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Geochemistry and Petrogenesis of Lower Jurassic Mafic Rock Suites in the
External Rif Belt, and Chemical Geodynamics of the Central Atlantic
Magmatic Province (CAMP) in NW Morocco
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Faouziya Haissen, 2 Mohamed Najib Zaghloul, 3Yildirim Dilek, 4 Oriol Gimeno-Vives,
4
Geoffroy Mohn, 5Aitor Cambeses, 4Dominique Frizon de Lamotte & 6Valerie Bosse
Département de Géologie, Faculté des Sciences Ben M’sik, Université Hassan II de
Casablanca, B.P. 7955, Casablanca, Morocco; ORCID Number: 0000-0002-0980-7535
2
Département de Géologie, Faculté des Sciences et Techniques, Université Abdelmalek Essaadi,
Tanger, Morocco
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Department of Geology and Environmental Earth Science, Miami University, Oxford, OH
45056, USA; ORCID Number: 0000-0003-2387-9575
4
Département de Géosciences et Environnement (GEC), Université Cergy CY Paris, 1 rue
Descartes 95000 Neuville/Oise Cedex, France
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Department of Mineralogy and Petrology, Faculty of Sciences, University of Granada, Campus
Fuentenueva s/n, 18002 Granada, Spain
6
Laboratoire des Magmas et Volcans (UMR6524), CNRS, Université Blaise Pascal, ClermontFerrand, France
1
Corresponding author:
Professor Faouziya Haissen
E-mail: [email protected]
Phone: +212 6 61253824
Submitted to: The Journal of Geology
Special Issue: Plate Tectonics Anniversary
Revised Ms: 20 December 2020
Revision #3: 28 May 2021
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This is the author’s accepted manuscript without copyediting, formatting, or final corrections.
It will be published in its final form in an upcoming issue of Journal of Geology, published by ​The University of Chicago Press.
Include the DOI when citing or quoting: https://doi.org/10.1086/716499 Copyright 2021 The University of Chicago Press.
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Abstract
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Jurassic mafic rock suites within a >200-km-long curvilinear belt in the Rif orogenic belt in
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northern Morocco, and show that these rock assemblages formed as part of the Central Atlantic
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Magmatic Province (CAMP). The CAMP represents a large igneous province that straddles the
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edges of the modern, peri–Atlantic continents. It developed ~200 Ma, following the initiation of
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the breakup of Pangea. Main magmatic rocks in the Rif External Zone include basaltic lavas,
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massive dolerite, and isotropic and cumulate gabbros, all intruded by dolerite and trondhjemite
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dikes and sills. Available U/Pb zircon ages from dolerite, gabbro, and trondhjemite dike rocks are
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200±4 Ma, 196±4 Ma and 192±Ma, respectively. Based on their geochemical affinities and
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isotopic compositions, the analyzed rocks define basalt–dolerite and gabbro–cumulate gabbro–
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trondhjemite groups. The basalt–dolerite group samples are sub-alkaline in nature and have low–
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TiO2 contents, whereas the gabbro–cumulate gabbro–trondhjemite group samples are alkaline and
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display high–TiO2 values. Most samples are tholeiitic in character and show large–ion lithophile
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(LILE) and light rare earth (LREE) enrichments, and high field strength element (HFSE) depletion
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compared to N–MORB. Samples of both groups display low 143Nd/144Nd201Ma (0.51182–0.51262)
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and high
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group rocks have E–MORB compositions, compatible with the Low–Ti CAMP suites, whereas the
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gabbro–cumulate gabbro–trondhjemite group rocks have OIB compositions reminiscent of High–
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Ti CAMP suites in other continents. Geochemical features of the OIB–like gabbro–cumulate
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gabbro–trondhjemite group suggest that their magmas underwent differentiation through tholeiitic
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fractionation. Magmas of the rocks of both groups included melt components, originated from
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partial melting of a previously subduction–modified subcontinental lithospheric mantle. Our
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results indicate that the Early Jurassic CAMP magmatism in northern Morocco was more extensive
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than its previously recognized manifestations in Morocco, and that it marked a major episode of
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continental magmatism prior to the opening of the Maghrebian Tethys between Africa and Iberia
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in the latest Jurassic.
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Key words: External Zone, Rif orogenic belt (Morocco); Early Jurassic continental magmatism
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in NW Africa; E–MORB and OIB magmas; Central Atlantic Magmatic Province (CAMP) in
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Morocco; Low–Ti versus High–Ti CAMP rocks; Maghrebian Tethys.
We present new field evidence, geochemical and isotopic data, and age constraints on Lower
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Sr/86Sri ratios with Nd values ranging from (-1.51) to (+4.85). The basalt–dolerite
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It will be published in its final form in an upcoming issue of Journal of Geology, published by ​The University of Chicago Press.
Include the DOI when citing or quoting: https://doi.org/10.1086/716499 Copyright 2021 The University of Chicago Press.
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INTRODUCTION
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Rifting and breakup of Pangea during the Late Triassic–Early Jurassic was associated with
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a large igneous province known as the Central Atlantic Magmatic Province (CAMP; Figure 1A
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Marzoli et al., 1999). The CAMP magmatism was important for two major geodynamic reasons:
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(1) This aerially extensive magmatic event and the attendant extensional deformation were
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followed by the opening of the Central Atlantic Ocean at about 175 Ma, after a phase of
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hyperextension (see review in Biari et al., 2017), and (2) it was also a precursor to the opening of
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the Maghrebian and Ligurian Tethys basins to the west of the Apulia promontory (Figure 2) during
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the Late Jurassic (Favre et al, 1991; Sallarès et al., 2011; Dilek and Furnes, 2019; Gimeno-Vives
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et al., 2019). These two Mesozoic Tethyan seaways were directly connected to the opening of the
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Central Atlantic Ocean, rather than to the evolution of the main trunk of Neotethys to the east
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(Ziegler, 1988; Frizon de Lamotte et al., 2011; Dilek and Furnes, 2019; Tugend et al., 2019).
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The CAMP is widely recognized as one of the major Phanerozoic Large Igneous Provinces
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(LIPs) (Marzoli et al., 2018). However, understanding of the mantle melt source(s) of different
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CAMP domains and pulses is still limited. In Morocco, remnants of CAMP magmatism have been
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reported from the Anti Atlas, High Atlas, Middle Atlas, and Meseta (Figure 1B), spurring
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extensive field–based geochemical, geochronological, and stratigraphic research during the past
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twenty years to better constrain the areal extent of the CAMP suites in this country (Bertrand et
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al., 1982; Sebai et al., 1991; Youbi et al., 2003; Marzoli et al., 2004; Knight et al., 2004; Verati et
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al., 2007; Nomade et al., 2007; Deenen et al., 2010; Bensalah et al., 2011; Dal Corso et al., 2014;
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Marzoli et al., 2018; 2019). However, any occurrence of mafic rock suites associated with CAMP
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magmatism in the Rif orogenic Belt farther north in the northernmost Morocco (Figure 1B) has
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not yet been fully documented. The Early Jurassic mafic rock assemblages in northern Morocco
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provide significant insights into the mode and nature of CAMP magmatism in NW Africa.
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In this paper we document the field occurrence of the Lower Jurassic mafic rock sequences
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in the External Zone of the Rif Belt (Mesorif and Prerif) in northern Morocco (Figures 1B & 3),
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and present new geochemical and isotopic data from these spatially and temporally related
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magmatic rock associations. Distributed along a >200–km–long curvilinear zone in the Rif Belt,
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these mafic rock sequences are compositionally, geochemically and geochronologically similar to
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some of the CAMP igneous rock suites in the peri–Atlantic continents and indicate the existence
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of previously unrecognized products of CAMP magmatism in northernmost Morocco. Our data
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and findings are significant in showing that: (a) CAMP magmatism during the breakup of NW
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Africa from Europe and Iberia during the dispersal of Pangea was associated with partial melting
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of the continental lithospheric mantle (CLM), and not plume related, unlike for many other LIPs;
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and (b) shallow–depth emplacement of gabbroic and hypabyssal dolerite intrusions was more
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extensive than basaltic volcanism at the surface during the CAMP magmatism in northern
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Morocco. These observations and interpretations provide important insights for the mode and
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nature of continental rift magmatism, as discussed at the end of the paper. In the first part of the
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paper, we summarize the definition, aerial distribution, geochronology and geochemical
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characteristic of CAMP in order to provide a geological context for the Early Jurassic mafic
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magmatism in northern Morocco. Next, we discuss the regional geology of the Rif Belt with a
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particular focus on the External Zone, where the Early Jurassic mafic rock sequences are exposed.
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Then, we present our data on the lithological distributions of the Early Jurassic mafic rock
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assemblages in the field, their petrography, major–trace element geochemistry, and isotopic
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compositions. In the last part of the paper, we discuss the melt source and evolution of the CAMP
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suites in the External Zone of the Rif Belt in comparison to the other CAMP occurrences and
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present a tectonomagmatic model for their petrogenetic development. Our findings make
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important contributions to our understanding of the chemical geodynamics of the CAMP.
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CENTRAL ATLANTIC MAGMATIC PROVINCE (CAMP)
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Aerial Distribution and Geological Features of CAMP
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The breakup of the Pangea supercontinent and the opening of the Central Atlantic Ocean
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were accompanied by widespread continental magmatism during the Late Triassic-Early Jurassic
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(Schlische et al., 2003; Sahabi et al. 2004; Labails et al., 2010). Products of this magmatic event
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cover areas of different sizes and shapes in four major circum-Atlantic continents today (Figure
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1A) and constitute the Central Atlantic Magmatic Province (CAMP; Marzoli et al, 1999).
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Geochemical affinities of the CAMP rock suites exposed in different regions have been shown to
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display significant similarities, and their crystallization ages display a narrow period of time for
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their emplacement and eruption as briefly summarized below. The northernmost manifestation of
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CAMP magmatism found to date is the Kerkoune dike in France (Caroff et al., 1995; Jourdan et
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al., 2003), whereas the westernmost, the southernmost and the easternmost occurrences exist in
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Texas (Baksi and Archibald, 1997), in Bolivia and Mali (Bertrand, 1991; Bertrand et al., 2005;
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Verati et al., 2005), and in Algeria (Chabou et al., 2010; Meddah et al., 2017), respectively (Figure
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1A). Hence, the CAMP magmatic event is considered as the aerially most extensive Large Igneous
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Province (LIP) in the geological record of the earth (Blackburn et al., 2013; Callegaro et al., 2014;
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Marzoli, 2018; 2019).
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Lithological units in the CAMP occurrences consist largely of shallow intrusions (dikes,
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sills, laccoliths, and layered gabbros) and lava flows, reported from Brazil, Mali, Guinea, central
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and southern Morocco, Spain, Liberia, Guyana, and the USA (Figure 1A; Sebaï et al., 1991;
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Deckart et al., 1997; Marzoli et al., 1999, 2004; Hames et al., 2000; Verati et al., 2007; Nomade
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et al., 2007; Jourdan et al., 2009; Callegaro et al., 2014, 2017; Marzoli et al., 2018; 2019). The
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most extensive and thickest exposures of CAMP lava flows occur in Canada, USA, central
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Morocco and NE Brazil (Figure 1A & B; Marzoli et al., 2019, and references therein).
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Geochronology of CAMP Magmatism
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The large 40Ar/39Ar geochronological database available from different CAMP occurrences
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indicates a well-constrained time period of 202–200 Ma for the development of the CAMP, with
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a peak magmatic activity around 201 Ma (Marzoli et al., 2018 and references therein; Schoene et
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al., 2010; Blackburn et al., 2013; Davies et al., 2017). The most recent U-Pb zircon dating of
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suitable CAMP rocks constrains the main CAMP magmatic activity around 201 Ma (Marzoli et
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al., 2018, and references therein), consistent with the previously obtained
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(Renne et al., 1998; Renne, 2000; Min et al., 2001; Nomade et al., 2004; Schoene et al., 2006;
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Schaltegger et al., 2008). Younger ages between ~196 Ma (Sinemurian) and 192 Ma (and even
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younger) have been also reported from CAMP outcrops in Brazil, Morocco and the USA (Sutter,
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1988; Hames et al., 2000; Sebai et al., 1991; Deckart et al., 1997; Marzoli et al., 1999, 2004, 2011;
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Knight et al., 2004; Nomade et al., 2007; Verati et al., 2007; Ruhl et al., 2016; Jourdan et al., 2009;
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White et al., 2017). This age span suggests a protracted nature of the CAMP magmatic activity for
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nearly 10 million years after the main event, until ca.192 Ma (Marzoli et al., 2018, and references
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therein).
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Geochemical Features of CAMP Sequences
6
40
Ar–39Ar age data
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CAMP magmatic products are made typically of tholeiitic continental flood basalts or
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basaltic andesites, classified in two compositional groups (De Min et al., 2003): Low-Ti (TiO2 < 2
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wt. %) and High-Ti (TiO2 ≥ 2 wt. %) groups. Low-Ti rocks represent the Prevalent Group, whereas
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High-Ti sequences are restricted to small areas in Suriname, French Guyana and Northern Brazil
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in South America, and in Liberia and Sierra Leone located in the southern margin of the West
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African Craton (Figure 1A; Dupuy et al., 1988; Bertrand, 1991; Chalokwu, 2001; Nomade et al.,
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2002; De Min et al., 2003; Deckart et al., 2005; Merle et al., 2011). A relatively reduced volume
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of High-Ti rocks and the lack of acidic and alkaline rocks is a distinctive feature of CAMP
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compared to other LIPs (Marzoli et al., 2018; Svensen et al., 2020). Low-Ti tholeiites of the CAMP
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display enriched compositions in comparison to normal mid-ocean ridge basalts (N-MORB) with
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higher concentrations of light rare earth elements (LREE) and large-ion lithophile elements (LILE)
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together with a more enriched isotopic signature. Isotopic signature of the small volume, High-Ti
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group is similar to that of enriched MORB (E-MORB) (Maroli et al. 2018, and references therein).
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A mantle plume origin was initially proposed for the origin of the CAMP (May, 1971; Hill
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1991; Oyarzun et al., 1997; Courtillot et al., 1999; Leitch et al., 1998; Wilson, 1997). However,
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the lack of characteristic geochemical signatures of plume magmas and the absence of any textural
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and geochemical evidence of very high mantle temperatures in the record of the CAMP lavas
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(McHone, 2000; Pegram, 1990; Puffer, 2001; Callegaro et al., 2013; Hole, 2015; Whalen et al.
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2015) have weakened this plume origin hypothesis. Alternative models such as plate boundary
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forces (Bott, 1982), edge-driven convection (Anderson, 1982; King and Anderson, 1995), global
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warming of the mantle (Coltice et al., 2007, 2009; Oyarzun et al. 1999; De Min et al., 2003;
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McHone, 2000) or lithospheric delamination as a mechanism of shallow-level emplacement
7
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(Lustrino, 2005) have been proposed. Thus, the origin of CAMP is still a matter of debate (Marzoli
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et al., 2018).
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Evidence for CAMP Magmatism in Morocco
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In Morocco, megadikes (up to 200-km-long in continuous length) and sill intrusions in
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Foum Zguid (Figure 1B), Ighrem, and Draa Valley (Anti Atlas) have been correlated with similar
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dike occurrences in southwestern Europe (mainly in Iberia). Extrusive rock suites of CAMP origin
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are lacking in the Anti Atlas but are widespread in all other tectonic domains (i.e., High Atlas,
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Middle Atlas, Meseta). The best preserved and thickest CAMP lava sequences are found in the
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Central High Atlas (Figure 1B; Marzoli et al., 2019, and references therein). These lavas were
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erupted in subaerial conditions in the once-contiguous extensional basins of eastern North America
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(Coastal NE Magmatic Province in Figure 1A) and Morocco, where they were intercalated with
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Triassic-Jurassic fluvial sediments (Marzoli et al., 2019, and references therein). CAMP lavas
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exposed in the Central High Atlas (to the south of 32°N latitude in Figure 1B) have been
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subdivided, based on their stratigraphic positions and geochemical fingerprints, into four
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distinctive flow units: Lower, Intermediate, Upper, and Recurrent Lavas from the oldest at the
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bottom to the youngest on top (Bertrand et al., 1982). These different lava flows are separated by
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layered sedimentary rocks, indicating that volcanism and deposition were synchronous.
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The CAMP magmatic rock units exposed in the Central High Atlas Mountains were dated
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using both 40Ar/39Ar and U-Pb zircon methods (Marzoli et al; 2019, and references therein). The
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available data indicate at least two temporally overlapping pulses of CAMP magmatism (Sebai et
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al., 1991; Marzoli et al., 2004; Knight et al., 2004; Verati et al., 2007; Nomade et al., 2007; Palencia
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Ortas et al., 2011; Blackburn et al., 2013). Of the four main CAMP units defined by Bertrand et
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al. (1982) in the basalt flows of the Central High Atlas, three of them, the Lower, Intermediate and
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Upper basalts, have indistinguishable 40Ar/39Ar plateau ages that range from 202.7±1.6 Ma to
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199.3±0.6 Ma with a clear age peak at 201.3 Ma. These radiometric dates are consistent with the
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U-Pb zircon ages reported from the similar units by Blackburn et al. (2013). However, the fourth
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unit, Recurrent Unit basalts, are younger with 40Ar/39Ar plateau ages ranging from 199.6±2.3 Ma
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to 196.3±2.4 Ma (Verati et al., 2007). Intrusive rocks of the Foum Zguid dike in the Anti Atlas
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(Figure 1B) yielded a U-Pb zircon age of 201.11±0.07 Ma (Davis et al. 2017; Marzoli et al., 2019),
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also consistent with extant ages from the CAMP units in the Central High Atlas and in other
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countries.
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REGIONAL GEOLOGY OF THE RIF BELT
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The Rif Belt constitutes the northernmost tectonic zone in Morocco and represents the
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western termination of the Peri-Mediterranean Alpine orogenic chain (Figure 2; Chalouan et al.,
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2008). It connects with the Betic Cordillera of the SE Iberian Peninsula to the north through the
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Gibraltar Arc (Figure 2), which is a major oroclinal belt that includes the Rif belt in NW Africa
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and the Betic belt in the southern Iberian Peninsula (Durand-Delga & Fontboté, 1980). The Rif
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Belt continues to the east into the Tell Mountains in Algeria and Tunisia (Wildi, 1983; Leprêtre et
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al., 2018) and farther east into the Southern Apennines through Sicily (Figure 2; Henriquet et al.,
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2020). The Rif segment of the Betic–Rif orogenic belt developed as a result of the collision
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between the AlKaPeCa (Alboran-Kabylies-Peloritan-Calabria) ribbon continent to the north and
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the rifted margin of North Africa to the south (Figure 2; Guerrera et al., 2005; Puga et al., 2017;
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Leprêtre et al., 2018; Gimeno-Vives et al., 2019, 2020b). The Rif Belt consists of W– to SW–
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vergent, large thrust sheets and nappes, exposed in three paleogeographic zones (Figure 3). From
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the north to the south these zones include the Internal Rif (or Alboran Domain), the Flysch Zone
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(Maghrebian Tethys Domain), and the External Rif (Figure 3).
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Internal Rif (Alboran Domain) Zone
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This zone is part of the AlKaPeCa ribbon continent and comprises several
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tectonostratigraphic units (Figure 3; Bouillin, 1986). It has been studied extensively, particularly
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its large upper mantle peridotite massif (Beni Bousera Massif, Figure 3) and high-grade
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metamorphic rock units of both Variscan and Alpine origins. The Internal Rif also includes weakly
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metamorphosed units, the Ghomaride nappe, containing the Dorsale calcaire, which represents
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the eastern part of the paleo northern margin of the Maghrebian Tethys (Figure 3; Chalouan et al.,
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2008 and Rossetti et al., 2010).
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Flysch Zone
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This zone constitutes the sedimentary cover of the Maghrebian Tethys (Figure 3), which
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initially developed as a left-lateral oceanic wrench fault system between Africa and Iberia. The
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Maghrebian Tethys was connected to the Central Atlantic Ocean in the west and to the Ligurian
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Tethys in the north during the latest Jurassic (Leprêtre et al., 2018). The main tectono-stratigraphic
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units of the Flysch Zone range in age from the Late Jurassic to the late Burdigalian (De Capoa et
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al., 2007; Zaghloul et al., 2007). The terminal closure of the Maghrebian Tethys during the
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Langhian-Serravallian (Vitale et al., 2014; Vitale et al., 2015) resulted in the collision of the
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AlKaPeCa microcontinent with the Mesozoic North African rifted margin and in the formation of
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major, south-vergent nappe stacks (Figure 3).
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External Rif Zone
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The External Rif comprises a nappe stack of Upper Triassic to Cenozoic rock units,
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including the Lower Jurassic mafic rock assemblages, which we investigated in this study. The
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current structural architecture of this zone developed as a result of the inversion of Mesozoic rock
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sequences and structures of the rifted continental margin of North Africa (Chalouan et al., 2008;
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Durand-Delga et al., 1960; Gimeno-Vives et al., 2020b; Abbassi et al., 2020; Suter, 1980a, 1980b).
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The External Rif is divided into three sub-domains (Figure 3), which include from the north to the
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south: Intrarif, Mesorif, and Prerif (Suter, 1965; Suter, 1980a, 1980b).
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The occurrence of large intrusive complexes in the External Rif was reported from the
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Mesorif sub-domain before and during the 1960s (Lacoste, 1934, Marçais in Durand-Delga et al.,
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1960; Suter, 1964a, 1964b, 1965). However, these magmatic rocks were originally mapped as
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granites and were interpreted as part of the crystalline basement rocks. Their mafic nature was
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recognized subsequently by Leblanc (1979), and they were interpreted by Vidal (1983) as Lower
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Cretaceous intrusive bodies within a “mélange”. The first complete field description of some of
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these mafic rock bodies was published by Benzaggagh (2011), followed by the first geochemical
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study of Harrara extrusions and Bou Adel (H and BA in Figure 3, respectively) intrusions
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(Benzaggagh et al., 2014). Different interpretations and ages have been proposed for these mafic
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rocks (Benzaggagh et al., 2014; Michard et al., 2014, 2018). Gabbroic rocks have been interpreted
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as an ophiolite complex, representing the Mesorif suture (Michard et al., 2014). A trondhjemitic
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intrusion in Bou Adel revealed a U-Pb zircon age of 190±2 Ma, pointing to an Early Jurassic
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episode of magmatism in the Mesorif (Michard et al., 2018).
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A more recent study has proposed that remnants of a hyper–extended, rifted continental
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margin of North Africa are preserved in the Mesorif (Gimeno–Vives et al., 2019). In this model,
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the southern sub–domain, represented by the Prerif, constitutes a proximal rifted margin of NW
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Africa, whereas the northern sub–domain, Intrarif, makes up a distal rifted margin of the continent.
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The Mesorif in the center, represents a transitional, lithospheric–scale necking zone between these
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two sub–domains. Gabbros and basaltic lavas in the Mesorif are envisaged in this model as artifacts
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of an episode of Late Triassic-Early Jurassic continental magmatism, which was followed by a
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main pulse of rifting in the Middle Jurassic that led to the opening of the Maghrebian Tethys (Favre
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et al., 1991; Gimeno-Vives et al., 2019; 2020a; 2020b). This Middle Jurassic rifting event was also
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responsible for the exhumation of subcontinental mantle at an ocean–continent transition (OCT)
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zone, which is currently exposed in the Beni Malek ultramafic body located in the Intrarif sub-
277
domain (BMP in Figure 3; Michard et al., 1992, 2007).
278
279
MAFIC ROCK UNITS IN THE EXTERNAL ZONE OF THE RIF BELT
280
Field Occurrences and Lithologies
281
For this study, we investigated all known mafic rock sequences exposed in the External
282
Rif, and in particular in the Mesorif sub-domain. Major mafic rock assemblages in the Mesorif
283
include the Tainest, Kef el Ghar, Zaitouna, Bou Adel, Laklaiaa, Ain Chejra, Jbel Bayo, Harrara
284
and Jbel Aghbar massifs (marked as red stars, labeled T, KG, Z, BA, KL, AC, JB, H, and JA in
285
Figure 3). We also examined and sampled the Dar Alami and Jorf Melha massifs in the Prerif sub-
286
domain to the south (black stars labeled DA and JM in Figure 3). Sizes of mafic rock outcrops in
287
these massifs are highly variable, ranging from hundreds of meters to decameters. Main lithologies
288
in these massifs include gabbro, cumulate gabbro, hypabyssal massive dolerite, and doleritic and
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289
leucocratic (trondhjemite) dike and sill intrusions (Figure 4). We summarize below the structure
290
and stratigraphy of the main exposures of the Early Jurassic mafic rock assemblages in the External
291
Rif.
292
293
Bou Adel Massif
294
The Bou Adel massif (BA in Figure 3) is the biggest and the best exposed mafic massif in
295
the Mesorif. It is stratigraphically overlain by a ~10-m-thick breccia, composed of angular and
296
subangular clasts (up to 10 cm in length) of gabbro and limestone in a volcaniclastic matrix. This
297
breccia is tectonically overlain along a thrust fault by a Middle Liassic, thick–bedded limestone
298
(Figures 4 & 5a) and Toarcian–Bajocian marl deposits. The Bou Adel massif consists of fine-
299
grained dolerite, porphyritic olivine gabbro, foliated layered gabbro, and trondhjemite (Figures 4
300
& 5a–b). Leucocratic trondhjemite rocks occur as cm-sized veins, dikes and sills, crosscutting the
301
gabbro and dolerite outcrops or also as small inrusions in the gabbros. Thus, they make up the
302
youngest intrusive rocks in the massif.
303
304
Zaitouna Massif
305
The Zaitouna Massif (Z in Figure 3) is exposed to the east of Taounate City in the east –
306
central Mesorif (Figure 3), where a nearly 500-m-wide, gabbro outcrop is directly overlain
307
stratigraphically by a ~3–m–thick breccia unit (Figures 5c–d), which contains clasts (3-cm to 30-
308
cm-long) of dolerite and limestone in a silty–muddy matrix. The gabbro here is intruded by
309
decameter-thick, NNW–SSE–striking and steeply E-dipping dolerite dikes (Figure 5e). Dikes
310
display coarse–grained microgabbro centers (Figure 5f).
311
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Kef El Ghar Massif
313
The Kef El Ghar Massif is located 30 km east of Bou Adel (KG in Figure 3), and its mafic
314
rocks are exposed in two outcrops: the first one is observed north of the village of Kef El Ghar and
315
the second one several kilometers farther to the west, called the Dar Bou Azza outcrop. The main
316
magmatic lithologies are gabbro–layered gabbro, and dolerite with trondhjemite dike and sill
317
intrusions. These rocks are stratigraphically overlain by breccia and conglobreccia units composed
318
mainly of gabbro clasts; this breccia phases upwards into red shale and pebbly sandstone, which
319
are in turn tectonically overlain along a ENE–dipping normal fault by a nearly 200–m–thick
320
succession of calcareous turbidites and a black shale, making up the Ferrysch Sequence in the
321
Mesorif (Figures 4 & 6a). The gabbro unit starts at the bottom with a coarse-grained biotite–
322
gabbro, phasing upward into a thin (~1 m) layered gabbro, which is depositionally overlain by a
323
~1.5–m–thick breccia, composed of 1 cm to 14 cm-long clasts of gabbro in a silty–sandy matrix.
324
This breccia is stratigraphically overlain by a ~5–m–thick, fine-grained reddish sandstone and silty
325
shale sequence (Figure 6b). Laminated sandstone includes cm-long, red-green chert lenses and is
326
crosscut by small-scale normal faults indicating NE–SW extension.
327
328
Tainest Massif
329
330
Mafic rocks crop out extensively in the Tainest Massif (T in Figure 3), and consist of
dolerite, microgabbro, gabbro, basaltic lavas, and trondhjemite intrusions (Figure 4).
331
332
Jbel Bayo Massif
333
The Jbel Bayo Massif is located near the Nekor fault zone and is the northeasternmost
334
outcrop investigated in this study (JB in Figure 3). A 100–m–wide massive gabbro outcrop
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335
displays meter-sized trondhjemite segregations and dolerite dike intrusions. This gabbro is
336
stratigraphically overlain by a 2–m–thick conglobreccia and volcaniclastic rocks with a medium-
337
grained, greenish sandstone matrix and clasts of gabbro and dolerite. Both the gabbro and the
338
conglobreccia are thrust over by an Upper Jurassic sandstone–mudstone sequence.
339
340
Laklaaia Massif
341
The Laklaaia Massif (KL in Figure 3) is located 4 to 5 km east of Ghafsai Village and 20
342
km NW of the City of Taounate, and is composed mainly of dolerite and layered gabbro, showing
343
variable grain sizes (Figure 4). In Section–I where we sampled the gabbros, a nearly 14–m–thick
344
layered gabbro is stratigraphically overlain by a 20–m–thick conglobreccia with angular clasts of
345
gabbro and recrystallized limestone in siliciclastic and volcaniclastic matrix. A Triassic red
346
mudstone unit tectonically overlies the gabbro and the conglobreccia along a thrust fault. In a
347
different section (Section–II to the north of the Village of Laklaaia) we sampled a nearly 70–m–
348
thick layered gabbro (REO 06, 07 and 08), which is directly overlain by a 50–cm–thick basaltic
349
lava unit. Coarse-grained limestone and volcaniclastic rocks depositionally overlie the gabbro and
350
lava units.
351
352
Ain Chejra Massif
353
This small massif (AC) is located close to the Laklaaia Massif and to the ESE of the Village
354
of Ouratzagh (Figure 3). It is made of 10–m–thick gabbro and dolerite, stratigraphically overlain
355
by ~ a 5- to 10–m–thick breccia, composed mainly of dolerite and gabbro clasts. Doleritic rocks
356
also occur as olistostromal blocks within mudstone and evaporite deposits (Figure 4).
357
Compositionally, the Ain Chejra dolerite is similar to the dolerite in the Laklaaia Massif.
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359
Jbel Aghbar Massif
360
The Jbel Aghbar massif (JA in Figure 3) is exposed along the Chefchaouen-Mokhrisset
361
Road, nearly 2 km North of the village of Mokhrisset. It is ~500–m–long and 100–m–wide and is
362
composed mainly of massive dolerite and basaltic lavas with an aphanitic texture. The dolerite–
363
basalt unit is thrust over a 50–m–thick conglobreccia, which is made of cm- to dm-long clasts of
364
gabbro, dolerite, limestone, and red sandstone in a red-coloured mudstone matrix (Figure 4).
365
366
Harrara massif
367
The Harrara massif is located (H in Figure 3) ~40 km N of Ouezzane (BenYaich, 1991;
368
Benzaggagh, 2000, 2011). It is composed mainly of fine-grained gabbro and massive dolerite,
369
stratigraphically overlain by fine-grained volcaniclastic rocks, made of broken fragments (clasts)
370
of basaltic lavas in a siliciclastic matrix, and a conglobreccia, which consists of dolerite, limestone,
371
mudrock, and red sandstone clasts in a fine-grained clastic matrix. Stratigraphically upward above
372
this conglobreccia is a calcareous turbiditic sandstone intercalated with reddish mudrock (Figure
373
4).
374
375
Dar Alami Massif
376
The Dar Alami Massif occurs within the Prerif sub-domain and is exposed on the
377
Ouezzane-Fes Road, nearly 22 km south of Ouezzane (DA in Figure 3). It is composed mainly of
378
massive dolerite and basaltic lavas with an aphanitic texture (Figure 4).
379
380
Jorf Melha massif
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Jorf Melha Massif (JM in Figure 3) is exposed nearly 43 km SE of Ouazzane on the
382
Ouazzane–Fes Road, and consists mainly of large pillow lava flows. These basaltic pillow lavas
383
are locally embedded within a vari-coloured, clayey matrix, and are overlain stratigraphically by
384
a 1- to 2-m-thick, brecciated lava rocks within a sandstone-claystone matrix (Figure 4). Pillow
385
lavas are made of basalt with microlithic or ophitic textures.
386
387
Mineralogy and Textures of Lower Jurassic Magmatic Rocks
388
We focused our observations on the mineralogy and textures of the mafic and leucocratic
389
rock assemblages in the Bou Adel, Zaitouna, Kel El Ghar, Tainest, Jbel Bayo, Laklaiaa, Ain
390
Chejra, Jbel Aghbar, Harrara, Dar Alami and Jorf Melha massifs as shown in Figure 4. Gabbros
391
are the dominant lithology in the Mesorif massifs, whereas they are nearly absent within the Prerif
392
sub-domain (DA and JM), where dolerite and basaltic lavas are the dominant lithologies.
393
394
Bou Adel Massif
395
Gabbroic rocks in Bou Adel display different textures based on their grain size and
396
cumulate versus isotropic nature. Layered gabbros have a heteroadcumulate texture with poikilitic,
397
pale pink–brownish, and anhedral clinopyroxene oikocrysts, enclosing olivine phenocrysts and
398
euhedral plagioclase grains (Figure 7a). Clinopyroxene is, thus, a post-cumulus phase in these
399
gabbros, filling in the inter-cumulus space. These textural relationships suggest a crystallization
400
sequence of olivine > plagioclase > clinopyroxene > Fe-Ti oxides > biotite + apatite, typical of
401
troctolites (sample REO-17). A second group of cumulate gabbros is made of olivine gabbro,
402
containing subautomorphic clinopyroxene grains (instead of oikocrysts) (samples RE-23 and RE-
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403
24; Figure 7b). In this second type gabbro, pyroxene is the most abundant mineral, whereas in
404
troctolites olivine is the main mineral phase. In either gabbro types (olivine gabbro or troctolite),
405
biotite occurs as discrete grains and is spatially associated with oxide phases (Figure 7b), rather
406
than appearing as overgrowths or reaction rims around primary minerals. Minor amphibole occurs
407
as rims around clinopyroxene grains in the olivine gabbro. Iddingsite, epidote, talc, and muscovite
408
are secondary phases in all gabbros. Other types of gabbros can be observed as pegmatitic gabbro
409
with rare olivine and abundant clinopyroxene and zoned plagioclase phases. Sericite, chlorite and
410
sphene are common secondary minerals in the gabbros, and idiomorphic apatite is also widespread
411
in them. Dolerite in this massif shows medium to coarse-grained, intergranular to sub-ophitic
412
textures, and its mineralogy is dominated mainly by clinopyroxene and plagioclase, less abundant
413
orthopyroxene, and rare olivine and opaque minerals. Chlorite, epidote, and calcite occur as
414
secondary minerals.
415
Leucocratic rocks occur as cm–sized veins, dikes and sills in the gabbros and massive
416
dolerite. Their mineralogy includes mainly plagioclase and quartz, with small amounts of alkali
417
feldspar, clinopyroxene pseudomorphs, amphibole, biotite, oxide minerals and accessory apatite.
418
419
Zaitouna Massif
420
In this massif, the dominant lithology is massive dolerite; gabbro can also be locally
421
observed beneath the massive dolerite unit or as pegmatitic enclaves in dolerite dikes. Dolerite and
422
gabbro rocks contain the same primary mineral assemblage of clinopyroxene, plagioclase, some
423
orthopyroxene, rare olivine, and opaque minerals (Figure 7c); only the grain size of minerals is
424
different in these rocks. Secondary minerals include chlorite + amphibole + serpentine +epidote ±
425
zeolite± prehnite.
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427
Kef El Ghar Massif
428
Mineral phases in dolerite and gabbro rocks in this massif are nearly the same as in other
429
massifs, although the modal percentage of plagioclase reaches 80% in some samples and
430
clinopyroxene and olivine are rarely preserved. Chlorite, talc, calcite, apatite, sericite, and
431
serpentine are widely present as secondary phases. Plagioclase and clinopyroxene represent the
432
main phases in phaneritic gabbros, whereas olivine makes up small grains. Biotite is rare and is
433
commonly observed together with sub-automorphic to skeletal Fe-Ti oxides. Trondhjemite is
434
composed of K-feldspar, plagioclase, quartz, clinopyroxene pseudomorphs, Fe-Ti oxides, and
435
secondary amphibole and chlorite.
436
437
Tainest Massif
438
The mineral assemblages in mafic rocks of the Tainest massif include clinopyroxene,
439
olivine, orthopyroxene, plagioclase and oxide minerals, although their olivine contents are higher
440
than those in other massifs (Figure 7d). Olivine and clinopyroxene phenocrysts are resorbed
441
extensively and commonly occur as euhedral pseudomorphs. Gabbroic rocks also include large
442
grains of primary biotite and oxide minerals. A basaltic dike intruded into a gabbro unit shows a
443
well-developed porphyritic texture with euhedral olivine and clinopyroxene phenocrysts in a
444
groundmass of olivine, pyroxene and oxide minerals.
445
446
Jbel Bayo Massif
447
Gabbroic rocks in this massif are rich in olivine phenocrysts, which show well-developed
448
core and rim segments, indicating their two-stage growth history. Secondary mineral phases
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449
include iddingsite and serpentine after olivine, serpentine, chlorite and amphibole after
450
clinopyroxene, and sericite after plagioclase.
451
452
Laklaiaa Massif
453
Dolerite in this massif consists mainly of clinopyroxene, plagioclase (up to 50%) and
454
olivine, minor orthopyroxene and oxide minerals (showing skeletal textures), and secondary
455
phases of iddingsite, serpentine, epidote, amphibole, chlorite, talc, smectite and prehnite. Gabbro
456
contains euhedral phenocrysts of olivine and clinopyroxene. Olivine phenocrysts exhibit a coarse-
457
grained core and an outer rim suggesting its two-stage growth mechanism (Figure 7e).
458
Clinopyroxene and olivine phenocrysts are surrounded by a groundmass of small olivine,
459
plagioclase, epidote, secondary amphibole and oxide minerals.
460
461
Other Mafic Massifs
462
Compositionally, the Ain Chejra Massif dolerites (Figure 7f) are similar to those in the
463
Laklaiaa Massif. Dolerite and microgabbro rocks in the Jbel Aghbar and Harrara massifs in the
464
western Mesorif have similar mineralogy as the dolerites in the other massifs. Dolerite in the Dar
465
Alami massif in the Prerif sub-domain is also identical but its grain size is smaller in comparison
466
to doleritic rocks in the other massifs. Pillow lavas in the Jorf Melha massif are made of microlithic
467
basalt.
468
MAJOR AND TRACE ELEMENT GEOCHEMISTRY
469
Except for the Bou Adel and Harrara massifs, the data on the geochemistry and isotopic
470
compositions of all the other massifs (i.e., Zaitouna, Ain Chejraa, Laklaiaa) are presented and
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471
interpreted for the first time in this paper. We selected thirty-two (32) rock samples from different
472
massifs to be analyzed in the Bureau Veritas Commodities Laboratory in Canada. These analyses
473
were done with calibrations against certified rock standard reference materials using glass discs
474
for major elements, and powder pellets for trace elements (LF200 analytical method). Major
475
element compositions of SiO2, Al2O3, Fe203, MgO, CaO, Na2O, K2O, TiO2, P2O5, MnO, Cr2O3 and
476
the elements of Ba, Ni, Sc were analyzed by X-ray fluorescence (XRF). Trace elements of Be, Co,
477
Cs, Ga, Hf, Nb, Rb, Sn, Sr, Ta, Th, U, V, W, Zr, and Y were analyzed by Inductively Coupled
478
Plasma-Mass Spectrometry (ICP-MS). The analyzed rock types included dolerite, basaltic lavas,
479
microgabbro, isotropic gabbro, cumulate gabbro, troctolites and trondhjemite dikes cross-cutting
480
all these lithologies. The whole-rock major and trace element data for the analyzed samples are
481
presented in Table 1.
482
483
Alteration effects
484
485
The analyzed rock samples display slight to moderate alteration effects, as reflected in their
486
loss-on-ignition (LOI) values, ranging mainly between 0.7 and 6 wt. %, and reaching values as
487
high as 8 wt. % in the most altered samples (e.g., Kef el Ghar sample RE-61). Given this range of
488
variations in the LOI values, we re-calculated the major element oxides to total 100% on a volatile–
489
free basis before plotting geochemical diagrams and the HFSE for our classification. The cumulate
490
gabbro samples show the lowest LOI values compared to other lithologies (Figure I
491
Supplementary material), consistent with our petrographic observations, indicating that dolerite,
492
basalt and trondhjemite rock samples are in general more strongly altered than cumulate gabbros.
493
Post-crystallization alteration appears to have resulted in partial leaching of the most mobile
21
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494
elements (e.g., alkalis, Rb, Ba and Cs). CaO was also mobilized as inferred by the presence of
495
secondary calcite and carbonate veins in most rock samples. The primary plagioclase (the most
496
abundant mineral phase in these rocks) show variable degree of albitization, sericitization, and
497
saussuritization. Clinopyroxene is locally altered to chlorite or amphibole, and olivine is totally or
498
partially replaced by iddingsite and serpentine minerals. Albitization process of plagioclase
499
resulted in a decrease of CaO and Al2O3, and in an increase of Na2O in the whole-rock
500
(Supplementary Figure I). This post-crystallization alteration was thus responsible for
501
redistribution of alkalis and Ca in the altered samples and of widespread presence of secondary
502
phases in the analyzed rocks.
503
504
Major element geochemistry
505
506
The analyzed rock samples show scattered values of Na2O, MgO, Al2O3, FeO, CaO, K2O,
507
Sr, Rb and Cs in binary diagrams (Supplementary Figure I). Such major elements are known to
508
be mobile during surface weathering and post-crystallization alteration, and hence cannot be used
509
effectively for tectonic discrimination of their magma compositions (Xia and Li, 2019). Their SiO2
510
contents range from 46 to 56 wt.%, with the highest values shown by trondhjemite dikes. MgO
511
values are between 2.59–12.74 wt.%, with a trondhjemite dike in the Bou Adel Massif (RE-75)
512
having the lowest value (2.59 wt%), whereas the cumulate gabbro samples showing variable MgO
513
contents. The MgO contents of troctolites are 11.61 wt. %, and of coarse-grained olivine gabbros
514
around 7 wt. %. Pegmatitic gabbro samples from the Laklaiaa Massif have MgO contents similar
515
to those from the Bou Adel troctolites. The highest Mg-number [(Mg-number =
516
100xMg/(Mg+Fe2+), with Fe2+ being 87% of total Fe] belongs to a suite of basalt and dolerite
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517
samples with values up to 82 (56–82). Isotropic gabbro samples display a range of Mg numbers
518
between 48 and 55, and cumulate gabbro samples between 44 and 70. These Mg numbers of the
519
analyzed samples reflect a measure of the degree of differentiation of their mafic magmas.
520
The Al2O3 contents of the gabbro and cumulate gabbro samples are ~18 wt.%, and likely
521
indicate strong plagioclase accumulation. Figure II in Supplementary Material shows how
522
cumulate gabbro compositions are related mainly to the accumulation of plagioclase (Diagrams
523
A and C), and not to the crystallization of Fe-Mg phases (Diagram B). The main cumulus mineral
524
is plagioclase, as observed in thin sections (Figures 7a, b) and inferred from variable Eu
525
anomalies, and the correlation between the Pl-Ap parameter and trace elements, which are
526
compatible in plagioclase (such as Sr). The Al2O3 and CaO contents in the gabbro-cumulate gabbro
527
group are determined by plagioclase accumulation. Apatite accumulation is also apparent in rocks
528
with positive Eu anomaly and high P2O5 contents. The highest Al2O3 content (about 21 wt. %) is
529
shown by trondhjemite dikes in the Bou Adel massif (sample RE-75), whereas the dolerite and
530
basalt samples have the lowest Al2O3 contents (12–15 wt.%).
531
Based on their lithologies, textures and major–element compositions, we have subdivided
532
our rock samples into two groups: the basalt–dolerite and the gabbro–cumulate gabbro–
533
trondhjemite groups. These two groups display major differences in terms of their TiO2, Hf, Y and
534
Lu contents (Figure I in supplementary material). The gabbro and cumulate gabbro samples
535
(except troctolites) show the highest values of TiO2 (>2 wt.%), particularly the coarse-grained
536
olivine gabbros from the Bou Adel and isotropic gabbros from the Kef El Ghar massifs. The TiO2
537
and P2O5 contents of these rocks are positively correlated. The higher TiO2 values in the gabbro
538
and cumulate gabbro samples are likely due to differentiation and are mainly associated with the
539
abundant occurrence of Fe-Ti oxide phases (ilmenite and/or titanomagnetite). High P2O5 contents
23
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540
of the gabbro-cumulate-gabbro-trondhjemite group are related to high apatite contents in these
541
rocks.
542
543
Trace element geochemistry
544
545
In order to see through the post-crystallization alteration effects, we have used immobile
546
trace elements rather than major elements to further classify the mafic rock suites in our study. We
547
have applied, for example, the Zr/TiO2 vs. Nb/Y diagram (Winchester and Floyd, 1977) for a more
548
effective classification of extensively altered oceanic rocks (Xia and Li, 2019). In this diagram
549
(Figure 8), most of the analyzed dolerite and basalt rock samples plot in the sub-alkaline basalt
550
field, whereas cumulate gabbro, isotropic gabbro and trondhjemite dike rocks plot mainly in the
551
alkaline field.
552
In a multi-element diagram normalized to Silicate Earth (Figure 9A–D; McDonough and
553
Sun, 1995), all analyzed samples display enrichments in large-ion lithophile elements (LILE as
554
Ba, Rb, K, Th, U) and light rare earth elements (LREE) compared to N-MORB, and depletions in
555
high-field strength elements (HFSE, as Nb, Ta). The Chrondrite–normalized rare earth element
556
(REE) diagrams (McDonough and Sun, 1995) display similar and sub-parallel patterns with
557
enrichment in LREE compared to heavy rare earth elements (HREE), characteristic of E-MORB–
558
type magmas (Figures 9E–H). The gabbro–cumulate gabbro group (Figure 9G) shows more
559
fractionated LREE-HREE ratios (e.g. La/YbN= 4.14 to 6.79) than the basalt–dolerite group, which
560
exhibits moderately sloping patterns (Figure 9B–F) with moderate values of LREE/HREE
561
(La/YbN= 2.13 - 3.96). The cumulate gabbro samples are more depleted in some MREE and HREE
562
compared to N-MORB and E-MORB, and they display noticeable positive Eu anomalies (up to
24
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563
1.83) indicating that plagioclase was a major fractionating phase (Figure 9G). Weak negative Eu
564
anomaly is generally observed in isotropic gabbro samples (Figure 9G). The MREE/HREE ratios
565
of the gabbro–cumulate gabbro–trondhjemite group (with steeper patterns) are higher than those
566
of the basalt–dolerite group that can be related probably to different depths of partial melting of
567
their mantle source. This inference is also supported by the low Y and Yb contents of the gabbro
568
and cumulate gabbro samples.
569
570
ISOTOPIC COMPOSITIONS
571
The Sr and Nd isotope ratios were determined in TIMS Laboratory of Granada University
572
in Spain, using Thermal Ionization Mass Spectrometry (TIMS) with a Finnigan Mat 262.
573
Normalization values were 86Sr/88Sr = 0.1194 and 146Nd/144Nd= 0.7219. Blanks were 0.6 and 0.09
574
ng for Sr and Nd, respectively. The external precision (2σ), estimated by analyzing 10 replicates
575
of the standard WS-E (Govindaraju et al., 1994), was better than ± 0.003% for
576
0.0015% for 143Nd/144Nd. The 87Rb/86Sr and 147Sm/144Nd values were determined directly by ICP-
577
MS following the method developed by Montero and Bea (1998), with a precision better than ±
578
1.2% and ± 0.9% (2σ) respectively. Measured radiogenic isotope ratios, associated errors and age-
579
corrected isotope ratios (at 201 Ma) for five selected rock samples (dolerite, gabbro and
580
trondhjemite rocks) are presented in Table 1.
581
All samples show low
143
Nd/144Ndi (0.5118–0.51262) and high
87
87
Sr/86Sr, and ±
Sr/86Sri ratios (0.7042–
582
0.7066) with εNd values ranging from -1.51 to +4.85. These values overlap with those from Low-
583
Ti CAMP rocks [(87Sr/86Sri (0.705–0.707),
25
143
Nd/144Ndi (0.5125–0.5122, εNd:-4 to +1)] but are
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584
slightly more depleted, resembling High-Ti CAMP suites (De Min et al. 2003; Deckart et al. 2005;
585
Merle et al. 2011; Klein et al. 2013 in Marzoli et al. 2018). This overlap is clear in Figure 10,
586
wherein the isotopic values of all our samples plot within the Sr-Nd isotopic range of the published
587
CAMP data. Only two rock samples (RE-47 and RE-27) in our study show positive εNd values of
588
+0.9 and +4.8, respectively, whereas the other two rock samples fall within the enriched domain
589
of the εNd vs. 87Sr/86Sr diagram.
590
591
COMPARISON WITH OTHER CAMP OCCURRENCES
592
The geochronology, lithology, geochemistry, and isotopic compositions of mafic rock
593
suites from the External Zone of the Rif belt in northern Morocco strongly resemble those of the
594
CAMP rock sequences exposed in other countries. In the section below, we discuss different
595
aspects of this comparison.
596
597
Zircon ages and geochronology
598
The results of zircon geochronology analysis of some of the dolerite, gabbro and
599
trondhjemite samples discussed in the current text from several massifs in the External Zone have
600
been published recently by Michard et al. (2018) (one sample) and Gimeno-Vives et al. 2019 (four
601
samples). Therefore, we do not present these data and the related interpretations in detail here.
602
Instead, we give a brief summary of the obtained ages from a dolerite sample from the Laklaiaa
603
Massif (sample RE-07), a microgabbro from Kef el Ghar (sample RE-61), a cumulate gabbro from
604
Bou Adel (sample RE-27), and one trondhjemite dike sample from Kef el Ghar (sample REO-09)
605
that were collected during our fieldwork.
26
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606
Dolerite and gabbro samples yielded concordant206Pb/238U ages ranging from 200 to 195
607
Ma (Gimeno-Vives et al., 2019). A gabbro sample (RE-27) from Bou Adel gave an age of 196±4
608
Ma, whereas a microgabbro sample from Kef el Ghar (RE-61) yielded an age of 195±4 Ma. A
609
dolerite sample from the Laklaiaa Massif (RE-07) revealed an age of 200±4 Ma. A trondhjemite
610
dike sample from the Kef el Ghar Massif gabbro (REO-9) gave a younger age of 192±4 Ma. This
611
age is consistent with a U-Pb zircon age of 190±2 Ma obtained previously from a similar
612
trondhjemite intrusion in the Bou Adel Massif (Michard et al., 2018).
613
Some of the analyzed zircon grains revealed much older ages. A microgabbro sample (RE-
614
61) from the Kef el Ghar massif provided an age of 462±9 Ma, and a single zircon grain gave an
615
age of 2 Ga (Michard et al., 2018). These inherited zircons either existed in the mantle melt source
616
and were incorporated into mafic melts during partial melting, or were picked up by magmas
617
ascending through the old continental crust prior to their emplacement in the Rif Belt.
618
In summary, the crystallization ages obtained from various mafic rock units in the Mesorif
619
sub-domain of the External Zone range between 200-195 Ma, whereas the late-stage trondhjemite
620
intrusions are 192 Ma in age (Michard et al. 2018, Gimeno-Vives et al. 2019). These ages correlate
621
well with the available ages reported in the literature from different CAMP domains, and indicate
622
that the main CAMP magmatic event occurred at 200–199 Ma, followed by two minor pulses
623
around 195 Ma and 192 Ma (Marzoli et al., 2018 and references therein).
624
625
Lithological Comparison of CAMP suites
626
The Lower Jurassic mafic rock sequences in the External Rif belt show strong similarities
627
to the contemporaneous rock suites reported from other tectonic zones in northern Morocco and
628
from other countries where CAMP rock units occur. Pillow lavas observed in the Jorf el Melha
27
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629
massif (JM) of the Prerif are lithologically similar to pillow lava occurrences reported from the
630
Western Meseta (Cogney et al, 1971; Cogney and Faugeres, 1975; Cogney et al, 1974) and from
631
the Central High-Atlas Mountains in Morocco. Some of these pillow lavas might have erupted in
632
fresh-water bodies and are locally spatially associated with phreato-magmatic deposits, such as
633
those pillow lava accumulations in the Algarve basin of southern Portugal (Figure 2; Youbi et al.
634
2003; Martins et al., 2008). Doleritic dikes with ophitic to pegmatitic textures and basaltic lava
635
flows that are spatially associated with Triassic sedimentary rocks have been also reported from
636
the Sub-Betic Zone (External Betic Zone) in southern Spain (Puga, 1987; Puga and Portugal
637
Ferreira (1989); Puga et al. (1989) and Morata et al. 1997). These intrusive and extrusive rock
638
associations in the Sub-Betic Zone have been interpreted as a result of magmatism that was
639
associated with the opening of the Central Atlantic Ocean (Comas et al., 1986; Puga et al., 1989;
640
Morata et al., 1997).
641
642
DISCUSSION
643
Geochemical Characterization of CAMP Suites
644
Based on their Nb/Y and Zr/P2O5 ratio values (Floyd and Winchester, 1975), most of our
645
analyzed samples are classified as tholeiitic in character (Figure 11). The range of MgO contents
646
of the CAMP rock suites (3-14 wt.%) shows a significantly evolved character of their magmas
647
(Marzoli et al., 2018). These geochemical features suggest high percentages (10–50 wt.%) of
648
fractional crystallization of primary mantle melts of the CAMP rock suites. Incompatible trace–
649
element contents of most of our samples from the basalt-dolerite group are similar to the Low–Ti
650
CAMP (CAMP Ti–poor) rocks with E–MORB compositions from other CAMP suites in different
651
countries (Figure 11); these rocks display depletion in HFSE and enrichment in LILE (Marzoli et
28
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652
al., 2018). The gabbro–cumulate gabbro–trondhjemite group plots mainly in the High–Ti and in
653
the in the OIB field (Figure 11). This geochemical shift was likely a result of differentiation of
654
basaltic magmas through tholeiitic fractionation. Thus, the rocks in this group show patterns that
655
are similar to those generally observed in high–Ti CAMP rock suites (Figure 11).
656
In geochemically characterizing our rock samples in line with the CAMP nomenclature,
657
we have used the classification of Marzoli et al. (2018). This classification (Figure 12) is based
658
on a large number of major–trace element and Sr-Nd-Pb isotopic data available from different
659
CAMP domains, and consider all CAMP basaltic lava flows, dikes and sills into six main groups
660
(Figure 12): (1) Tiourdjal group (TiO2 = 1.3–1.5 wt.%; MgO = 6–8 wt.%; La/Yb = 6–8); (2)
661
Prevalent group (TiO2 = 1.0–1.3 wt.%; MgO = 6–8 wt.%; La/Yb = 3.5–5.5); (3) Holyoke group
662
(TiO2 = 0.8–1.0 wt.%; MgO = 6–8 wt.%; La/Yb = 2.5–3.5); (4) Recurrent group (TiO2 = 1.4–1.6
663
wt.%; MgO = 4–6 wt.%; La/Yb = ~ 2); (5) Carolina group (TiO2 = 0.5 wt.%; MgO = up to 13
664
wt.%; La/Yb = 1–3); and, (6) High Ti group (TiO2> 2.1 wt.%; MgO = 3–8 wt.%; La/Yb = 2–8).
665
.Our basalt–dolerite group plots with the Prevalent Group of Marzoli et al. (2018) (Figure 12),
666
which includes the Intermediate and Upper Unit lava flows in Morocco (Bertrand et al. (1982)
667
together with the great majority of CAMP lavas from Portugal, USA, Canada, South America, and
668
most dike and sill occurrences in NW Africa and in NE North America (New England to Canada).
669
However, the most differentiated rocks of the gabbro–cumulate gabbro–trondhjemite group fall
670
outside the field of the Prevalent Group (Figure 12), and some of these rocks (samples RE-25,
671
RE-27, RE-68) show an affinity with the High-Ti Group (TiO2> 2.1 wt.%, 3–8 wt.% MgO, La/Yb
672
= 2–8) of Marzoli et al. (2018) (Figure 12). This High-Ti Group includes high-Ti lava flows from
673
the Parnaiba basin (Brazil) and the high-Ti dikes from Liberia, French Guiana, Suriname, and NE
674
Brazil, as well as the Freetown Layered Intrusion in Sierra Leone. Other samples (RE-61, RE-75)
29
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675
display a large scatter and do not show any affinity with any other group in the classification of
676
Marzoli et al. (2018).
677
678
Melt Source and Evolution of CAMP SUITES
679
In evaluating the mantle melt source and the melt evolution patterns of the CAMP suites
680
we have investigated in this study, we have utilized some of the widely used trace – element
681
discrimination diagrams in order to compare our samples to those reported from other CAMP
682
suites from the peri – Atlantic regions. We realize that such diagrams have limitations in accurately
683
discriminating among basalts produced in different tectonic settings (see, for example, Li et al.,
684
2015), particularly when significant overlaps exist between the different types of basalts. However,
685
the well–established continental rifting related origin of the CAMP rock suites provides an
686
important geological constraint in our interpretations of these diagrams here. We have plotted our
687
samples together with the two main types of CAMP occurrences (i.e., Ti–rich and Ti–poor) on
688
three different discrimination diagrams in order to evaluate their mantle melt source. Our samples
689
plot on the mixing line overlapping the CAMP compositions from Morocco and other countries in
690
the Y/Nb vs. Yb/Nb tectonic discrimination diagram (Figure 13 A). The basalt–dolerite group
691
plots in the typical E-MORB field, whereas the differentiated group of gabbro–cumulate gabbro–
692
trondhjemite intrusions plot near the OIB end-member. In a Th/Yb vs. Nb/Yb discrimination
693
diagram (Figure 13B; Pearce, 2008), the basalt–dolerite group and the Low-Ti CAMP rock suites
694
plot above the mantle array. We interpret this shift as a manifestation of crustal contribution in the
695
melt evolution of this group. The most differentiated group of rocks straddle the OIB and E-MORB
696
fields in both the Th/Yb vs. Nb/Yb and TiO2/Yb vs. Nb/Yb diagrams (Figures 13B & C).These
697
observed geochemical trends and the Nb and Ta (in Ti) anomalies are characteristic of most CAMP
30
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698
basalts, and suggest that their magmas were either derived from partial melting of a subduction-
699
modified mantle source, or from melts whose magmas interacted with continental crust prior to
700
their eruption and emplacement at shallow crustal depths.
701
CAMP basalts from Europe, NW Africa, eastern North America and South America
702
display major differences in their Sr-Nd-Pb isotopic compositions. These isotopic variations
703
suggest melt contributions from different mantle sources, as well as varied degrees of crustal
704
contamination at different depths during their magmatic evolution (Marzoli et al., 2018, and
705
references therein). The enriched Sr-Nd isotopic compositions of the majority of the CAMP
706
basalts, their depleted Nb values, and relatively high LILE (such as Rb, Ba) and light REE (such
707
as La) contents may collectively suggest crustal contamination–assimilation effects. Modeling of
708
the whole-rock and mineral chemistry of rock units from various CAMP domains (Dorais and
709
Tubrett, 2008; Callegaro et al., 2013, 2014; Merle et al., 2011, 2014; Marzoli et al., 2014) have
710
shown, however, that the maximum extent of crustal assimilation could not exceed 10 wt.% of the
711
primary magma volume (Marzoli et al., 2018).
712
We infer, therefore, that the relatively elevated Th/La ratios of up to 0.342 in our basalt–
713
dolerite group (Jochum et al. 1991; Dostal et al., 2016), when compared to the mantle values of
714
0.12 (Sun and McDonough,1989), indicate that their parental magmas were likely derived from a
715
previously subduction–modified mantle (Shellnutt et al., 2018). Slightly enriched Nd isotopic
716
signatures of the gabbro and trondhjemite dike rocks (εNd = -1.5) or marginally depleted (+0.9) to
717
highly depleted εNd values (+4.8, sample RE-47) point to the involvement of a subduction–
718
modified sub-continental lithospheric mantle (SCLM) source in the melt evolution of this group.
719
The basalt–dolerite group plots within the EMII mantle domain in the Nd vs.87Sr/86Sr diagram
720
(Figure 10), indicating an enriched mantle source for their melts. This enrichment of the CAMP
31
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721
mantle source was likely caused by subducted crustal material and / or subducted oceanic
722
sediments (Pegram, 1990; Puffer, 2001; Dorais and Tubrett, 2008; Callegaro et al., 2013, 2014;
723
Merle et al., 2011, 2014; Whalen et al., 2015). This crustal material and oceanic sediments were
724
introduced into the shallow mantle via the Paleozoic or Proterozoic subduction events during the
725
assembly of West Gondwana (Marzoli et al., 2018).
726
727
TECTONOMAGMATIC MODEL FOR THE MELT–MAGMA EVOLUTION OF THE
728
CAMP IN THE MESORIF
729
We present a tectonomagmatic model here, integrating our field observations and new
730
geochemical and petrological data from the mafic rock assemblages in the Mesorif with the extant
731
geochronological and geochemical data from the other CAMP rock suites in Morocco and in other
732
peri–Atlantic countries. The short time (~10 million years) span between the crystallization ages
733
of the hypabyssal dolerite–gabbro and the trondhjemite dike intrusions indicates a relatively short–
734
lived magmatic episode (201–198 Ma) for the CAMP magmatism. Initially slow lithospheric
735
stretching and crustal thinning of the Supercontinent Pangea during the latest Triassic–Early
736
Jurassic led to the development of the CAMP (Figure 14A). This extensional event promoted
737
asthenospheric upwelling and associated decompression melting with mildly elevated mantle
738
potential temperatures (~1350°C), as modeled for other domains of the CAMP (Marzen et el.,
739
2020), that in turn produced mafic magmas emplaced at shallow depths in the continental crust of
740
NW Africa (Figure 14B). Basaltic lavas were erupted in playa lakes and fluvial–lacustrine
741
environments intercalating with syn–rift siliciclastic and volcaniclastic sediments and evaporites
742
in terrestrial basins (Figures 14B & C). The occurrence of more widespread and thicker gabbro
743
and massive dolerite rocks in comparison to limited lava flows in the Mesorif hint that CAMP
32
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744
magmatism produced mainly hypabyssal intrusions, but not extensive volcanic eruptions at the
745
surface. The absence of any subcontinental mantle peridotites and high–grade lower crustal rocks
746
in the Mesorif indicates that lithospheric–crustal thinning was not advanced enough to result in
747
their exhumation, and that the zone of lithospheric necking and crustal thinning was restricted to
748
a short distance of 100 km or less in the Mesorif. A palinspastically restored width of the Mesorif
749
sub-domain after removing its Cretaceous and Miocene contractional deformation and shortening
750
effects is still less than 100 km, and is thus consistent with this interpretation.
751
The two geochemical groups we have identified in this study, the gabbro–cumulate gabbro–
752
trondhjemite and the basalt–dolerite groups, display geochemical fingerprints that are
753
characteristic of moderately to highly enriched (E–MORB) mantle sources and also a mantle that
754
had an OIB metasomatic imprint (Figures 10–12). Therefore, we posit that the asthenospheric
755
mantle beneath NW Africa was highly heterogeneous with these geochemically and isotopically
756
discrete domains, and that progressive partial melting of these mantle domains during the
757
upwelling of the asthenosphere contributed to the melt evolution beneath the narrow rift axis
758
(Figure 14D). Furthermore, the continental lithospheric mantle beneath NW Africa was
759
metasomatized by slab derived fluids and melts derived from subducted oceanic sediments during
760
the assembly of West Gondwana, as we deduce from high Th/La ratios of the basalt–dolerite group
761
rocks and enriched Nd isotopic signatures of the gabbro–trondhjemite group rocks from the
762
Mesorif. Asthenospheric heat as well as decompression melting of this subduction–metasomatized
763
subcontinental lithospheric mantle produced magmas that were coalesced within the melt column
764
beneath the rift axis; they were then channeled upwards within a plumbing system (Figure 14C).
765
Distinctly absent mantle plume geochemical and isotopic signatures in the petrogenesis of the
766
CAMP rocks in the Mesorif indicate that plume magmatism did not play a recognizable role in the
33
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767
evolution of this segment of the CAMP, unlike in the evolution of many other LIPs around the
768
globe. This interpretation is consistent with the findings from other peri–Atlantic CAMP domains
769
(Marzoli et al., 2018, and references therein).
770
The gabbro–cumulate gabbro–trondhjemite rocks that are 4 to 8 million year younger than
771
the dolerite–basalt group rocks in the Mesorif deviate geochemically from typical E–MORB
772
compositions with their high Ti and V values, suggesting that their magmas underwent tholeiitic
773
fractionation. The wide range of MgO contents and the significant positive Eu anomalies of the
774
cumulate gabbros support extensive fractionation of the CAMP magmas (Figure 14E). These
775
features were associated with the fractionation of olivine, Fe–Ti oxide phases (ilmenite and
776
titanomagnetite) and plagioclase, as our petrographic observations suggest (Figure 7). Based on a
777
review of a comprehensive geochemical dataset from different CAMP domains, Marzoli et al.
778
(2018) have proposed that the primary mantle melts of the CAMP experienced high percentages
779
(~10–50%) of fractional crystallization. We think that this extent of fractional crystallization may
780
have taken place in several different stages at different depths in the middle to upper continental
781
crust in NW Africa. As the magmas ascended through the crust within a plumbing system, they
782
experienced some degree of assimilation and fractional crystallization (Figure 14E). The depleted
783
Nb values, enriched Sr–Nd isotopic compositions, and relatively high LILE and LREE values of
784
the basaltic lavas may, in fact, be a result of crustal contamination and assimilation. However,
785
modeling of the whole–rock and mineral chemistry data from basaltic lavas of the other CAMP
786
domains has shown that the degree of crustal assimilation was no higher than 10% of the primary
787
magma volume (Marzoli et al., 2018, and references therein). Thus, we posit that fractional
788
crystallization occurred in magma pools and pathways in the crust, and that as the magma travelled
789
upwards within the plumbing system more fractionation of olivine, plagioclase and clinopyroxene
34
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790
took place (Figure 14E). The existence of large and zoned olivine phenocrysts, surrounded by
791
smaller olivine and clinopyroxene crystals in the hypabyssal dolerite, microgabbro and basalt rock
792
samples support this multiple cooling and crystallization stages for their magmatic development.
793
794
Our study of the mafic CAMP assemblages in the Mesorif has shown that their mantle melt
795
source was heterogeneous, and subduction influenced. The depth of partial melting of this mantle
796
source has not been well constrained for the CAMP–Morocco. A (Ce/Yb)N versus (Dy/Yb)N
797
discrimination diagram, showing the plots of our Mesorif samples and the CAMP–Morocco and
798
Ti–poor and Ti–rich CAMP basaltic rocks indicates that partial melting of the CAMP mantle
799
source started in the garnet stability field (G–MORB) and continued into the N–MORB spinel field
800
(Figure 14F). This inference suggests that mantle depths of initial partial melting must have been
801
75 to 85 km (Figure 14C; Saccani, 2015; Saccani et al., 2015).
802
Our data and interpretations indicate that the Early Jurassic CAMP rifting in the Mesorif
803
zone was limited in terms of its areal extent (<100 km wide), the degree of its mantle partial
804
melting, and time span (<10 million years). The regional geology shows that continued continental
805
rifting in NW Africa shifted farther to the north later in the Late Jurassic and Early Cretaceous,
806
resulting in the development of the Maghrebian Tethys (today’s Flysch Zone in Figure 3), which
807
was connected with the coeval Ligurian and Western Tethys seaways to the north and the Central
808
Atlantic Ocean to the west (Dewey et al., 1973; Dercourt et al., 1986; Ziegler, 1988; Smith and
809
Livermore, 1991; Dilek and Furnes, 2019). Thus, the CAMP in Morocco was a significant
810
precursor to the Mesozoic ocean basin development episodes along the northern edge of Western
811
Gondwana and to the birth of an equatorial (E–W–trending) Neotethyan oceanic realm (Dilek and
812
Furnes, 2019; Furnes et al., 2020).
35
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813
814
CONCLUSIONS
815
816
(1) We have defined and described the field occurrence of newly recognized, 200–192 Ma mafic
817
rock assemblages within a >200–km–long curvilinear zone in the Mesorif and Prerif tectonic sub-
818
domains of the Rif orogenic belt in northern Morocco as part of the Early Jurassic Central Atlantic
819
Magmatic Province (CAMP).
820
821
(2) Major lithologies in these two sub-domains include cumulate and isotropic gabbros, massive
822
dolerite, trondhjemite dikes and sills, and basaltic lavas. These rocks make up compositionally and
823
isotopically two distinct groups: The basalt–dolerite group with a subalkaline affinity, low TiO2
824
contents, and E–MORB compositions, and the gabbro–cumulate gabbro–trondhjemite group with
825
an alkaline affinity, high TiO2 contents, and OIB compositions. The former group is compatible
826
with the Low–Ti CAMP suites, whereas the latter group is analogous to High–Ti CAMP suites
827
documented from the other CAMP occurrences in different peri–Atlantic continents.
828
829
(3) Geochemical fingerprints of the two geochemical groups point to a highly enriched (E–MORB)
830
mantle source and an OIB metasomatic imprint, indicating a highly heterogeneous mantle source
831
beneath NW Africa. Progressive partial melting of these mantle domains during the upwelling of
832
the asthenosphere contributed to the melt evolution beneath the narrow rift axis. In addition, the
833
magmas of both rock groups were influenced by melts derived from partial melting of a previously
834
subduction–metasomatized continental lithospheric mantle beneath modern NWAfrica. These
36
This is the author’s accepted manuscript without copyediting, formatting, or final corrections.
It will be published in its final form in an upcoming issue of Journal of Geology, published by ​The University of Chicago Press.
Include the DOI when citing or quoting: https://doi.org/10.1086/716499 Copyright 2021 The University of Chicago Press.
835
magmas underwent tholeiitic fractionation–related differentiation during their ascent to shallow
836
crustal depths and to the surface.
837
838
(4) The Early Jurassic CAMP magmatism in the Rif Belt in northern Morocco was an important
839
episode of mafic intrusions at shallow depths within the West Gondwana continental crust that
840
lasted for ~10 million years. This magmatic pulse and the associated extensional deformation were
841
a precursor to the opening of the Maghrebian Tethys to the north of NW Africa that occurred
842
during the Late Jurassic.
843
844
Acknowledgements
845
This paper is a contribution to the IGCP Project #683 (igcp683.org). O.G.V. thanks
846
Université Cergy-Pontoise (France) for a PhD scholarship. TOTAL R & D “les marges de
847
convergence” project (Sylvain Calassou) is gratefully acknowledged for financial support towards
848
the organization of our field campaigns in the Rif Belt of Morocco. We are grateful to the Journal
849
reviewers, Professors Cathy Busby (University of California-Davis, USA) and Paola Tartarotti
850
(University of Milan, Italy) for their constructive reviews and insightful comments that helped us
851
improve the organization and the science in the paper. We thank Editor David Rowley for his
852
insightful comments on various aspects of the paper that helped us improve it.
853
854
37
This is the author’s accepted manuscript without copyediting, formatting, or final corrections.
It will be published in its final form in an upcoming issue of Journal of Geology, published by ​The University of Chicago Press.
Include the DOI when citing or quoting: https://doi.org/10.1086/716499 Copyright 2021 The University of Chicago Press.
855
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1320
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Youbi, N., Martíns, L.T., Munha, J.M., Ibouh, H., Madeira, J., Chayeb, A. and El Boukhari, A.
(2003). The Late Triassic–Early Jurassic volcanism of Morocco and Portugal in the geodynamic
framework of the opening of the Central Atlantic ocean. In: Hames WE, McHone JG, Renne PR,
Ruppel C (eds), The Central Atlantic magmatic province: Insights from fragments of Pangaea. Am
Geophys Un Geophys Monogr, 136,179–207.
1324
1325
Winchester, J. A., Floyd, P.A. (1977). Geochemical discrimination of different magma series and
their differentiation products using immobile elements. Chemical Geology, 20, 325-343.
1326
1327
Xia, L., Li, X. (2019). Basalt geochemistry as diagnostic indicator of tectonic setting. Gondwana
Research, 65, 43-67.
1328
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1331
Zaghloul, M.N., Di Staso, A., de Capoa, P., Perrone, V. (2007). Occurrence of upper Burdigalian
silexite beds within the Beni Ider Flysch Fm. in the Ksar-es-Seghir area (Maghrebian Flysch Basin,
Northern Rif, Morocco): stratigraphic correlations and geodynamic implications. Bollettin della
Società Geologica Italiana 126, 223–239.
1332
1333
Ziegler, P.A. (1988). Evolution of the Arctic – North Atlantic and the Western Tethys. American
Association of Petroleum Geologists (AAPG) Memoir 43, doi: https://doi.org/10.1306/M43478.
51
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1334
Figure captions
1335
1336
Figure 1. A – Simplified reconstruction of the Central Atlantic Magmatic Province (CAMP) in a
1337
Pangea realm around 202–192 Ma (gray dashed line outlines the CAMP). Major occurrences of
1338
CAMP dike and sill intrusions and lava flows are shown (modified after Davis et al., 2017, and
1339
Marzoli et al. 2019). Coloured fields represent different CAMP groups according to the
1340
geochemical classification of Marzoli et al. (2018). Outside these fields, all other CAMP
1341
occurrences marked by the pink color belong to the Prevalent Group (TiO2 = 1.0–1.3 wt.%; MgO
1342
= 6–8 wt.%; La/Yb = 3.5–5.5). Outlined box in NW Africa marks the location of Figure 3. B –
1343
Simplified tectonic map of Northern Morocco, showing different tectonic domains and the
1344
locations of CAMP intrusions and lava flows (data from Marzoli et al. 2018, and this study). Stars
1345
mark the newly identified CAMP occurrences in the Rif belt.
1346
1347
Figure 2. Tectonic map of the Western Mediterranean region, showing the Betic Cordillera (Spain)
1348
and the Rif Belt (Morocco) as part of the Gibraltar Arc and the major modern marine basins in the
1349
region. Also shown are the remnants of the Mesozoic Maghrebian Tethys (in black) and the main
1350
exposures of the Variscan crystalline basement units.
1351
1352
Figure 3. Tectonic map of the Rif Belt in northern Morocco, showing different tectonic zones with
1353
distinct tectonostratigraphic units. Notice the generally S–directed thrust and nappe sheets within
1354
the Rif Belt. The External Zone includes, from the north to the south, Intrarif, Mesorif and Prerif
1355
sub-domains. Mafic rock suites investigated in this study occur mainly in the Mesorif sub-domain.
1356
Red stars and black stars mark the locations of the major mafic massifs and sampling sites in the
1357
Mesorif and Prerif sub-domains, respectively. Sites in the Mesorif: AC = Ain Chejra; BA= Bou
1358
Adel; H = Harrara; JA = Jbel Aghbar; JB = Jbel Bayo; KG = Kef El Ghar; Kl = Laklaiaa; T =
1359
Taineste; Z = Zaitouna. Sites in the Prerif: DA = Dar Alami; JM = Jorf Melha.
1360
1361
Figure 4. Stratigraphic columnar sections of the igneous and sedimentary rock units in different
1362
mafic massifs in the Rif Belt. The 20–meter scale bar applies to all columns, except where we
1363
show specific scale intervals within certain lithological units. See the text for the description of
1364
different massifs and their lithological components as displayed in this figure.
52
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1365
1366
Figure 5. Field images of different lithological units in the CAMP massifs in the Rif Belt. a.
1367
Isotropic gabbro with irregular trondhjemite intrusions in the Bou Adel massif. This gabbro unit
1368
is thrust over by Lower Jurassic (Liassic) carbonate rocks; b. Foliated layered gabbro outcrop in
1369
the Bou Adel massif. Magmatic layers dip to the SW, whereas the spaced foliation dips to the NE.
1370
c. Isotropic gabbro in the Zaitouna massif, stratigraphically overlain by a mafic breccia containing
1371
angular clasts of massive dolerite and limestone in a muddy–silty matrix. Person for a scale. d.
1372
Close–up outcrop image of the breccia unit in c. See the vertical scale bar e. NNW–SSE–striking
1373
and steeply E–dipping dolerite dikes that are intrusive into the isotropic gabbro. Person for a scale.
1374
f. Coarse–grained microgabbro in the center of dolerite dikes. Marker for a scale.
1375
1376
Figure 6. Field images of different lithological units in the Kef El Ghar Massif. a. Gabbro–layered
1377
gabbro unconformably overlain by the Upper Jurassic Ferrysch sequence, and they are both
1378
tectonically overlain by Liassic carbonates along a W–SW–directed thrust fault. b. Layered gabbro
1379
unit, overlain by a brecciated gabbro, which is in turn overlain by a red shale and pebbly sandstone
1380
sedimentary unit (~5–m–thick). An ENE–dipping normal fault separates this unit below from
1381
gently ENE–dipping calcareous turbiditic rocks above. Person in a blue shirt for a scale.
1382
1383
Figure 7: Microphotographs of different lithological units from the mafic rock suites in the
1384
Mesorif. a. Typical troctolite from layered gabbros in the Bou Adel massif, showing a
1385
heteroadcumulate texture with poikilitic, anhedral clinopyroxene (Cpx) oikocrysts enclosing
1386
olivine (Ol) and plagioclase (Pl). Cpx is a post–cumulus phase. The main mineral phases include:
1387
Cpx + Pl + Ol + Fe-Ti oxides + Bt (biotite), and secondary minerals. b. Cumulate gabbro from the
1388
Bou Adel massif, composed of Ol+ Cpx + Bt + Fe-Ti oxides. Biotite is commonly spatially
1389
associated with the oxide minerals. c. Gabbro from the Zaitouna massif; mineral phases include:
1390
Cpx+Opx (orthopyroxene) + Pl + rare Ol + Fe-Ti oxides and secondary minerals (chlorite,
1391
serpentine, amphibole). d. Dolerite from the Tainest massif. The main mineral phases include Ol
1392
+ Cpx + Opx + Pl + Fe-Ti oxides + Bt ± epidote. This massive dolerite rock contains abundant
1393
olivine compared to dolerites in the other massifs. e. Microgabbro from the Laklaiaa massif,
1394
showing phenocrysts of Cpx and euhedral Ol in a groundmass of small Ol, Pl, epidote, Fe-Ti oxide
1395
grains and secondary amphibole. Ol phenocrysts display zoning, suggesting possible two–stage
53
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1396
crystallization history. f. Dolerite from the Ain chejra massif with a mineral association of Cpx +
1397
Pl + Ol + Fe-Ti oxides + Bt, and secondary minerals.
1398
1399
Figure 8. Nb/Y vs. Zr/TiO2 plot (after Winchester and Floyd, 1977) of the samples analyzed in
1400
this study in comparison to the Ti–poor and Ti–rich CAMP units and representative CAMP–
1401
Morocco samples, and to the representative N–MORB, E–MORB and OIB fields. The N–MORB,
1402
E–MORB and OIB fields are constructed based on the data from Arevalo and McDonough (2010),
1403
Willbold and Stracke (2010), and Gale et al. (2013). See text for discussion.
1404
1405
Figure 9. A–D: Silicate Earth–normalized multi-element plots, and E–H: Chondrite–normalized
1406
REE diagrams for mafic rocks from the Mesorif (this study). Normalizing values are from
1407
McDonough and Sun 1995. Representative CAMP and CAMP–Morocco fields are shown for
1408
comparison. Blue, yellow and red patterns display the OIB, E–MORB and N–MORB patterns
1409
CAMP rocks from different localities of this province and from Morocco for comparison (see
1410
references used in text). OIB, N-MORB and E-MORB patterns are also plotted for comparison.
1411
Data sources: Cebria et al. (2003); De Min et al. (2003); Villaseca et al. (2004); Verati et al. (2005);
1412
Deckart et al. (2005); Mahmoudi et al. (2007); Martins et al. (2008); Cuppone et al. (2009); Chabou
1413
et al. (2010); Bensalah et al. (2011); Merle et al. (2011); Marzoli et al. (2011, 2014, 2019);
1414
Callegaro et al. (2014); Cirricione et al. (2014); Meddah et al. (2017); Heimdal et al. (2019).
1415
1416
Figure 10. ƐNd201Ma vs
1417
Mesorif (this study), plotted against the majority of the mafic rocks from the External Zone in the
1418
Rif Belt, Ti–poor and Ti–rich CAMP rocks, and the representative fields of the N–MORB, E–
1419
MORB and OIB fields. Data sources are the same as in Figure 9. The data for the N–MORB, E–
1420
MORB and OIB fields are from Arevalo & McDonough (2010), Willbold & Stracke (2010), and
1421
Gale et al. (2013). See the text for discussion.
87
Sr/86Sr201Ma variation diagram of the analyzed rock samples from the
1422
1423
Figure 11. Zr/(P2O5*10000) vs. Nb/Y discrimination diagram (after Floyd and Winchester, 1975).
1424
The majority of our samples and the CAMP units from other countries show a tholeiitic affinity.
1425
See text for discussion.
1426
54
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1427
Figure 12. TiO2 (wt.%) vs. La/Yb correlation diagram of the analyzed rock samples from the
1428
Mesorif plotted against different geochemical groups of the CAMP magmatic products (from
1429
Marzoli et al. 2018). Also plotted here are the Ti–poor and Ti–rich CAMP units, and mafic rock
1430
samples from other tectonic domains in Morocco (CAMP Morocco). Representative E–MORB,
1431
N–MORB, and OIB fields are shown for comparison. The majority of the mafic rock samples
1432
(mainly dolerite-basalt group) from the Mesorif-Prerif plot in the Prevalent Group field, which
1433
includes the Intermediate and Upper Lava Units from Morocco, and the majority of dikes and sills
1434
from Africa and other CAMP occurrences elsewhere. The OIB, N–MORB and E–MORB patterns
1435
are also plotted for comparison (data source from Sun & McDonough, 1989).
1436
1437
Figure 13. A. Y/Nb vs. Yb/Nb diagram. B. Th/Yb vs. Nb/Yb diagram (after Pearce 2008). C.
1438
TiO2/Yb vs. Nb/Yb diagram (after Pearce 2008). Samples representing CAMP rocks from
1439
Morocco and different localities in the whole province are also plotted (see references used in
1440
text). Also shown are the representative CAMP unites from other countries and Morocco. The N–
1441
MORB, E–MORB and OIB fields are constructed based on the data from Arevalo & McDonough
1442
(2010), Willbold & Stracke (2010), and Gale et al. (2013). See the text for discussion.
1443
1444
Figure 14. Integrated tectonomagmatic model, depicting the structural, tectonic, and melt
1445
evolution of Lower Jurassic mafic rock associations in a continental rift zone, represented by the
1446
CAMP in NW Africa. A- The Pangea Supercontinent and the general outline of the CAMP around
1447
200 Ma (modified from Trond et al., 2012). The white ellipse marks the Rif orogenic belt in NW
1448
Africa and depicts the approximate location of the tectonic cross-section in Panel B. B- Interpretive
1449
tectonic cross-section from NW Africa, showing the mode of continental rifting and associated
1450
magmatism in the Early Jurassic (~200–192 Ma). Lithospheric necking and crustal thinning led
1451
into asthenospheric upwelling and decompression melting, which in turn caused partial melting of
1452
the previously subduction–metasomatized lithospheric mantle. Grabens and half–grabens
1453
produced by extensional normal faulting were underlain by hypabyssal mafic intrusions and filled
1454
by fluvial and lacustrine sediments and basaltic lavas. Key to acronyms: BDTZ = Brittle – ductile
1455
transition zone; CLM = Continental lithospheric mantle. C- Inferred schematic cross-section (not
1456
to scale), showing the heterogeneous mantle structure beneath the rift axis of the CAMP in NW
1457
Morocco. Note that partial melting of various mantle domains in the garnet and spinel stability
55
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1458
fields. See text for further discussion. Key for acronyms: CLM = Continental lithospheric mantle;
1459
Gt – Garnet; Sp = Spinel. D- Simplified, chondrite–normalized REE diagram, indicating that
1460
magmas of the cumulate gabbro in the Mesorif that compositionally overlapped with E–MORB
1461
melts experienced significant plagioclase fractionation as evidenced by large positive Eu
1462
anomalies. E- Plumbing system of the CAMP magmas at different crustal levels, whereby
1463
tholeiitic fractionation of olivine, clinopyroxene, orthopyroxene and plagioclase took place in
1464
magma pools and pathways (modified from Heinonen et al., 2019). F- A (Ce/Yb)N versus
1465
(Dy/Yb)N discrimination diagram of the mafic rock suites from the Mesorif (this study) and various
1466
CAMP domains. The density distribution of the data shows progression of partial melting from
1467
the garnet stability field to the spinel stability field (pink arrow). See text for further discussion.
1468
Normalization values are from Sun and McDonough (1989).
1469
1470
Table 1: Major-Trace Elements Geochemistry and Sr-Nd data from mafic rock suites in the
1471
Mesorif.
1472
1473
Supplementary Materials:
1474
1475
Figure I: Binary diagrams of mobile and immobile elements vs. LOI (wt.%) in analyzed samples,
1476
showing the effects of alteration on mobile elements.
1477
1478
Figure II: Parameters that indicate accumulation of plagioclase and ferromagnesian minerals in
1479
mafic samples. (A) Eu/Eu* [Eu*=EuN/√(SmN*GdN)] vs. MgO (wt %); (B) Cr (ppm) vs MgO
1480
(wt. %); (C) Ni (ppm) vs. MgO (wt.%); (D) Eu/Eu* vs. Sr (ppm); (E) Eu/Eu* vs. P2O5 (wt.%).
1481
Abbreviations for minerals: Ol – Olivine; Opx – Orthopyroxene; Cpx – Clinopyroxene; Amp –
1482
Amphibole; Bt – Biotite; Plg – Plagioclase; Ap – Apatite.
56
Figure 1
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Figure 2
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Figure 3
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Figure 4
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Figure 5a&b
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Figure5c-f
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Figure 6
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Table 1
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Table 1: whole- rock major and trace elements for representative samples of mafic rocks from External Rif with Sr-Nd isotopes for specific
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samples
Basalt
RE-08
Major elements (wt %)
SiO2
53.61
TiO2
1.92
Al2O3
18.78
FeOT
5.38
MnO
0.13
MgO
4.93
CaO
2.53
Na2O
7.27
K 2O
0.63
P2O5
0.45
LOI
3.60
Total
99.23
Trace elements (ppm)
Rb
18
Cs
Be
3.0
Sr
195
Ba
58
Sc
27.0
V
138.0
Co
19.5
Ni
65.0
Ga
10.5
Y
24.9
Nb
12.2
Ta
0.8
Zr
240
Hf
4.3
Sn
3.0
U
0.70
Th
1.0
La
10.4
Ce
27.3
Pr
3.7
Nd
17.3
Sm
3.75
Eu
1.49
Gd
4.45
Tb
0.75
Dy
4.68
Ho
1.01
Er
3.02
Tm
0.44
Yb
2.87
Lu
0.43
Isotopes
87
Rb/86Sr
87
Sr/86Sr
87
Sr/86Sr201 Ma
147
Sm/144Nd
143
Nd/144Nd
143
Nd/144Nd201 Ma
εNd201 Ma
Basalt, Pillow lava
Jorf el Melha
Basalt
RE-09
Basalt
RE-10
49.62
1.79
18.68
6.04
0.12
7.69
2.15
5.39
1.09
0.41
6.10
99.08
46.04
1.53
16.53
6.98
0.46
10.74
4.14
2.68
3.30
0.33
5.80
98.53
50.29
0.97
14.45
7.86
0.11
10.83
4.25
3.54
1.92
0.07
4.50
98.79
50.72
1.49
13.42
9.23
0.10
8.52
5.88
3.65
2.00
0.15
3.50
98.66
48.84
1.02
15.16
9.02
0.17
8.52
7.34
3.33
1.60
0.09
3.60
98.69
49.94
1.04
14.83
8.54
0.21
9.54
4.86
3.92
1.57
0.10
4.20
98.75
32
0.1
5.0
173
57
22.0
151.0
31.0
76.0
17.0
29.0
12.6
0.7
226
4.1
2.0
0.50
0.9
15.1
34.9
4.4
18.6
4.01
1.55
5.04
0.79
5.12
1.06
3.12
0.45
2.85
0.43
46
1.4
2.0
493
1073
33.0
155.0
35.6
171.0
15.1
23.5
10.6
0.7
180
3.4
2.0
0.40
0.6
11.0
25.6
3.2
14.0
3.34
1.34
4.14
0.68
4.24
0.87
2.67
0.38
2.34
0.37
21
26
22
0.3
118
202
37.0
297.0
40.4
72.0
15.3
15.7
3.3
0.2
53
1.5
2.0
0.30
0.9
7.6
15.8
1.9
8.8
2.33
0.95
3.03
0.51
3.30
0.63
1.71
0.23
1.49
0.21
219
161
38.0
358.0
33.1
42.0
16.2
30.3
7.3
0.4
121
3.5
3.0
0.50
2.5
15.9
34.0
4.2
17.9
4.84
1.67
5.80
0.94
5.64
1.13
3.31
0.44
2.94
0.42
31
0.2
3.0
231
200
37.0
282.0
45.9
99.0
15.0
20.2
5.1
0.5
78
2.3
3.0
0.30
1.4
8.8
18.8
2.5
11.5
2.89
0.98
3.77
0.62
3.73
0.83
2.26
0.31
2.09
0.31
Zitouna
Dolerite
Gabbro
REO-14
REO-15
Dolerite & basalt
Bou Adel
Dolerite
RE-26
Laklaai
Dolerite
MRE-02
255
215
38.0
287.0
41.2
83.0
14.5
18.1
4.4
0.4
79
2.3
21.0
0.30
1.5
9.4
20.3
2.5
11.4
2.77
0.91
3.36
0.57
3.57
0.74
2.23
0.33
2.01
0.29
This is the author’s accepted manuscript without copyediting, formatting, or final corrections.
It will be published in its final form in an upcoming issue of Journal of Geology, published by ​The University of Chicago Press.
al Rif with Sr-Nd isotopes for specific
continued
Include theTable
DOI1when
citing or quoting: https://doi.org/10.1086/716499 Copyright 2021 The University of Chicago Press.
e & basalt
Laklaai
Dolerite
MRE-05
50.84
0.99
14.14
8.56
0.18
8.61
6.73
3.86
1.59
0.09
3.20
98.79
28
0.3
268
260
37.0
275.0
41.1
87.0
11.8
17.5
4.2
0.3
76
2.1
1.0
0.20
1.4
7.6
16.5
2.1
9.4
2.28
0.64
2.86
0.51
3.16
0.71
2.08
0.28
1.84
0.27
Dolerite
MRE-06
Major elements (wt %)
SiO2
52.48
TiO2
1.39
Al2O3
15.27
FeOT
7.92
MnO
0.10
MgO
8.70
CaO
3.02
Na2O
2.91
K 2O
1.48
P2O5
0.15
LOI
5.50
Total
98.92
Trace elements (ppm)
Rb
15
Cs
Be
Sr
109
Ba
42
Sc
39.0
V
334.0
Co
36.2
Ni
37.0
Ga
16.5
Y
26.6
Nb
7.0
Ta
0.4
Zr
121
Hf
3.4
Sn
3.0
U
0.70
Th
2.4
La
15.6
Ce
29.6
Pr
3.8
Nd
16.2
Sm
3.87
Eu
1.24
Gd
4.68
Tb
0.75
Dy
4.80
Ho
1.01
Er
2.99
Tm
0.42
Yb
2.67
Lu
0.39
Isotopes
87
Rb/86Sr
87
Sr/86Sr
87
Sr/86Sr201 Ma
147
Sm/144Nd
143
Nd/144Nd
143
Nd/144Nd201 Ma
εNd201 Ma
Dolerite & basalt
Taineste
Dolerite
Dolerite
RE-33
RE-35
Laklaai
Dolerite
MRE-07
Gabbro
MRE-13
52.16
2.34
13.75
8.58
0.10
6.80
6.77
1.86
0.90
0.22
5.30
98.78
47.01
1.01
13.90
9.74
0.11
11.91
5.49
2.94
1.18
0.09
5.20
98.58
48.40
1.04
15.14
9.35
0.12
9.26
4.30
3.27
2.90
0.09
4.80
98.67
9
14
1.0
373
35
44.0
505.0
32.8
24.0
17.8
35.8
12.0
0.7
184
5.2
2.0
1.20
4.0
18.5
37.7
4.8
21.0
4.79
1.57
6.01
1.03
6.69
1.35
4.20
0.57
3.64
0.58
79
76
43.0
275.0
52.8
126.0
11.8
14.4
4.1
0.3
74
2.1
2.0
0.30
1.2
6.6
13.7
1.9
8.1
2.11
0.66
2.65
0.47
2.87
0.60
1.63
0.22
1.44
0.23
1.679
0.724850
0.720052
0.1270
0.511997
0.511830
-10.7
Basalt
RE-36
Harrara
Dolerite
RE-02
49.24
0.91
14.70
8.29
0.14
8.22
8.43
3.39
1.93
0.07
3.40
98.72
46.35
0.94
12.87
9.49
0.16
8.64
14.10
0.96
2.38
0.09
2.60
98.58
50.53
1.10
14.30
9.78
0.27
7.27
8.90
2.49
1.42
0.11
2.50
98.67
36
0.2
23
59
1.3
38
0.3
142
243
39.0
290.0
45.4
71.0
16.4
17.9
4.8
0.3
73
2.2
2.0
0.30
1.2
7.1
16.8
2.2
10.1
2.72
0.97
3.21
0.54
3.26
0.70
1.94
0.27
1.94
0.26
311
156
38.0
282.0
37.8
79.0
14.9
15.5
3.8
0.3
57
1.5
1.0
0.30
0.9
6.6
14.1
1.7
8.3
2.28
0.76
2.60
0.48
2.94
0.61
1.71
0.24
1.57
0.23
486
908
34.0
258.0
41.3
76.0
13.4
19.0
4.7
0.2
77
2.1
9.0
0.50
1.8
9.1
19.6
2.8
10.6
2.89
0.82
3.24
0.55
3.45
0.74
1.95
0.28
1.83
0.27
189
180
38.0
296.0
43.8
69.0
19.3
21.7
5.9
0.3
88
2.4
1.0
0.40
1.5
9.6
20.7
2.7
12.3
3.02
1.04
3.73
0.63
3.93
0.84
2.47
0.32
2.26
0.34
This is the author’s accepted manuscript without copyediting, formatting, or final corrections.
It will be published in its final form in an upcoming issue of Journal of Geology, published by ​The University of Chicago Press.
continued
Include theTable
DOI1when
citing or quoting: https://doi.org/10.1086/716499 Copyright 2021 The University of Chicago Press.
Dar Alami
Basalt
RE-04
55.59
0.94
12.96
7.68
0.05
10.17
1.72
3.28
0.15
0.13
6.20
98.87
3
84
23
34.0
253.0
33.5
52.0
13.3
19.5
4.7
0.2
73
1.9
1.0
0.20
1.2
6.3
16.9
2.4
10.6
3.02
0.97
3.58
0.60
3.46
0.77
2.21
0.29
1.88
0.28
Dar Alami
Basalt
RE-06
Major elements (wt %)
SiO2
48.85
TiO2
1.04
Al2O3
14.78
FeOT
8.04
MnO
0.09
MgO
12.00
CaO
2.90
Na2O
3.94
K 2O
0.16
P2O5
0.09
LOI
6.90
Total
98.79
Trace elements (ppm)
Rb
2
Cs
Be
3.0
Sr
104
Ba
21
Sc
38.0
V
296.0
Co
40.4
Ni
76.0
Ga
14.9
Y
17.2
Nb
4.8
Ta
0.3
Zr
77
Hf
2.3
Sn
1.0
U
0.40
Th
1.4
La
8.3
Ce
17.7
Pr
2.2
Nd
10.0
Sm
2.59
Eu
0.90
Gd
3.11
Tb
0.53
Dy
3.30
Ho
0.71
Er
2.03
Tm
0.28
Yb
1.71
Lu
0.26
Isotopes
87
Rb/86Sr
87
Sr/86Sr
87
Sr/86Sr201 Ma
147
Sm/144Nd
143
Nd/144Nd
143
Nd/144Nd201 Ma
εNd201 Ma
Dolerite & basalt
Jbel Aghbar
Basalt
Basalt
RE-12
RE-13
Basalt
RE-15
Brawa
Basalt
RE-20
53.45
1.54
13.23
9.71
0.18
7.49
4.00
3.75
1.58
0.23
3.50
98.66
51.29
1.19
13.88
8.76
0.18
8.25
5.46
2.49
3.50
0.12
3.60
98.72
49.22
1.03
14.35
9.21
0.16
11.31
3.44
2.90
2.12
0.09
4.80
98.63
49.18
1.03
14.00
8.31
0.12
13.00
2.71
2.61
1.89
0.10
5.80
98.75
49.77
1.27
18.17
5.16
0.03
10.93
1.85
4.78
0.23
0.13
6.80
99.12
19
58
1.5
22
0.3
25
1.0
5
0.2
219
593
38.0
316.0
36.3
54.0
14.2
18.5
5.6
0.4
95
2.8
1.0
0.30
1.6
8.5
19.0
2.4
10.7
2.62
0.82
3.41
0.59
3.86
0.79
2.29
0.32
2.16
0.33
128
179
36.0
288.0
47.6
78.0
18.3
17.7
5.0
0.4
84
2.4
1.0
0.50
1.7
7.6
16.5
2.1
9.5
2.40
0.72
2.93
0.48
3.30
0.70
2.15
0.30
2.00
0.31
145
165
36.0
281.0
41.4
80.0
13.2
18.2
5.0
0.5
82
2.2
96
8
35.0
272.0
31.1
67.0
10.9
19.2
5.8
0.5
98
2.6
0.30
1.3
6.6
15.0
2.0
9.1
2.48
0.94
3.17
0.54
3.60
0.75
2.14
0.30
2.10
0.30
0.40
1.6
4.5
11.7
2.1
10.1
3.04
1.18
3.82
0.63
3.84
0.74
2.29
0.35
2.13
0.34
Ain Chejra
Dolerite
MRE-01
3.0
101
142
36.0
270.0
35.3
38.0
13.7
31.5
10.7
0.6
174
4.9
1.0
0.70
3.3
18.3
41.4
5.2
21.9
4.90
1.36
6.23
1.01
6.36
1.31
3.92
0.54
3.33
0.52
This is the author’s accepted manuscript without copyediting, formatting, or final corrections.
It will be published in its final form in an upcoming issue of Journal of Geology, published by ​The University of Chicago Press.
continued
Include the DOITable
when1 citing
or quoting: https://doi.org/10.1086/716499 Copyright 2021 The University of Chicago Press.
Cummulitic gabbro
Bou Adel
Ol-Gabbro
RE-23
48.00
2.21
16.53
9.43
0.15
7.09
10.33
3.17
0.93
0.21
0.60
98.65
9
0.2
1.0
538
102
35.0
270.0
45.2
48.0
15.4
16.0
13.5
0.8
95
2.3
1.0
0.40
1.1
9.4
20.0
2.5
11.3
2.89
1.23
3.45
0.54
3.08
0.62
1.67
0.22
1.44
0.20
Cummulitic gabbro
Bou Adel
Dolerite
Troctolite
RE-25
REO-17
Major elements (wt %)
SiO2
47.04
47.57
TiO2
4.97
0.66
Al2O3
16.70
17.99
FeOT
9.26
8.57
MnO
0.15
0.12
MgO
5.69
11.61
CaO
10.45
7.79
Na2O
3.33
2.91
K 2O
0.76
0.77
P2O5
0.13
0.13
LOI
0.20
0.60
Total
98.68
98.72
Trace elements (ppm)
Rb
7
7
Cs
0.1
Be
Sr
561
525
Ba
93
79
Sc
38.0
8.0
V
401.0
62.0
Co
42.0
60.1
Ni
241.0
Ga
13.6
13.5
Y
14.9
7.1
Nb
23.7
7.5
Ta
1.4
0.4
Zr
105
57
Hf
2.7
1.1
Sn
1.0
2.0
U
0.30
0.30
Th
0.9
0.9
La
8.5
6.3
Ce
15.7
12.6
Pr
2.1
1.4
Nd
10.0
6.1
Sm
2.61
1.37
Eu
1.25
0.87
Gd
3.10
1.53
Tb
0.49
0.24
Dy
2.72
1.36
Ho
0.55
0.27
Er
1.57
0.69
Tm
0.20
0.11
Yb
1.33
0.65
Lu
0.20
0.09
Isotopes
87
Rb/86Sr
87
Sr/86Sr
87
Sr/86Sr201 Ma
147
Sm/144Nd
143
Nd/144Nd
143
Nd/144Nd201 Ma
εNd201 Ma
Bou Adel
Ol-Gabbro
RE-27
Gabbro
Kef el Ghar
Gabbro
Gabbro
RE-61
RE-68
Gabbro
REO-10
47.14
4.62
16.86
8.90
0.17
4.74
9.24
3.49
1.54
0.23
1.70
98.63
47.25
1.69
15.33
4.11
0.06
11.00
6.39
4.32
0.47
0.39
8.20
99.21
47.64
4.23
15.97
8.91
0.28
5.01
8.10
3.95
1.79
0.32
2.50
98.70
49.47
1.63
16.00
8.47
0.14
6.02
8.65
3.90
0.84
0.22
3.50
98.84
21
0.9
2.0
847
215
35.0
384.0
38.1
44.0
16.1
23.7
38.8
2.1
248
5.7
3.0
0.80
2.7
16.7
31.6
3.8
15.5
3.91
1.39
4.84
0.79
4.77
0.95
2.73
0.37
2.18
0.32
9
0.3
6.0
151
56
19.0
166.0
16.7
209.0
21.7
30.2
29.6
1.5
200
4.3
4.0
0.40
4.6
37.5
75.5
8.0
32.2
5.94
1.29
6.53
1.01
5.86
1.10
2.93
0.44
2.60
0.37
16
0.4
4.0
455
326
34.0
334.0
26.4
10
0.1
0.089
0.705678
0.705425
0.2293
0.512731
0.512429
1.0
0.045
0.706642
0.706513
0.1265
0.512468
0.512302
-1.5
18.3
17.5
29.4
1.9
146
3.5
1.0
0.80
2.0
14.3
29.0
3.5
15.8
4.04
1.41
4.31
0.65
3.61
0.66
1.90
0.25
1.55
0.22
328
110
23.0
147.0
29.3
50.0
16.4
16.5
12.3
0.7
98
2.3
3.0
0.50
1.1
8.9
18.6
2.4
10.3
2.90
1.14
3.80
0.58
3.27
0.64
1.79
0.23
1.46
0.22
This is the author’s accepted manuscript without copyediting, formatting, or final corrections.
It will be published in its final form in an upcoming issue of Journal of Geology, published by ​The University of Chicago Press.
continued
Include the DOI Table
when1citing
or quoting: https://doi.org/10.1086/716499 Copyright 2021 The University of Chicago Press.
Leucocratic facies
Bou Adel
Trondhjemite
RE-75
50.86
3.00
21.03
5.86
0.10
2.59
8.87
4.47
1.45
0.22
0.70
99.15
11
0.3
678
161
18.0
195.0
19.9
19.0
11.7
13.6
1.0
98
2.2
1.0
0.60
1.4
11.1
20.5
2.4
9.9
2.48
1.38
2.65
0.43
2.54
0.49
1.22
0.18
1.11
0.15
Leucocratic facies
Kef el Ghar
Jbel Bayou
Trondhjemite
Trondhjemite
REO-9
RE-47
Major elements (wt %)
SiO2
52.46
TiO2
2.37
Al2O3
17.78
FeOT
4.89
MnO
0.03
MgO
6.71
CaO
2.49
Na2O
6.41
K 2O
0.55
P2O5
0.55
LOI
5.00
Total
99.24
Trace elements (ppm)
Rb
8
Cs
0.1
Be
1.0
Sr
245
Ba
99
Sc
19.0
V
152.0
Co
7.4
Ni
63.0
Ga
17.7
Y
28.0
Nb
37.8
Ta
2.5
Zr
223
Hf
4.8
Sn
5.0
U
2.10
Th
3.7
La
15.9
Ce
33.8
Pr
4.0
Nd
16.5
Sm
4.82
Eu
1.52
Gd
6.04
Tb
1.05
Dy
6.07
Ho
1.14
Er
3.00
Tm
0.39
Yb
2.46
Lu
0.32
Isotopes
87
0.084
0.102
Rb/86Sr
87
0.706921
0.704534
Sr/86Sr
87
0.706681
0.704242
Sr/86Sr201 Ma
147
0.1397
0.0874
Sm/144Nd
143
0.512556
0.512743
Nd/144Nd
143
0.512628
Nd/144Nd201 Ma 0.512372
εNd201 Ma
-0.1
4.9