern (see I
Fig. 9) is characterized by a concave bell shape influenced by a low
sand/shale ratio with values fluctuating between 100 and 200 gAPI. AF4a is
identified in the “Talus à Tigillites” Formation and the Emsian
sequence of the Illizi Basin (Eschard et al., 2005) and in the
Lochkovian “Oued Samene” Formation (Wendt et al., 2006). AF4b
is interpreted as a shoreface environment. The presence of swaley
cross-bedding produced by the amalgamation of storm beds (Dumas
and Arnott, 2006) and other cross-stratified beds is indicative of upper
shoreface environments (Loi et al., 2010). The
gamma-ray pattern (see J and K Fig. 9) displays concave erratic egg shapes
with a very regularly decreasing-upward trend and ranging from offshore
shale with mid values (80 to 60 gAPI) to clean sandstone with lower values
at the top (40 to 60 gAPI). AF4b is observed in the “Atafaitafa” Formation
corresponding to the “Zone de passage” Formation of the Illizi Basin
(Eschard et al., 2005). AF4c is interpreted as a lower
shoreface environment (Dumas and Arnott, 2006; Suter, 2006).
The gamma-ray pattern displays the same features as the upper shoreface
deposits with higher values (i.e., muddier facies) ranging from 100 to 80 gAPI (see J and K Fig. 9).
Offshore marine environments
This depositional environment is composed of AF5a and AF5b facies
associations (Table 1). AF5a is mainly defined by wavy to planar-bedded
heterolithic silty-shales interlayered with fine-grained sandstones. It also
contains bundles of skeletal wackestones and calcareous mudstones. The main
sedimentary structures are lenticular sandstones, HCS
, mud mounds, low-angle cross-bedding, tempestite
bedding, slumping, and deep groove marks. Sediments can present rare
horizontal bioturbation such as Zoophycos (Z), Teichichnus (Te), and Planolites (Pl). AF5b is
characterized by an association of black silty shales with occasional
bituminous wackestones and packstones. It is composed of graptolites,
gonitaties, orthoconic nautiloids, pelagic pelecypods, limestone nodules,
tentaculitids, ostracods, and rare fish remains. Rare bioturbation such as
Zoophycos (Z) is visible.
In AF5a, the occurrence of HCS, the decrease in sand thickness and grain
size, and the bioturbation and the floro–faunal associations
indicate a deeper marine environment under the influence of storms
(Aigner, 1985; Dott and Bourgeois, 1982; Reading and Collinson, 2009). AF5a is interpreted as upper offshore
deposits (i.e., offshore transitional). The gamma-ray pattern is serrated and
erratic with values well grouped around high values from 120 to 140 gAPI
(see L Fig. 9). Positive peaks may indicate siltstone to sandstone ripple
beds. AF5b is interpreted as lower offshore deposits
(Aigner, 1985; Stow et al., 2001; Stow and Piper, 1984). Here again the gamma-ray
signature is serrated and erratic with values well grouped around 140 gAPI
(see L Fig. 9). Hot shales with anoxic conditions are characterized by
gamma-ray peaks (> 140 gAPI). These gamma-ray patterns are
typical of offshore environments dominated by shales
(Rider, 1996; Serra and Serra, 2003; Wagoner
et al., 1990). AF5a and AF5b are observed in the Silurian “Argiles à
Graptolites” Formation and the Emsian “Orsine” Formation of the Illizi
Basin (Beuf et al., 1971;
Eschard et al., 2005; Legrand, 1986, 2003b). The “Argiles de Mehden Yahia”
and “Argiles de Temertasset“ shales have the same facies
(Wendt et al., 2006).
Sequential framework and unconformities
The high-resolution facies analysis, depositional environments, stacking
patterns, and surface geometries observed in the Devonian succession reveal
at least two different orders of depositional sequences (large and medium
scale, Fig. 8) considered as transgressive/regressive (T/R)
(Catuneanu et al., 2009). The sequential framework proposed in Fig. 8b results from the
integration of the vertical evolution of the main surfaces (Fig. 8a) and the
gamma-ray pattern (Fig. 9). The Devonian series under focus exhibits 9
medium-scale sequences (D1 to D9, Fig. 8; Figs. 10, 11, 12, and 13) bounded
by 10 major sequence boundaries (HD0 to HD9), and 9 major flooding
surfaces (MFS1 to MFS9). The correlation of the different sequences at the
scale of the different basins and arches is used to build three cross sections – two E–W (Figs. 10, 11 and 12) and one N–S (Fig. 13).
SE–W cross-section between the Reggane Basin, the Azzel Matti Arch,
the Ahnet Basin, the Arak-Foum Belrem Arch, the Mouydir Basin, and the Amguid El Biod Arch
(well locations in Fig. 3). The well W1 biozone calibration is from
Hassan (1984) and the internal report is based on the Magloire (1967) classification: biozone G3-H (Wenlock–Ludlow, Upper Silurian),
biozone I-K (Lochkovian–Emsian, Lower Devonian), biozone L1-3
(Eifelian–Givetian, Middle Devonian), biozone L4 (Frasnian, Upper
Devonian), biozone L5-7 (Famennian, Upper Devonian), and biozone M2
(Tournaisian–Lower Carboniferous). The well W7 biozone calibration is from
Azzoune (1999) and the internal report is based on the Boumendjel (1987) classification: biozone 7–12 (Lochkovian, Lower Devonian) and biozone 15
(Emsian, Lower Devonian). Interpretation of the basement is based on Figs. 1, 2, and 15. Well location is in Fig. 2.
The result of the analysis of the general pattern displayed by the
successive sequences reveal two major patterns (Figs. 10, 12 and 13) limited
by a major flooding surface MFS5. The first pattern extends from the Oued
Samene to Adrar Morrat formations and is dated from the Lochkovian to
Givetian. D1 to D5 medium-scale sequences indicate a general proximal
clastic depositional environment (dominated by fluvial to transitional and
shoreface facies) with intensive lateral facies evolution. This first
pattern is thin (from 500 m in the basin depocenter to 200 m around the
basin rim) and with successive amalgamated surfaces on the edge of the
arches between the “Zone de passage” and “Oued Samene” formations (e.g.,
Figs. 10 and 13). It is delimited at the bottom by the HD0 surface
corresponding to the Silurian–Devonian boundary. D1 to D3 are composed of
T/R sequences with a first deepening transgressive trend indicative of a
transition from continental to marine deposits bounded by a major MFS and
evolving into a second shallowing trend from deep marine to shallow marine
depositional environments. D1 to D3 thin progressively toward the edge and
the continental deposits, in the central part of the basin, pass laterally
into a major unconformity. The amalgamation of the surfaces and lateral
variations of facies between the Ahnet Basin and Azzel Matti and Arak-Foum
Belrem arches demonstrate a tectonic control related to the presence of
subsiding basins and paleo-highs (i.e., arches).
D4 and D5 display the same T/R pattern with a reduced continental influence
and upward decrease in lateral facies variations and thicknesses where the
MFS4 marks the beginning of a marine-dominated regime in the entire area. It
is identified as the Early Eifelian transgression defined by
Wendt et al. (2006). The D5 sequence is mainly composed of
shoreface carbonates. Evidence of mud mounds preferentially located along
faults are well-documented in the area for that time
(Wendt et al., 1993, 1997, 2006; Wendt and Kaufmann, 1998). This change in the
general pattern indicates reduced tectonic influence.
MFS5, at the transition between the two main patterns, represents a major
flooding surface on the platform and is featured worldwide by deposition of
“hot shales” during the early Frasnian (Lüning et al., 2003, 2004; Wendt et al., 2006).
SE–W cross section between the Arak-Foum Belrem Arch, the
Mouydir Basin, and the Amguid El Biod Arch. Outcrop cross-section
correlations and biostratigraphic calibrations are based on the compilation
of published papers (Wendt et al., 2006, 2009b).
Interpretation of the basement is based on Figs. 1, 2, and 15. Outcrop
location is in Fig. 2.
The second pattern extends from the “Mehden Yahia”, “Temertasset” to the
“Khenig” formations dated Frasnian to Lower Tournaisian. This pattern is
composed of part of D5–D9 medium-scale sequences. It corresponds to
homogenous offshore depositional environments with no lateral facies
variations. However, local deltaic (fluvio–marine) conditions are observed
during the Frasnian at the Arak Foum Belrem Arch (“Grès de Mehden
Yahia” in Fig. 12). A successive alternation of shoreface and offshore
deposits is organized into five medium-scale sequences (part of D5, and D6
to D9; Figs. 10, 11 and 12). They, in particular, show some regressive phases
with the deposition of both “Grès de Mehden Yahia” and “Grès du
Khnig” sandstones (bounded by HD6 and HD9). This pattern (i.e., part of D5
to D9) corresponds to the general maximum flooding (Lüning et al., 2003,
2004; Wendt et al., 2006) under eustatic control with no tectonic
influences.
NE–W cross section between the Reggane Basin, the Azzel Matti Arch,
the Ahnet Basin, the Arak-Foum Belrem Arch, the Mouydir Basin, and the Amguid El Biod Arch.
The well W18 biozone calibration is based on Kermandji
et al. (2009): biozone (Tm) tidikeltense microbaculatus (Lochkovian, Lower Devonian), biozone (Es)
emsiensis spinaeformis (Lochkovian–Pragian, Lower Devonian), biozone (Ac) arenorugosa caperatus (Pragian, Lower
Devonian), biozone (Ps) poligonalis subgranifer (Pragian–Emsian, Lower Devonian), biozone (As)
annulatus svalbardiae (Emsian, Lower Devonian), biozone (Mp) microancyreus protea (Emsian–Eifelian, Lower to Middle
Devonian), and biozone (Vl) velatus langii (Eifelian, Middle Devonian). The well W19 and W20
biozones calibration from internal reports (Abdesselam-Rouighi,
1991; Khiar, 1974) is based on the Magloire (1967)
classification: biozone H (Pridoli, Upper Silurian), biozone I (Lochkovian,
Lower Devonian), biozone J (Pragian, Lower Devonian), biozone K (Emsian,
Lower Devonian), and biozone L1-5 (Middle Devonian to Upper Devonian).
Interpretation of the basement is based on Figs. 1, 2, and 15. Outcrop and
well location is in Fig. 2.
Subsidence and tectonic history: an association of low rate extensional
subsidence and positive inversion pulses
The backstripping approach (Fig. 14) was applied to five wells (W1, W5, W7,
W17, and W21). The morphology of the backstripped curve and subsidence rates
can provide clues as to the nature of the sedimentary basin (Xie
and Heller, 2009).
In intracratonic basins, reconstructed tectonic
subsidence curves are almost linear to gently exponential in shape, similar
to those of passive margins and rifts (Xie and Heller, 2009). The
compilation of tectonic backstripped curves from several wells in
peri-Hoggar basins (Fig. 14a, see Fig. 1 for location) and from wells in the
study area (Fig. 14b) display low rates of subsidence (from 5 to 50 m Myr-1)
organized in subsidence patterns of: inversion of the low rate subsidence
(ILRS type c, red line, Fig. 14c), deceleration of the low rate subsidence
(DLRS type b, black line), and acceleration of the low rate subsidence (ALRS
type a, blue line).
N–S cross section in the Ahnet Basin between Azzel Matti Arch
and Arak-Foum Belrem Arch; well W7 biozone calibration from
Azzoune, (1999) internal report based on the Boumendjel (1987) classification: biozones 7–12 (Lochkovian, Lower Devonian), biozone
15 (Emsian, Lower Devonian). Well W18 biozone calibration is based on
Kermandji et al. (2009): biozone (Tm) tidikeltense microbaculatus (Lochkovian,
Lower Devonian), biozone (Es) emsiensis spinaeformis (Lochkovian–Pragian, Lower Devonian), biozone
(Ac) arenorugosa caperatus (Pragian, Lower Devonian), biozone (Ps) poligonalis subgranifer (Pragian–Emsian, Lower
Devonian), biozone (As) annulatus svalbardiae (Emsian, Lower Devonian), biozone (Mp)
microancyreus protea (Emsian–Eifelian, Lower to Middle Devonian), and biozone (Vl) velatus langii (Eifelian,
Middle Devonian). The well W12 biozone calibration from
Abdesselam-Rouighi (1977) internal report is based on the
Boumendjel (1987) classification: biozone J (Pragian, Lower
Devonian), biozone K (Emsian, Lower Devonian), biozone L1 (Eifelian, Middle
Devonian), and biozone L7-3, L7-9 (Frasnian–Famennian, Upper Devonian).
Interpretation of the basement is based on Figs. 1, 2, and 15. Outcrop and
well location is in Fig. 2.
Each period of ILRS, DLRS, and ALRS may be synchronous among the different
wells studied (see B1 to J, Fig. 14b) and some wells from published data (see
D to J Fig. 14a).
The Saharan Platform is marked by a rejuvenation of basement structures,
around arches (Figs. 1, 2, and 3), linked to regional geodynamic pulses
during Neoproterozoic to Paleozoic times (Fig. 14). A compilation of the
literature shows that the main geodynamic events are associated with
discriminant association of subsidence patterns:
(a) Tectonic backstripped curves of the Paleozoic North Saharan
Platform (peri-Hoggar basins) compiled from literature. 1: HAD-1 well in
Ghadamès Basin (Makhous and Galushkin, 2003b); 2: well
RPL-101 in Reggane Basin (Makhous and Galushkin, 2003b); 3: L1-1
well in Murzuq Basin (Galushkin and Eloghbi, 2014); 4: TGE-1 in
Illizi Basin (Makhous and Galushkin, 2003a); 5: REG-1 in
Timimoun Basin (Makhous and Galushkin, 2003b); 6: Ghadamès-Berkine Basin (Allen and
Armitage, 2011; Yahi, 1999); 7: well in Sbâa Basin
(Tournier, 2010); and 8: well B1NC43 in Al Kufrah Basin
(Holt et al., 2010). (b) Tectonic backstripped curves of
wells in the study area 1: well W17 in Ahnet Basin; 2: well W5 in Ahnet
Basin; 3: well W7 in Ahnet Basin; 4: well W21 in Mouydir Basin; and 5: well W1
in Reggane Basin. (c) Typologies of subsidence curves morphologies. A: Late
Pan-African compression and collapse (type a, b, and c subsidence), B: undifferentiated Cambrian–Ordovician (type a, b, and c subsidence),
B1: Cambrian–Ordovician tectonic quiescence (type a subsidence), B2: Cambrian–Ordovician extension (type b subsidence), C: Late Ordovician
glacial and isostatic rebound (type c subsidence), D: Silurian extension
(type b subsidence), E: Late Silurian Caledonian compression (type c
subsidence), F: Early Devonian tectonic quiescence (type a subsidence), G–H: Middle to late Devonian extension with local compression (i.e., inversion
structures, type b and c subsidence), I: Early Carboniferous extension with
local tectonic pre-Hercynian compression (type c and b subsidence), and J: Middle Carboniferous tectonic extension (type b subsidence).
Late Pan-African compression and collapse (patterns a, b, and c, A Fig. 14a).
The infra-Cambrian (i.e., top Neoproterozoic) is characterized by horst
and graben architecture associated with wedge-shaped unit DO0 in the
basement (Fig. 7). This structuring, probably related to Pan-African
post-orogenic collapse, is illustrated by intracratonic basins infilled with
volcano-sedimentary molasses series
(Ahmed and Moussine-Pouchkine, 1987; Coward and Ries, 2003; Fabre et al., 1988;
Oudra et al., 2005).
Cambrian–Ordovician geodynamic pulse (Fig. 14). Highlighted by the
wedge-shaped units DO1 (Figs. 6a and 7), the horst and graben system is
correlated with deceleration (DLRS pattern a, B1) and with local
acceleration of the subsidence (ALRS pattern b, B2). The
Cambrian–Ordovician extension is documented on arches (Arak-Foum Belrem,
Azzel Matti, Amguid El Biod, Tihemboka, Gargaf, Murizidié, Dor El Gussa,
etc.) of the Saharan Platform by synsedimentary normal faults, reduced
sedimentary successions
(Bennacef
et al., 1971; Beuf et al., 1968a, b, 1971; Beuf and Montadert, 1962;
Borocco and Nyssen, 1959; Claracq et al., 1958; Echikh, 1998; Eschard et
al., 2010; Fabre, 1988; Ghienne et al., 2003, 2013; Zazoun and Mahdjoub,
2011), and by stratigraphic hiatuses
(Mélou
et al., 1999; Oulebsir and Paris, 1995; Paris et al., 2000; Vecoli et al.,
1995, 1999).
Late Ordovician geodynamic pulse (i.e., Hirnantian glacial and isostatic
rebound; Fig. 14). Late Ordovician incisions mainly situated at the hanging
walls of normal faults (Fig. 7c and d) are interpreted as Hirnantian
glacial-paleovalleys (Le Heron, 2010; Smart, 2000) and
followed by local inversion of low rate subsidence (ILRS of type c, C in
Fig. 14).
Silurian extensional geodynamic pulse (D, Fig. 14). The Silurian
post-glaciation period is featured by the reactivation and sealing of the
inherited horst and graben fault system (i.e., wedge-shaped unit DO2; Figs. 6b, c, 7a and b).
It is linked to an acceleration of the subsidence (ALRS of pattern b in Fig. 14). This tectonic extension is documented in seismic
records (Najem et al., 2015) and is associated with the Silurian major
transgression on the Saharan Platform
(e.g., Eschard et al., 2005;
Lüning et al., 2000).
Late Silurian to Early Devonian geodynamic pulse (Caledonian
compression; E Fig. 14). Late Silurian times are marked by reactivation and
local positive inversion of the former structures (Figs. 6c and 7b); this occurs due to
truncations located at fold hinges (Figs. 6c and 7) and due to a major shift
from marine to fluvial/transitional environments (e.g., Fig. 10).
Backstripped curves register an inversion of the subsidence (ILRS of pattern
c, in Fig. 14). The Caledonian event is mentioned as related to large-scale
folding or uplifted arches (e.g., the Gargaff, Tihemboka, Ahara,
Murizidé-Dor el Gussa and Amguid El Biod arches) and it is associated
with breaks in the series and with angular unconformities
(Beuf
et al., 1971; Biju-Duval et al., 1968; Boote et al., 1998; Boudjema, 1987;
Boumendjel et al., 1988; Carruba et al., 2014; Chavand and Claracq, 1960;
Coward and Ries, 2003; Dubois and Mazelet, 1964; Echikh, 1998; Eschard et
al., 2010; Fekirine and Abdallah, 1998; Follot, 1950; Frizon de Lamotte et
al., 2013; Ghienne et al., 2013; Gindre et al., 2012; Legrand, 1967a, b;
Magloire, 1967).
Early Devonian tectonic quiescence (F Fig. 14). This is characterized by
a deceleration of the low rate subsidence (DLRS of pattern a, F in Fig. 14).
During this period, we have detected Emsian truncation from satellite images
(Fig. 6d and e) and erosion and pinch out of Upper Emsian to Eifelian
series from well cross sections (Figs. 10, 12 and 13). In previous works,
these hiatuses/gaps (i.e., Upper Lochkovian, Lower Pragian, Upper Pragian,
Upper Emsian, Lower Eifelian) are observed in the Ahnet Basin
(Kermandji, 2007;
Kermandji et al., 2003, 2008, 2009; Wendt et al., 2006), in the Illizi
(Boudjema, 1987) and in the Reggane (Jäger et al., 2009).
Middle to late Devonian geodynamic pulse (extension and local
inversions, G and H Fig. 14). The Middle to Late Devonian period is
characterized by large wedge hiatuses and truncations associated with the
reactivation of horst and graben structures and local positive inversion
(OD3 in Figs. 6d, e, f, 7 and 10 to 13). This period is characterized by
inversion and acceleration of low rate subsidence (patterns c and b: ILRS –
ALRS, Fig. 14). Some of the Middle to Late Devonian syntectonic structures
and hiatuses (e.g., Givetian/Frasnian) are noticed in the Ahnet Basin
(Wendt et al., 2006), on the Amguid Ridge
(Wendt et al., 2009b), in the Illizi Basin
(Boudjema, 1987; Chaumeau et al.,
1961; Eschard et al., 2010; Fabre, 2005; Legrand, 1967a), on the Gargaf
(Carruba et al., 2014; Collomb, 1962;
Fabre, 2005; Massa, 1988) and elsewhere on the platform (Frizon
de Lamotte et al., 2013).
Pre-Hercynian to Hercynian geodynamic pulses (I and J Fig. 14).
This period is organized in Early Carboniferous pre-Hercynian (I, Fig. 14)
to Late Carboniferous–Early Permian Hercynian compressions limited by Mid
Carboniferous tectonic quiescence/extension (J, Fig. 14). The Carboniferous
period is characterized by a normal reactivation and local positive
inversion of the previous structural patterns involving reverse faults,
overturned folds, transpressional flower structures along strike-slip fault
zones (Figs. 6f, 7b, c and d). The major Carboniferous tectonic event on
the Saharan Platform impacted all arches and it is mainly controlled by
near-vertical basement faults with a strike-slip component
(Boote
et al., 1998; Caby, 2003; Carruba et al., 2014; Haddoum et al., 2001, 2013;
Liégeois et al., 2003; Wendt et al., 2009a; Zazoun, 2001, 2008).
According to these authors basement fabric features exerted a very strong
control on the structural evolution during the Hercynian deformation. Two
major hiatuses (i.e., Mid Tournaisian to Mid Visean–Serpukhovian) are
recognized (Wendt et al., 2009a).
The geodynamic pulses attest to the reactivation of the terranes and
associated lithospheric fault zones. This observation questions the nature
of the Precambrian basement and associated structural heritage.
(a) Interpreted aeromagnetic anomaly map
(https://www.geomag.us/, last access: 1 December 2016) of the Paleozoic
North Saharan Platform (peri-Hoggar basins) showing the different terranes
delimited by N–S, NW–SE, and NE–SW lineaments and megasigmoid structures
(SC – shear fabrics). (b) Bouguer
anomaly map (from the International Gravimetric Bureau:
http://bgi.omp.obs-mip.fr/, last access: 1 December 2016) of the North
Saharan Platform (peri-Hoggar basins) presenting evidence of positive
anomalies under arches and negative anomalies under basins.
Geochronological data show that the different terranes were reworked during
several main thermo-orogenic events. The two main events deduced from
geochronological data are the Neoproterozoic (i.e., Pan-African) and
Paleoproterozoic (i.e., Eburnean) episodes
(Bertrand and Caby, 1978).
Aeromagnetic anomaly surveys are commonly used to analyze geological
features such as rock types and fault zones (e.g.,
Turner et al., 2007). A similar study was led in the meantime showing
similar interpretations (Bournas
et al., 2003; Brahimi et al., 2018). In this study, these data highlight the
geometries and the extension of the different terranes under the sedimentary
cover. Four main domains can be identified from the aeromagnetic anomaly
map, delimited by contrasted magnetic signatures and interpreted as suture
zones (thick black lines, Fig. 15a). The study area is bounded to the south
by the Tuareg Shield (TS), to the north by the south Atlassic Range, to the
west by the West African Craton (WAC), and to the east by the East Saharan
Craton (ESC) or Saharan Metacraton (Abdelsalam et al., 2002).
The magnetic disturbance features (Fig. 15a) show three main magnetic
trends. A major N–S sinuous fabric and two minor sinuous 130–140∘ E and N45∘ E trends. The major N–S lineaments coincide with terrane
boundaries and mega-shear zones (e.g., 4∘50′, 4∘10′,
WOSZ, EOSZ, 8∘30′, RSZ shear zones; Fig. 1). Sigmoidal-shaped
terranes 200 to 500 km long and 100 km wide are characterized (red lines in
Fig. 15a). The whole assemblage forms a typical SC-shaped shear fabric
(Choukroune et al., 1987) associated with vertical
mega-shear zones and suture zones (e.g., WOSZ, EOSZ, 4∘10′,
4∘50′ or 8∘30′ Hoggar shear zones in Fig. 1). The SC
fabrics combined with subvertical lithospheric shear zones (Fig. 16b and c)
are typical features of the Paleoproterozoic accretionary orogens
(Cagnard et al., 2011; Chardon et al., 2009).
This architecture is concordant with the Neoproterozoic collage of the
Tuareg Shield (i.e., mobile belt) between the West African Craton and the
East Saharan Craton (i.e., cratonic blocks) described by
(Coward and Ries, 2003; Craig et al., 2008).
The gravimetric anomaly map (Fig. 15b) shows a correlation between
gravimetric anomalies and tectonic architecture (intracratonic
syncline-shaped basin and neighboring arches). Positive anomalies
(> 66 mGal) are mainly associated with arches, whereas negative
anomalies are related to intracratonic basins (< 66 mGal).
Nevertheless, negative anomaly disturbance is found in the Hoggar Massif
probably due to Cenozoic volcanism and the Hoggar swell
(Liégeois et al., 2005) or to Eocene
Alpine intraplate lithospheric buckling
(Rougier et al., 2013).
(a) Interpreted map of basement terranes according to their age
(compilation of datasets in Fig. 1 and Supplement data 1); (b) Satellite
images (7ETM+ from USGS: https://earthexplorer.usgs.gov/, last access: 29 October 2016) of
Paleoproterozoic Issalane-Tarat terrane, central Hoggar (see C for
location). (c) Interpreted satellite images of Paleoproterozoic
Issalane-Tarat terrane showing sinistral sigmoid mega-structures associated
with transcurrent lithospheric shear fabrics (SC).
The Precambrian structural heritage is characterized by accreted
lithospheric terranes limited by vertical strike-slip mega-shear zones (Fig. 16b and c). A zonation is observed between the Paleozoic basins and arches
configurations and the different terranes (thermo-tectonic age). Arches are
linked to Archean to Paleoproterozoic continental terranes in contrast to
syncline-shaped basins which are associated with Meso-Neoproterozoic
terranes (Figs. 1, 2 and 16a).
(a) Different structural model styles identified from the
analysis of seismic profiles and from interpretation of the satellite
images. (b) A conceptual model of the architecture of an intracratonic low rate
subsidence basin and synthesis of the tectonic kinematics during the
Paleozoic. Note that the differential subsidence between arches and basins
is controlled by terrane heterogeneity (i.e., thermo-chronologic age,
rheology, etc.).
Low subsidence rate intracratonic Paleozoic basins of the central Sahara
provide a basis for an integrated modeling study
Paleozoic intracratonic basins with similar characteristics (architecture,
subsidence rate, stratigraphic partitioning, alternating episodes of
intraplate extension, and short duration compressions with periods of
tectonic quiescence, etc.) have been documented in North America
(e.g., Allen and Armitage, 2011; Beaumont et al., 1988; Burgess, 2008; Burgess et
al., 1997; Eaton and Darbyshire, 2010; Pinet et al., 2013; Potter, 2006;
Sloss, 1963; Xie and Heller, 2006), South America
(Allen and Armitage, 2011; de Brito Neves et al., 1984; Milani and Zalan, 1999; de
Oliveira and Mohriak, 2003; Soares et al., 1978; Zalan et al., 1990), Russia
(Allen and Armitage, 2011; Nikishin et al., 1996) and Australia
(Harris, 1994; Lindsay and Leven, 1996; Mory et al., 2017). However, the nature of the potential
driving processes (lithospheric folding, far-field stresses, local increase
in the geotherm, mechanical anisotropy from lithospheric rheological
heterogeneity, etc.) associated with the formation of intracratonic
Paleozoic basins remains highly speculative
(Allen and Armitage, 2011; Armitage and Allen, 2010; Braun et al., 2014; Burgess
and Gurnis, 1995; Burov and Cloetingh, 2009; Cacace and Scheck-Wenderoth,
2016; Célérier et al., 2005; Gac et al., 2013; Heine et al., 2008;
Leeder, 1991; Vauchez et al., 1998).
The multiscale and multidisciplinary analysis performed in this study enable
us to document a model of Paleozoic intracratonic central Saharan basins
that couples basin architecture and basement structures (Fig. 17). While we do
not provide any quantitative explanations for the dynamics of these basins,
our synthesis highlights that their subsidence is not the result of a single
process and we attempt to make a check-list here of the properties that a
generic model of formation of such basins must capture:
The association of syncline-shaped wide basins and neighboring arches
(i.e., paleo-highs). The structural framework shows a close association of
syncline-shaped basins, interbasin principal to secondary arches, and
intra-basin secondary arches (see Fig. 2).
By local horst and graben architecture linked to steep-dipping planar
normal faults and associated with normal fault-related fold structures (i.e.,
forced folds; a, Fig. 17a). Locally, the extensional structures are
disrupted by positive inversion structures (b, Fig. 17a) or transported
normal faults (c, Fig. 17a).
A low rate of subsidence ranging between 5 to 50 m Myr-1 (Fig. 14).
Long periods of extension and tectonic quiescence are interrupted by
brief periods of compression or glaciation/deglaciation events
(Beuf et al., 1971; Denis et al., 2007;
Le Heron et al., 2006). These periods of compression are possibly related to
intraplate compression linked to distal orogenies (i.e., Late Silurian
Caledonian event, Late Carboniferous Hercynian, (Frizon de
Lamotte et al., 2013) or to intraplate arch uplift related to magmatism
(Derder et al., 2016; Fabre, 2005; Frizon de Lamotte et al., 2013; Moreau et al., 1994).
Synsedimentary divergent onlaps and local unconformities are identified
from integrated seismic data, satellite images, and borehole data (Figs. 5,
6, 7 and 10 to 13). The periods of tectonic activity are characterized by
normal to reverse reactivation of border faults, emplacement of wedge-shaped
units, and erosional unconformities neighboring the arches.
The stratigraphic architecture displays a lateral facies variation and
partitioning between distal marine facies infilling the intracratonic basins
(i.e., offshore deposits) and proximal amalgamated facies (i.e.,
fluvio–marine, shoreface) associated with prominent stratigraphic hiatus and
erosional unconformities in the vicinity of the arches.
A close connection is evidenced between the period of tectonic
deformation and the presence of erosional unconformities (i.e., 2, 3, 6, 8,
10 geodynamic events in Fig. 17b). By contrast, the periods of tectonic
quiescence and extension are characterized by low lateral facies variations,
thin deposits, and the absence of erosional surfaces.
The Precambrian heritage corresponds to Archean to Paleoproterozoic
terranes identified in the Hoggar Massif and reactivated during the
Meso-Neoproterozoic Pan-African cycle (Fig. 1). The Precambrian lithospheric
heterogeneity illustrated by the different characteristics of Precambrian
terranes (wavelength, age, nature, fault zones) spatially control the
emplacement of the syncline-shaped intracratonic basins underlain by
Meso–Neoproterozoic oceanic terranes and the arches underlain by Archean to
Paleoproterozoic continental terranes (Figs. 1, 2 and 16). Many authors
suggest control of the basement fabrics is inherited from the Pan-African
orogeny in the Saharan basins (Beuf et al., 1968b, 1971; Boote et al., 1998; Carruba et al., 2014; Coward and
Ries, 2003; Eschard et al., 2010; Guiraud et al., 2005; Sharata et al.,
2015).
Conclusions
Our integrated approach using both geophysical (seismic, gravity,
aeromagnetic, etc.) and geological (well, seismic, satellite images, etc.)
data has enabled us to decrypt the characteristics of the intracratonic
Paleozoic Saharan basins and the control of the heterogeneous lithospheric
heritage of the horst and graben architecture, low rate subsidence,
and the association of long-lived broad synclines and anticlines (i.e., arches
swells, domes, highs or ridges) with very different wavelengths (λ)
(tens to hundreds of kilometers). A coupled basin architecture and basement
structures model is proposed (Fig. 17).
This study highlights a tight control of the heterogeneous lithosphere
zonation over the structuring of the intracratonic central Saharan Basin.
This particular type of basin is characterized by a low rate of subsidence
and fault activation controlling the homogeneity of sedimentary facies and
the distribution of the main unconformities. The low rate activation of
vertical mega-shear zones bounding the intracratonic basin during Paleozoic
times contrasts markedly with classic rift kinematics and architecture.
Three different periods of tectonic compressional pulses (i.e., Caledonian,
Middle to Late Devonian, and pre-Hercynian), extension, and quiescence are
identified and controlled the sedimentary distribution (Fig. 17). An
understanding of tectono-sedimentary interaction is key to understanding the
distribution of the Paleozoic petroleum reservoirs of this first-order oil
province.
Seismic and well log data analysed in this study are part
of the Neptune Energy/SONATRACH internal database. Unfortunatly, they are not
publicly available. Nonetheless, satellite images and geophysical data are
all available (see data and methods).
The Supplement related to this article is available online at: https://doi.org/10.5194/se-9-1239-2018-supplement.
The strucutral seismic and
photogeology interpretation as well as the basement interpretation and
analyses throughout this study were mainly undertaken by PP and MG.
Interpretations of the well logs, the sedimentology, and the sequence
stratigraphy were primarily carried out by PP and EV. Backstripping was led
by PP and controlled by IM and EP. The paper was written by PP, with
additional input and scientific editing from MG and EV. All authors
contributed to the technical interpretation, extensive discussions, and ideas
throughout the study and the writing of the paper.
The authors declare that they have no conflict of
interest.
Acknowledgements
We are most grateful to Neptune Energy and ENGIE which provided the database used in this paper and funded the
work. Special thanks to the data management service of Neptune Energy
(especially Aurelie Galvani) for their help with the database. Thanks also go
to Jobst Wendt, Réda Samy Zazoun, and Fabio Lottaroli for detailed
reviews/comments , along with a short comment from Alexander Peace, which
considerably enriched and improved this paper. Edited by: Mark Allen Reviewed by: Reda Samy
Zazoun, Fabio Lottaroli, and Jobst Wendt
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