# Initial Pangean rifting north of the West African Craton: Insights from late Permian U-Pb and 40Ar/39Ar dating of alkaline magmatism from the Eastern Anti-Atlas (Morocco)

Amine Najih, Pilar Montero, Chrystèle Verati, Moulley Charaf Chabou, Abdelilah Fekkak, Lahssen Baidder, Hassan Ezzouhairi, Fernando Bea, André Michard

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## ► To cite this version:

Amine Najih, Pilar Montero, Chrystèle Verati, Moulley Charaf Chabou, Abdelilah Fekkak, et al. Initial Pangean rifting north of the West African Craton: Insights from late Permian U-Pb and 40Ar/39Ar dating of alkaline magmatism from the Eastern Anti-Atlas (Morocco). Journal of Geodynamics, 2019, 132, pp.101670. 10.1016/j.jog.2019.101670. hal-03575924

HAL Id: hal-03575924

https://hal.science/hal-03575924v1

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# Initial Pangean rifting north of the West African Craton: Insights from late Permian U-Pb and ⁴⁰Ar/³⁹Ar dating of alkaline magmatism from the Eastern Anti-Atlas (Morocco)

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Amine Najih¹, Pilar Montero², Chrystèle Verati³, Moulley Charaf Chabou⁴, Abdelilah

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Fekkak¹, Lahssen Baidder⁵, Hassan Ezzouhairi¹, Fernando Bea², André Michard⁶*

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¹Chouaib Doukkali University, Faculty of Sciences, Earth Sciences Department, El-Jadida, Morocco. E-mails najih.a@ucd.ac.ma (A.N.); abfekkak@gmail.com (A.F.); ezzouhairi@hotmail.com (H.E.)

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²Granada University, Faculty of Sciences, Department of Mineralogy and Petrology, Granada, Spain. E-mail pilar_montero@me.com

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³ Université Côte d'Azur, Nice Sophia Antipolis University, Observatoire de la Côte d'Azur, UMR CNRS Géoazur, 06560 Sophia Antipolis, France. E-mail chrystele.verati@univ-cotedazur.fr

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⁴Université Ferhat Abbas, Sétif 1, Institut d'Architecture et des Sciences de la Terre, Département des Sciences de la Terre, Campus El-Bez, 19000, Sétif, Algeria. E-mail charaf.chabou@univ-setif.dz

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⁵ Hassan II University, Faculty of Sciences Ain Chock, Geosciences Laboratory, Casablanca, Morocco.

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⁶ Em. Pr. Paris-Sud University, Earth Sciences Department, 91405 Orsay Cedex, France. E-mail andremichard@orange.fr

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* Corresponding author

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## Abstract

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Numerous mafic dykes, sills and laccoliths crop out in the southern part of the Tafilalt basin (Eastern Anti-Atlas, Morocco). These rocks intrude the mildly folded Ordovician to Early Carboniferous formations, consisting mainly of lamprophyric dolerites and camptonites with minor gabbros and syenodiorites. Previous geochemical studies have shown that the Tafilalt magmatism of sodic-alkaline affinity has been produced by low degrees of partial melting from an enriched deep mantle source within the garnet stability field. However, the age and the geodynamic context of these rocks were presently unknown since no isotopic dating had
so far been made of the Tafilalt dolerites. To resolve this issue, we present here the first
$^{40}$Ar/$^{39}$Ar biotite and U-Pb zircon dating from the Tafilalt alkaline magmatism. Three samples
(biotite separates) yielded well-defined $^{40}$Ar/$^{39}$Ar plateau ages of 264.2 ± 2.7 Ma, 259.0 ± 6.3
Ma and 262.6 ± 4.5 Ma whereas $^{206}$Pb/$^{238}$U dating of zircon from one of these samples yielded
an age of 255 ± 3 Ma. These ages coincide within the dating error, and indicate that this
magmatism occurred in the late Permian. Considering geochronological and geochemical
data, we propose that the Tafilalt magmatism reflects an early-rift magmatic activity that
preceded the Triassic rifting heralded by the Central Atlantic mMagmatic Province. This
magmatic activity is recorded in both sides of the future Atlantic Ocean by small-volume
alkaline magmatism that started in the late Permian and extends into the Triassic. The alkaline
magmas are probably generated in response to an increase in the mantle potential temperature
underneath the Pangea supercontinent.

41 **Keywords:** Gondwana margin; pre-CAMP magmatism; Late Permian; Alkaline dolerites;
42 Morocco

1. Introduction

Different magmatic events generally record each step of the evolution of collisional belts
from birth to collapse and to continental breakup. We present in the following an example of
magmatism postdating the latest steps of the Variscan evolution at the NW margin of
Gondwana. This example is taken from the Late Devonian-Early Carboniferous South Tafilalt
Basin in the eastern Anti-Atlas belt, which is the foreland belt of the Variscan Meseta Orogen
of Morocco (Fig. 1A). During the Paleozoic and Mesozoic, Morocco has been the site of
widespread magmatism associated with various geodynamic settings (Ezzouhairi et al., 2008;
Ikenne et al., 2017; Álvaro et al., 2018) : (i) Cambrian-Ordovician magmatism related to the
rifting of the NW Gondwana margin (Nance et al., 2012); (ii) typical Variscan magmatism
(Essaifi et al., 2019), including synsedimentary to syntectonic Late Devonian-Early Carboniferous bimodal magmatism, syn- to post-tectonic Serpukhovian-Pennsylvanian granite emplacement, and late orogenic, Early Permian (Cisuralian) granite and ignimbrites emplacement (Youbi et al., 1995; Roddaz et al., 2002; Saidi et al., 2002; Hoepffner et al., 2005; El Hadi et al., 2006); (iii) the dykes, sills and lava flows of the Central Atlantic Magmatic Province (CAMP) emplaced at the Triassic/Jurassic boundary during the Pangean rifting that led to the opening of the Central Atlantic Ocean (Marzoli et al., 1999; Youbi et al., 2003; Marzoli et al., 2018, 2019).

In the Eastern Anti-Atlas, precisely in the southern part of the Tafilalt basin (Fig. 1B), several occurrences of dolerites and gabbros have been recognized for long, hereafter labeled the South Tafilalt magmatic rocks. They form dykes, sills and laccoliths hosted in the mildly folded Ordovician-Visean series. They were firstly mapped in the historical Georges Choubert's map, scale 1:500,000 (1950) as Paleozoic, "Hercynian dolerites" (δ₁H), due to their appearance of being folded together with the hosting sedimentary rocks. Conversely, in the geological map of Morocco, scale 1:200,000, sheet Tafilalt-Taouz (Destombes and Hollard, 1986), they were noted as Triassic-Liassic dolerites by comparison with the dolerites of the Tindouf basin (Zemoul area), which have been just dated at 180-187 Ma (K-Ar method; Hailwood and Mitchell, 1971) when the draft of the map was achieved (1975) and recognized as pseudo-folded intrusions emplaced in previously folded beds (Hollard, 1973). In the new geological maps of the South Tafilalt region, scale 1:50,000, sheets Al-Atrous, Irara, Marzouga, Mfis, Taouz (2014), they are left undated, but the corresponding explanatory memoirs suggest, although cautiously, that these rocks emplaced during the Late Devonian and the Visean (e.g., Benharref et al., (2014a, 2014b, 2014c)). Lastly, Pouclet et al., (2017), based on a synthesis of the geochemical characters of the Tafilalt magmatic rocks, reaffirmed that they not emplaced during the Late Triassic, but during two pre-Variscan pulses, i.e.,
during the Famennian-Tournaisian and the Early Visean. Contrastingly, and based on a geochemical discussion, Chabou et al., (2017a, 2017b) stressed, too, that the Tafilalt lamprophyric dolerites differ from the CAMP tholeiitic rocks, but they emphasized that these intrusions also differ from the Devonian-Carboniferous intrusions of the surrounding areas (Moroccan Meseta, southern Hoggar). They suggested that “the Tafilalt mafic rocks could represent the Permian-Triassic (?) alkaline precursor of the CAMP magmatism itself as observed elsewhere (Spain, Nova Scotia).”

Our aim in the present work is essentially to report on the first isotopic dating of zircon and biotite minerals from the South Tafilalt intrusions. We show in the following that these geochronological results confirm Chabou et al. (2017a, 2017b) hypothesis and demonstrate that the South Tafilalt magmatism is a late Permian event. Eventually, we discuss the significance of this magmatism in the geodynamic frame of the Pangea amalgamation and breakup.

## 2. Geological setting

The Tafilalt region extends at the eastern tip of the Anti-Atlas, where the latter connects with the Ougarta belt (Fig. 1A). The Anti-Atlas and Ougarta belts both involve gently folded Upper Ediacaran-Carboniferous sequences overlying unconformably a faulted Pan-African basement that crops out in faulted antiforms. The Neoproterozoic Pan-African belt, and the superimposed Anti-Atlas and Ougarta Paleozoic belts are molded against the northern part of the West African Craton (WAC; Ennih and Liégeois, 2008; Gasquet et al., 2008; Michard et al., 2017). The Anti-Atlas can be regarded as the thick-skinned foreland belt of the Variscan Orogen, which is exposed in two distinct domains, i) the Mauritanide thrust-fold belt cropping out south of latitude 28° along the Atlantic margin, and ii) the Moroccan Meseta Paleozoic belt outcropping in the Atlas-Meseta basement massifs north of the South Atlas Fault
102 (Burkhard et al., 2006; Michard et al., 2010). The Anti-Atlas Paleozoic sequence begins with
103 the Early Cambrian in the Western Anti-Atlas, but not before the late Early Cambrian or
104 Middle Cambrian in the Eastern Anti-Atlas (Destombes et al., 1985; Geyer and Landing,
105 1995) (Fig. 2). This results from the activity of synsedimentary ENE-trending normal faults,
106 also responsible for alkaline basalt outpours during the Middle Cambrian (Raddi et al., 2007;
107 Soulaimani et al., 2014; Pouclet et al., 2017). The Ordovician-Silurian period is characterized
108 by sandy to argillaceous deposits, only perturbed during the Hirnantian glacial events (
109 Destombes et al., 1985; Clerc et al., 2013; Ghienne et al., 2014). The Middle-Upper Devonian
110 deposits show dramatic thickness and facies variations illustrating the coeval disintegration of
111 the northern Saharan platform (Wendt, 1985; Baidder et al., 2008; Frizon de Lamotte et al.,
112 2013). The South Tafilalt basin formed at that time south of the Ougnat-Erfoud high, and east
113 of the Ougnat-Ouzina Axis (Fig. 1B). The youngest terms of the folded sequence are late
114 Visean in age (Benharref et al., 2014). Post-folding, molasse-type subaerial deposits,
115 Bashkirian and Pennsylvanian in age (from ca. 320 to 300 Ma; Cavaroc et al., 1976; Fabre,
116 2005; Cózar et al., 2014) are preserved in the Tindouf cratonic basin south of the Anti-Atlas.
117 In contrast, marine deposits accumulated up to the late Moscovian (ca. 305 Ma) in the
118 Bechar-Abadla Basin to the east of the Ougarta belt. The Abadla basin received subaerial red
119 bed deposits during the late Westphalian-Autunian (Fabre, 1976, 2005; Bouabdallah et al.,
120 1998), suggesting diachronic folding of the western Anti-Atlas and eastern Anti-Atlas-
121 Ougarta belts.

122 During the Triassic, the eroded Anti-Atlas Paleozoic fold belt remained emerged as the
123 common shoulder of the Atlantic and Tethyan (Atlas) rifts (Frizon de Lamotte et al., 2008).
124 However, the Anti-Atlas and adjacent areas (Bechar Basin, Atlas and Meseta domains) were
125 intruded by numerous dykes and sills and covered (locally or extensively) by lava piles of the
126 Central Atlantic Magmatic Province, dated c. 201 Ma (Sebai et al., 1991; Derder et al., 2001;
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Youbi et al., 2003; Chabou et al., 2007; Meddah et al., 2007; Verati et al., 2007; Dal Corso et al., 2014). Continental or shallow marine sediments have covered the Anti-Atlas basement during the Cretaceous-Neogene, and now form the Kem-Kem and Guir Hamada plateaus (Zouhri et al., 2008).

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**3. Sampling and methods**

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**3.1. Mapping and petrographic analysis**

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Mapping of the South Tafilalt area was firstly performed by one of us (A.F.) in the frame of the South Tafilalt mapping program of the Moroccan Ministry of Energy and Mines, which resulted in the five sheets, scale 1:50,000, cited above. Detail mapping and sampling by the first author (A.N.) was then directed towards the intrusive rocks south of Rissani, the Tafilalt capital (Fig. 1B). Numerous dolerite dykes can be observed there crosscutting down to the Ordovician formations of the Ougnat-Ouzina axis between Taouz and Ouzina (Fig. 3A). The first-order, kilometer-long dykes follow the main fault zones and their magmatic material is rich in xenoliths and phenocrysts, especially biotite, and frequently brecciated. The second-order dykes that ensure vertical communication between sills and laccoliths may contain fragments of reworked first-order dykes.

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Dolerite sills are located in the most argillaceous, shaly or silty levels of the Silurian, Devonian and Carboniferous, up to the Upper Visean (Figs. 2, 3A). Although they are mostly localized in the weakness zones between incompetent and competent layers, sills may change laterally their position in the stratigraphic pile. Laccolith architecture is observed in two areas i) in the mushroom-shaped Znaigui magmatic body, at the base of the lower Famennian marls, and ii) at Jbel Dboa in the core of the Mfis anticline. In the latter area, five interconnected laccoliths intrude the Famennian marls and shales, each one being shifted laterally above the
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lowest (Figs. 3B, 4). These tens of meter-thick magmatic bodies consist of coarse grained gabbros or syeno-diorites.

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The Tafilalt magmatic rocks were first identified by petrographic analysis under the optical microscope. In order to precise the geochemical characters of these rocks, they were the object of 98 analyses during the South Tafilalt mapping program cited above (Álvaro et al., 2014; Benharref et al., 2014a, 2014b, 2014c; Pouclet et al., 2017), and of 33 new analyses (Najih et al., 2018).

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## 3.2. U-Pb zircon dating (SHRIMP)

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For U-Pb zircon dating we selected the more leucocratic, syeno-dioritic sample collected in the core of the second Dboa laccolith (Fig. 3B). This sample (KS-LA2) is particularly rich in pinky K-feldspar, besides of plagioclase, biotite and kaersutite (see Appendix). Zircon was separated using panning, first in water and then in ethanol. After eliminating the magnetic fraction from the concentrates with a neodymium magnet, zircon grains were handpicked under a binocular microscope. About 130 zircon grains from sample L-2 plus several grains of standards were mounted on a 3.5 cm diameter epoxy SHRIMP megamount (Ickert et al., 2008). Once polished, zircon grains were studied by optical (reflected and transmitted light) and scanning electron microscopy (secondary electrons and cathodoluminescence), coated with a 13 nm thick gold layer, and analyzed for U-Th-Pb using a SHRIMP IIe/mc ion microprobe at the IBERSIMS laboratory of the CIC- University of Granada, Spain. The SHRIMP U-Th-Pb analytical method roughly followed that described by Williams and Claesson, (1987), and is described in detail at www.ugr.es/ibersims. Each selected spot is rastered with the primary beam for 120 s prior to the analysis, and then analysed by 6 scans, following the isotope peak sequence ¹⁹⁶Zr₂O, ²⁰⁴Pb, ²⁰⁴.¹background, ²⁰⁶Pb, ²⁰⁷Pb, ²⁰⁸Pb, ²³⁸U, ²⁴⁸ThO, and ²⁵⁴UO. Every peak of every scan is measured sequentially 10 times with the

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following total counting times per scan: 2 s for mass 196; 5 s for masses 238, 248, and 254; 15 s for masses 204, 206, and 208; and 20 s for mass 207. Uranium concentration was calibrated using the SL13 reference zircon (U: 238 ppm; Claoué-Long et al., 1995). U/Pb ratios were calibrated using the TEMORA-II reference zircon (417 Ma; Black et al., 2004) which was measured every 4 unknowns. All calibration procedures are performed on the standards included on the same mount. Mass calibration is done on the REG zircon (ca. 2.5 Ga, very high U, Th and common lead content). Data reduction was done with the SHRIMPTOOLS software (downloadable from www.ugr.es/~fbea) using the STATA™ programming language.

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### 3.3. $^{40}$Ar-$^{39}$Ar analysis

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Samples KSLA-01 (biotite gabbro), KSLA-02 (biotite syeno-diorite) and KSTA69 (biotite gabbro; see Appendix) were crushed and 200-500 µm size fraction was cleaned in ultrasonic bath. Five grains of biotite for each sample were carefully handpicked under a binocular microscope to select only grains without evidence of alteration or inclusions. Selected grains were packaged in aluminum foils and were irradiated for 90 h in the McMaster Nuclear Reactor together with Fish Canyon sanidine grains as flux monitor (28.030 ± 0.056 Ma, Jourdan and Renne, 2007). The argon isotopic interferences on K and Ca were determined by the irradiation of KF and CaF₂ pure salts from which the following correction factors were obtained: ($^{40}$Ar/$^{39}$Ar)K = 2.97 x 10⁻² ± 10⁻³ at 1σ, ($^{38}$Ar/$^{39}$Ar)K = 1,24 x 10⁻² ± 5 x 10⁻⁴ at 1σ, ($^{39}$Ar/$^{37}$Ar)Ca = 7.27 x 10⁻⁴ ± 4 x 10⁻⁵ at 1σ, and ($^{36}$Ar/$^{37}$Ar)Ca = 2.82 x 10⁻⁴ ± 3 x 10⁻⁵ at 1σ.

$^{40}$Ar/$^{39}$Ar step heating analyses were performed at Geoazur Nice (France) using a continuous 100 W PhotonMachine CO₂ (IR) laser used at 5-15% during 60 s. Argon isotopes were measured in static mode using an ARGUS VI mass spectrometer from Thermo-Fischer.

Measurements were carried out in multi-collection mode using four Faraday cups equipped with 10¹² ohm (masses 40, 39, 38 and 37) and one low-background compact discrete dynode
ion counter to measure mass 36. Collector gain calibration is performed by the computer-controlled application of predetermined voltages to each collector. Mass discrimination for the mass spectrometer was monitored by regularly analyzing air pipette volumes. The raw data were processed using the ArArCALC software (Koppers, 2002), and ages were calculated using the decay constants given by Steiger and Jäger (1977). Blanks were monitored after every three sample analyses. All parameters and relative abundance values are provided in compiled files in online supplementary data and have been corrected for blanks, mass discrimination, and radioactive decay. Atmospheric $^{40}$Ar was estimated using a value of the initial $^{40}$Ar/$^{36}$Ar of 298.56 (Lee et al., 2006). Our criteria for the determination of a plateau are as follows: a plateau must include at least 70% of $^{39}$Ar released, over a minimum of three consecutive steps agreeing at 95% confidence level. Plateau ages are given at the 2σ error level, and the plateau age uncertainties include analytical and J-value errors. All the errors on the inverse isochron, total fusion ages, and initial $^{40}$Ar/$^{36}$Ar ratios are quoted at the 2σ error.

## 4. Petrographic and geochemical outlines of the South Tafilalt magmatism

### 4.1 Petrography

Petrographic analysis indicates broadly constant lamprophyric facies for the dykes, most of the sills and the laccoliths margins, and gabbro and syeno-diorite facies for the inner parts of the laccoliths. The dykes are composed mainly of plagioclase and kaersutite phenocrysts in a doleritic groundmass made of plagioclase, alkali feldspar, biotite, pyroxene, and brown hornblende (Figs. 5A). Sills of olivine dolerites are hosted only in the lowest, Silurian to Emsian beds of the pile. They show intersertal porphyritic texture with olivine and clinopyroxene phenocrysts and scarce kaersutite microcrystals at their core. The plagioclase crystals (> 1.25 mm) are transformed into epidote and white mica ("sericite"), whereas olivine and clinopyroxene grains are completely altered into epidote and chlorite. Other sills
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consisting of kaersutite lamprophyres with olivine percentage less than 2% are found at all levels of the Paleozoic pile. Olivine and clinopyroxene minerals are partially or totally transformed into calcite, epidote and chlorite. Kaersutite (<1 mm) crystals are the most abundant ferromagnesian minerals. They are usually altered into chlorite and actinolite.

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The laccoliths show fine-grained border zones and medium to coarse-grained cores. The fine-grained borders exhibit lamprophyric facies, whereas microgranular to coarse grained gabbroic and syeno-dioritic textures are seen at the center of the laccoliths. The gabbroic rocks are composed of olivine, pyroxene, kaersutite, biotite, alkali feldspar, hornblende and plagioclase (Figs. 5B-D, F). The secondary minerals are "sericite", calcite and opaque minerals. The syeno-diorite rocks (Fig. 5E) are composed of alkali feldspar, biotite, kaersutite, plagioclase and altered pyroxene crystals. The secondary minerals are calcite, chlorite, opaque minerals, and elongated crystals of apatite.

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## 4.2. Geochemistry

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Detailed geochemical works on Tafilalt lamprophyric dolerites published recently (Pouclet et al., 2017; Najih et al., 2018) provide a good indication of the chemical composition of these rocks. Globally, compositions are basanitic (except for some differentiated rocks) and fall within the alkali lamprophyres field drawn by Rock (1987) on the TAS diagram (Fig. 6). The rocks are sodic ($K_2O/Na_2O < 1$) and undersaturated (olivine-nepheline normative), with the exception of a few differentiated rocks sampled in the core of the Dboa laccolith. The high HFSE, LILE and REE concentrations along with the relatively high Nb/Y ratios (2-5) and La/Sm ratios (5-13) are indicative of the alkaline character of these rocks (Kay and Gast, 1973; Pearce and Cann, 1973). The chondrite-normalized REE patterns of the Tafilalt lamprophyric dolerites (Fig. 7) display strong light REE (LREE) enrichment relative to heavy REE (HREE) and high REE fractionation with $(La/Yb)_N$ ratios varying from 14 to 64. The

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patterns have no Eu anomalies. Plots of elements normalized to primitive mantle (Fig. 8)
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show enrichment in large ion lithophile elements (LILE) relative to HREE, strong Ba
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enrichment, and a small negative Sr and Ti anomalies (except for olivine dolerites which
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display a weak positive Sr anomaly). The patterns do not show negative anomalies in Ta-Nb
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relative to the LILE. The chondrite- and mantle-normalized patterns of the Tafilalt
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lamprophyric dolerites are similar to those of alkaline lamprophyres (Rock, 1991) (Fig. 7, 8)
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and differ from those of calc-alkaline lamprophyres which typically display a negative Ta-Nb
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anomaly.
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5. Age of the Tafilalt magmatism

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5.1. *U-Pb data*

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For U-Th-Pb SHRIMP analyses we selected 130 crystals from sample L-2 (part of the KSLA-02 sample). Zircons are small, usually between <100 to 150 µm, translucent or opaque yellow to brown in color, and stubby bi-pyramidal A-type crystals with no prismatic faces, typical of being crystallized in alkaline magmas (Fig. 9). They are usually low cathodoluminescent or completely black, with no internal structure or with a slight patchy zonation. Only three among them show truncated prismatic forms with well-defined oscillatory zoning, two of them being inherited, older than crystallization zircon grains. Some other crystals are partially influenced by fluids and appear full of small irregular inclusions and fractures.

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All zircons except those three with oscillatory zonation are extremely rich in Th, with
contents ranging from >1000 to about 7000 ppm (average = 2930 ppm). U content is however
normal, ranging from 216 to 750 ppm (average =411 ppm). As a consequence Th/U ratios are
very high, although very variable, ranging from 3.3 to 18.8 (average=8.2). There is no
correlation between U and Th contents.
271 Thirty-eight analyses have been performed in 38 grains. Four of them had a discordance
272 >10% and were rejected. The other 34 analyses are concordant or nearly concordant, with a
273 small or negligible content of common-lead. All together plot in concordia in a Whetherill
274 diagram, 32 of them group in a cluster and two scatter along the concordia at 617 and 701 Ma
275 (Fig. 10A) The cluster is composed of 32 analyses yielding an average (weighted mean)
276 $^{206}$Pb/$^{238}$U age of 255 ± 3 Ma (MSWD=2.0), which is considered the age of crystallization of
277 the rock (Fig. 10B). The scarcity of inherited zircons among the grains here analyzed is
278 noticeable, and contrasts with their almost exclusive occurrence in the more mafic, gabbro-
279 dioritic samples analyzed by Pouclet et al. (2017) from the same area.

280 5.2. $^{40}$Ar-$^{39}$Ar data

281 Our measurements were performed on three samples from distinct units of the J. Dboa
282 laccolith system (Figs. 3B, 4). The results are presented below and summarized in Fig. 11.

283 KSLA-01 biotite displays well-defined plateau age at 264.2 ± 2.7 Ma (MSWD=1.89),
284 concordant with the inverse isochron age of 264.9 ± 2.9 Ma, representing more than 97% of
285 $^{39}$Ar released (Fig. 11). The isochron age yields atmospheric initial $^{40}$Ar/$^{36}$Ar ratio. Therefore,
286 this plateau age is considered as laccolith emplacement age during Capitanian-Wuchiapingian
287 period. KSLA-02 sample yields a less constrained plateau age at 259 ± 6.3 Ma
288 (MSWD=8.46); it shows typical loss of radiogenic $^{40}$Ar in biotite in the first heating steps,
289 and displays disturbed spectra suggesting recoil effect in the biotite during neutron irradiation.
290 The inverse isochron age is concordant with the calculated plateau age, with nevertheless an
291 erroneous MSWD value. KSLA-02 plateau age of 259.0 ± 6.3 Ma is considered valid and
292 represents a minimum age for the biotite crystallization. KSTA-69 biotite displays a plateau
293 age at 262.6 ± 4.5 Ma (MSWD=7.59) concordant with the inverse isochron age, which shows
294 however an initial $^{40}$Ar/$^{36}$Ar intercept lower than the atmospheric ratio of 298.56. The plateau
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age of KSTA-69 sample is concordant with KSLA-01 and KSLA-02 suggesting that these
rocks belong to the same late Permian magmatism.

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# 6. Discussion

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## 6.1. Age of the Tafilalt magmatism

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Table 1 summarizes the new isotopic ages presented in sections 5.1 and 5.2 above. Referring to the International Chronostratigraphic Chart 2018/8 (Fig. 12), all samples plot in the Middle-Upper Permian (Guadalupian-Lopingian), referable unformally to the late Permian. The SHRIMP age of sample KS-LA1, 255 ± 3 Ma, is consistent with its plateau age at 259.0 ± 6.3 Ma in the limits of the error bar. ⁴⁰Ar/³⁹Ar ages of the 3 analysed samples are indistinguishable at the 2 sigma level. Agreeing with the essential homogeneity of the Tafilalt magmatism as described by Pouclet et al. (2017) and Najih et al. (2018), we may conclude that the South Tafilalt intrusions emplaced during the late Permian. Therefore, the Tafilalt magmatism occurred about 30 My after the end of the mild Variscan folding of the area, which can be ascribed to the Late Pennsylvanian or Early Permian at a maximum (Fabre, 2005; Michard et al., 2010; Baidder et al., 2016). Conversely, the South Tafilalt magmatism emplaced about 60 My before the CAMP, which is well-represented west and east of the Tafilalt (Zemoul area in northern Tindouf Basin; Bechar Basin and Ksour Mountains; Chabou et al., 2007, 2010; Meddah et al., 2007, 2017). Our U-Pb and Ar-Ar results demonstrate for the first time the occurrence of a late Permian alkaline magmatism at the edge of the Sahara Platform where only CAMP intrusions were known.

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## 6.2. Nature and source of the Tafilalt magmatism

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In their recent review of the South Tafilalt magmatic rocks, Pouclet et al. (2017) describe these rocks as a system of basaltic dolerites, lamprophyric dolerites, analcite-bearing
camptonites, gabbros and diorites, sharing a sodic alkaline affinity. Additionally, Pouclet et al. (2017) infers that the intrusive bodies were linked to sub-water volcanic activities with spilitic lava flows and pyroclatites. In the absence of isotopic data these Authors conclude that the Tafilalt alkaline rocks emplaced during two main pulses in the Famennian-Tournaisian and the Early Visean. However, the isotopic dates presented here (Table 1, Fig. 12) discard the interpretation of the brecciated spilitic dolerites as lava flows or pyroclastic beds and oblige to consider them as sills emplaced in under-compacted, water-rich sediments. The probable volcanic manifestations of the Tafilalt magmatism occurred at the surface of the Variscan belt during the late Permian and were subsequently eroded, most likely during the Late Triassic-Middle Jurassic (Gouiza et al., 2016).

The geochemical characteristics of the Tafilalt lamprophyric dolerites indicate that these rocks are cogenetic (i.e., derived from a common mantle source) and were produced by low degrees of partial melting from an enriched deep mantle source within the garnet stability field, (Pouclet et al., 2017; Najih et al., 2018). The Tafilalt magmatism does not show any subduction signature.

### 6.3. *Correlations and geodynamic significance*

In the Variscan (Hercynian) belt of Western Europe, two dominantly felsic magmatic events have been known for a long time during the Late Carboniferous-Early Permian, with calc-alkaline then alkaline affinities, respectively. They record crustal thickening during the collision climax (i.e., the Pangea birth) from 325-310 Ma, and the subsequent, late orogenic extension from 310-270 Ma (Timmerman, 2004; Ballouard et al., 2018; and references therein). This evolution predates by ~70 Ma the well-known Central Atlantic Magmatic Province (CAMP) tholeiitic magmatism that heralds the Triassic rifting of Pangea (Fig. 13; Marzoli et al., 1999; Frizon de Lamotte et al., 2015). However, this time span was punctuated
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by distinct magmatic events recognized locally in central Pangea in both sides of the future Atlantic Ocean (i.e., Morocco, south-western Europe and eastern North America). These events have been known for years (Table 2).

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In the southwestern Meseta of Morocco, microdiorites emplaced at the end of the Permian or in the early Triassic times are known in the Rehamna (Hoepffner, 1982; Baudin et al., 2003; Razin et al., 2003) and Jebilet massifs (Huvelin, 1971; Bouloton and Gasquet, 1995; Youbi et al., 2001; Bouloton et al., 2019). The magmatic rocks, also referred to as quartz-monzodiorites, show transitional characteristics from calc-alkaline to alkaline affinities (Youbi et al., 2001; Bouloton et al., 2019). The precise age of this magmatism is not well constrained. In the eastern Rehamna massif, several quartz-monzodiorite dykes crosscut the Autunian red beds (Razin et al., 2003). In the Jebilet massif, a $^{40}$K-$^{40}$Ar age of 255 Ma is reported by Youbi et al. (2001) and interpreted as a minimum estimate for the timing of emplacement of this magmatism, whereas Dostal et al., (2005) obtained a SHRIMP U-Pb age of 235 Ma, interpreted as a maximum age. Finally, Bouloton et al. (2019) proposed an age of 240 ± 10 Ma on the basis of $^{40}$K-$^{40}$Ar dating on kaersutite phenocrysts. Elsewhere in Morocco, magmatism of presumed late Permian-Early/Middle Triassic age showing the same geochemical transitional affinities as the Jebilet and Rehamna microdiorites are described in the Khenifra Massif (Central Morocco) (Youbi et al., 1995), and in the southeast border of the Argana basin (Aït Chayeb et al., 1998).

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In the Iberian Peninsula, several camptonitic lamprophyre dykes of late Permian age are present in the Spanish Central system. The dykes gave $^{40}$Ar/$^{39}$Ar ages of 264.6 ± 1.6 Ma, 264.3 ± 0.9 Ma and 264.4 ± 0.9 Ma in the Sierra de Gredos (Scarrow et al., 2006) and 264 ± 3 Ma in the Sierra Guadarrama (Perini and Timmerman, 2008). Similar alkaline lamprophyres reported in the Subbetic zone (Southern Spain) gave $^{40}$Ar/$^{39}$Ar ages of 217 ± 2.5 Ma (Puga et al., 2010). Finally, in the western part of the Pyrenean Axial zone, K-Ar datings on kaersutite

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367 phenocrysts of four mafic dykes of alkaline affinity yielded ages varying from 266 ± 10 Ma to
368 271 ± 9 Ma (Debon and Zimmermann, 1993). This suggests that the post-Hercynian and pre-
369 CAMP alkaline magmatism of that part of the central Pangaea spread over a long period of
370 time from the late Permian to Upper Triassic.

371 On the other side of the Central Atlantic rift, Triassic magmatic rocks of alkaline affinity
372 predating the CAMP are widespread in New England and Atlantic Canada (McHone, 1992).
373 These include the lamprophyric sill of Malpeque Bay (Prince Edward Island, Canada) dated at
374 247 Ma (Greenough et al., 1988); a camptonite dyke located in Cambridge, Massachusetts
375 (USA), which gave a precise <sup>40</sup>Ar/<sup>39</sup>Ar age on kaersutite of 246 ± 4 Ma (Ross, 2010), and
376 lamprophyre dykes near Plymouth, Nova Scotia (Canada), which yielded <sup>40</sup>Ar/<sup>39</sup>Ar
377 hornblende-biotite ages ranging between 231 ± 3 Ma and 222 ± 3 Ma (Pe-Piper and
378 Reynolds, 2000).

379 The <sup>40</sup>Ar/<sup>39</sup>Ar plateau ages and U-Pb zircon age of the Tafilalt lamprophyric dolerites (this
380 study) are identical with the precise ages obtained for lamprophyres of the Sierra de Gredos
381 and the Sierra Guadarrama areas (Spanish Central system) (Scarrow et al., 2006; Perini et al.,
382 2008) and close to the age obtained for the lamprophyric sill of Malpeque Bay (Canada) and
383 the camptonite dyke of Cambridge, Massachusett (Greenough et al., 1988; Ross, 2010).
384 Conversely, the Nova Scotia lamprophyres are younger, early Upper Triassic (Pe-Piper and
385 Reynolds, 2000). Due to the lack of precise age determination for the other late Permian to
386 Late/Middle Triassic magmatic events in Morocco, geochronological correlations between the
387 South Tafilalt lamprophyric dolerites and the Jebilet-Rehamna microdiorites or the SE Argana
388 mafic rocks are currently not possible.

389 Geochemically, the South Tafilalt dolerites resemble lamprophyres of the Sierra de
390 Gredos area (Spanish Central system) and the late Permian-Triassic alkaline rocks
(lamprophyres and dolerites) of New England (USA), Atlantic Canada, Subbetic zone (Spain) and southeastern Iberian Chain (Lago San Jose et al., 2000). By contrast, they differ significantly from the Jebilet microdiorites. Such similarities and differences are best highlighted in a discriminant diagram (Chabou et al., 2017a) based on ratios of incompatible elements (Zr/Ti vs. Zr/Nb). Those elements are deemed to be immobile during weathering, post-consolidation alteration and metamorphic processes (Winchester and Floyd, 1977; Pearce, 2014). In this diagram (Fig. 14), the Tafilalt lamprophyric dolerites correlate well with the late Permian-Norian alkaline magmatism (dolerites and lamprophyres) of Spain (Lago San José et al., 2000; Orejana et al., 2008; Puga et al., 2010) and Nova Scotia (Canada) (Greenough et al., 1988; Pe-Piper and Reynolds, 2000). This diagram shows that the Jebilet microdiorites do not share the chemical compositions displayed by the South Tafilalt mafic rocks. The geochemical fingerprinting of the Jebilet microdiorites assign them to the Variscan post-orogenic magmatism rather than to the pre-rifting stage of Central Atlantic (Youbi et al., 2001) or to any other magmatic event independent of the Variscan orogeny (Bouloton et al., 2019).

In the post-orogenic magmatic belts of the Variscan Bohemian massif, a well-dated (292 Ma) quartz-monzodiorite was reported in association with post-orogenic granites (Laurent et al., 2014). This monzodiorite share the same geochemical characteristics with the Jebilet microdiorites (Figs. 14, 15) indicating that both are probably formed in the same geodynamic context (i.e., post-orogenic context). The poorly constrained late Permian to Early Triassic ages assigned to the Jebilet microdiorites are probably younger than the time of igneous crystallization, due to the alteration of the rocks and the dating method used (K-Ar). Likewise, the Permian/Lower-Middle Triassic Argana mafic rocks and Pyrenean mafic dykes mentioned above do not share chemical compositions displayed by the Tafilalt magmatism; they plot between the Tafilalt and CAMP fields in (Fig. 14).
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Recently, Mekkaoui et al., (2017) reported on two types of geochemical signatures in dolerites from the Daoura-Ougarta Range (Algeria), one similar to that of the Tafilalt alkaline rocks (Fig. 14) and the other one to that of the CAMP tholeiite. They would be therefore derived from two different mantle sources, through different melting conditions and possibly at two distinct periods.

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Several geodynamic models have been invoked to explain the emplacement of the late Permian-Triassic magmatism of the central Pangaea province. Some authors attributed this magmatism to the post-orogenic magmatic activity that marked the final stage of the evolution of the Variscan belt (Bea et al., 1999; Lago et al., 2004; Scarrow et al., 2006, 2011), while others have interpreted it as witnessing the pre-rifting stage of the Central Atlantic (Aït Chayeb et al., 1998; Youbi et al., 2001; Ross, 2010). More recently, Bouloton et al. (2019) have suggested that the microdioritic magmatism of the Jebilet Massif (Morocco) neither represent the ultimate stage of the Variscan orogeny, nor a precursor to the breakup of Pangea. They interpreted this magmatism to reflect an aborted rifting event, or a consequence of the Late Paleozoic right-lateral shear zone developed between Europe and North Africa. In the Pyrenees, Debon and Zimmermann (1993) considered the alkaline mafic dykes as marking the transition from the Variscan cycle to the Atlantic-Alpine one.

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The Tafilalt magmatism cannot be considered as post-orogenic since the Tafilalt area was far southward of the domain of crustal thickening due to the Variscan collision, i.e., the Meseta Orogen. Furthermore, geochemical signatures and petrographic suite of the Tafilalt magmatism do not correspond to those of post-orogenic rocks.

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The geodynamic interpretation proposed by Bouloton et al. (2019) is speculative in the absence of reliable dating of microdioritic magmatism of the Jebilet Massif. As has been pointed out, this magmatism is likely to be linked to the latest stage of the Variscan orogeny.

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Considering geochronological and geochemical data: (i) The geochemistry of the late Permian South Tafilalt magmatism is similar to the lamprophyres of the same age and other alkaline dolerites and lamprophyres of Triassic age emplaced in central Pangea in both sides of the Atlantic Ocean. This alkaline magmatism was therefore spread over a long period of time from the late Permian to Upper Triassic as mentioned above, and ended before the emplacement of the huge Central Atlantic Magmatic Province (CAMP) at the Triassic/Jurassic boundary and the subsequent opening of the Central Atlantic Ocean in the Early Jurassic (Labails et al., 2010). (ii) Kaersutite-bearing, alkaline, OIB-type geochemical characteristics of mafic rocks such as those of the South Tafilalt are often indicative of an intraplate volcanism formed at the onset of lithospheric extension associated with initial rifting (eg. Zeng et al., 2019), we propose that the Tafilalt magmatism reflects an early-rift magmatic activity that preceded the opening of the Central Atlantic Ocean. This proposal is consistent with the location of this alkaline magmatism in both sides of the Atlantic Ocean as evidenced by the Upper Triassic lamprophyres of the Spanish Subbetic zone (Puga et al., 2010), alkaline dolerites of the South-East Iberian Chain (Lago San José et al., 2000) and the alkaline dolerite dykes of the New England (McHone, 1992) and Atlantic Canada (Pe-Piper et al., 2000) that share the same geochemical characteristics as the Tafilalt mafic rocks. These early magmatic events occurred at about the same locations as the CAMP manifestations, and then can be regarded as early-rift manifestations, although they precede the Late Triassic rifting by ~60 My. We acknowledge that 60 My is a huge amount of time, which makes also correct to envision the peri-Atlantic late Permian magmatic manifestations as an aborted rifting event (cf. Bouloton et al., 2019). However, the apparent delay between the Tafilalt magmatism and the onset of rifting in the Central Atlantic area is reduced to ~10-15 Ma considering that syn-rift continental deposits accumulated as early as the Middle Triassic (247-237 Ma) in the Argana and High Atlas basins of Morocco (Medina et al., 2001; El Arabi
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et al., 2006; Voigt et al., 2011) and the Fundy Basin of Canada (Fig. 13; Leleu and Hartley, 2010). At the early-rift stage, reactivated zones of weakness appear to control the emplacement of the late Permian alkaline magmatism. In the specific case of the Tafilalt basin we may observe that the late Permian magmatism developed on a mosaic of faulted basement blocks located at the junction between two branches of the Variscan Orogen, i.e., the ENE-trending Anti-Atlas belt and the NW-trending Ougarta belt (Baidder et al., 2016).

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Small-volume alkaline igneous rocks are temporally and spatially associated with many, if not most Large Igneous Provinces (LIPs) on Earth (Ernst, 2014). Kent et al. (1992) reported evidence for alkali volcanism from five major LIPs (Siberian Trapps, Karoo, Parana Basin, Rajmahal-Kerguelen and Deccan provinces), preceding (and in some case, post-dating) the onset of the main tholeiitic magmatism by several million years (< 10 My for Deccan province to 32 My for Rajmahal). In some LIPs, alkaline magmatism emplaced in the main phase of flood volcanism (e.g., Etendeka Province, Schmitt et al., 2000). In older LIPs, there is evidence of punctual alkaline magmatic events that preceed the main rifting magmatism by a long amount of time. In the Keweenawan LIP, the age of alkaline dolerite and lamprophyre dykes (1147 -1140 Ma) is some 35 My older than the age of the main rift magmatism (1115-1106 Ma) (Queen et al., 1996; Heaman et al., 2007). In the Central Iapetus Magmatic Province (CIMP), the oldest alkaline igneous rocks were dated at 615 Ma, 60 My before the onset of major ocean crust formation at ca. 555 Ma (eg. Ernst and Bell, 2010, and references therein). Finally, the onset of alkaline magmatism in the youngest intra-continental rifting system on Earth (East-African rift) began at ca. 40 Ma (Bailey, 1992) and continue to be active today (Ernst and Bell, 2010).

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In the LIPs whose genesis is attributed to a mantle plume, the alkaline volcanism that precedes the onset of the main pulse of tholeiitic basalts is attributed to the earliest melts associated with the initial impact of the mantle plume starting-head against the lithosphere

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(e.g., Gibson et al., 2006). In this case, alkaline igneous activity predates the main tholeiitic volcanism within a short time-span (< 10 Ma). In the case of the CAMP, the global warming mantle better explain the large delay between the Tafilalt magmatism (~260 Ma) and the rifting-related magmatic climax (~200Ma).

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The CAMP is often cited as an example of LIP whose origin is not linked to a mantle plume (McHone, 2000), although recent studies invoke a passive role of a plume in the genesis of this province (Marzoli et al., 2018; Tegner et al., 2019). Some characteristics of the CAMP are consistent with a mantle global warming model (Coltice et al., 2007; 2009; Hole, 2015). In this model, the aggregation of Pangea causes changes to the temperature and flow of underlying mantle through thermal insulation and the enlargement of flow wavelength, triggering a large scale melting in the asthenosphere and the lithosphere leading to emplacement of the CAMP. This model suggests subcontinental increase in temperature as large as 100 °C for time scales of ~100 My (Coltice et al., 2009), and predicts, amongst others, that heating occurs over an area comparable to the supercontinent and that melt extraction is controlled by tectonics (Coltice et al., 2009). In line with this model, it is conceivable that the mantle global warming associated with the Pangea supercontinent can account for the occurrence of punctual alkaline events over a long time-span before the onset of the CAMP. The heating of the subcontinental lithosphere and asthenosphere by 50°C at the time of the emplacement of the first alkaline magmas and the gradual increase of temperature due to mantle insulation can trigger low partial melting of metasomatic layer of the subcontinental lithospheric mantle and provide the thermal source of this magmatism. When the temperature of the subcontinental mantle increases by 100°C, a large scale melting in the asthenosphere and the lithosphere may occur leading to the emplacement of the brief, geographically large and high-volume CAMP episode (Fig. 16).

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This model can explain why these early magmatic events occurred at about the same
locations as the CAMP manifestations and were controlled by tectonic zones of weakness.
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Furthermore, it is also worth noting that some LIPs cited as an example of magmatic
provinces that developed following a mantle global warming processes show a long lived
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and of the CIMP linked to the amalgamation of the Pannotia supercontinent (Coltice et al.,
2007; Doblas et al., 2002).
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7. Conclusions

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The Late Devonian-Early Carboniferous South Tafilalt Basin (eastern Anti-Atlas belt, Morocco) represents the foreland belt of the Variscan Meseta Orogen of Morocco. The mildly folded Paleozoic series (from Ordovician to Early Carboniferous) were intruded by numerous mafic dykes, sills and laccoliths of lamprophyric dolerites and camptonites with minor gabbros and syeno-diorites. Coarse-grained gabbro and syeno-diorite occur in the inner part of the laccoliths. Geochemically, the Tafilalt magmatism shows sodic-alkaline affinity, being produced by low degrees of partial melting from a metasomatized deep mantle source. Biotite ⁴⁰Ar/³⁹Ar and zircon ²⁰⁶Pb/²³⁸U dating of gabbros and syeno-diorite provide the first robust geochronological evidence of the age of the South Tafilalt magmatism. Results of ⁴⁰Ar/³⁹Ar dating of biotite give plateau ages of 264.16 ± 2.72 Ma, 262.61 ± 4.5 Ma and 259.04 ± 6.31 Ma), whereas ²⁰⁶Pb/²³⁸U dating of zircon yields a mean age of 255 ± 3 Ma). These ages coincide within the dating error, and indicate that this magmatism occurred during the late Permian. The geochemistry of the Tafilalt magmatism is similar to the lamprophyres of the same age and other alkaline dolerites and lamprophyres of Triassic age emplaced in central Pangea in both sides of the Atlantic Ocean (i.e., Morocco, south-western Europe and eastern North America). These rocks, including the Tafilalt ones, would reflect an early-rift magmatic activity likely related to the global mantle warming underneath the Pangea supercontinent.
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This precursory magmatism was controlled locally by the extensional/transtensional reactivation of zones of crustal weakness within and around the future CAMP area.

**Acknowledgments.** We would like to thank the Editor and the two anonymous Reviewers for their comments and feedback, which helped to improve the initial version of the manuscript. This research did not receive any specific grant from funding agencies in the public, commercial, or not-for-profit sectors. AM acknowledges logistic support from the Geoscience Laboratory, Faculty of Sciences Ain Chock, Hassan II University of Casablanca, Morocco. This is the IBERSIMS publication N° 62.

## Data Availability

Datasets related to this article can be found at https://earthref.org/ERDA/2407/?database_name=erda&search_start=main

file name: tafilalt.magmatism.analytical.data.zip

## Appendix: Petrography and Geochemistry of the dated samples (Najih et al., 2018)

<table><thead><tr><th>Rock / Sample</th><th>KSLA1</th><th>KSLA2</th><th>KSTA69</th></tr></thead><tbody><tr><td>Type of rock</td><td>Biotite gabbro</td><td>Syeno-diorite</td><td>Biotite gabbro</td></tr><tr><td>Texture</td><td>Fine grained, microgranular to granular</td><td>Microgranular to granular</td><td>Coarse microgranular to granular</td></tr><tr><td>Primary minerals</td><td>Olivine, plagioclase, K-feldspar, pyroxene, brown hornblende, biotite</td><td>Perthitic K-feldspar, plagioclase, brown hornblende, pyroxene, biotite, apatite.</td><td>As KS-LA1, except a greater content in K-feldspar</td></tr><tr><td>Secondary minerals</td><td>Mg-chlorite, actinolite, calcite, epidote, Fe-oxides</td><td>Mg-chlorite, calcite, epidote, Fe-oxides</td><td>As KS-LA1</td></tr><tr><td>Remarks</td><td></td><td>Zircon grains, sometimes in biotite</td><td></td></tr></tbody></table>
<table><thead><tr><th>Geochronological ID</th><th>KSLA1</th><th>KSLA2</th><th>KSTA69</th></tr></thead><tbody><tr><td>Geochemical ID</td><td>TA55</td><td>TA53</td><td>TA69</td></tr><tr><td>Setting</td><td>Laccolith 1</td><td>Laccolith 2</td><td>Laccolith 4</td></tr><tr><td>Hostrock<br>(wt%)</td><td>Early Famennian</td><td>Early Famennian</td><td>Middle Famennian</td></tr><tr><td>SiO<sub>2</sub></td><td>43,08</td><td>50,29</td><td>42,82</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>15,16</td><td>18,79</td><td>14,23</td></tr><tr><td>Fe<sub>2</sub>O<sub>3</sub></td><td>10,43</td><td>6,34</td><td>10,94</td></tr><tr><td>MgO</td><td>6,8</td><td>4,03</td><td>6,84</td></tr><tr><td>CaO</td><td>9,14</td><td>4,89</td><td>8,25</td></tr><tr><td>Na<sub>2</sub>O</td><td>4,04</td><td>3,65</td><td>4,13</td></tr><tr><td>K<sub>2</sub>O</td><td>1,61</td><td>4,57</td><td>1,1</td></tr><tr><td>TiO<sub>2</sub></td><td>2,62</td><td>1,89</td><td>2,4</td></tr><tr><td>P<sub>2</sub>O<sub>5</sub></td><td>2,12</td><td>0,59</td><td>2,44</td></tr><tr><td>MnO</td><td>0,14</td><td>0,13</td><td>0,26</td></tr><tr><td>Cr<sub>2</sub>O<sub>3</sub></td><td>0,01</td><td></td><td>0,013</td></tr><tr><td>LOI</td><td>4,1</td><td>4</td><td>5,8</td></tr><tr><td>Sum<br>(ppm)</td><td>99,26</td><td>99,21</td><td>99,22</td></tr><tr><td>Ni</td><td>63</td><td></td><td>69</td></tr><tr><td>Sc</td><td>14</td><td>6</td><td>13</td></tr><tr><td>Ba</td><td>2837</td><td>1933</td><td>3389</td></tr><tr><td>Be</td><td>3</td><td>4</td><td>3</td></tr><tr><td>Co</td><td>35,6</td><td>12,8</td><td>36,7</td></tr><tr><td>Cs</td><td>0,4</td><td>0,3</td><td>0,8</td></tr><tr><td>Ga</td><td>19,3</td><td>20,3</td><td>18,5</td></tr><tr><td>Hf</td><td>6,3</td><td>8</td><td>6,5</td></tr><tr><td>Nb</td><td>97,2</td><td>113,3</td><td>83,7</td></tr><tr><td>Rb</td><td>18,4</td><td>43,8</td><td>14,7</td></tr><tr><td>Sn</td><td>2</td><td>2</td><td>1</td></tr><tr><td>Sr</td><td>1932,9</td><td>3188,1</td><td>1626</td></tr><tr><td>Ta</td><td>4,7</td><td>5,9</td><td>4</td></tr><tr><td>Th</td><td>8,4</td><td>9,8</td><td>8,4</td></tr><tr><td>U</td><td>2,4</td><td>3</td><td>2,4</td></tr><tr><td>V</td><td>165</td><td>92</td><td>157</td></tr><tr><td>W</td><td>21</td><td>62,5</td><td>27,8</td></tr><tr><td>Zr</td><td>315,7</td><td>392,5</td><td>313,9</td></tr><tr><td>Y</td><td>23,6</td><td>20,7</td><td>26</td></tr><tr><td>La</td><td>127,2</td><td>116,6</td><td>125,8</td></tr><tr><td>Ce</td><td>239,3</td><td>214,4</td><td>247,4</td></tr><tr><td>Pr</td><td>24,68</td><td>21,02</td><td>25,92</td></tr><tr><td>Nd</td><td>90,5</td><td>70,3</td><td>94,5</td></tr><tr><td>Sm</td><td>13,21</td><td>10,55</td><td>14,45</td></tr><tr><td>Eu</td><td>3,84</td><td>2,98</td><td>3,91</td></tr><tr><td>Gd</td><td>10,46</td><td>7,85</td><td>10,12</td></tr><tr><td>Tb</td><td>1,22</td><td>0,95</td><td>1,27</td></tr><tr><td>Dy</td><td>5,79</td><td>4,62</td><td>6,3</td></tr><tr><td>Ho</td><td>0,91</td><td>0,76</td><td>0,92</td></tr><tr><td>Er</td><td>2,15</td><td>1,85</td><td>2,38</td></tr><tr><td>Tm</td><td>0,29</td><td>0,28</td><td>0,3</td></tr><tr><td>Yb</td><td>1,82</td><td>1,76</td><td>1,77</td></tr><tr><td>Lu</td><td>0,25</td><td>0,24</td><td>0,24</td></tr></tbody></table>
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843 Rock, N.M.S., 1987. The nature and origin of lamprophyres: an overview. Geol. Soc. London, Spec. Publ. 30, 191-226.
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864

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867
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873 Sun, S., McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. Geol. Soc. London, Spec. Publ. 42, 313–345.

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Petrogenesis and $^{40}$Ar/$^{39}$Ar geochronology of the Brandberg complex, Namibia: evidence
for a major mantle contribution in metaluminous and peralkaline granites. J. Petrol., 41,
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886 Verati, C., Rapaille, C., Féraud, G., Marzoli, A., Bertrand, H., Youbi, N., 2007. $^{40}$Ar/$^{39}$Ar ages
and duration of the Central Atlantic Magmatic Province volcanism in Morocco and
Portugal and its relation to the Triassic–Jurassic boundary. Palaeogeogr. Palaeoclimatol.
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890 Wendt, J., 1985. Disintegration of the continental margin of northwestern Gondwana: Late
891 Devonian of the eastern Anti-Atlas (Morocco). Geology 13, 815.

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893 Caledonian metamorphism of high grade paragneisses from the Seve Nappes,

894 Scandinavian Caledonides. Contrib. Mineral. Petrol. 97, 205-217.

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896 and their differentiation products using immobile elements. Chem. Geol. 20, 325-343.

897 Youbi, N., Bellon, H., Marzin, A., Piqué, A., Cotten, J., Cabanis, B., 2001. Du cycle

898 orogénique hercynien au pré-rifting de l'Atlantique central au Maroc occidental : les

899 microdiorites des Jebilet sont-elles des marqueurs magmatiques de ce passage ? C. R.

900 Acad. Sci. Paris - Ser. IIA - Earth Planet. Sci. 333, 295-302.

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902 massif permien de Khénifra (Sud-Est du Maroc Central). Geodinamica Acta 8, 158-172.

903 Youbi, N., Martins, L.T., Munhá, J.M., Ibouh, H., Madeira, J., Aït Chayeb, E.H., El Boukhari,

904 A., 2003. The Late Triassic-Early Jurassic volcanism of Morocco and Portugal in the

905 framework of the Central Atlantic Magmatic Province: An overview. In: W.E. Hames,

906 J.G. McHone, P.R. Renne, C. Ruppel, (Eds). The Central Atlantic Magmatic Province:

907 Insights from Fragments of Pangea. Am. Geophys. Union, Geophys. Mono. 136, 179-

908 207.

909 Zeng, Y.C., Xu, J.F., Ducea, M.N., Chen, J.L., Huang, F., Zhang, L., 2019. Initial rifting of

910 the Lhasa Terrane from Gondwana: insights from the Permian (~262 Ma) amphibole-rich

911 lithospheric mantle-derived Yawa basanitic intrusions in southern Tibet. J. Geophys.

912 Res., Solid Earth. doi.org/10.1029/2018JB016281

913 Zouhri, S., Kchikach, A., Saddiqi, O., Haïmer, F.Z. El, Baidder, L., Michard, A., 2008. The
914 Cretaceous-Tertiary Plateaus, in: Michard, A., Saddiqi, O., Chalouan, A., Frizon de

915 Lamotte, D. (Eds.), Continental evolution: the geology of Morocco, Lect. Notes Earth

916 Sci. 116,331–358.

917

918

919
920 **Figure captions**
921

922 Fig. 1: A: Location of the Tafilalt Basin at the northern border of the Saharan platform, in the
923 gently folded foreland of the Variscan Meseta belt (after Michard et al., 2010). B: Distribution
924 of the alkaline magmatism in the South Tafilalt Basin.

925 Fig. 2: Paleozoic stratigraphy of the South Tafilalt region, after Álvaro et al. (2014), and
926 location of the gabbroic sills and laccoliths in the lithological column, after the published
927 geological maps and this work.

928 Fig. 3: Geological maps of the South Tafilalt region, showing the location of the Tafilalt
dykes, sills and laccoliths (A) and the detail of the Jbel Dboa laccolith system (B), after the
929 Geological map of Morocco, scale 1:50,000, Al Atrous, Marzouga, Mfis and Tawz sheets,
930 and this work.
931

932 Fig. 4: Field view of the southern part of the Jbel Dboa laccolith system viewed from the
933 Upper Famennian sandstones (d7c) in the southeast. The sampling sites of samples KS-
934 LA1/2 are shown.

935 Fig. 5: Microphotographs of Tafilalt intrusive rocks. A: Kaersutite lamprophyre dyke; Znaigui
936 laccolitic intrusion. - B: Gabbro, Jbel Dboa laccolith 1. C-F: Microphotographs of the
937 analyzed samples showing the fresh biotite crystals used for $^{40}$Ar/$^{39}$Ar dating. C and D: Thin
938 section of sample KS-LA1 sample, plane polarized light, not analyzed and analyzed,
939 respectively. E and F: thin sections of samples KS-LA2 and KSTA-69, respectively. Act:
940 actinolite; Ap: apatite; Bi: biotite; Chl: chlorite; Kaer: kaersutite; Hb: brown hornblende; Ol:
941 olivine; Pl: plagioclase; Pyr: pyroxene. Opaque grains: Fe-Ti oxide.
943

Fig. 6. Total alkali-silica classification diagram (LeBas et al., 1986) for Tafilalt lamprophyric dolerites (grey area) (Pouclet et al., 2017; Najih et al., 2018). The studied samples are reported (red triangles), as well as the field of alkaline lamprophyres (AL, after Rock, 1987). Boundary line between alkaline and sub-alkaline series after Irvine and Baragar (1971).

944

945

946

947

Fig. 7. Chondrite-normalized (McDonough and Sun, 1995) REE contents of the studied samples (red triangles) and the South Tafilalt lamprophyric dolerites (grey field). Data from Pouclet et al. (2017) and Najih et al. (2018).

948

949

950

Fig. 8. Primitive mantle-normalized (Sun and McDonough, 1989) trace element contents of the studied samples (red triangles) and of the South Tafilalt lamprophyric dolerites (grey field). Data from Pouclet et al. (2017) and Najih et al. (2018).

951

952

953

Fig. 9: CL images of zircons from sample L-2. All crystals are low cathodoluminescent, with a faint zonation (grains a-i) or with no internal structure (grains j, k). Three zircons show a well-defined oscillatory zonation (grain m, restitic). Some other zircons are very dark or black and appear plenty of inclusions (grain l) and could not be analyzed.

954

955

956

957

Fig. 10: Whererill concordias for analyzed zircons from sample L-2. A: plot of 34 analyses including two older than crystallization inherited zircons. B: cluster of the remaining 32 analyses yielding an average $^{206}$Pb/$^{238}$U age of 255 ± 3 Ma.

958

959

960

Fig. 11: $^{40}$Ar/$^{39}$Ar age spectra as a function of $^{39}$Ar released by biotite from two gabbros (KS-LA-01 and KS-TA-69) and syeno-diorite (KS-LA-02). The error boxes of each step are at the 2σ level. Arrows show the plateau age fraction. Error of the plateau ages is given at the 2σ level. Detailed analytical data are reported in online Supplementary Data.

961

962

963
964

Fig. 12: Plot of the new isotopic ages of the Tafilalt magmatism on the International
Chronostratigraphic Chart of the Middle-Upper Permian series. Zr: SHRIMP age of zircon;
Ar: $^{40}$Ar/$^{39}$Ar ages of biotite.

965

966

967

Fig. 13: Distribution of the alkaline magmatic intrusions of post-Autunian, pre-Rhetian age (red stars) in the Central Atlantic area, within the CAMP domain. See the corresponding ages in Tables 1 (Tafilalt) and 2 (other occurrences). Background map after Knight et al., 2004, modified. C: Cambridge (Massachusetts); Kh: Khenifra; KT: Kahel Tabelbala; MP: Malpeque Bay; P: Plymouth (Nova Scotia); Py: Pyrenees; SB: Subbetic Zone; T: Tafilalt. The Jebilet star includes the Rehamna occurrences too (SW Moroccan Meseta). The Avila sector of the Iberian Meseta includes the Sierra de Gredos and Sierra de Guadarrama.

968

969

970

971

972

973

974

Fig. 14. Nb/Y vs. Zr/Ti diagram showing compositional field of Tafilalt mafic rocks, Jebilet microdiorites, other late Permian/Triassic alkaline rocks (lamprophyres and dolerites) of the central Pangaea province and CAMP tholeiites from North West Africa 1: Tafilalt mafic rocks (Pouclet et al., 2017; Najih et al., 2018). 2: Jebilet microdiorites (Youbi et al., 2001; Bouloton et al., 2019). 3: Middle Permian alkaline mafic rocks of the Pyrenean Axial zone (Debon and Zimmermann, 1993). 4: lamprophyres of the Sierra de Gredos area (Spanish Central system) (Orejana et al., 2008). Lamprophyres of Atlantic Canada are also included in this field (Greenough et al., 1988; PePiper and Reynolds, 2000). 5: Lamprophyres of the Subbetic zone (Southern Spain) (Puga et al., 2010). 6: Triassic alkaline dolerites (SE Iberian Chain) (Lago San José et al., 2000). 7: Alkaline dolerites of the Ougarta Range (Southwestern Algeria) (Mekkaoui et al., 2017). 8: Permian Argana mafic rocks (Ait Chayeb et al., 1998). 9: CAMP tholeiites from North West Africa (Marzoli et al., 2004; Verati et al., 2005; Meddah et al., 2007; Chabou et al., 2010). 10: Quartz-monzodiorite of the Bohemian massif (Laurent et al., 2014).

975

976

977

978

979

980

981

982

983

984

985

986

987
Fig. 15. Primitive mantle-normalized (Sun and McDonough, 1989) trace element contents of the Tafilalt lamprophyric dolerites (data from Pouclet et al., 2017, and Najih et al., 2018), compared to the Jebilet microdiorite (data after Bouloton et al., 2019), and to the quartz-monzodiorite of the Bohemian massif data after Laurent et al., 2014).

Fig. 16: Magmatic events related to the successive stages of the Pangea rifting in the Central Atlantic area. Sketch of lithospheric profiles after Reston (2009) and Leleu et al. (2016). A: Early magmatic events. TAM: Tafilalt alkaline magmatism. B: CAMP magmatism. 1: Greenough et al. (1988), Pe-Piper and Reynolds (2000), Ross (2010). 2: Aït Chayeb et al. (1998). 3: This work. 4: Marzoli et al. (2019).

## Tables

<table><caption>Table 1: Summary of the new datings of the Jbel Dboa laccoliths. See samples location in Figs. 3, 4.</caption><thead><tr><th>Samples</th><th>Laccolith Numbers</th><th>Rock types</th><th>SHRIMP age (Ma)</th><th><sup>40</sup>Ar-<sup>39</sup>Ar biotite age (Ma)</th></tr></thead><tbody><tr><td>KS-LA1</td><td>1</td><td>Biotite gabbro</td><td>-</td><td>264.2±2.7</td></tr><tr><td>KS-LA2 = L-2</td><td>2</td><td>Kaersutite syeno-diorite</td><td>255.0±3.0</td><td>259.0±6.3</td></tr><tr><td>KS-TA69</td><td>4</td><td>Biotite gabbro</td><td>-</td><td>262.6±4.5</td></tr></tbody></table>

<table><caption>Table 2: Age of the alkaline dolerites reported in the literature from the CAMP domain (See Fig. 13 for location).</caption><thead><tr><th></th><th>Location and rock type</th><th>U/Pb age (Ma)</th><th>K/Ar age (Ma)</th><th>Ar/Ar age (Ma)</th><th>Reference</th></tr></thead><tbody><tr><td rowspan="3">Morocco</td><td rowspan="3">Meseta - Jebilet</td><td>-</td><td>255</td><td>-</td><td>Youbi et al. (2001)</td></tr><tr><td>235</td><td>-</td><td>-</td><td>Dostal et al. (2005)</td></tr><tr><td>-</td><td>240 ± 10</td><td>-</td><td>Bouloton et al. (2019)</td></tr></tbody></table>
<table><tbody><tr><td></td><td>Anti Atlas<br/>-<br/>Tafilalt</td><td>Lamprophyres<br/>dolerites and gabbros</td><td>-<br/>255 ±3</td><td>-</td><td>264.2 ± 2.7</td><td>This study</td></tr><tr><td></td><td></td><td></td><td>-</td><td>-</td><td>259 ± 6.3</td><td>This study</td></tr><tr><td></td><td></td><td></td><td>-</td><td>-</td><td>262.6 ± 4.5</td><td>This study</td></tr><tr><td></td><td></td><td></td><td>-</td><td>-</td><td>264.6 ± 1.6</td><td></td></tr><tr><td rowspan="3">Spain</td><td>Sierra de Gredos</td><td>Lamprophyres</td><td>-</td><td>-</td><td>264.3 ± 0.9</td><td rowspan="3">Scarrow et al. (2006)</td></tr><tr><td></td><td></td><td>-</td><td>-</td><td>264.4 ± 0.9</td></tr><tr><td>Sierra Guadarrama</td><td>Camptonites</td><td>-</td><td>-</td><td>264 ± 3</td></tr><tr><td></td><td>Subbetics</td><td>Lamprophyres</td><td>-</td><td>-</td><td>217 ± 2.5</td><td>Perini and<br/>Timmerman (2008)</td></tr><tr><td></td><td>Pyrenees</td><td>Alkaline dykes</td><td>-</td><td>266</td><td>266 ± 10<br/>-<br/>271 ± 9</td><td>Puga et al. (2010)<br/>Debon and<br/>Zimmermann (1993)</td></tr><tr><td rowspan="2">Canada</td><td>Malpeque Bay, Prince<br/>Edwards</td><td>Lamprophyres</td><td>-</td><td>247</td><td>-</td><td>Greenough et al.<br/>(1988)</td></tr><tr><td>Plymouth, Nova<br/>Scotia</td><td>Lamprophyres</td><td>-</td><td>-</td><td>222 ± 3<br/>-<br/>231 ± 3</td><td>Pe-Piper and<br/>Reynolds (2000)</td></tr><tr><td>USA</td><td>Cambridge,<br/>Massachussetts</td><td>Camptonites</td><td>-</td><td>-</td><td>246 ± 4</td><td>Ross (2000)</td></tr></tbody></table>

1005

1006

1007

1008
A

500 km

<figure><img src="image_3.png" /></figure>

<figure><img src="image_4.png" /></figure>

Pan-African belts

<figure><img src="image_5.png" /></figure>

Paleozoic (platform)

<figure><img src="image_6.png" /></figure>

Mesozoic to Recent

<figure><img src="image_7.png" /></figure>

Paleoproterozoic

<figure><img src="image_8.png" /></figure>

Neoproterozoic

<figure><img src="image_9.png" /></figure>

Variscan front

<figure><img src="image_10.png" /></figure>

Archean

<figure><img src="image_11.png" /></figure>

-Cambrian

Variscan belts

<figure><img src="image_12.png" /></figure>

Mauritanides front

B

4°20′′W

4°10′′W

4°0′′W

3°50′′W

<figure><img src="image_13.png" /></figure>

31°100′′N

31°00′′N

30°50′′N
Silurian
Devonian

Ordovician

Cambrian

Ediacaran

<figure><img src="image_14.png" /></figure>

Devonian

Carboniferous

<figure><img src="image_15.png" /></figure>

<figure><img src="image_16.png" /></figure>

Pelite, sandstones

<figure><img src="image_17.png" /></figure>

Olistostromes

<figure><img src="image_18.png" /></figure>

Pelites, turbidites

<figure><img src="image_19.png" /></figure>

Limestones

<figure><img src="image_20.png" /></figure>

Pelites, limestones

<figure><img src="image_21.png" /></figure>

Mudstones, silts
black shales (S1-4)

<figure><img src="image_22.png" /></figure>

Mudmounds, reefs

<figure><img src="image_23.png" /></figure>

Tilloids, diametites

<figure><img src="image_24.png" /></figure>

Iron oolith

<figure><img src="image_25.png" /></figure>

Quartzites

<figure><img src="image_26.png" /></figure>

Sandstones, arkoses
red silts

<figure><img src="image_27.png" /></figure>

Alkaline basalts

<figure><img src="image_28.png" /></figure>

Granites (Ougnat)

<figure><img src="image_29.png" /></figure>

Tafilit alkaline
intrusives

500 m
<figure><img src="image_30.png" /></figure>

<figure><img src="image_31.png" /></figure>

<figure><img src="image_32.png" /></figure>
<figure><img src="image_33.png" /></figure>

SW

NE
<figure><img src="image_34.png" /></figure>

<figure><img src="image_35.png" /></figure>

<figure><img src="image_36.png" /></figure>

<figure><img src="image_37.png" /></figure>

<figure><img src="image_38.png" /></figure>

<figure><img src="image_39.png" /></figure>
Na₂O + K₂O (wt %)

<figure><img src="image_40.png" /></figure>

Picrobasalt

<figure><img src="image_41.png" /></figure>

Basanite

<figure><img src="image_42.png" /></figure>

Hawaiite

Basalt

50

Basaltic andesite

Andesite

Sub-alkaline series

SiO₂(wt %)

60

Dacite

70

Rhyolite

Trachy- andesite

Basaltic trachy- andesite

Phono- tephrite

Tephri- phonolite

Phonolite

Trachyte
Trachydacite

Alkaline series

Foidite
Rock / chondrite

<figure><img src="image_43.png" /></figure>

La

Pr

Sm

Gd

Dy

Er

Yb

Ce

Nd

Eu

Tb

Ho

Tm

Lu
a

257

<figure><img src="image_44.png" /></figure>

b

253

<figure><img src="image_45.png" /></figure>

c

262

<figure><img src="image_46.png" /></figure>

d

257

<figure><img src="image_47.png" /></figure>

f

258

<figure><img src="image_48.png" /></figure>

g

263

<figure><img src="image_49.png" /></figure>

h

259

<figure><img src="image_50.png" /></figure>

i

252

<figure><img src="image_51.png" /></figure>

k

246

<figure><img src="image_52.png" /></figure>

l

<figure><img src="image_53.png" /></figure>

m

701

<figure><img src="image_54.png" /></figure>

100 µm

e

252

<figure><img src="image_55.png" /></figure>

j

253

<figure><img src="image_56.png" /></figure>
Rock / Primitive mantle

<figure><img src="image_57.png" /></figure>
Wetherill plot. Common-lead uncorrected

<figure><img src="image_58.png" /></figure>

Sample L-2

Wetherill plot. Common-lead uncorrected

<figure><img src="image_59.png" /></figure>

$206\text{Pb}/^{238}\text{U age} = 255 \pm 3 \text{ Ma}$

($\text{Mswd} = 2.0, n=32$)
Ma

249

Lower Triassic

254

Middle-Upper Permian
Guadalupian

Lopingian

Changsingian
Wuchiapingian

Zr

LA2

Ar

<figure><img src="image_60.png" /></figure>

LA1
Ar

Capitanian

Wordian

Roadian

259

264

269

TA69
Ar

<figure><img src="image_61.png" /></figure>

<figure><img src="image_62.png" /></figure>
~200 Ma

<figure><img src="image_63.png" /></figure>

Atlantic seaward -dipping reflectors

CAMP lava flows

CAMP dykes

CAMP sills

Onshore Triassic basins

Archean craton nuclei

Late Permian - Middle Triassic dykes
<figure><img src="image_64.png" /></figure>

1: Tafilalt

2: Jebilet

3: Pyrenees Axial Zone

4: Sierra de Gredos

5: Subbetics

6: SE Iberian Chain

7: Ougarta

8: Argana (pre-CAMP)

9: CAMP

10: Bohemia
Rock / Primitive mantle

<figure><img src="image_65.png" /></figure>

Tafalt dolerites

Jebilet microdiorites

Bohemia monzodiorites
## A: Early magmatism (260 Ma to ~210 Ma)

<figure><img src="image_66.png" /></figure>

## B: CAMP magmatism (~201 Ma)

<figure><img src="image_67.png" /></figure>
# Geochemistry and age of magmatism during the Pangean breakup - a Moroccan vision

<figure><img src="image_68.png" /></figure>