# Petrochronology of Alkaline Intrusive Magmatism in the Central High-Atlas (Morocco) and Relationships With Jurassic Salt Diapirism

Youssef Ahechach, Muhammad Ouabid, Françoise Roger, Fleurice Parat, Otmane Raji, Jean-Marie Dautria, Valérie Bosse, Cheikh Elwali Malainine, Hicham El Messbahi, Hamza Skikra, et al.

► To cite this version:

Youssef Ahechach, Muhammad Ouabid, Françoise Roger, Fleurice Parat, Otmane Raji, et al. Petrochronology of Alkaline Intrusive Magmatism in the Central High-Atlas (Morocco) and Relationships With Jurassic Salt Diapirism. Geochemistry, Geophysics, Geosystems, 2025, 26 (6), pp.e2025GC012178. 10.1029/2025GC012178. hal-05399239

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# Geochemistry, Geophysics, Geosystems®

## RESEARCH ARTICLE

10.1029/2025GC012178

**Key Points:**

* First U-Pb zircon ages constrained the precise timing and tectonic setting of the Moroccan Central High Atlas Jurassic alkaline intrusions

* An alkaline intrusive province closely associated with Triassic sediments exhumed by diapirism in the Central High Atlas of Morocco

* The relationship of the Central High Atlas Jurassic alkaline intrusions to diapiric structures and the magmatic and hydrothermal processes

**Supporting Information:**

Supporting Information may be found in the online version of this article.

**Correspondence to:**

M. Ouabid, muhammad.ouabid@um6p.ma

**Citation:**

Ahechach, Y., Ouabid, M., Roger, F., Parat, F., Raji, O., Dautria, J.-M., et al. (2025). Petrochronology of alkaline intrusive magmatism in the Central High-Atlas (Morocco) and relationships with Jurassic salt diapirism. *Geochemistry, Geophysics, Geosystems*, 26, e2025GC012178. https://doi.org/10.1029/2025GC012178

Received 14 JAN 2025

Accepted 9 MAY 2025

© 2025 The Author(s). Geochemistry, Geophysics, Geosystems published by Wiley Periodicals LLC on behalf of American Geophysical Union. This is an open access article under the terms of the Creative Commons Attribution License, which permits use, distribution and reproduction in any medium, provided the original work is properly cited.

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## Petrochronology of Alkaline Intrusive Magmatism in the Central High-Atlas (Morocco) and Relationships With Jurassic Salt Diapirism

Youssef Ahechach<fnref n="1" />, Muhammad Ouabid<fnref n="1" />✉, Françoise Roger<fnref n="2" />, Fleurice Parat<fnref n="2" />, Otmane Raji<fnref n="1" />, Jean-Marie Dautria<fnref n="2" />, Valérie Bosse³, Cheikh Elwali Malainine<fnref n="1" />, Hicham El Messbahi⁴, Hamza Skikra<fnref n="1" />✉, and Jean-Louis Bodinier⁵

**Abstract** A striking feature of the Imilchil alkaline province, in the Moroccan Central High Atlas, is the close association of magmatic intrusions with Triassic sediments exhumed by diapirism. This study reports field, petrographic, geochemical and Sr-Nd isotopic data, as well as new U-Pb zircon ages, which allow us to constrain the timing of the intrusions and to discuss their relationship to diapiric structures and the magmatic and hydrothermal processes. Our results indicate a restricted range of Late Jurassic ages (ca. 163 ± 1—ca. 149 ± 2 Ma) for the paroxysm of magmatic activity and confirm the diachronicity between diapirism and magmatism, albeit with a short interval of only a few My between the end of diapirism and the magmatic paroxysm. Deep-seated faults associated with transtensional tectonics, accommodating the increasing spreading rate of the Central Atlantic during the Jurassic, played a key role in the close association of magmatic intrusions with diapirs. The faults first triggered active diapirism in the Early-Jurassic and then guided OIB-type partial melts formed by local decompression melting of the mantle. At shallower depths, the diapiric structures acted as corridors for magma migration, probably using evaporites as preferential pathways. Magmas underwent fractional crystallization and interacted with evaporites to produce pervasive Na-rich hydrothermal fluids.

**Plain Language Summary** This study focuses on the central Moroccan High Atlas of Imilchil, characterized by an association between alkaline igneous intrusions and Triassic sediments exhumed by salt diapirism. Our new field, petrographic, geochemical, and isotopic data and U-Pb ages for five main intrusions indicate for the first time that these Imilchil intrusions are Late Jurassic in age (ranging from 163 Ma to 149 Ma) and provide a better understanding of the magmatic and hydrothermal processes involved in their formation. They also show that the regional major paleo-faults reactivated by transtensional tectonics played a significant role in the association of these magmatic intrusions with the diapirs. These faults first generated active salt diapirism in the Early Jurassic, followed by passive diapirism associated with filling sedimentary mini basins in the Middle Jurassic. Subsequently, diapiric structures also facilitated the migration of mantle-derived magma, which interacted with evaporites to generate Na-rich hydrothermal fluids that locally modified differentiated igneous rocks during their crystallization.

## 1. Introduction

When emplaced in continental environments, magmas of mantle origin exhibit greater geochemical and isotopic diversity than their oceanic counterparts, due to the multiple sources of differentiation and contamination resulting from the need to rise through—generally thick—lithospheric chemical and conductive thermal boundary layers. The extent of change that magmas undergo strongly depends on how easily they can penetrate these boundaries, and hence on the local tectonic setting. In turn, deciphering their geochemical signature, combined with field observations, can reveal the geodynamic scenario in which the magmatic activity took place (Fischer et al., 2020; Rollinson & Pease, 2021 and references therein). Along the northern margin of the West African Craton (WAC), the intracontinental High-Atlas belt of Morocco is an inverted rift that was uplifted in the Late Cretaceous and Cenozoic, following several stages of continental rifting during the Mesozoic, coeval with the opening of the Atlantic Ocean (Figure 1a) (Frizon De Lamotte et al., 2008; Moussaid et al., 2023; Skikra

<footnote n="1">¹Geology & Sustainable Mining Institute (GSMI), University Mohammed VI Polytechnic (UM6P), Benguerir, Morocco,</footnote>

<footnote n="2">²Géosciences Montpellier, Université de Montpellier & CNRS, Montpellier, France, ³Laboratoire Magmas et Volcans,</footnote>

<footnote>Université Clermont Auvergne, Campus Universitaire des Cézeaux, Clermont-Ferrand, France, ⁴Department of Geology,</footnote>

<footnote>Polydisciplinary Faculty, Sidi Mohamed Ben Abdallah University, Taza, Morocco, ⁵Institute for Advanced Studies,</footnote>

<footnote>University Mohammed VI Polytechnic (UM6P), Benguerir, Morocco</footnote>
<figure><img src="image_3.png" /><figcaption>**Figure 1.** (a) Sketch map showing the location of the Central High-Atlas Mountains in NW WAC (adapted from Ouabid, Garrido, et al., 2021). (b) Simplified geological map and (c) cross-section showing the spatial distribution of the studied Jurassic-Cretaceous magmatism and their relationships with the Mesozoic sedimentary series (adapted from Michard et al. (2011) and references therein). (d) Detailed geological map (adapted from 1:100,000 geological maps of Imilchil, and Tounfite) and (e) cross-section showing the spatial distribution and relationships among the lithological igneous mafic and felsic units of the studied Tirrhist intrusions.</figcaption></figure>
et al., 2021 and references therein). A key feature of the belt is a Jurassic to Eocene magmatic activity in the Central part of the High Atlas, including in particular the Jurassic to Cretaceous Imilchil alkaline province, expressed by a network of plutonic and hypabyssal intrusions (dykes, sills and laccoliths of mafic and felsic rocks) and secondarily by some basaltic lava flows (Figures 1b–1d) (e.g., Allouban et al., 2021; Armando, 1999; Bensalah et al., 2013; Essaifi & Zayane, 2018; Guezal et al., 2011; Ouabid, Raji, et al., 2021; Studer, 1987). Compared to the rest of the High Atlas segments (Western and Eastern), the Central High Atlas also experienced maximum subsidence and sedimentation in the Mesozoic and was pervasively structured in the Jurassic by syn- to post-rift halokinesis into a system of minibasins separated by diapir ridges (Martín-Martín et al., 2017; Michard et al., 2011; Saura et al., 2014). Of particular note is the close association of igneous intrusions with diapiric structures, suggesting that both are associated with the same tectonic structures at crustal/lithospheric scales, and that magmas used the diapirs as pathways through the Mesozoic cover (Saura et al., 2014). This paper presents and discusses new integrated field, petrographic, mineral and bulk rock geochemical, and Sr-Nd isotopic data for the major alkaline intrusions in the Central High-Atlas. In addition, we report for the first time zircon U-Pb ages to determine the precise timing of the magmatic event. Combined with literature data, these new results allow us to constrain the tectonic framework of the intrusions and their relationship to diapirs, with implications for the rifting evolution of the Moroccan High Atlas.

## 2. Geological Context of the Central High Atlas Imilchil Alkaline Province

The Moroccan High Atlas is an intra-continental alpine chain with heights in the range of 3–4 km, which developed on the NW margin of the WAC, on the North African plate (Figure 1a). The development and structuring of the belt is linked to rifting processes that began in the Early Triassic and resulted in the formation of a series of sedimentary basins throughout most of the Mesozoic (Calvín et al., 2018; Frizon De Lamotte et al., 2008; Piqué et al., 2002; Skikra et al., 2021; Teixell et al., 2003 and references therein). Rifting occurred in response to the opening of the Atlantic Ocean and resulted in the break-up of the Variscan/Pan-African basement of Pangea, controlled by the reactivation of Paleozoic/Precambrian paleofaults (Figure 1b) (e.g., Domènech et al., 2015; Frizon De Lamotte et al., 2008; Laville et al., 2004; Ouabid & Garrido, 2023; Piqué et al., 2002; Teixell et al., 2003; and references therein). Subsidence and sedimentation reached a maximum in the Central High Atlas, providing a comprehensive record of the orogenic evolution of the High Atlas. Furthermore, since the early recognition of diapirism in the Central High Atlas (Bouchouata et al., 1995; Ettaki et al., 2007), there has been a growing consensus that Jurassic syn- and post-rift halokinesis is responsible for most of the structuring and deformation of the sedimentary cover in this region (Martín-Martín et al., 2017; Michard et al., 2011; Saura et al., 2014). The resulting architecture is framed by diapir ridges and elongated mini-basins with a dominant NE-SW trend, and exhibits distinctive halokinetic structures similar to those described in salt basins (Mart & Ross, 1987; Rowan & Vendeville, 2006). The diapirism of the Late Triassic evaporite-bearing deposits is thought to have been triggered by Paleozoic/Precambrian basement faulting associated with the transtensional to extensive tectonic regime that dominated the Central High Atlas in the Mesozoic (Bouchouata, 1994; Laville, 1988; Martín-Martín et al., 2017; McClay et al., 2004; Saura et al., 2014). Allochthonous blocks of the Early Jurassic cover (Hettangian-Early Sinemurian) were also transported by the diapirs, dating the onset of halokinesis from the Sinemurian to the earliest. Subsequently, passive diapiric growth occurred during the Early Jurassic as a result of sedimentation in mini-basins, with the most recent activity recorded in the Middle Jurassic in the Central High Atlas Imilchil region based on deposition of post-diapiric Bathonian-Callovian sediments (Saura et al., 2014). The structure has been only slightly modified by orogenic compression and uplift into a series of faulted anticlines and flat-bottomed synclines, mainly oriented ENE-WSW (Figures 1b and 1c) (Mattauer et al., 1977; Michard et al., 2011; Piqué et al., 2002; Teixell et al., 2003).

During the Triassic, argillites and salts (mainly halite) were first deposited in association with voluminous tholeiitic lava flows of the Central Atlantic Magmatic Province (CAMP) (e.g., Frizon de Lamotte et al., 2008; Michard et al., 2011; Marzoli et al., 2019 and references therein). Subsequently, during the Early Jurassic, a regional marine transgression led to the accumulation of a carbonate sequence (mostly limestones and marls). This sequence was then sealed, between the Middle Jurassic and the Early Cretaceous, by a red bed of continental sandstones formed as a consequence of basin filling or marine regression (e.g., Cavallina et al., 2018; Frizon de Lamotte et al., 2008, 2009; Haddoumi et al., 2002, 2010; Michard et al., 2011; Oussou et al., 2024 and references therein). During the Middle Cretaceous, a new marine platform sedimentary series composed mainly of carbonates developed (e.g., Cavallina et al., 2018; Froitzheim et al., 1988 and references therein), which was
subsequently sealed in the Late Cretaceous by red terrigenous deposits. The latter are considered to be the first
markers of the Africa-Europe convergence and the onset of the Alpine orogeny (e.g., Ettachfini & Andreu, 2004;
Ettachfini et al., 2005; Froitzheim et al., 1988; Laville et al., 1977; Michard et al., 2011).

Basaltic lava flows occur interspersed with Cretaceous red beds and are therefore considered to be of the same age (e.g., Bensalah et al., 2013, 2016; Frizon de Lamotte et al., 2008; Haddoumi et al., 2010; Michard et al., 2013), as confirmed by Ar-Ar dating of lavas south of the Imilchil alkaline province (Moratti et al., 2018). However, the most obvious manifestation of igneous activity in the Central High-Atlas is a large suite of alkaline to transitional hypabyssal and plutonic intrusions located in the core of anticlinal ridges, in elliptical to rhomboidal structures, where they are intimately associated with diapiric Triassic sediments in variable proportions (Figures 1b–1e, Table 1) (Ahechach et al., 2024; Allouban et al., 2021; Armando, 1999; Essaifi & Zayane, 2018; Lhachmi et al., 2001; Studer, 1987; Zayane et al., 2002; and references therein). The contacts between the composite—igneous and sedimentary—units and the Mesozoic sedimentary country rocks are occasionally subvertical and faulted. The subvertical intrusive contacts may alternatively represent inherited diapiric boundaries (e.g., Teixell et al., 2024). In the field, they are underlain by breccias composed of fragments of various sizes (up to 50 cm) of igneous intrusive rocks (mainly mafic) and limestones disseminated in a white (albite-rich) powdery matrix (Figures 1e and 2a–2b) (e.g., Essaifi & Zayane, 2018). The country rocks show moderate schistous deformation (e.g., Laville & Pique, 1992) but lack intrusive contacts and contact metamorphism. Some rare chilled margins are locally observed at the margins of igneous intrusions within the composite structures (e.g., Figure 2b).

The geochronological data available in the literature for the Imilchil alkaline province (mainly whole-rock and mineral K/Ar or Ar/Ar data) indicate ages ranging from 173 ± 4 to 107 ± 3 Ma, suggesting that igneous activity occurred over a long period of about 60–70 Myr (e.g., Armando, 1999; Hailwood & Mitchell, 1971; Martin et al., 1978; Rolley, 1978; Westphal et al., 1979). However, some authors have questioned certain age estimates, especially those obtained for lava flows, due to inconsistencies with their stratigraphic positions (e.g., Bensalah et al., 2013, 2016; Michard et al., 2013). As for the igneous rocks, their intrusion into diapiric sediments and the faulted contacts with the host sedimentary sequences preclude any age interpretation based on cross-cutting relationships (Figures 1b–1e, 2a, and 2b). The precise timing of intrusive magmatism in the Imilchil alkaline province is therefore still poorly constrained.

3. The Alkaline Intrusions: Internal Structures and Petrography

This study focuses on five of the major structures hosting the alkaline igneous intrusive rocks, which are char-
acterized by their larger size (Table 1; Figure 1b): (a) Tassent (~30 km long and 1–3.5 km wide), (b) Tasraft (~22 km long and 0.5–2.5 km wide), (c) Anemzi (~12 km long and 1–2 km wide), (d) Tirrhist (~10 km long and 1–2.5 km wide), and (e) Anefgou (~8 km long and 0.5–4 km wide). In addition to sedimentary rocks, all these structures contain two types of igneous rocks: mafic (mostly gabbroic) on the one hand, and felsic (mostly syenite) on the other (Table 1; Figure 1b). The sedimentary rocks in these structures or massifs are present in highly variable proportions, ranging from 10% to 70% (Figures 1d and 1e, Table 1), in the form of slabs of variable thickness up to 100 m. They are composed mainly of sandy rocks and clays from the Triassic red bed formation, and subordinately of Jurassic limestones, sometimes recrystallized at temperatures of about 300–500°C (Armando, 1999; Essaifi & Zayane, 2018; Lhachmi et al., 2001). When observed, the contacts between the intrusive igneous units and the sedimentary slabs are sharp. However, these contacts are often brecciated, more or less powdery, and correspond to local zones of brittle deformation. Within the composite units or structures, the igneous intrusions occur as a set of intrusive layered bodies that were injected parallel to each other. They are subvertical to vertical in several instances. Such a structural pattern is particularly visible in the Tirrhist unit, where vertical layers of cumulate troctolite, troctolite, and monzogabbros, ranging between 30 and 400 m in thickness, alternate, however, without showing any rhythmic repetition (Figures 1e and 2c). Sub-horizontal layering is also observed, at least locally, in the Tassent, Tasraft, Anemzi, and Anefgou units. Armando (1999) and Zayane et al. (2002) have discussed the implication of flat layering in those units for the emplacement mechanisms of the igneous intrusions. However, in their recent paleomagnetic study of the gabbros, Calvín et al. (2017) propose that all the intrusions originally occurred vertically and that the current dips result from tilting due to post-emplacement tectonics; Table 1 summarizes the obtained rotation data.

Most mafic bodies exhibit magmatic bedding characterized by alternating white and dark layers (Figures 2d and 2e). The white beds are generally cm-thick and almost mono-mineral (mainly plagioclase). In contrast, the dark
<table><caption>**Table 1**

General Characteristics of the Studied Igneous Intrusive Structures of the Central High-Atlas and List of Samples Studied</caption><thead><tr><th>Structure</th><th>Geometric shape/sedimentary unit (vol.% of massif)</th><th>Rotation</th><th>Magmatic unit (vol.% of massif)</th><th>Sample</th><th>Petrography</th><th>Major &amp; trace elements</th><th>Mineral chemistry</th><th>U-Pb zircon dating</th><th>Sr-Nd isotopes</th></tr></thead><tbody><tr><td>Tassent</td><td>5.69/32.21</td><td>Elongated-folded<br>ellipse (area ~72 km²), ~30 km length and 1–3.5 km wide, with sedimentary unit of about ~45 vol.%</td><td>~80–160° in the core of igneous bodies</td><td>Mafic (~10 vol.%)<br>AP01<br>AP28<br>AP29<br>Felsic (~45 vol.%)<br>AP30<br>AP27</td><td>OM<br>OM<br>OM<br>OM<br>OM</td><td>XRF<br>XRF<br>XRF<br>XRF<br>XRF*<br>ICPMS</td><td>ICPMS<br>ICPMS<br>ICPMS<br>ICPMS<br>ICPMS</td><td>EPMA<br>ICPMS<br>EPMA<br>EPMA<br>EPMA</td><td>LA-ICPMS<br>LA-ICPMS<br>LA-ICPMS<br>LA-ICPMS<br>LA-ICPMS</td><td>TIMS</td></tr><tr><td>Tasraft</td><td>5.86/32.19</td><td>Elongated<br>ellipse (area ~80 km²), 22 km length and 0.5–2.5 km wide, with sedimentary unit of about ~70 vol.%</td><td>~130–175° in the core of the igneous bodies, 40° in Mafic (~75 vol.%)<br>AP40<br>AP37<br>Felsic (~10 vol.%)<br>AP38</td><td>OM<br>OM<br>OM<br>OM</td><td>XRF<br>XRF<br>XRF<br>XRF</td><td>ICPMS<br>ICPMS<br>ICPMS<br>ICPMS</td><td>EPMA<br>EPMA<br>EPMA<br>EPMA</td><td>LA-ICPMS<br>ICPMS<br>EPMA<br>EPMA</td><td>TIMS<br>TIMS</td></tr><tr><td>Anemzi</td><td>5.24/32.31</td><td>Elongated<br>ellipse (area ~25 km²), ~12 km length and 1–2 km wide with sedimentary unit of about ~20 vol.%</td><td>Mafic (~65 vol.%)<br>AN4<br>Felsic (~25 vol.%)<br>AP22<br>AP23<br>AN1<br>AN5<br>AN2</td><td>OM<br>OM<br>OM<br>OM<br>OM<br>OM</td><td>XRF<br>XRF<br>OM<br>OM<br>XRF<br>XRF*<br>ICPMS</td><td>ICPMS<br>ICPMS<br>ICPMS<br>ICPMS<br>ICPMS<br>ICPMS</td><td>EPMA<br>EPMA<br>EPMA<br>EPMA<br>EPMA<br>EPMA</td><td>LA-ICPMS<br>LA-ICPMS<br>LA-ICPMS<br>LA-ICPMS<br>LA-ICPMS<br>LA-ICPMS</td><td>TIMS<br>TIMS<br>TIMS<br>TIMS<br>TIMS<br>TIMS</td></tr><tr><td>Tirrthist</td><td>5.45/32.20</td><td>Elongated<br>rhombus (area ~34 km²), ~10 km length and</td><td>~10–11° in the core of intrusions, and 45–49°<br>In contact with the country sediment rocks</td><td>Mafic (~70 vol.%)<br>IN01<br>AP32<br>IN34<br>IN19</td><td>OM<br>OM<br>OM<br>OM</td><td>XRF<br>XRF<br>XRF<br>XRF</td><td>ICPMS<br>ICPMS<br>EPMA<br>EPMA<br>EPMA<br>EPMA</td><td>LA-ICPMS<br>LA-ICPMS<br>LA-ICPMS<br>LA-ICPMS<br>LA-ICPMS<br>LA-ICPMS</td><td>TIMS</td></tr></tbody></table>
<table><caption>**Table 1**
*Continued*</caption><thead><tr><th>Structure</th><th>Geometric shape/sedimentary unit (vol.% of massif)</th><th>Rotation</th><th>Magmatic unit (vol.% of massif)</th><th>Sample</th><th>Petrography</th><th>Major &amp; trace elements</th><th>Mineral chemistry</th><th>U-Pb zircon dating</th><th>Sr-Nd isotopes</th></tr></thead><tbody><tr><td rowspan="10">G.Cooor (WGS84)</td><td rowspan="10">up to ~2.5 km wide, with sedimentary unit of about ~10 vol.%</td><td rowspan="10"></td><td rowspan="10"></td><td>AP19</td><td>OM</td><td>XRF</td><td>ICPMS</td><td>EPMA</td><td>LA-ICPMS</td></tr><tr><td>IN2</td><td>OM</td><td>XRF</td><td>ICPMS</td><td></td><td>LA-ICPMS</td><td>TIMS</td></tr><tr><td>IN3</td><td>OM</td><td>XRF</td><td></td><td></td><td></td><td></td></tr><tr><td>AP18</td><td>OM</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>IN11</td><td>OM</td><td>XRF</td><td></td><td></td><td></td><td></td></tr><tr><td>IN14</td><td>OM</td><td>XRF</td><td></td><td></td><td></td><td></td></tr><tr><td>IN32</td><td>OM</td><td>XRF</td><td></td><td></td><td></td><td></td></tr><tr><td>IN40</td><td>OM</td><td>XRF</td><td></td><td></td><td></td><td></td></tr><tr><td>Felsic (~10 vol.%)</td><td>IN4</td><td>OM</td><td>XRF</td><td>ICPMS</td><td></td><td>LA-ICPMS</td><td>TIMS</td></tr><tr><td rowspan="10">Anefgou 5.43/32.27</td><td rowspan="10">~20–35° in the core of the igneous bodies</td><td rowspan="10"></td><td rowspan="10"></td><td>AP14</td><td>OM</td><td>XRF</td><td>ICPMS</td><td>EPMA</td><td>LA-ICPMS</td></tr><tr><td>AP620A</td><td>OM</td><td>XRF</td><td>ICPMS</td><td></td><td>LA-ICPMS</td><td></td></tr><tr><td>AP12</td><td>OM</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>AP16</td><td>OM</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>IN20</td><td>OM</td><td>XRF</td><td>ICPMS</td><td>EPMA</td><td>LA-ICPMS</td><td>TIMS</td></tr><tr><td>AP10</td><td>OM</td><td>XRF*</td><td>ICPMS</td><td>EPMA</td><td>LA-ICPMS</td><td>LA-ICPMS</td></tr><tr><td>IN44/43</td><td>OM</td><td></td><td></td><td>EPMA</td><td>LA-ICPMS</td><td>LA-ICPMS</td></tr><tr><td>IN13</td><td>OM</td><td>XRF</td><td></td><td></td><td></td><td></td></tr><tr><td>Mafic (~75 vol.%)</td><td>APVI</td><td>OM</td><td>XRF</td><td>ICPMS</td><td>EPMA</td><td></td><td></td></tr><tr><td rowspan="10">Cylindrical to rough rhombus (area ~20 km²), ~8 km length and 0.5–4 km wide, with sedimentary unit of about ~15 vol.%</td><td rowspan="10"></td><td rowspan="10"></td><td rowspan="10"></td><td>IN22</td><td>OM</td><td>XRF</td><td>ICPMS</td><td>EPMA</td><td>LA-ICPMS</td></tr><tr><td>APV</td><td>OM</td><td>XRF</td><td>ICPMS</td><td>EPMA</td><td></td><td></td></tr><tr><td>IN27</td><td>OM</td><td>XRF</td><td>ICPMS</td><td></td><td></td><td></td></tr><tr><td>Felsic (~10 vol.%)</td><td>AP43</td><td>OM</td><td>XRF</td><td>ICPMS</td><td></td><td></td><td></td></tr><tr><td>APIV</td><td>OM</td><td>XRF</td><td>ICPMS</td><td>EPMA</td><td></td><td></td><td></td></tr><tr><td>IN23</td><td>OM</td><td>XRF</td><td>ICPMS</td><td></td><td></td><td></td><td></td></tr><tr><td>IN21</td><td>OM</td><td>XRF</td><td>ICPMS</td><td></td><td></td><td></td><td></td></tr><tr><td>AP44</td><td>OM</td><td>XRF</td><td>ICPMS</td><td>EPMA</td><td>LA-ICPMS</td><td></td><td></td></tr><tr><td>IN28</td><td>OM</td><td>XRF</td><td></td><td></td><td></td><td></td><td></td></tr></tbody></table>

*Note.* Optical petrographic microscope (OM), X-ray Fluorescence (XRF). Inductively coupled plasma spectrometry (ICPMS). Electron probe micro-analyzer (EPMA). Laser ablation-ICPMS (LA-ICPMS). Thermal ionization mass spectrometer (TIMS). *Analyses from Ouabid, Raji, et al. (2021).*
<figure><img src="image_4.png" /><figcaption>**Figure 2.** Field photographs showing (a), (b) the nature of the main contacts between sedimentary and magmatic units, marked by the breccia zone (e.g., Tirrhist intrusion), (c) alternating intrusive bodies (e.g., Anefgou intrusion), (d), (e) magmatic bedding (Tirrhist and Tassent, respectively), (f) spheroidal weathering affecting the studied mafic magmatic unit (Tirrhist intrusion).</figcaption></figure>

layers are rich in ferromagnesian minerals, mainly clinopyroxenes (±olivine). They are more diffuse, up to a few cm-thick, and contain a significant proportion of plagioclase. This magmatic bedding is mainly sub-vertical (Figure 2d), parallel to the layer but often presents undulations, criss-cross structures, and even local disturbances, with zones of discordant bedding and sub-horizontal bedding as well (Figure 2e). In the outcrop, the mafic units are more altered than the felsic units with widespread supergene spheroidal weathering (Figure 2f). The mafic units seem to be cut by felsic intrusive units varying in size and shape. The largest have lengths ranging from 0.5 to 2 km long and widths never exceeding 700 m, and their shapes, in most cases, seem little influenced by the gabbroic intrusions they crosscut (Figures 1d and 1e). Where observable, the contact between felsic and mafic units is sharp (Figures 3a and 3b). In addition to the largest felsic unit structures described above, numerous smaller felsic cm-sized pockets, and dykes or veins up to 40 cm thick crosscut the mafic units (Figures 3c and 3e).
<figure><img src="image_5.png" /><figcaption>**Figure 3.** Field photographs of the studied Central High-Atlas Imilchil igneous intrusive massifs ((a) Anefgou and (b)–(f) Tirrhist intrusions), showing the main structures of felsic units and their relationships with mafic units.</figcaption></figure>

The felsic rocks are mainly coarse-to medium-grained monzodiorite-monzonite-syenite rock associations (Figures 1e and 3). In the field, the boundaries between these different felsic rock types are often hidden, and their relationships are difficult to understand. However, syenite is always the dominant rock type. All the rocks belonging to the felsic units are leucocratic rocks, always forming high reliefs, and they form most of the main crests of the intrusive massifs (e.g., J. Agra, and J. Hayim; Figures 1d, 3e, and 3a; elevation ~2.7–2.8 km). Another particularity of the felsic units is the presence of pegmatite pockets (Figure 3f), whose diameters vary between few cm–to–few m. These pockets are generally grouped at the apex of felsic unit bodies. The mineralogy of these pockets is very peculiar and consists of albite-feldspar, garnet, zircon, amphibole, magnetite, titanite, and apatite. All these phases reach unusual sizes, always greater than 1 cm (some can even be 10 cm), and the crystals (some of gem-quality) are often subhedral and interlocked. The most beautiful Moroccan apatite gems are extracted from veins associated with these felsic unit rocks (Ouabid & Raji, 2023; Ouabid, Raji, et al., 2021).

The petrographic characteristics of the main lithological facies of the mafic and felsic units are quite similar throughout the studied major Central High-Atlas Imilchil intrusive massifs (e.g., Figures 1b–1e; Tables 1 and 2). Based on their mineral paragenesis and textures, the magmatism produces four main lithologies for the mafic unit,
<table><caption>**Table 2**
Petrographic Features of the Selected Samples From the Central High-Atlas Intrusive Massifs</caption><thead><tr><th>Massif</th><th>Unit</th><th>Facies</th><th>Sample</th><th>Aspect & texture</th><th>Major minerals</th><th>Accessories</th><th>Secondary minerals</th></tr></thead><tbody><tr><td rowspan="3">Tassent</td><td rowspan="3">Mafic</td><td rowspan="3">Cumulate troctolite gabbro<br>Hydrothermal altered gabbros</td><td>AP01</td><td>Dark green rock, coarse-grained, poikilitic</td><td>O1, Fsp, Cpx, Bt, Opq</td><td>Ap, Zrn</td><td>Srp, Chl</td></tr><tr><td>AP28</td><td>Dark gray to gray, coarse-grained</td><td>Fsp, Cpx, Bt, Opq</td><td>Ilm</td><td>Chl</td></tr><tr><td>AP29</td><td>Dark gray to gray rock, coarse-grained</td><td>Fsp, Ap, Cpx, Amp, Opq</td><td>Rt, Ttn</td><td>Prh, Chl</td></tr><tr><td rowspan="3">Felsic</td><td rowspan="3">Monzonitic rock<br>Syenite</td><td>AP30</td><td>Gray to white-gray rock, coarse-grained</td><td>Fsp, Amp, Cal</td><td>Opq, Ttn, Ap</td><td>Cal</td></tr><tr><td>AP27</td><td>White rock, coarse-grained</td><td>Fsp, Amp, Bt</td><td>Ap</td><td>Prh, Chl, Cal</td></tr><tr><td>AP40</td><td>Dark green rock, coarse-grained, intersertal</td><td>Fsp, Ol, Cpx.Bt</td><td>Opq, Ap</td><td>Srp</td></tr><tr><td rowspan="3">Tasraft</td><td rowspan="3">Mafic</td><td rowspan="3">Troctolite gabbro<br>Hydrothermal altered gabbro</td><td>AP37</td><td>Dark gray rock, coarse-grained</td><td>Fsp, Amp, Cpx</td><td>Ttn, Opq, Ap</td><td>Prh, Cal, Chl</td></tr><tr><td>AP38</td><td>White rock, medium to coarse-grained</td><td>Fsp, Amp, Cal</td><td>Opq, Qz</td><td>Qz, Cal, Ep</td></tr><tr><td>AN4</td><td>Dark green rock, poikilitic</td><td>O1, Fsp, Cpx, Bt</td><td>Opq, Ap</td><td>Srp, Chl</td></tr><tr><td rowspan="3">Anemzi</td><td rowspan="3">Mafic</td><td rowspan="3">Troctolite gabbro<br>Felsic Monzonitic rocks</td><td>AP22</td><td>Gray to white-gray rock, coarse-grained</td><td>Fsp, Amp, Bt</td><td>Opq, Ap, Zrn</td><td>Ep</td></tr><tr><td>AP23</td><td>Gray to white-gray rock, coarse-grained</td><td>Fsp, Amp, Bt</td><td>Opq, Ap, Zrn</td><td>Ep</td></tr><tr><td>AN1</td><td>White rock, coarse-grained</td><td>Fsp, Cpx, Amp</td><td>Opq, Qz, Rt, Zrn</td><td>Cal, Chl, Ep</td></tr><tr><td rowspan="3">Syenitic rocks</td><td rowspan="3">Mafic</td><td rowspan="3">Troctolite gabbro</td><td>AN2</td><td>White rock, coarse-grained</td><td>Fsp, Cpx, Amp</td><td>Opq, Qz, Rt, Zrn</td><td>Chl, Cal, Ep</td></tr><tr><td>AN5</td><td>White rock, coarse-grained</td><td>Fsp, Cpx, Amp</td><td>Opq, Zr, Qz, Zrn</td><td>Ser, Chl</td></tr><tr><td>AP32</td><td>Dark green rock, medium-grained, intersertal and locally poikilitic</td><td>Fsp, Ol, Cpx, Opq</td><td>Bt, Ap</td><td>Chl, Srp</td></tr><tr><td rowspan="3">Tirrhist</td><td rowspan="3">Mafic</td><td rowspan="3">Troctolite gabbros</td><td>IN32</td><td>Dark green rock, medium-grained, intersertal and locally poikilitic</td><td>Fsp, Ol, Cpx, Opq</td><td>Bt, Ap</td><td>Chl, Srp</td></tr><tr><td>IN34</td><td>Dark green rock, medium-grained, intersertal and locally poikilitic</td><td>Fsp, Ol, Cpx, Opq</td><td>Ap</td><td>Chl</td></tr><tr><td>IN19</td><td>Dark green rock, medium to coarse-grained, intersertal</td><td>Fsp, Cpx, Ol, Bt, Opq</td><td>Ap</td><td>Chl, Ep, Prh</td></tr><tr><td rowspan="3">Hydrothermal altered gabbros</td><td rowspan="3">Mafic</td><td rowspan="3">Troctolite gabbro</td><td>IN40</td><td>Dark gray rock, coarse-grained</td><td>O1, Fsp, Cpx, Bt, Opq</td><td>Ap</td><td>Chl</td></tr><tr><td>IN01</td><td>Dark gray rock, medium-grained, intersertal</td><td>Fsp, Ol, Cpx, Bt, Opq</td><td>Ap</td><td>Chl</td></tr><tr><td>IN03</td><td>Dark gray rock, medium to fine grained</td><td>Fsp, Ol, Cpx, Bt, Opq</td><td>Ap</td><td>Chl, Ep, Prh</td></tr><tr><td rowspan="3">Monzogabbros</td><td rowspan="3">Mafic</td><td rowspan="3">Troctolite gabbro</td><td>AP19</td><td>Dark gray rock, medium-grained, intersertal</td><td>Fsp, Cpx, Ol, Bt, Opq</td><td>Ap</td><td>Ser</td></tr><tr><td>IN2</td><td>Dark gray rock, coarse-grained</td><td>O1, Fsp, Cpx, Bt, Opq</td><td>Ap, Zrn</td><td>Ser</td></tr><tr><td>AP18</td><td>Dark gray rock, coarse-grained</td><td>Fsp, Ol, Cpx, Opq, Bt</td><td>Ap</td><td>Chl, Srp</td></tr><tr><td rowspan="3">Tasraft</td><td rowspan="3">Mafic</td><td rowspan="3">Troctolite gabbro</td><td>IN11</td><td>Dark gray rock, coarse-grained</td><td>O1, Fsp, Cpx, Bt, Opq</td><td>Ap</td><td>Chl, Srp</td></tr><tr><td>IN14</td><td>Dark gray rock, coarse-grained</td><td>Fsp, Cpx, Ol, Bt, Opq</td><td>Ap</td><td>Chl, Srp</td></tr></tbody></table>
<table><caption>Table 2<br>Continued</caption><thead><tr><th>Massif</th><th>Unit</th><th>Facies</th><th>Sample</th><th>Aspect & texture</th><th>Major minerals</th><th>Accessories</th><th>Secondary minerals</th></tr></thead><tbody><tr><td rowspan="10">Felsic</td><td rowspan="10">Monzonitic rocks</td><td rowspan="10">Syenitic rocks</td><td>IN4</td><td>Gray to white gray rock, coarse-grained</td><td>Fsp, Amp</td><td>Ap, Zrn, Opq</td><td>Ep, Chl</td></tr><tr><td>AP14</td><td>White-gray rock, coarse-grained</td><td>Fsp, Amp</td><td>Ap, Opq</td><td>Cal</td></tr><tr><td>AP12</td><td>White rock, coarse-grained</td><td>Fsp, Amp</td><td>Ap, Opq</td><td>Amp, Cal, Chl</td></tr><tr><td>AP16</td><td>White rock, coarse-grained</td><td>Fsp, Amp</td><td>Ap, Opq</td><td>Cal, Amp</td></tr><tr><td>IN20</td><td>White rock, coarse-grained</td><td>Fsp, Amp, Cpx</td><td>Ap, Opq, Qz, Zrn</td><td>Ep, Ser</td></tr><tr><td>AP620A</td><td>White rock, coarse-grained</td><td>Fsp, Amp</td><td>Ap, Opq, Qz, Zrn</td><td>Ep, Ser</td></tr><tr><td>AP10</td><td>White rock, coarse-grained</td><td>Fsp, Grt, Cpx</td><td>Ap, Opq, Zrn</td><td>Ep, Ser, Chl</td></tr><tr><td>IN43</td><td>White rock, coarse-grained/pegmatite</td><td>Fsp, Grt, Amp</td><td>Opq, Zrn, Cpx, Ap</td><td>Ser, Ep</td></tr><tr><td>IN44</td><td>White rock, coarse-grained/pegmatite</td><td>Fsp, Grt, Amp</td><td>Opq, Zrn, Cpx, Ap</td><td>Ep, Ser, Chl</td></tr><tr><td>IN13</td><td>White rock, coarse-grained/pegmatite</td><td>Fsp</td><td>Zrn, Ap, Opq</td><td>Chl</td></tr><tr><td rowspan="10">Anefgou</td><td rowspan="10">Mafic</td><td rowspan="10">Cumulate troctolite gabbro<br>Hydrothermal altered gabbros</td><td>APVI</td><td>Dark green rock, medium-grained, Poikilitic</td><td>Ol, Fsp, Cpx, Opq, Bt</td><td>Ap</td><td>Chl, Srp</td></tr><tr><td>IN22</td><td>Dark green rock, medium-grained intersertal</td><td>Fsp, Ol, Cpx, Opq</td><td>Bt, Zrn</td><td>Chl</td></tr><tr><td>APV</td><td>Dark gray rock, medium-grained intersertal</td><td>Fsp, Cpx, Amp, Opq, Bt</td><td>Ap</td><td>Ep</td></tr><tr><td>IN27</td><td>Dark gray rock, coarse-grained</td><td>Fsp, Amp, Opq</td><td>Ttn, Ap</td><td>Ep</td></tr><tr><td>APIV</td><td>Gray to white-gray rock, medium to coarse-grained</td><td>Fsp, Cpx, Opq</td><td>Ttn, Ap</td><td>Ept</td></tr><tr><td>IN21</td><td>White rock, coarse-grained</td><td>Fsp, Amp</td><td>Opq, Ap, Zrn</td><td>Ep, Ser</td></tr><tr><td>IN23</td><td>White rock, coarse-grained</td><td>Fsp, Amp</td><td>Opq, Ap, Opq, Qz, Zrn</td><td>Ep, Ser</td></tr><tr><td>AP43</td><td>White rock, coarse-grained</td><td>Fsp</td><td>Opq, Rt, Qz</td><td>Chl, Cal</td></tr><tr><td>AP44</td><td>White rock, coarse-grained</td><td>Fsp</td><td>Opq, Rt, Qz</td><td>Chl, Cal</td></tr><tr><td>IN28</td><td>White-gray rock, coarse-grained</td><td>Fsp, Am</td><td>Opq, Qz, Ap</td><td>Chl, Cal</td></tr></tbody></table>

Note. Abbreviation of common rock-forming minerals are following Warr (2021) and Whitney and Evans (2010).
including cumulate troctolite-gabbro, troctolite-gabbros, monzogabbros, and hydrothermally altered gabbros. In addition, the felsic units consist mainly of monzodiorite, monzonite, and syenite rocks, with the latter always being the dominant rock type. Detailed petrographic features of the mafic and felsic unit rocks are summarized in Table 2 and Figures 4 and 5 and described below.

## 3.1. Mafic Rock Unit

Petrographic observations reveal that the mafic unit is composed of four distinct facies: cumulative troctolite-, troctolite-, hydrothermally altered-, and monzo-gabbros. The cumulate troctolite gabbros are dark green, medium to coarse-grained rocks with a tendency to a poikilitic-to-intertedral texture; they present a relatively rough magmatic foliation (Figure 4). The petrographic lithology is predominantly plagioclase, olivine, clinopyroxene, opaque minerals (mainly Fe-Ti-oxides), and biotite (Figure 4, Table 2). Apatite and zircon are the main accessory minerals and serpentine and chlorite occur as the main secondary phases. Plagioclase (~40–55 vol. %) is always the dominant phase. It occurs as subhedral to euhedral laths up to 6 mm in length and seems to be the earliest phase during fractional crystallization together with olivine. The olivine (up to 30 vol.%) forms subhedral to ovoid shape crystals up to 1–2 mm in size. The crystals exhibit numerous microfractures. Moreover, the olivine grains exhibit slight serpentinization and ferruginous alteration along the microfractures. Both plagioclase and olivine crystals include an early generation of small opaques of Fe-Ti oxides (<100 µm). Clinopyroxene (mainly augite) crystallizes in the form of anhedral crystals. Finally, interstitial Fe-Ti oxides crystallize. The last phase to crystallize seems to be biotite: it is systematically associated with opaque minerals and can locally form a continuous corona around them. Moreover, the biotite periphery is partially transformed into chlorite. In some samples, olivine contains micro-inclusions of Fe-Ti oxides, generally ovoid in shape and about 100 µm in size, and aligned along sub-parallel directions oblique to the elongation crystals. Few euhedral zircons (up to 200 µm) and apatite as inclusions are observed as major accessory minerals.

With petrographic characteristics similar to cumulate troctolitic gabbro lithology, the troctolite gabbros are also dark gray green rocks, and they display coarse-to-medium grained textures, often with a tendency to intersertal-to-poikilitic texture, and most of them show a relatively clear magmatic fluidity. Troctolite gabbros are composed of plagioclase, olivine, clinopyroxene, Fe-Ti-oxides, and biotite (e.g., Figure 4, Table 2), but they have more clinopyroxene and less olivine than cumulate troctolite gabbro. Plagioclase remains the dominant mineral (40–50 vol.%) and appears to be the earliest in the order of crystallization. Olivine crystals are subhedral to ovoid in shape and can reach 2 mm in size. As in the cumulate gabbro, olivines are also highly microfractured and filled by serpentinite. Clinopyroxene is the third phase to crystallize and occurs as subhedral to anhedral grains often poikilitic. Finally, the oxides (mainly ilmenite) appeared whose habit is always interstitial. Biotite is late and it systematically forms coronas, sometimes continuous, around Fe-Ti oxides. In addition, the biotite is partially chloritized. Euhedral apatite needles (up to ~500 µm long) and euhedral zircon (up to 300 µm long) are always present as accessory minerals but rare; they are probably the first minerals to have crystallized in troctolite and cumulate gabbros.

The monzogabbro lithologies are more evolved and lighter in color than cumulate and troctolite gabbros, being generally pale green to dark gray rocks (e.g., Figures 2 and 4, Table 2). They are medium-to coarse-grained, and their textures are typically gabbroic, but unlike troctolites, magmatic foliation is almost absent. As with previous gabbro types, they are composed of plagioclase, olivine, clinopyroxene, and biotite, but they differ in the proportions of mineral phases (e.g., olivine is minor). Accessory and secondary mineral phases consist mainly of oxides, apatite, zircon, titanite, rutile, chlorite, sericite, serpentine, prehnite, and epidote. Plagioclase is the dominant phase (~40–55 vol.%) and it occurs as subhedral to euhedral laths (up to 8 mm). As in the troctolites, the olivine crystals are ovoid (up to 3 mm in size). The second phase in the order of crystallization seems olivine, but it is minor, constituting less than 10 vol.% of the rock. Clinopyroxene (10–20 vol.%) appears as subhedral crystals (size, up to 3 mm, up to 6 mm in length). Biotite (~15 vol.%) occurs as anhedral grains and is often associated with iron oxide (mainly ilmenite). Apatite needles (50–500 µm long) are always present and are more abundant than in troctolitic rocks. In addition, euhedral zircons (up to 500 µm) are present but rare.

Finally, the hydrothermally altered gabbro lithologies have a greenish and subhedral-granular texture. They consist of both relic primary assemblages and an important secondary mineral association, including green amphibole, epidote, calcite, chlorite prehnite, albite, and titanite. This lithology seems to be the product of hydrothermal alteration of the previous gabbro types.
<figure><img src="image_6.png" /><figcaption>**Figure 4.** Representative petrographic microphotographs of the main lithological facies of the mafic units of the studied Central High-Atlas Imilchil intrusive magmatism.</figcaption></figure>
## 3.2. Felsic Rock Unit

The rocks of the felsic units—including monzonitic facies (monzodiorite–monzonite) and syenite facies (dominant) association—of the studied Central High-Atlas Imilchil intrusive magmatism are even lighter in color (Table 2, Figures 3–5). The monzodiorite samples occur as greyish and coarse to medium-grained rocks composed primarily of plagioclases (up to 60–70 vol.%) occurring as subhedral to euhedral lathes, 2 mm in size with a sericitized core (Table 2, Figure 5). Amphibole is abundant (up to 30 vol.%) and the crystals display strong and discontinuous optical zoning with generally green cores and brown rims. It also exhibits a strong alteration with chlorite or second amphibole in contact with iron oxides. Minor µm-sized biotite flakes are also observed in or close to the amphibole. Clinopyroxene is occasionally observed with grains up to 0.2 mm in size. Titanite and apatite are abundant but randomly distributed throughout the rocks, and epidote and chlorite are the most common secondary phases together with iron oxides. Akin to monzodiorite rocks, monzonite samples are coarse-grained and white gray to gray rocks. They are mostly composed of plagioclase (up to 70 vol.%). The plagioclase is mostly albite. These rocks also contain a significant amount of amphibole, up to 30 vol.% of the rock. The amphibole occurs as green sub-euhedral to euhedral crystals with sizes up to 1 cm long. Clinopyroxene grains are rarely observed scattered throughout the rocks with diameters up to 0.2 mm. The main accessory phases are iron oxides, titanite, apatite and zircon. Secondary mineral phases such as calcite, epidote, and chlorite are observed.

Toward, the more evolved facies of the studied felsic unit rock association, the syenites are leucocratic to hol-
oleucocratic white rocks with coarse-grained to pegmatitic textures. They are composed of albitic plagioclase
(~70–90 vol.%), amphibole (~3–15 vol. %), clinopyroxene (<5 vol.%), opaques (<5 vol.%) (e.g., Figure 5,
Table 2). Accessory phases include magnetite, titanite, zircon, and apatite. In some syenites, quartz occurs but
always in very small amounts. Amphibole and clinopyroxene coexist in these rocks and the crystals of these two
phases are subhedral and they have sizes up to 0.7 cm long. The first one is bright green possibly of tremolite and
edinite composition, while the second is beige in color like diopside.

The apexes of syenite intrusions are pegmatitic and consist of up to 80 vol.% of pluricentimeter laths of near pure, strongly twinned albite (e.g., Figure 3f). These laths are oriented and often contain apatite inclusions. They also present a diffuse clay alteration organized in tiling. Some include dissolution cells filled with clumps of chlorite and sericite. Anhedral magnetite crystals occupy the interstices between the albite crystals. Subhedral to anhedral cm-sized amphibole crystals also occupy the interstices between the albite crystals. In some samples, subhedral to anhedral crystals of clinopyroxene (~5–10 vol.%) occur highly fractured and their rims are reactional and transformed to actinote. They also contain iron-oxide inclusions. Zircon, apatite, rutile, titanite, prehnite, and epidote, in addition to magnetite, are the main accessory and secondary phases in pegmatites. In some localities (Tirrhist and Tassent), subhedral crystals up to 8 mm in diameter of garnet are abundant and can reach up to 5 vol.%.

The quartz-syenites rocks are leucocratic coarse-grained rocks, consisting primarily of euhedral to subhedral albite crystals (up to 70 vol.%); size reach up to 4 mm, subhedral amphibole crystals (up to 20 vol.%) with a possible ferro-actinolite affinity, and display a distinct zoning from core to rim under microscope. They contain several acicular microcrystals of apatite, which occur in both subhedral and anhedral amphibole crystals. Locally, these amphiboles fill the spaces between the albite crystals, sometimes associated with rare quartz grains. When in contact with quartz, the amphiboles have euhedral forms, indicating that these spaces likely originally contained residual liquid. These ferro-actinolite typically cluster the multi-millimeter aggregates, reaching up to 4 mm in size and containing subhedral to euhedral titanomagnetite crystals, surrounded by light brown areas containing biotite and ferriferous chlorite. These titanomagnetites often contain apatite inclusions. Additionally, rare epidote and zircon crystals are scattered throughout the rock.

# 4. Materials and Methods

## 4.1. Mineral Chemistry

In situ major element micro-analyses for minerals were carried out at the Geosciences Montpellier laboratory (Montpellier, France) using the Cameca SX100 Electron probe micro-analyzer (EPMA). Major element analyses were performed using a µm-focus beam, an accelerating voltage of 15 kV and a probe current of 20 nA. Analytical settings were used with a count time of 20 s for Si, Al, Fe, Mg, Ca, Ti, Mn, Na, K, Cr, and Ni. A combination of natural and synthetic standards is used for calibration, including albite for Na, orthoclase for K, Al₂O₃ for Al,
<figure><img src="image_7.png" /><figcaption>**Figure 5.** Representative petrographic microphotographs of the main lithological facies of the felsic units of the Central High-Atlas magmatism studied.</figcaption></figure>
wollastonite for Si and Ca, TiO₂ for Ti, forsterite for Mg, rhodochrosite for Mn, Fe₂O₃ for Fe, and native metal for Ni. The major element compositions of analyzed main minerals are listed in the supplementary materials (Table S1 in Supporting Information S1 for olivines, Table S2 in Supporting Information S1 for feldspars, Table S3 in Supporting Information S1 for clinopyroxenes, Table S4 in Supporting Information S1 for amphiboles and Table S5 in Supporting Information S1 for iron oxides) and described below in the Result section.

## 4.2. Zircon LA-ICP-MS U-Th-Pb Dating

Separated zircon grains from a total of 12 rock samples (Table 1) were obtained by standard crushing and heavy liquid and magnetic separation using Frantz isodynamic techniques. The grains were handpicked under a binocular microscope and then mounted in epoxy resin discs and polished to expose mid-grain sections. Zircon crystals were also analyzed in situ on thin section. Cathodoluminescence (CL) and back-scattered electrons (BSE) images were used to select points for analysis. U-Th-Pb isotope data on the zircons were obtained at the Laboratoire Magmas & Volcans (LMV), Clermont-Ferrand, France by Laser-Ablation Inductively Coupled Plasma Mass-Spectrometry (LA-ICP-MS). The analytical method for isotope dating with LA-ICP-MS is similar to that reported by Paquette and Tiepolo (2007) and Hurai et al. (2010). The analyses involved the ablation of minerals with a Resonetics M-50 laser system operating at a wavelength of 193 nm (Müller et al., 2009). The ablated material was carried in helium and then mixed with nitrogen and argon before injection into the plasma source of a Thermo Element XR Sector Field high-resolution ICP-MS. The alignment of the instrument and mass calibration were performed before every analytical session using the NIST SRM 612 references glass by inspecting the signal of ²³⁸U and minimizing the ThO⁺/Th⁺ ratio (<1%). The signals of ²⁰⁴(Pb + Hg), ²⁰⁶Pb, ²⁰⁷Pb, ²⁰⁸Pb, ²³²Th and ²³⁸U masses were acquired. The ²³⁵U signal was calculated from ²³⁸U on the basis of the ratio ²³⁸U/²³⁵U = 137.88. Single analyses consisted of 30 s of background integration with the laser off, followed by 60 s integration with the laser firing and a 20 s delay to wash out the previous sample and prepare for the next analysis. The occurrence of common Pb in the sample was monitored by the evolution of the ²⁰⁴(Pb + Hg) signal intensity, but due to the non-negligible mercury content of the Ar gas creating an isobaric interference between ²⁰⁴Pb and ²⁰⁴Hg, no common Pb correction was applied prior to data reduction using the GLITTER® software package (van Achterbergh et al., 2001). Raw data are corrected for U-Pb fractionation during laser sampling and for instrumental mass bias by standard bracketing with repeated measurements of a primary standard (GJ-1 zircon, Jackson et al., 2004). The reproducibility and accuracy of the corrections were controlled by repeated analyses of secondary standards (91,500 zircon, Wiedenbeck et al., 1995) treated as unknown. For each analysis, the time resolved signals of single isotopes and isotope ratios were monitored and carefully inspected to verify the presence of perturbations related to inclusions, fractures, mixing of different age domains or common Pb. The detailed analytical procedures are presented in Table S6 in Supporting Information S1. The calculated ratios were exported and ages and diagrams generated using the Isoplot/Ex v. 2.49 software package (Ludwig, 2001). The decay constants used for the U-Pb system are those determined by Jaffey et al. (1971) and recommended by the IUGS (Steiger & Jäger, 1977). In the text, tables and figures of U-Pb zircon geochronology, all uncertainties are given at 2σ level. The concentrations of U, Th and Pb were calibrated relative to the certified values for primary standards (Jackson et al., 2004). In general, only data with a concordance upper than ±95% are considered in the calculation of ages. Discordant data are excluded because of doubtful interpretation. Indeed, in laser ablation-ICP-MS analyses, several factors that cannot be easily detected from the inspection of the time-resolved signals might contribute to discordance (e.g., common Pb, losses Pb, mixing of different age domains, small cracks, inclusions). However, discordant data can only be taken into account in the linear regression calculation if they allow defining possible discordia lines on the U-Th-Pb diagrams and if one or both intercepts between the discordia line and the concordia curve have a geological meaning. All obtained U-Th-Pb data are reported in Table S7 in Supporting Information S1, and presented and summarized in Tera-Wasserburg diagrams and Table 3, and described below.

## 4.3. Bulk Rock Major and Trace Element Analyses

For whole-rock geochemical analysis of major and trace elements of a total 39 samples (Table 1), each rock sample was crushed and powdered using a jaw crusher and agate ring mill. Rock powders were analyzed for major elements using a Rigaku ZSX Primus IV wavelength dispersive X-ray fluorescence (WD-XRF) spectrometer from the GeoAnalytical laboratory (UM6P, Benguerir, Morocco) and XRF at the RRXG platform at the University of Montpellier (France). Loss on ignition (L.O.I) was determined by drying the samples at 900°C and is
<caption>**Table 3**</caption>

Summary of Geographic Location, Rock Types, and U-Th-Pb Geochronological Results for Samples From Studied Central High-Atlas Igneous Intrusive Structures

<table><thead><tr><th>Location/Rock type</th><th>Sample</th><th>G.Coor. (WGS84)</th><th>U-Th-Pb age (Ma)</th><th></th></tr></thead><tbody><tr><td colspan="5">Mafic unit rocks</td></tr><tr><td rowspan="2">Tassent/Cumulate troctolite gabbro</td><td rowspan="2">AP01</td><td rowspan="2">32.21/-5.69</td><td>Zircon population 1</td><td>158.6 ± 1.3</td></tr><tr><td>Zircon population 2</td><td>153.1 ± 1.5</td></tr><tr><td rowspan="2">Anefgou/Hydrothermally altered gabbro</td><td rowspan="2">IN22</td><td rowspan="2">32.27/-5.41</td><td>Zircon population 1</td><td>162.8 ± 1.4</td></tr><tr><td>Zircon population 2</td><td>152.0 ± 2.1</td></tr><tr><td>Tirrhist/Monzogabbro</td><td>IN2</td><td>32.19/-5.46</td><td></td><td>160.8 ± 0.7</td></tr><tr><td colspan="5">Felsic unit rocks</td></tr><tr><td>Tirrhist/Monzonitic rock</td><td>IN4</td><td>32.19/-5.46</td><td></td><td>154.6 ± 0.7</td></tr><tr><td>Tirrhist/Syenite</td><td>IN20</td><td>32.21/-5.43</td><td>Zircon population 1</td><td>156.0 ± 1.1</td></tr><tr><td></td><td></td><td></td><td>Zircon population 2</td><td>149 ± 2</td></tr><tr><td>Tirrhist/Syenite</td><td>AP10</td><td>32.2/-5.45</td><td></td><td>157.2 ± 0.9</td></tr><tr><td>Tirrhist/Syenite pegmatite</td><td>IN44</td><td>32.20/-5.45</td><td></td><td>158.8 ± 0.9</td></tr><tr><td>Anefgou/Syenite</td><td>IN23</td><td>32.27/-5.43</td><td></td><td>152.1 ± 0.8</td></tr><tr><td>Anefgou/Syenite</td><td>IN21</td><td>32.27/-5.41</td><td></td><td>159.0 ± 1.3</td></tr><tr><td>Anemzi/Syenite</td><td>AN1</td><td>32.3/-5.27</td><td>Zircon population 1</td><td>160.6 ± 1.7</td></tr><tr><td></td><td></td><td></td><td>Zircon population 2</td><td>151.6 ± 1.4</td></tr><tr><td>Anemzi/Syenite</td><td>AN2</td><td>32.28/-5.27</td><td></td><td>159.6 ± 0.8</td></tr><tr><td>Anemzi/Syenite</td><td>AN5</td><td>32.29/-5.26</td><td></td><td>159 ± 1</td></tr></tbody></table>

generally low to moderate in the felsic and mafic unit rocks (0.04–3.4 wt.%, except one sample AP38 with L.O.I ~8.4 wt.%, linked to intense secondary calcite amygdales).

For whole-rock trace element analyses, samples were analyzed using an Agilent 7700x ICPMS at the AETE Platform (OSU-OREME, Montpellier, France). Prior to analyses, 100 mg of rock powder went through a preparation protocol consisting of two acid digestions at 110°C for 24 hr each (1 ml HNO₃ + 2.5 ml HF and 0.5 ml HNO₃ + 1.5 ml HCl) and three evaporation to dryness at increasing temperatures from 150°C to 190°C (0.5 ml HClO₄, 0.5 ml HNO₃, and 0.25 ml HNO₃). The obtained solid residues were subsequently diluted before analysis by adding 2 ml of Milli-Q water and then 1.5 ml of ultrapure HNO₃. In addition to the studied rock samples, international standard samples (UB-N-1, BHVO-1, UB-N-2, BHVO-2, MAN, MRG1, and BIR) were also prepared to control the accuracy of the analyses. The obtained whole-rock major and trace element data are reported in Table S8 in Supporting Information S1 and described below.

## 4.4. Bulk Rock Sr-Nd Isotopic Analyses

For Sr-Nd radiogenic isotopic analyses of selected samples (Table 1), about 100 mg of rock powder from each sample was weighed in a Teflon beaker and dissolved in 1:1 a mixture of HF/HNO₃. After dissolution, the samples were diluted in 1 ml of 2% HNO₃ and Nd-Sr was extracted from the matrix (2N HNO₃) using a combination of Sr-Spec and Thru-Spec Eichrom resins. Mixed Sr and rare earth elements (REE) were loaded onto a Re filament and were run sequentially (first Sr then Nd) using a double Re filament protocol. Monitoring of ⁸⁷Rb and ¹⁴⁴Sm interferences was performed according to the protocol of Li et al. (2012) and the quality and reproducibility of the measurements were controlled using a sequential measurement of isotopic standards (SRM 987 and JNdI-1), doped isotopic standards (NBS 987+ Rb and JNdI + Sm) and laboratory-dedicated Sr + REE artificial solutions. Sr-Nd isotopic chemistry and data acquisition were performed at GET-OMP, University of Toulouse (France), using the Thermo Scientific TRITON + solid source mass spectrometer (TIMS, procedures are detailed in Li et al. (2011, 2012)). Standard reproducibilities are for SRM-987 (pure and doped) and fall within the recommended values. Measured blanks are for Nd and for Sr. ⁸⁷Sr/⁸⁶Sr and ¹⁴³Nd/¹⁴⁴Nd ratios were normalized to ⁸⁶Sr/⁸⁸Sr = 0.1194 and ¹⁴⁶Nd/¹⁴⁴Nd = 0.7219, respectively, after corrections from isobaric
interferences using $^{87}$Rb/$^{85}$Sr = 0.387041 for $^{87}$Sr and combination of $^{147}$Sm/$^{149}$Sm = 1.08583 and $^{147}$Sm/$^{144}$Sm = 4.87090 for $^{144}$Nd. Results of Sr-Nd radiogenic isotopic analyses are reported in Table S9 in Supporting Information S1, and described below.

# 5. Results

## 5.1. Mineral Compositions

### 5.1.1. Olivines

Olivine is the common mineral in all mafic unit rocks of the studied Central High-Atlas intrusive massifs, including cumulate troctolite gabbro, troctolite gabbro and monzogabbro samples. All analyses of olivine grains from analyzed cumulate troctolite gabbro exhibit homogenous concentrations of SiO$_2$ (~37–39 wt.%), FeO (~22–28 wt.%), and MgO (~35–40 wt.%). They have MgO-rich affinities with forsterite (Fo) pole on about ~Fo$_{70-77}$. In some olivine grains, they show slight chemical zoning from core to rims, with forsterite percentage range from ~Fo$_{70}$ in the rims, and about ~Fo$_{73}$ in the core. In troctolite gabbro, olivine is characterized by a decrease in the MgO content (~26–36 wt.%), and an increase in FeO (25–37 wt.%), with variable Fo values ranging from 56% to 72%. On the other hand, olivine grains from monzogabbros are Fe-rich olivine, consisting of high FeO (~48–50 wt.%), and low MgO (~16–18 wt.%) with Fo values of about 26%–40%. In addition to the decline in Fo values from primitive gabbros (cumulate troctolite and troctolite gabbros) to evolved gabbros (monzogabbros), we also observed a slight increase in CaO and MnO concentrations, respectively, from 0.1 to 0.4 wt.% and 0.3 to 1.1 wt.%.

### 5.1.2. Feldspars

Feldspars are the predominant minerals found in all rocks from both mafic and felsic units. They are mainly feldspar plagioclase compositions in mafic unit rocks. In cumulate troctolite gabbros, they span a slight to moderate range of chemical compositions and can be classified as labradorite ($An_{52-71}Ab_{28-47}$). Plagioclases in troctolite gabbros are more sodic than those in the cumulate troctolite gabbros. They vary mainly between labradorite and andesine plagioclases ($An_{32-63}Ab_{36-64}$) with some oligoclase rims ($An_{11-27}Ab_{71-85}$). In monzogabbros, the plagioclases have more evolved chemical compositions with andesine ($An_{37-47}Ab_{51-60}$) to oligoclase ($An_{12-28}Ab_{69-80}$) affinities. In hydrothermally altered gabbros, they confirm a strong destabilization of the rock with a zoning from Ca-cores with labradorite compositions ($An_{50-56}Ab_{43-49}$), and Na-rich rims with andesine to albite affinities ($An_{8-37}Ab_{63-89}$). On the other hand, the chemical compositions of feldspars are more evolved and felsic in the monzonitic (monzodiorite-monzonite) and syenitic rock association of the felsic unit of the Central High-Atlas intrusive massifs, confirming the petrographic observations and showing that they are Si- and Na-rich. They mainly display albite compositions $Ab_{92-100}$. Some K feldspar grains are found in these rocks with compositions of about $Ab_{2-58}Or_{39-98}$.

### 5.1.3. Clinopyroxenes

Clinopyroxenes are also one of the common major minerals in the studied mafic and felsic unit rocks (gabbros and syenites). They have quite similar compositions in all gabbroic rocks with mainly augite to diopside compositions ($Wo_{43-49}$, $En_{34-43}$, $Fs_{12-18}$), with a slight decrease in Mg from cumulative troctolite gabbros to monzogabbros (Mg# ~0.78 for cumulate troctolite gabbros and ~0.5–0.8 in monzogabbros). Compared to gabbroic rocks, the syenite clinopyroxenes have variable compositions from diopside to hedenbergite ($Wo_{46-49}$, $En_{16-38}$, $Fs_{13-36}$).

### 5.1.4. Amphiboles

Amphiboles have been observed in felsic unit rocks and the obtained compositions from studied monzonitic and syenitic rocks show variable Ca content (~1.3–1.9 atoms per formula unit = a.p.f.u), Si (~7–7.8 a.p.f.u) and Mg (Mg#: Mg/(Mg + Fe) Mg# ~0.2 to 0.9), low to moderate (Na + K) values (0.1–1 a.p.f.u) and low Al concentrations (~0.01–1 a.p.f.u), characteristics of calcic amphiboles (e.g., Leake et al., 1997) with high Fe (mostly ~1–4 a.p.f.u) and slight high values of Na (up to 0.9 a.p.f.u) are mainly observed in amphiboles of syenitic rocks.

### 5.1.5. Iron Oxides

Iron oxides are the common opaque minerals from mafic and felsic rocks of the studied Central High-Atlas intrusive massifs. In mafic lithological facies (cumulate troctolite gabbro, troctolite gabbros, and
monzogabbros), iron oxides are ilmenite with FeO of about ~43–49 wt.% and ~48–53 wt.% TiO₂ with high values of MgO (up to ~1–6 wt.%). In felsic rock facies, ilmenite displays similar FeO and TiO₂ concentrations but with high MnO (up to ~3–5 wt.%) and low MgO (<0.03 wt.%). On the other hand, magnetite crystals are found in the felsic unit facies (monzonitic and syenitic rocks) with FeO content ranging between 84 wt.% and 92 wt.% and variable TiO₂ concentrations (~0.3–8 wt.%).

## 5.2. Zircon U-Th-Pb Ages

In this section, we present the results of LA-ICP-MS dating of zircons for 12 samples from different lithologies from Central High Atlas intrusive mafic and felsic units that constrain the Late Jurassic age of this magmatism (Table 3; Figures 6–9).

### 5.2.1. Mafic Unit Rocks

#### 5.2.1.1. Tassent Cumulate Gabbro AP01

A total of 7 zircon crystals or fragments were recovered from 8 kg of crushed samples of Tassent cumulate gabbro. Most of these crystals were sub-euhedral, transparent, and colorless. They exhibited CL-dark zones and well-developed concentric oscillatory and sector (hourglass) zoning typical of igneous zircons (e.g., Figure 6; Corfu et al., 2003; Hanchar & Miller, 1993; Hoskin & Schaltegger, 2003; Pidgeon, 1992). No evidence of inherited cores was observed. Fourteen analyses were conducted on all these zircon grains, one in the central zone of the grain and another on the rim (e.g., Table S7 in Supporting Information S1). Zircons display a high degree of variability in their concentrations, with values ranging from low to moderate for Pb (14–444 µg/g), moderate to

<figure><img src="image_8.png" /><figcaption>**Figure 6.** Cathodoluminescence images of zircons separated from the studied rock samples within the mafic and felsic units of the Central High-Atlas intrusive massifs.</figcaption></figure>
<figure><img src="image_9.png" /><figcaption>**Figure 7.** Tera-Wasserburg diagrams obtained on zircon of the mafic unit rocks of the Central High-Atlas intrusive massifs. Dotted ellipses are not taken into account for the age calculation and the histogram as they are discordant (see details in the text), while white ellipses that are concordant or sub-concordant are included in the histogram but not in the concordia age calculation. MSWD(C + E) is the mean square of weighted deviates for concordance and equivalence.</figcaption></figure>

high Th (~500 µg/g to ~2.3%), and U (~455 µg/g to ~1.1%), as well as high values for the Th/U ratios (~1–4). Plotted in the Tera-Wasserburg diagram, a single data point is discordant (spot#13, dotted ellipse), probably due to the presence of common Pb. The remaining 13 ellipses exhibited a range of dates between ca. 150 Ma and 160 Ma along the concordia curve (Figure 7a). The binned frequency histogram, coupled with probability, reveals that the data are distributed into two relatively similar-sized populations, with a peak around 158 Ma and a second around 153 Ma. The oldest population is predominantly composed of cores (4 cores and 2 rims) and yields a concordia age of 158.6 ± 1.3 Ma (MSWD(C+E) = 0.81; N = 6), while the youngest population is mainly comprised of rims (6 rims and 1 core) and gives a concordia age of 153.1 ± 1.5 Ma (MSWD(C+E) = 1.2; N = 7).

### 5.2.1.2. Anefgou Troctolite Gabbro IN22

A total of 14 zircon crystals were extracted from 8 kg of troctolite gabbro from the Anefgou massif. The crystals were anhedral, transparent and colorless to pink with a size range of ~80 to ~200 µm. The CL-images display either a dark pattern without internal structure or exhibit extremely patchy zoning, the latter being largely attributable to the presence of numerous inclusions (e.g., Figure 6). No evidence of inherited cores was observed. These zircons generally have highly variable concentrations of Pb 2–315 µg/g (most ca. 20–215 µg/g), Th of 6 µg/g–1.7% (most ca. 0.1–1.7%) and U of 95–9,690 µg/g–1.7% (most ca. 0.05%–1%), and with a high Th/U mainly between 0.67 and 3.68 (Table S7 in Supporting Information S1). As shown in the Tera-Wasserburg diagram, the 20 ellipses are located between 148 Ma and 165 Ma (Figure 7). Among the aforementioned ellipses, four (#3, 7, 10, 13) are evidently discordant (14%–17%), likely due to common Pb contamination. Consequently, these will not be considered in subsequent discussions. The 16 analyses are plotted in histograms, showing two populations of equivalent size, separated by a range of approximately 10 million years at ca. 162 and 152 Ma. The weighted mean age of the oldest population, calculated from the ²³⁸U/²⁰⁶Pb ratio, is 163.1 ± 1.2 Ma (MSWD(C+E) = 0.68; N = 7). Due to the slight discordance (5%–6%) of ellipses 8 and 9, likely due to low common Pb contamination, these two ellipses were excluded from the calculation of the concordia age, which is 162.8 ± 1.4 Ma (MSWD(C+E) = 0.81; N = 5). These two dates are equivalent, and it is therefore proposed that the best estimate is the concordia age. In the case of the youngest population, the seven ellipses are slightly more dispersed, resulting in a weighted average of 152.0 ± 2.1 Ma (MSWD = 2.4; N = 7). Ellipses 6 and 16 are presumably the consequence of mixing among these populations or of a minor common Pb contamination and/or a radiogenic Pb loss.

### 5.2.1.3. Tirrhist Monzogabbro IN2

For the Tirrhist monzogabbro, the zircon grains that have been separated from approximately 7.7 kg of sample material are essentially euhedral to subhedral with width-to-length ratios of about 1:2. They are transparent and colorless to opaque red, indicating their metamict nature. Dark CL-images show the presence of concentric oscillatory zoning typical of magmatism. Additionally, the CL-images exhibit the presence of fractures and inclusions, yet no evidence of an inherited core was observed (Figure 6). A total of 26 analyses were conducted on the 16 zircon grains, either at the center of the crystal or at the rim (Table S7 in Supporting Information S1). Concentrations of Pb ranged
from low to moderate (3–89 µg/g), while those of Th, and U exhibited a broader range, from low to high (ca. 44 µg/g to 0.4% and ca. 120 µg/g to 0.3%, respectively). In addition, the Th/U ratio spanned a wide range from 0.37 to 1.9. With the exception of one analysis (#22), all the data were concordant to sub-concordant and spread along the concordia curve around 174 Ma and 152 Ma (Figure 7). If we exclude the discordant datum (#22), which is likely to be contaminated by common Pb, and the datum (#12) that gives a ²³⁸U/²⁰⁶Pb date that is slightly older than the others at 170.6 ± 3.3 Ma, the weighted mean is 159.8 ± 1.1 Ma (MSWD = 3.2; N = 24). The 20 data give a concordia age of 160.8 ± 0.7 Ma (MSWD(C+E) = 1.3; N = 20). The four ellipses with the youngest ²³⁸U/²⁰⁶Pb dates (ca. 156–154 Ma) were excluded from the concordia age calculation because they are likely to have experienced a slight loss of Pb; the best estimate of the date is the concordia age at 160.8 ± 0.7 Ma.

### 5.2.2. Felsic Unit Rocks

#### 5.2.2.1. Tirrhist Monzodiorite IN4

A significant quantity of zircon was obtained from 5.5 kg of sample of Tirrhist monzodiorite. The grains are either fragments or sub-euhedral in shape with width-to-length ratios of about 1:1 with a size range of ~80 µm to ~200 µm. They are transparent, colorless to opaque, and red brown, indicating their metamict nature. Moreover, the crystals contain numerous fractures and inclusions. Dark CL-imaging occasionally reveals oscillatory zoning, which is characteristic of magmatism. However, no inherited core is evident (e.g., Figure 6). A total of 23 analyses were carried out on the 22 zircon crystals (e.g., Table S7 in Supporting Information S1), both at the core and at the rim (e.g., Table S7 in Supporting Information S1). The Pb concentrations ranged from low to moderate (2–176 µg/g), while those of Th, and U exhibited a broader range, from low to high (ca. 34 µg/g to ~0.9%, 87 µg/g to ~0.4%, respectively). Additionally, the Th/U ratio spanned a wide range, from 0.3 to 2. With the exception of one analysis (#15), which is likely to be contaminated by common Pb, all the data were concordant or sub-concordant and distributed along the concordia curve at around 172 and 149 Ma (Figure 8). If we exclude the discordant data point (#15), and the three data points (#5, 21 and 23) that give slightly older ²³⁸U/²⁰⁶Pb dates at ca. 168–160 Ma, which may reflect the presence of an inherited component, the remaining data points yield a concordia age of 154.6 ± 0.7 Ma (MSWD(C + E) = 1.19; N = 19).

#### 5.2.2.2. Tirrhist Syenite IN20

In 5.8 kg of Tirrhist syenite IN20, we separated 20 sub-euhedral grains of zircon: sizes ~100–250 µm, metamict, transparent pink to opaque pink. They show patchy zoning, which is disturbed by the presence of numerous inclusions (Figure 6). These zircons have generally concentrations 9–64 µg/g Pb, ~430 µg/g–0.5% Th, and ~268 µg/g–0.1% U with and moderate to high Th/U (1–4). All U-Pb isotope analyses are plotted in the Concordia diagram with an age of 153 ± 2 Ma with a very high unreasonable MSWD of ~9, clearly reflecting the presence also of discordant analyses (e.g., spots 3, 12, and 13 with discordance up to 9%) and a spread and old age for the concordant/near concordant analyses 11/17 analyses with <3% discordant with and mean ²³⁸U/²⁰⁶Pb date of ca. 155 ± 1 Ma with low MSWD of 0.97 and 6/17 generate Concordia age = 156 ± 1 Ma and high MSWD ~0.98 (e.g., Figure 8). The remaining 6/17 near concordant/concordant spots in 6 zircons (1,2,5, 9, 16, and 19) yield a younger mean ²³⁸U/²⁰⁶Pb date of ca. 149 ± 2 Ma with low MSWD of 2.5 of which similar 4/11 (6, 9, 16, and 18) Concordia age of about c. 149 ± 2 Ma with MSWD = 1.8 (Figure 8). Within error are these Late Jurassic Oxfordian dates (c. 147–157 Ma), which we interpret as the best estimate of the crystallization age of Tirrhist syenite IN20.

#### 5.2.2.3. Tirrhist Syenite AP10

The zircon grains that were separated from approximately 5.3 kg of the Tirrhist syenite sample material are mainly fragments, although a few euhedral crystals are also present. They are transparent to colorless, ranging in size from 100 to 300 µm. Zircon grains show either dark CL-images or clear oscillatory zoning (Figure 6). A total of 26 analyses were carried out on the 19 zircon grains, either at the crystal center or at the rim (Table S7 in Supporting Information S1). The Pb concentrations ranged from low to moderate (3–167 µg/g), while Th, and U contents exhibited a wider range, from low to high ca. 31 µg/g to ~1% and ca. 73 µg/g to ~0.4%, respectively. Furthermore, the Th/U ratios spanned a large range, from 0.32 to 2.78, with the exception of analysis #26 where Th/U = 0.08. With the exception of analysis #2, all the data are concordant to sub-concordant. They are
<figure><img src="image_10.png" /><figcaption>**Figure 8.** Tera-Wasserburg diagrams obtained on zircon of the felsic rocks of the Anefgou and Tirrhist massifs (Central High-Atlas). Dotted or white ellipses are not taken into account for the age calculation (see details in the text), while gray and dotted ellipses that are concordant or sub-concordant are included in the weighted average but not in the concordia age calculation. MSWD(C + E) is the mean square of weighted deviates for concordance and equivalence.</figcaption></figure>

positioned on the concordia curve between ca. 166 Ma and ca. 150 Ma (Figure 8). The weighted mean of $^{238}$U/$^{206}$Pb dates is 155.5 ± 0.9 Ma (MSWD = 1.9; N = 24/25) rejecting data #6, which has a slightly older date at 162.9 ± 3.1 Ma. Of these 24 data, 20 give a concordia age of 157.2 ± 0.9 Ma (MSWD(C+E) = 1.6). The best estimate of the date is the concordia age.
5.2.2.4. *Tirrhist Syenite Pegmatite IN44*

Zircon megacrysts up to ~0.8 cm in size were separated from syenite pegmatite associated with syenite outcrops of the syenite AP10. These crystals are more or less transparent to opaque with some displaying significant metamict alteration. They contain clinopyroxene inclusions. In CL imaging, sometimes these megacrysts exhibit slight oscillatory zoning, and no evidence of inherited cores (Figure 6). Twenty-one analyses were carried out on five crystals, and exhibit variable concentrations of Pb (~2–100 µg/g), Th (~100 µg/g to ~0.5%), and U (~58 µg/g to ~0.3%), with moderate Th/U (~1–4.5) (Table S7 in Supporting Information S1). With the exception of data #1, which is highly discordant (not shown in Figure 8 as it is out of frame), the remaining data are distributed along or in close proximity to the concordia curve between ca. 166 Ma and ca. 143 Ma in the Tera-Wasserburg diagram (Figure 8). A concordia age of 158.8 ± 0.9 Ma (MSWD<sub>(C+E)</sub> = 1.3; N = 14) was obtained. The four discordant data (#2, 4, 11, 12), and the two data (#14, 16) with the youngest <sup>238</sup>U/<sup>206</sup>Pb dates are not taken into account in this calculation due to a probable slight contamination in the common Pb in the first case, and a probable loss of radiogenic Pb in the second.

5.2.2.5. *Anefgou Syenite IN23*

Numerous zircon grains were found in 3.8 kg of Anefgou syenite IN23 that were crushed. Most of these crystals are either broken or sub-euhedral with a long stick shape with width-to-length ratios of about 1:4 to 1:7. They are transparent to opaque, indicating metamict alteration, and pinkish in color. CL imaging shows clear oscillatory zoning (Figure 6). A total of 19 analyses were conducted on 17 zircons (Table S7 in Supporting Information S1). These zircons generally have low concentrations of Pb 3–64 µg/g (most ca. 3–26 µg/g), high variable concentration of Th (124 µg/g to 0.7%), and U (75 to 589 µg/g). Additionally, Th/U ratios are high and vary widely from 0.9 to 28. All the data form a cluster between ca. 163 Ma and ca. 144 Ma (Figure 8). Except for data #1, which has a slightly older <sup>238</sup>U/<sup>206</sup>Pb date at ca. 160 Ma, the set data yield a weighted mean date of 151.4 ± 1.0 Ma (MSWD = 1.9; N = 18). Among these data, four data (#2, 10, 14, 16) are slightly discordant due a probably common Pb contamination, and the remaining data give a concordia age of 152.1 ± 0.8 Ma (MSWD<sub>(C+E)</sub> = 1.4; N = 14). The best estimate of the date is the concordia age.

5.2.2.6. *Anefgou Syenite IN21*

In syenite Anefgou IN21, the zircons grains are either fragments or sub-euhedral in shape, squat with width-to-length ratios of about 1:1 to 1:2. They are transparent, colorless to opaque, and red brown, indicating their metamict nature. Dark CL imaging exhibits clear oscillatory zoning and the occurrence of inclusions and fractures (Figure 6). A total of 17 analyses were conducted in 12 zircon crystals (Table S7 in Supporting Information S1), either at the center of the crystal or at the rim. The concentration of Pb ranges from low to moderate (31–124 µg/g), while those of Th and U exhibit large variations ranging from ca. 44 µg/g to 0.4% and ca. 120 µg/g to 0.3%, respectively. Moreover, the Th/U ratio spanned a wide range, ranging from 0.44 to 2. As shown in the Tera-Wasserburg diagram, the 17 ellipses are located between 143 and 175 Ma (Figure 8). Among the aforementioned ellipses, six (#3, 4, 5, 9, 10, 17) are evidently discordant (>4%), likely due to common Pb contamination and Pb losses (#4, 5, 9). Consequently, these will not be considered in subsequent discussions. The three ellipses (# 8, 13, 16) with the oldest mean <sup>238</sup>U/<sup>206</sup>Pb dates (ca. 163–171 Ma) were excluded from the concordia age calculation because they are likely to have experienced a slight loss of Pb. The 8 data given a concordia age of 159.0 ± 1.3 Ma (MSWD<sub>(C+E)</sub> = 1.3; N = 8). The best estimate of the date is the concordia age at 159.0 ± 1.3 Ma.

5.2.2.7. *Anemzi Syenite AN1*

The majority of the zircon grains extracted from 5.4 kg of the Anemzi syenite are euhedral, colorless, transparent to white, and opaque. Such grains are frequently metamict. CL imaging shows an oscillatory zoning that is strongly disrupted by the presence of inclusions (Figure 6). A total of 21 analyses were conducted on 19 zircon grains (e.g., Table S7 in Supporting Information S1). The concentrations of Pb (7–63 µg/g), Th (43 µg/g–0.3%), and U (321 µg/g–0.2%) exhibited variability, ranging from low to moderate. Th/U ratios were observed to span a range of 0.13–1.61, with the majority failing within 0.37–1.61. Of these analyses, five (#3, 4, 10, 12, 20) are discordant, probably due to common Pb contamination, one (#17) is concordant around 180 Ma and the remaining 15 ellipses are located between about 167 and 144 Ma (Figure 9). The binned frequency histogram, coupled with a probability-density distribution, shows that the data are distributed into two main populations, with a peak around
<figure><img src="image_11.png" /><figcaption>**Figure 9.** Tera-Wasserburg diagrams obtained on zircons of the felsic rocks of the Anemzi massif (Central High-Atlas). Dotted ellipses are not taken into account for the age calculation and the histogram as they are discordant (see details in the text). MSWD(C + E) is the mean square of weighted deviates for concordance and equivalence.</figcaption></figure>

151 Ma and a second around 160 Ma. The four oldest data (#1, 2, 5, 21, shown by gray ellipses in Figure 9) yield a concordia age of 160.6 ± 1.7 Ma (MSWD(C+E) = 1.03), while the youngest population (#6, 7, 8, 9, 11, 13, 14, 15, 16,18 and 19, shown by green ellipses in Figure 9) gives weighted mean ²³⁸U/²⁰⁶Pb of 151.0 ± 1.6 Ma (MSWD = 2.1; N = 11), of which five (#6, 9, 14, 16, 18) yield a concordia age of 151.6 ± 1.4 Ma (MSWD(C+E) = 0.99; N = 5). These two dates are similar within the error bars, and it is therefore proposed that the best estimate is the concordia age.

### 5.2.2.8. Anemzi Syenite AN2

For the Anemzi syenite AN2, the zircon grains that have been separated from approximately 5.5 kg of sample rock material are essentially euhedral, pinkish transparent to red opaque, with width-to-length ratios of about 1:2. They contain inclusions (e.g., Figure 6). They exhibited CL-dark zones and well-developed concentric oscillatory zoning and sector (hourglass) zoning, which are typical of magmatism zircon (Hoskin & Schaltegger, 2003). Twenty-five analyses were performed on 17 zircons, either at the center of the crystal or at the rim (Table S7 in Supporting Information S1). The Pb, Th, and U contents are very heterogeneous and are between 19 and 88 µg/g, 617 µg/g to ~0.4% and 617 µg/g to ~0.3%, respectively, with Th/U ratios ranging from 0.8 to 1.7. With the exception of two discordant data (#20, 21), which are likely the result of common Pb contamination, all the data form a concordant cluster that ranges between ca. 165 Ma and 152 Ma (Figure 9). The weighted mean is 159.6 ± 0.8 Ma (MSWD = 1.3; N = 23). No concordia age could be calculated.

### 5.2.2.9. Anemzi Syenite AN5

The majority of the extracted zircons, comprising 5.8 kg of Anemzi syenite AN5 are sub-euhedral to ovoid in shape. It is notable that a significant proportion of these grains are fragments of zircon. They are colorless, transparent to opaque white, with a substantial proportion being metamict. The dark CL images do not display the internal structures (zonation and inherited core) but the occurrence of a few inclusions (e.g., Figure 6). A total of 28 analyses were performed on 26 crystals (Table S7 in Supporting Information S1). The Pb concentrations exhibited a relatively narrow range, spanning from low to moderate (4–133 µg/g). In contrast, the Th, and U contents demonstrated a broader variation, encompassing levels from low to high (ca. 56 µg/g to ~0.6% and ca. 156 µg/g to ~0.3%, respectively). Additionally, the Th/U ratio exhibited a high range (~0.33–2.03). As shown in the Tera-Wasserburg diagram, the 24 ellipses are located between 154 Ma and 166 Ma (Figures 9c, Table 3). Except for the discordant data #1, which are likely the result of common Pb contamination, all the 27 data form a cluster around ca. 160 Ma (Figure 9) whose three data (#3, 11, 25) are slightly discordant. The remaining 24 analyses yield a concordia age of 159 ± 1 Ma (MSWD(C+E) = 1.09).

## 5.3. Whole-Rock Geochemistry

### 5.3.1. Major and Trace Elements

Major and trace element compositions of 39 samples of studied Jurassic mafic and felsic intrusive magmatism of the Central High-Atlas are listed in Table S8 in Supporting Information S1. To have more general information on this magmatism, all available literature data from the magmatism of the same studied intrusive structures are also used in different geochemical
<figure><img src="image_12.png" /><figcaption>**Figure 10.** Geochemical classification of the igneous lithological units of the studied Central High-Atlas intrusive massifs. (a) Na₂O + K₂O versus SiO₂ (TAS) classification diagram (Middlemost, 1994). (b) Zr/Ti versus Nb/Y (Winchester & Floyd, 1977). Ocean island basalts (OIB) (Sun & McDonough, 1989), Bulk continental crust (CC) (Rudnick & Gao, 2003), BAB (back-arc basin basalts) and MORB (Gale et al., 2013). (c) Na₂O versus. K₂O diagram (fields from Middlemost, 1975). (d) (K₂O + Na₂O) versus (K₂O/(K₂O + Na₂O))*100 plot (after Hughes, 1973). Mafic unit (Green Field) and felsic unit (Pink Field).</figcaption></figure>

classification diagrams (Figures 10 and 11; data from Armando (1999), Lhachmi et al. (2001), Zayane et al. (2002), Essaifi and Zayane (2018), Ouabid, Raji, et al. (2021) and Ahechach et al. (2024)). Major element compositions of the mafic and felsic units show SiO₂ content variation between ~41 wt.% and ~65 wt.%, and total alkalis (Na₂O + K₂O)% between ~1.5 wt.% and ~15 wt.%. In the Total Alkali Silica diagram (TAS) of Middlemost (1994) (Figure 10a), the samples draw an almost complete igneous plutonic series from the mafic to felsic members. They classify in the gabbroic rock fields for the mafic unit rock samples: peridot gabbro/gabbro field for mainly troctolitic rocks, gabbro–monzogabbro field for hydrothermally altered gabbro samples and monzogabbro field for monzogabbro samples. A few hydrothermally altered and troctolite gabbro samples fall in the foid gabbro field. On the other hand, felsic unit rocks plot mainly as monzodiorite–monzonite series for the studied samples of monzonitic rocks (monzodiorite and monzonite) and syenite (dominant)–quartz monzonite series for syenitic samples (Figure 10a). All mafic and felsic lithological units from the different massifs (including this study and those from the literature) are similar, plotting in the alkaline igneous series (Figure 10a) with slightly high values of Nb/Y (mostly felsic rocks, e.g., Figure 10b; Winchester & Floyd, 1977) and mostly sodic features (Figure 10c). Most samples are metaluminous (mol. Al₂O₃/CaO + Na₂O + K₂O (=ACNK) < 1 and Al₂O₃/Na₂O + K₂O (=ANK) > 1) for mafic rocks and felsic rocks with slightly peraluminous and peralkaline affinities for few felsic samples (ACNK > 1 and ANK < 1) (Table S8 in Supporting Information S1; e.g., Maniar & Piccoli, 1989). The assessment of the alteration effects using the discrimination diagram of Hughes (1973) in Figure 10d reveals that most mafic rocks belong to fresh-to-weakly altered igneous rock series, except for some hydrothermally altered gabbro samples plotted in Na- and K-metasomatism fields. On the other
<figure><img src="image_13.png" /><figcaption>**Figure 11.** Major and trace elements of studied Jurassic igneous intrusive mafic and felsic unit rocks from the Central High-Atlas. Literature data from Armando (1999), Lhachmi et al. (2001), Zayane et al. (2002), Essaifi and Zayane. (2018), Ouabid, Raji, et al. (2021) and Ahechach et al. (2024).</figcaption></figure>

hand, felsic unit rocks (monzonitic and syenitic rocks; Na₂O up to ~10–15 wt.%) plot mostly in the Na-metasomatism domain (≥80% of all analyses).

In the Harker variation diagrams (Figure 11), the High-Atlas Jurassic intrusive magmatism shows linear trends and progressive compositional variations, from the more primitive mafic unit rocks (gabbroic rocks) to the more evolved felsic rocks (syenitic rocks). Data show good to rough negative correlations between SiO₂ and TiO₂, MnO, Mg#, Ni, Cr, good positive correlations between SiO₂ and Na₂O, Zr, and Yb, and rough positive correlations between SiO₂ and Al₂O₃ and K₂O and Nb. Almost all analyses of the studied magmatism show fluctuate compositions from low-K series to shoshonitic series either for mafic or felsic unit rocks (e.g., Rickwood, 1989).

Figure 12 shows the REE concentrations normalized to chondrite abundances of McDonough and Sun (1995) and the multi-element spider diagrams normalized to the primitive mantle (PM) of Sun and McDonough (1989) of the mafic and felsic lithological facies of the major Central High Atlas Jurassic intrusive massifs. We have also shown the compositions of ocean island basalts (OIB) (Sun & McDonough, 1989) and bulk continental crust (CC) (Rudnick & Gao, 2003) for comparison purposes. In detail, all studied rock-forming mafic units (troctolitic rocks, hydrothermally altered gabbros, monzogabbros) have parallel and overlapped REE patterns with a slight variation in REE abundances (ΣREE ~ 37–200 ppm in cumulate troctolite gabbros, ~64–66 ppm in troctolite gabbros, ~60–100 ppm in monzogabbros and ~48–94 ppm hydrothermally altered gabbros). They generally show a high fractionation between light (L) and heavy (H) REEs (La/n/Lu_n ~ 5–8), and a moderate fractionation of LREEs (La/n/Sm_n ~ 2) and HREEs (Gd/n/Lu_n ~ 2–3), with negligible to weak negative Eu anomalies (Eu/Eu* = 0.7–1) (Figure 12).
<figure><img src="image_14.png" /><figcaption>**Figure 12.** (left) REE patterns normalized to chondrite values from McDonough and Sun (1995) and (right) spider diagrams normalized to the primitive mantle (PM) (Sun & McDonough, 1989) for the mafic and felsic magmatic rocks of the Central High-Atlas Jurassic alkaline intrusive structures.</figcaption></figure>

In PM normalized trace element diagrams (Figure 12; except one sample from cumulate troctolite gabbro with fluctuate concentrations), all studied mafic unit-forming rocks are characterized by uniform and parallel patterns with positive Pb anomaly (with slightly higher values in monzogabbros and hydrothermally altered gabbros), and negligible-to-weak Sr and Ti anomalies (mostly in monzogabbros and hydrothermally altered gabbro samples). They all have variable enrichment in K, Rb, Ba and Th and negligible depletion in Nb and Ta. Compared to OIB, the mafic samples have REE and multi-element patterns that are closely parallel and similar (with slight high values in OIB, Figure 12).

Similarly, the felsic unit-forming rocks (monzonitic and syenitic rocks) are also characterized by parallel REE patterns (Figure 12). They show variable REE concentrations (ΣREE ~20–209 ppm in monzonitic rocks, ΣREE ~28–252 ppm in syenitic rocks), variable fractionations between LREE and HREE ((La/Lu)~0.6–8), though most of this compositional variability corresponds to the syenitic rocks. Compared to mafic unit rocks, the felsic unit rocks are characterized by similar fractionation values of REE with ~4–7 La/Lu; ~2 La/Sm and ~2–3 Gd/Lu and ~0.6–1 Eu/Eu* mostly in monzonitic rocks, with somewhat fluctuating REE patterns in syenitic
rocks, on the other hand, have with ~0.6–8 La_n/Lu_n; ~1–3 La_n/Sm_n and ~0.5–3 Gd_n/Lu_n with high depletion in Eu (Eu/Eu* up to ~0.3). The multi-element spider diagrams of the monzonitic and syenitic rock patterns are also similar and parallel with high positive Pb anomaly (with slight fluctuating values in syenitic rocks), and negative Sr and Ti anomalies (mostly in syenites, and fluctuating values in monzonitic rocks). All have variable enrichment in Rb, Ba, Th, K, and P, and negligible Nb and Ta anomalies in most samples. Compared to OIB and CC, the studied Jurassic felsic unit-forming rock samples have REE and multi-element patterns that are closely parallel and similar to OIB in terms of element enrichment and CC in terms of element fluctuation.

### 5.3.2. Sr-Nd Isotopes

Sr-Nd isotopic data from mafic and felsic units of the studied Central High-Atlas Jurassic alkaline intrusive structures are listed in Table S9 in Supporting Information S1. Based on their Late Jurassic U-Pb zircon crystallization ages (Figures 7–9, Table 3), the initial Sr and Nd isotopic ratios were recalculated for a median age of 155 Ma using the parameters in DePaolo (1981), Goldstein and Jacobsen (1988), Faure (2001), and White (2014). The mafic unit rocks have relatively similar high ¹⁴⁵Nd/¹⁴⁴Nd(i) values of 0.51253–0.51256 with positive εNd(i) ~ 1.7–4.3 and low ⁸⁷Sr/⁸⁶Sr(i) (0.70384–0.70562), except one sample from hydrothermally altered gabbros, which yields a low Nd isotope signature (¹⁴⁵Nd/¹⁴⁴Nd(i) ~ 0.51202, and negative εNd(i) ~ -8). The felsic rocks (monzonitic and syenitic rocks) have Nd isotope signatures quite similar to the mafic rocks with high ¹⁴⁵Nd/¹⁴⁴Nd(i) values (0.51252–0.51267, positive εNd(i) ~ 1.6–4.5) but with variable, evolved and high ⁸⁷Sr/⁸⁶Sr(i) ratios (0.70502–0.70824).

# 6. Discussions

## 6.1. Revised Age of Central High-Atlas Jurassic Intrusive Magmatism

Among a hundred intrusions (dykes, sills and laccoliths) in the studied area, the Tassent, Tasraft, Anemzi, Tirrhist, and Anefgou massifs (Figure 1) constitute the main and key structures of the Central High Atlas geology as they have been widely considered evidence of a presumed Jurassic-Cretaceous intrusive series associated with anticlinal ridges. Over the last four decades, several attempts of geochronological and chronostratigraphic studies have been conducted on these major Central High-Atlas intrusive magmatic structures. Absolute dating was restricted to K/Ar and Ar/Ar age measurements on whole rock and mineral (through biotite) from the mafic rocks collected only in the Tassent and Anefgou massifs (e.g., Armando, 1999; Hailwood & Mitchell, 1971; Westphal et al., 1979). They suggest that the Central High-Atlas Tassent and Anefgou massifs recorded a prolonged magmatic intrusive activity from ca. 157 Ma to ca. 119 Ma, during about 40 Myr from the Late Jurassic to Early Cretaceous (Figure 13a). Variable imprints of the post-magmatic alteration observed in the studied Central High-Atlas igneous intrusive massifs (Figures 4, 5, 10) and the general thermal event that affected these intrusives during the Cretaceous period (e.g., Barbero et al., 2007) could have influenced the K-Ar/Ar-Ar dates resulting from uncertainties in crystallization age determination for magmatism (e.g., Kelley, 2002). In this study, we provide for the first time high temperature geothermometer, U-Pb zircon geochronological data (Table 3, Figures 7–9, and 13a), coupled with new exhaustive field observations, petrography, mineralogy and bulk rock geochemistry (major-trace elements and Sr-Nd isotopes) in different lithological units of the major Central High-Atlas igneous intrusive structures cited above to better understand the petrochronological and geodynamic issues related to this intrusive magmatism. The cumulate troctolite gabbro sample of the mafic unit from the Tassent structure yields two Jurassic age populations at ca. 159 Ma and ca 153 Ma. Similarly, the hydrothermally altered gabbro sample of the mafic unit from Anefgou massif and two syenitic rock samples of the felsic unit from Anemzi and Tirrhist massifs yield ages of ca. 163 Ma and ca. 152 Ma, ca. 161 Ma and ca. 152 Ma, and ca. 156 Ma and ca. 149 Ma, respectively (Table 3). Coeval dates are also found in monzogabbro sample from Tirrhist massif (ca. 161 Ma), in two other syenitic rock samples from Anemzi massif (ca 160–159 Ma), in two syenitic rock samples (between ca 159 and 157 Ma), one monzonitic rock sample (ca. 155 Ma) from Tirrhist massif and two syenitic rock samples from Anefgou massif (at ca. 159–152 Ma) (Figures 7–9, Table 3). The new U-Pb zircon geochronological data are also reported in Figure 13a, ranging from ca. 163 ± 1 Ma and ca. 149 ± 2 Ma (with error about 11 Myr), coinciding with most Late Jurassic dating measurements yielded in mafic rocks from the same intrusive structures (Tassent and Anefgou) by K-Ar and Ar-Ar biotite mineral techniques (Figure 13a). However, they contrast starkly with those of the Early Cretaceous obtained by the K-Ar whole-rock technique on the Tassent gabbros (Figure 13), which might be due to uncertainties in crystallization age determination by the technique. If we consider all obtained U-Pb zircon dates, the Central High-Atlas igneous intrusive structures yield
<figure><img src="image_15.png" /><figcaption>**Figure 13.** (a) Radiometric age synthesis regarding the Ediacaran magmatic period for the Moroccan Meseta belt based on our new data and literature data. Data projection is made with IsoplotR (Vermeesch, 2018). RPP (relative probability plot). Simplified geodynamic events occurring in the Central High Atlas are reported from Saura et al. (2014), Martín-Martín et al. (2017) and Calvín et al. (2018) and references therein. (b) Paleogeographic situation of the Central High-Atlas in the NW-Africa during the late Jurassic (adapted after Labails et al., 2010; Sibuet et al., 2012). (c) Upper Jurassic alkaline intrusive magmatism associated with post-rift diapirism and elongated mini basins.</figcaption></figure>

coeval Late Jurassic ages for mafic and felsic units, attesting that both units might confirm a common synchronous magmatic activity. More recently, U-Pb geochronological studies in gem-quality apatite (ca. 153 Ma; Apen et al., 2022) and titanite (ca. 152 Ma; Yuan et al., 2024) hosted in intrusive rocks from the studied area confirm also this Late Jurassic crystallization event.
As also observed in the outcrops, the sharp intrusive contact features between the mafic and felsic units and between the mafic bodies themselves (Figures 1d, 1e, 2c, and 3), could be produced by at least two main magma injections around 163–159 Ma and 155–149 Ma. The young magma injection event might rejuvenate some zircon grains as indicated by the two age populations in the same sample (e.g., Harlov et al., 2023; Pidgeon, 1992). However, the estimated ca. 11 Myr time gap at least between population dates can also be linked to the protracted crystallization time of the intrusive bodies. It should be noted that several samples yield only one age population with a range age that closely resembles that yielded by two age populations in the same sample, which we consider as a ca. 163–149 Ma date interval (Late Jurassic: Oxfordian–early Tithonian) is the best estimation for the crystallization age of the main Central High-Atlas intrusive massifs, not to mention possible magmatic precursors since the Early Jurassic as evidenced by few old zircons in studied rocks (ca. 180–170 Ma) (e.g., Figure 7–9).

## 6.2. A Geodynamic Scenario for the Alkaline Intrusive Magmatic Event

The timing of magmatic activity in the Imilchil alkaline province indicated by our data, from the subtle precursors in the Early and Middle Jurassic as recorded by a few zircon grains, to the paroxysmic activity in the Late Jurassic, coincides with a major change in the regional tectonic regime due to a significant acceleration of the Central Atlantic opening rate during the Jurassic. During this period, the oblique opening of the Central Atlantic accelerated from a half-spread rate of ~0.8 cm/yr in the Early Jurassic (ca. 200–175 Ma) to a maximum of about ~1.7–2.8 cm/yr in the Late Jurassic (ca. 165–145 Ma) (e.g., Labails et al., 2010). As a result, the motion of Africa relative to Iberia evolved from predominantly extensional—with a minor strike-slip component—to a predominantly left-lateral strike-slip regime that lasted until the Late Jurassic (ca. 145 Ma) and was accommodated in the Central High Atlas by extensive transtensional tectonics (Fernandez, 2019; Labails et al., 2010; Sahabi et al., 2004; Sibuet et al., 2012). Thus, the halokinetic structuring of the Central High Atlas, which is stratigraphically dated from ca. 190 Ma in the Early Jurassic to ca. 170–165 Ma in the Middle Jurassic (e.g., Figure 13; Saura et al., 2014, and references therein), developed mostly in a dominant extensional regime—a feature that is also supported by the similarity of the observed halokinetic architecture in elongated diapir ridges and mini-basins to that described in salt basins and passive margins (e.g., Figures 13b and 13c; Davison et al., 1996; Mart & Ross, 1987; Nalpas & Brun, 1993; Rowan & Vendeville, 2006; Vendeville & Jackson, 1992; Weston et al., 1993). Conversely, alkaline magmatism occurred mostly in a transtensional regime. The early, subtle magmatic pulses recorded in our data would coincide with the progressive change in the regional tectonic regime during the Early and Middle Jurassic, while the paroxysm of igneous activity during the Late Jurassic was coeval with the maximum reached by transtensional deformation accommodating the oblique opening of the Atlantic Ocean.

In several respects, the geodynamic scenario that emerges for the Central High Atlas is reminiscent of that suggested by studies of the North Pyrenean Zone, which in the Mesozoic accommodated extensive and strike-slip deformation associated with the movement of Iberia relative to Europe in response to the opening of the Central Atlantic Ocean (Choukroune & Mattauer, 1978; Jammes et al., 2009; Olivet, 1996). Like the Central High Atlas, the North Pyrenean Zone has been locally structured by salt tectonics (Ford & Vergés, 2020) and intruded by alkaline magmatism (Azambre et al., 1992; Miranda et al., 2009; Montigny et al., 1986; Rossy et al., 1992). However, apart from the fact that magmatism in the Pyrenees occurred in the Cretaceous rather than in the Late Jurassic, a major difference between the two regions is that the transtensional regime in the Pyrenees culminated in the opening of “hot” pull-apart basins, where hyperextension—coeval with magmatism—led to exhumation of mantle rocks directly beneath the Mesozoic sedimentary cover and HT-LP metamorphism in the latter (Clerc et al., 2012, 2015; Golberg & Leyreloup, 1990; Lagabrielle et al., 2010; Teixell et al., 2018). This event was associated with extreme thinning, boudinage and pull-apart extraction of the crustal basement (Clerc & Lagabrielle, 2014). Significant fragmentation of the lithospheric mantle into individual tectonic blocks would also explain the exhumation of mantle fragments from different depths in the North Pyrenean zone (Fabriès et al., 1991).

In the Central High Atlas, the hyperextensional phase is not recorded in the sedimentary cover, but transtensional tectonics probably also resulted in some pull-apart dislocation of the underlying lithospheric mantle and continental basement, probably partially decoupled from the sedimentary cover by the Triassic, as in the Pyrenees (Clerc et al., 2016). In this scheme, deformation at shallower levels was mainly accommodated by strike-slip movements concentrated along diapir ridges (e.g., Figure 13c). In the Atlas chain in general, strike-slip and
transtensional movements would be responsible for short-wavelength variations in lithospheric mantle thickness, as suggested by geophysical imaging (Bezada et al., 2014) and studies of mantle xenoliths (El Messbahi et al., 2015). These works propose piecewise delamination of the lithosphere at the short wavelength scale, combined with the development of small-scale instabilities of a few kilometers wide, involving feedback between (a) mantle upwelling into the rifted lithosphere, (b) decompression melting, and (c) thermal erosion of the lower lithosphere enhanced by melt infiltration and advective heat transport. In the topographic highs formed at the base of the lithosphere, a low degree of partial melting associated with decompression of the upwelling mantle may be the origin of the alkaline magmatism. According to this mechanism, it is likely that several unrelated partial melting zones formed beneath the Central High Atlas. However, as noted below (Section 6.4.), the geochemical and Nd isotopic homogeneity of the studied intrusions suggests that all the magmas were generated from a common mantle source, probably an underplated pre-Jurassic mantle plume with a geochemical signature close to that of OIB (e.g., Figures 10b, 12, and 14). Meanwhile, the pulling apart of the crustal basement probably created voids at the aplomb of major faults, which may have been filled, first by lateral flow of unstable Triassic sediments, and later by partial melts from the mantle. These areas would have served as feeder zones for both the salt diapirs and the magmas that infiltrated them. Injected through a highly fragmented crust, the magmas escaped crustal contamination at this stage, as indicated by the absence of symptomatic trace element and Nd-Sr isotopic enrichments in the mafic terms (Figures 10d and 14; Section 6.4.). However, they may have undergone an early stage of fractional crystallization, as suggested by the mildly evolved nature of the mafic magmas injected into the diapirs, most of which are characterized by Mg# values <0.7 (Figure 11) and are therefore not in equilibrium with the mantle, as noted by Lhachmi et al. (2001).

## 6.3. Relationships Between Magmatic Intrusions and Diapirs

The presence of magmatic intrusions within the diapiric structures of the Imilchil region (Figures 1 and 13c) is a distinctive feature that clearly distinguishes them from other Triassic diapirs in the Atlantic domain (Martín-Martín et al., 2017; Michard et al., 2011; Saura et al., 2014 and references therein). Similar magmatic intrusions have been identified in other areas such as J. Zerkoun (south of Z. Sidi Hamsa), Ait El Abbas, and other localities (Eduard et al., 2025; Teixell et al., 2024). The predominantly Upper Jurassic age of the intrusions revealed by our study confirms the diachronicity between diapiric upwelling and the paroxysm of magmatic activity, as suggested by previous work (e.g., Figures 13a; Fadile, 2003; Hailwood & Mitchell, 1971; Saura et al., 2014). However, our data also indicate that precursor magmatic injections probably occurred throughout the diapirism and that only a relatively short time gap separates the end of diapirism (ca. 170–165 Ma based on deposition of post-diapiric Bathonian-Callovian sediments–Saura et al., 2014) and the paroxysm of magmatic activity (ca. 163–149 Ma–this study). Therefore, while it is most likely that many of the intrusive bodies were emplaced late in already frozen diapiric structures, it is also likely that earlier magmatic injections played a role in sealing the diapirs, as noted by Magee et al. (2021) for the Santos Basin offshore Brazil based on geophysical evidence.

The location of Jurassic magmatic intrusions strictly within the diapirs of Triassic material (Figures 1 and 13c) likely reflects shared structural control by basement faults—possibly connected to major lithospheric faults (e.g., Ayarza et al., 2014)—, rooted in pull-apart lithospheric structures. This is also consistent with geochemical signatures suggesting that the Central High Atlas Jurassic magmatism originated and rose from a shallow mantle source of about 30–40 km (e.g., Albarede, 1992). However, the diapir walls, either still active or frozen, clearly served as preferential corridors for magma migration through the sedimentary cover (Martín-Martín et al., 2017; Saura et al., 2014). To date, only a few studies have addressed the complex interplay between magma migration and salt structures and tectonics, including for example, Schofield et al. (2010, 2014), Heimdal et al. (2019), and Magee et al. (2021). In short, the channeling of magma within evaporite layers and the replacement of the latter by igneous intrusions can be controlled by.

1. Partial or total assimilation of the evaporites by the magmas, facilitated by the low melting temperature of halite (~800°C), a process considered, for example, by Heimdal et al. (2019) to explain the contamination in Cl and radiogenic Sr in CAMPS basalts from the Amazonas and Solimões basins, in Brazil. However, the replacement of evaporites by magmatic rocks via an assimilation mechanism must result in significant magma contamination, especially in Na in the case of halite dissolution (Heimdal et al., 2019). Such contamination is not observed in the majority of the mafic rocks studied, which are essentially potassic. Sodium enrichment is only observed in evolved felsic unit rocks (monzonites and syenites) and some spatially associated, hydrothermally altered gabbros. However, as discussed below (Section 6.4.), the contamination in this case is more
likely to be related to Na-rich fluids derived (at least in part) from evaporite degassing than from direct assimilation of the latter.

2. Thermal decomposition of evaporite layers, including degassing (e.g., Pang et al., 2013) and enhanced creep deformation of certain lithologies (e.g., carnallite-rich layers, after Schofield et al., 2014), a process that may also trigger extrusion of low-viscosity layers and may in turn influence diapir dynamics (Magee et al., 2021; Schofield et al., 2014). These mechanisms may have occurred both downstream of melt intrusions and laterally, generating volatile-rich hydrothermal fluids and paving the way for further magmatic injections. This mechanism can explain the multiple injections of alkaline magmas and their crystallization as sub-vertical intrusions (e.g., Figure 1). In the case of efficient extrusion of partially decomposed, low-viscosity evaporites, magmatic circulation could also have been controlled by salt weld structures left after evaporite extrusion, a process well documented for hydrocarbon circulation in halokinetic contexts (e.g., Rowan et al., 2004).

One point to consider is the paucity of evaporites in the diapiric structures of the Central High-Atlas, except for small amounts of gypsum, and the complete absence of halite in the structures hosting the studied magmatic intrusions. This feature is not uncommon in fossil diapirs, especially when they have undergone orogenic compression (Ford & Vergés, 2020; Giles & Rowan, 2012; Graham et al., 2012; Harrison & Jackson, 2013; Kernen et al., 2012; Rowan & Vendeville, 2006; Saura et al., 2014, 2016). It is generally attributed to the preferential creep of evaporites (Jackson & Cramez, 1989; Wagner III & Jackson, 2011), which can lead to their extrusion toward the upper parts of diapirs and eventually to the surface. This phenomenon is likely to occur at all stages of diapiric evolution, from passive diapirism, during which extruded salt may be incorporated into sediments and/or dissolved in seawater, to the final squeezing of salt bodies during compressional orogenic stages (e.g., Ford & Vergés, 2020). Finally, salt may be dissolved by meteoric fluids during final exhumation. In the case of the Central High-Atlas, the absence of salt has been discussed by Saura et al. (2014) and Martín-Martín et al. (2017). As halite is known to be present in few diapiric structures, south of the Imilchil region (Martín-Martín et al., 2017; Teixell et al., 2003), where it is artisanal mined, Saura et al. (2014) did not rule out an “original configuration of the Late Triassic basin associated with a variable distribution of evaporites” and the possibility that “overpressured schists could also have played a significant role in the diapiric evolution of the Central High Atlas.” However, Martín-Martín et al. (2017) noted the presence of Triassic salt in the subsurface, based on Bouguer gravity anomaly data across the Central High Atlas (Ayarza et al., 2005). Finally, the authors considered extrusion and hydrothermal evaporite dissolution, favored by Jurassic plutonism, as a plausible factor to explain the disappearance of salt. Saura et al. (2014) suggested that the large amount of diapiric material replaced by igneous rocks, sometimes preserving only rare pelitic enclaves (cf. Anefgou and Tirrhist structures), could correspond to the missing evaporite rocks. Although more work remains to be done to elucidate the mechanisms of magma channeling within diapiric structures, the individualization and extensive circulation of Na-Cl hydrothermal fluids, probably derived from the decomposition of halite—as discussed in the following section—strongly argues for the presence of evaporites in the diapirs during the magmatism. The evaporites probably played a role in facilitating and focusing magma infiltration. Although significant amounts of evaporites may have been extruded from the diapirs before and after the Jurassic magmatic event, the intrusions probably played a role in explaining their current rarity in the Imilchil region. However, it is worth noting that evaporites are also absent from diapirs without extrusion (e.g., the High-Atlas Rich area), suggesting that compressional squeezing was probably the main controlling factor (Teixell et al., 2024).

## 6.4. Fingerprints of Magmatic and Hydrothermal Processes

The age differences observed in some complexes suggest that the injections during the paroxysm of magmatic activity occurred in at least two main episodes spread over a period of about ~11 Ma (Figure 13, Table 3; Section 6.1.). However, geochemical data such as parallel trace element patterns (Figure 12) and Nd-Sr isotopes for the gabbros (Figure 14) indicate a homogeneous composition of the injected primary magmas, both in time over this period, and spatially between the different structures. The distribution of the highly incompatible trace elements Th, Nb and Ta shows a within-plate signature, but these elements are significantly depleted in the gabbros compared to the OIB (Figure 14), a feature that cannot be explained solely by the cumulative nature of certain samples and must reflect the source composition, probably depleted as a result of previous melt extraction events. On the Nd-Sr isotope covariation diagram (Figure 14), our data and literature values for the same occurrences show compositions within the mantle array for the gabbros and some syenites ($ε_{Nd}$ ~ 2–4.3,
<figure><img src="image_16.png" /><figcaption>**Figure 14.** (a) Co-variation between the Sr and Nd isotopes of the studied Central High-Atlas Jurassic magmatism (OIB from Sun and McDonough (1989), BAB and MORB from Gale et al. (2013), continental crust from White (2014), Moroccan Triassic tholeiitic basalts (CAMP) from Marzoli et al. (2019) and Moroccan Eocene Tamazeght alkaline massif from Bernard-Griffiths et al. (1991) and Bouabdellah et al. (2010)). (a) Co-variation between the Sr isotopes and Na₂O (wt.%). Evaporite and crustal contamination trends in Figures 13a and 13b from Pang et al. (2013). (c) Nb versus Zr plot (Fractionation trend from Peltonen et al., 1996). (d, e) Th/Yb versus Ta/Yb and Th versus Ta diagrams (after Gorton & Schandl, 2000; Pearce, 1983).</figcaption></figure>

$^{87}$Sr/$^{86}$Sr < 0.706), consistent with an OIB source (Sun & McDonough, 1989) and slightly more enriched than the famous Eocene Tamazeght alkaline complex, in the eastern Central High Atlas (Bernard-Griffiths et al., 1991; Bouabdellah et al., 2010). Conversely, the Moroccan Triassic lavas from the CAMP, with negative ε$_{Nd}$ values (Marzoli et al., 2019), are much more isotopically enriched than the Imilchil alkaline intrusions, indicating a major change in the mantle source of magmatism from the Triassic syn-rift to the Jurassic post-rift events.

The magmas injected into the diapirs underwent cooling (e.g., Figures 13c and 15), as evidenced by the presence of frozen rims, and fractional crystallization, as indicated by the presence of cumulative gabbros and rock compositions aligned with geochemical differentiation trends (Figures 11 and 14). The individualization of the differentiated rocks (monzonitic and syenitic rocks) is also the result of this fractional crystallization process, as confirmed by similar crystallization ages and as generally agreed in previous works (Essaifi & Zayane, 2018;
<figure><img src="image_17.png" /><figcaption>**Figure 15.** (a) Cross-section through the diapiric ridge controlling the magmatism emplacement (adapted from Martín-Martín et al., 2017). (b) Cross-section summarizing the main components of the studied upper Jurassic intrusive magmatism of the Central High Atlas and possible reaction with host Triassic sedimentary rocks. Not to scale.</figcaption></figure>

Lhachmi et al., 2001; Zayane et al., 2002). The mass balance ratios of felsic to mafic rocks calculated from surface data (25/75 for Anemzi and 30/70 on average for the studied five major intrusions—Figure 1d, Table 1) agree well with the values predicted by Lhachmi et al. (2001) using a fractional crystallization model. The differentiated rocks were more pervasively contaminated by exogenous fluids than their mafic counterparts, as already noted by Essaifi and Zayane (2018) and confirmed by their selective enrichment in $^{87}$Sr (up to 0.708) on the Nd-Sr isotopic covariation diagram and the associated enrichment in sodium (Figure 14). Essaifi and Zayane (2018) attributed these enrichments to a contamination by a fluid emanating from limestones of the Jurassic cover, but we believe that an evaporite source, presumably present at the time of the intrusions, would better explain the sodium enrichment (e.g., Heimdal et al., 2019; Pang et al., 2013; Figures 14a and 14b). However, our petrographic observations and mineral chemical data, in particular the replacement of calcic plagioclase ($An_{12-71}Ab_{28-80}$) by albite ($Ab_{90-100}$), confirm that this contamination does not result from the direct assimilation of evaporites, but rather from the interaction of these rocks with Na-rich fluids derived from the thermal decomposition of evaporites (Figure 15). Although still poorly defined in terms of contours and intensity, the Central High Atlas HT-LP metamorphism, which selectively affects Triassic and Early Jurassic sediments (Ibouh, 1995) and locally peaks at about 500°C and 1.5 kbar (Lhachmi, 1992; Lhachmi et al., 2001), is probably related to this stage.

The generation of hydrothermal fluids as a result of thermal evaporite degassing has been proposed in studies of the North Pyrenean Zone (Corre et al., 2018), where this process would be responsible for an extensive Na metasomatic event coeval with alkaline magmatism and a HT-LP metamorphism locally exceeding 600°C at 2–5 kbar (Clerc et al., 2015; Clerc & Lagabrielle, 2014; Golberg & Leyreloup, 1990; Poujol et al., 2010)—and possibly also for advective heat transfer through the sedimentary cover (Dauteuil & Ricou, 1989). In the Central
High Atlas, exogenous fluids enriched in Na, Cl and S probably mixed in variable proportions with endogenous fluids resulting from the final crystallization of the magmas, which were most likely enriched in other elements such as P and F (Ouabid, Raji, et al., 2021). Due to the low fluid/rock ratios during their migration, such fluids can also be modified by reaction with the rocks they pass through, either igneous or sedimentary (Moragas et al., 2022), with their alkaline element content remaining largely buffered by the high Na composition of the evaporite-derived component. As a result, the hydrothermal fluids have had the capacity to modify host rocks, producing Na-metasomatized wall rocks. They are also probably responsible for the observed dispersion in the distribution of several trace elements in syenites, compared to gabbros (Figures 12 and 14a–14e). Finally, the hydrothermal fluids give rise to a variety of mineralized ores, such as the gem-quality apatites observed in the igneous rocks (Ouabid, Raji, et al., 2021) and the Ba- and Sr-rich veins developed in the sedimentary country rocks.

7. Conclusion

Our study focuses on five major igneous intrusive complexes of the Imilchil alkaline province in the Central High Atlas of Morocco. The alkaline suites were intruded into the diapiric walls of the halokinetic structures that developed in the Central High Atlas during the Jurassic as a result of rifting associated with the opening of the Central Atlantic. The combination of field and petrographic observations, mineral and whole-rock geochemical data, Nd-Sr radiogenic isotopes and the first U-Pb zircon ages obtained for the magmatic province allows the timing of the intrusions, their relationship to diapiric structures and the magmatic and hydrothermal processes to be specified. The main conclusions of this study and the resulting evolutionary scenario can be summarized as follows.

1. U-Pb radiometric data on zircon allow us to revise the age of the igneous rocks and to restrict the paroxysm of magmatic activity to the Upper Jurassic (ca. 163–149 Ma), whereas literature data suggest a range from the Middle Jurassic to the Cretaceous.

2. The new ages confirm the existence of a diachronicity between diapirism and magmatism, but also indicate that only a relatively short time gap of a few million years separates the end of the halokinetic event (~170–165 Ma) and the paroxysm of magmatic activity, not to mention possible magmatic precursors since the Early Jurassic (ca 180–170 Ma). It is therefore plausible that alkaline magmatism played a role in sealing and fossilizing the diapiric structures.

3. The location of Jurassic magmatic intrusions strictly within the diapirs of Triassic material probably reflects shared structural control by basement faults—possibly associated with major lithospheric faults rooted at about 30–40 km in pull-apart lithospheric structures where local decompression of the upwelling mantle with an OIB signature generated the alkaline magmatism.

4. The diapiric structures clearly served as corridors for magma migration through the sedimentary cover. The magmas may have used thermally decomposed and fluidized evaporites as preferred pathways.

5. Magmas were injected into the diapirs in several episodes and underwent cooling and fractional crystallization at different levels of the diapiric/magmatic structures. These processes led to the individualization of gabbros, sometimes with a cumulate character, and the differentiation of monzonites and syenites, with a mass ratio of felsic to mafic rocks of ~30/70.

6. Finally, thermal degassing of evaporites as a result of magma injection produced Na-rich hydrothermal fluids, which locally mixed in variable proportions with volatile-rich fluids produced during final magma solidification and reacted variably with the host rocks. The hydrothermal fluids modified some of the differentiated igneous rocks during the final stages of their crystallization and also gave rise to mineralized veins, such as those hosting gem-quality apatites cutting the igneous rocks, and the Ba- and Sr-rich veins developed within the sedimentary country rocks.

Data Availability Statement

Ahechach et al. (2024). [Dataset]. Armando (1999). [Dataset]. Bernard-Griffiths et al. (1991). [Dataset]. Bouabdellah et al. (2010). [Dataset]. Essaifi and Zayane (2018). [Dataset]. Lhachmi et al. (2001). [Dataset]. Marzoli et al. (2019). [Dataset]. Ouabid, Raji, et al. (2021). [Dataset]. Zayane et al. (2002). [Dataset].
Acknowledgments

We would like to thank Dr. Paul D. Asimow for editorial handling, and Dr. Antonio Teixell and the anonymous reviewer for their reviews that helped to clarify and improve some aspects of the paper. This paper is a contribution to the PhD research project of Y. Ahechach, supported by the GSMI-UM6P research projects on the geological exploration of mineral resources associated with igneous alkaline complexes. We would also like to thank all those who contributed and facilitated this contribution, in particular the teams from the GSMI-UM6P/OCP GeoAnalytical Lab (Morocco) and Geosciences Montpellier(France).

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