Petrological and geochemical constraints on the origin of apatite ores from Mesozoic alkaline intrusive complexes, Central High-Atlas, Morocco

1 1Petrological and geochemical constraints on the origin of apatite 2ores from Mesozoic alkaline intrusive complexes, Central High3Atlas, Morocco 4Muhammad Ouabid1,2,*, Otmane Raji1, Jean-Marie Dautria3, Jean-Louis Bodinier1,3, 5Fleurice Parat3, Hicham El Messbahi4, Carlos J. Garrido2, Youssef Ahechach1,5 61. Mohammed VI Polytechnic University, Geology & Sustainable Mining, Hay Moulay Rachid, 43150, 7Benguerir, Morocco 82. Instituto Andaluz de Ciencias de la Tierra, Consejo Superior de Investigaciones Científicas– 9Universidad de Granada, Avenida de las Palmeras 4, 18100 Armilla, Granada, Spain 10 3. Géosciences Montpellier, Université de Montpellier, CNRS, Cc 60, Place Eugène Bataillon, 34095 11 Montpellier Cedex 05, France 12 4. Département de Géologie, Faculté Polydisciplinaire, Université Sidi Mohamed Ben Abdellah, Route 13 d'Oujda, P.O. Box 1223, Taza, Morocco 14 5. Département de Géologie, Faculté des Sciences, Université Moulay Ismail, P.O. Box 11201 Zitoune, 15 Meknès, Morocco 16 *Corresponding author: Muhammad Ouabid, [email protected] 2 18 Abstract 19 Five main apatite deposits from the Moroccan Central High-Atlas —including 20 Anemzi, Tirrhist-Inouzane, Tassent, Ait-Daoud-Toumliline and Tasraft— have been 21 investigated through integrated field, petrographic, mineralogical, and geochemical 22 studies. The apatite ore occurs as veins, few mm to several cm thick, spatially 23 associated with syenite‒quartz monzonite (dominant) to monzodiorite‒gabbrodiorite 24 bodies emplaced during the Atlas Jurassic-Cretaceous alkaline magmatic event. Gem25 quality apatite crystals, up to 15 cm in size, are associated with albite, clinopyroxene 26 (hedenbergite–augite–diopside), amphibole (hornblende–edenite–actinolite), K-feldspar, 27 quartz, magnetite, titanite, epidote, prehnite, and calcite. Two types of apatite have 28 been distinguished according to their halogen contents: F-rich and Cl-rich apatites. Both 29 types are enriched in REE (up to 2 wt.% ΣREE), and have major and trace element 30 contents consistent with a magmatic origin, corroborated by the chemical similarity with 31 accessory apatite in the host alkaline intrusions, as well as with other reference igneous 32 apatites. However, hydrothermal fluids percolating through country sedimentary rocks 33 were also involved, as indicated by the striking Na enrichment observed in the Cl-rich 34 apatites, likely inherited from Triassic evaporites, showing close spatial relationships 35 with the alkaline intrusions. These fluids would also account for pervasive Na36 metasomatism of alkaline magmatic rocks, as observed in the wall rocks of the apatite 37 veins. 38 Keywords: Moroccan Central High-Atlas, alkaline magmatic event, vein-type apatite 39 ores, F-rich and Cl-rich apatites, magmatic origin/hydrothermal fluids. 3 41 1. Introduction 42 Apatite is the most ubiquitous accessory phosphorus(P)-bearing mineral found in 43 sedimentary, igneous and metamorphic rocks. It is also sometimes a major rock-forming 44 mineral in some peculiar rock types: e.g. sedimentary phosphorites and some igneous 45 rocks of alkaline, carbonatitic and anorthositic affinities (e.g., McConnell 1973; Glenn et 46 al., 1994; Dymek and Owens 2001; Martin and Rakovan, 2013; Duchesne and Liégeois, 47 2015; Pufahl and Groa et al., 2017). Apatite is the only major potential resource of P 48 used in the manufacturing of e.g., nitrogen-phosphorus-potassium (NPK) fertilizers, 49 animal feed supplements, ceramics, pharmaceuticals, textiles and batteries. The 50 sedimentary phosphorites make the largest contribution to the world production of 51 phosphate, whereas igneous apatite rocks account for only ~ 10 % (e.g., Jasinski, 2016; 52 Pufahl and Groa et al., 2017). The igneous phosphate ores are nevertheless 53 economically valuable since they provide high quality phosphate concentrates with low 54 contents of unwanted contaminants such as Cd, As, Pb, Si and Al, and they are rich in 55 Rare Earth Elements (REE) (e.g., Webster and Piccoli, 2015; Chakhmouradian et al 56 2017). In terms of world resources, Morocco is the largest holder and producer of 57 sedimentary phosphates, with reserves estimated to represent about ~ 76 % of world 58 reserves (e.g., Jasinski, 2016, 2019; Pufahl and Groa et al., 2017; www.ocpgroup.ma). 59 However, Morocco also bears igneous phosphate ores, especially in the High-Atlas 60 mountain range. Over the last two decades, gem-quality apatite crystals well61 appreciated by mineral collectors and known as "Moroccan High-Atlas Imilchil-Anemzi 62 apatite" were made available on the international mineral market (e.g., 63 https://www.minfind.com/mineral-569173.html; https://www.mchminerals.com/fine- 4 64 mineral-specimens/apatite-morocco; Rakovan, 2015). Despite their economic and 65 scientific interests, only one recent study (Dumańska-Słowik et al., 2018) was devoted 66 to the Anemzi apatite, chiefly dealing with organic inclusions, and a detailed 67 characterization of the apatite ores remains to be done. This paper provides the first 68 description of the main apatite deposits in the High-Atlas region. We carried out an 69 exhaustive petrographic, mineralogical, and geochemical study (major, volatile, and 70 trace elements) of the apatites and associated minerals, and of the host lithologies. 71 These data are used to identify suitable sources for the phosphate mineralization. 72 2. Geological setting 73 The High-Atlas (up to 4000 m elevation) constitutes a typical intra-continental 74 Alpine orogen, forming a thrust chain resulting from the Africa–Eurasia convergence. Its 75 growth started during the early Mesozoic with the reactivation of Pangea paleo-faults 76 which delimited subsidence sedimentary basins, leading to the Atlantic Ocean opening 77 at the NW Africa margin (e.g., Stampfli and Borel, 2002; Piqué et al., 2002; Teixell et al., 78 2003; Laville et al., 2004; Frizon de Lamotte et al., 2008, 2009; Domènech et al., 2015). 79 This chain includes several segments, such as the Western, Central and Eastern High80 Atlas (Fig. 1a). In contrast with the two other segments, the Central High-Atlas 81 corresponds to the deepest part of the orogen (Fig. 1b; e.g., Teixell et al., 2003; Frizon 82 de Lamotte et al., 2008, 2009; Gouiza et al., 2010; Michard et al., 2011). Several works 83 have described the lithostratigraphical sequence of this belt (Fig. 1b‒c; Charrière et al., 84 2005, 2009; Haddoumi et al., 2002, 2010; Ettaki et al., 2007; Frizon de Lamotte et al., 85 2008, 2009; Cavin et al., 2010; Saura et al., 2014; Charrière and Haddoumi, 2016, 86 2017; Cavallina et al., 2018). The first Mesozoic sediments are Triassic red argilites and 5 87 evaporites (mostly halite and gypsum) forming more than 1000 m-thick series which 88 deposited within subsident basins and which cover unconformably a leveled 89 Precambrian-Paleozoic basement (e.g., Du Dresnay, 1987; Piqué et al., 2002; Eddif et 90 al., 2007; Ouabid et al., 2017, 2020; Boukerrou et al., 2018; Karaoui et al., 2021). These 91 deposits are generally interbedded with tholeiitic lava flows linking to the Central Atlantic 92 Magmatic Province (e.g., Marzoli et al., 2019 and references therein). Subsequently, a 93 regional marine transgression led to the deposit of a series of Early Jurassic carbonates 94 (mostly limestones and marls). During the Middle Jurassic and Lower Cretaceous, 95 conformable red beds of continental sandstone settled, indicating the filling of basins 96 and a marine regression. Overall, these post-rift basins were controlled by sinistral 97 strike-slip faults, and separated by Triassic diapiric anticlinal ridges with dominant NE 98 trending (Fig. 1b‒c; e.g., Laville and Pique, 1992; Saura et al., 2014). In the Middle 99 Cretaceous, a new marine carbonated platform sets up (e.g., Froitzheim et al. 1988; 100 Cavallina et al., 2018 and references therein), and at the end of the Cretaceous, the 101 sediments consist of terrigenous red beds, expressing the early manifestation of the 102 Alpine orogeny (e.g., Laville et al. 1977; Froitzheim et al. 1988; Ettachfini and Andreu, 103 2004; Ettachfini et al., 2005; Michard et al., 2011; Cavallina et al., 2018). The Central 104 High-Atlas Mesozoic ridges are often unconformably overlain by Paleocene non-marine 105 red bed formation which draws the synclinal geometries attributed to Late 106 Cretaceous‒Cenozoic shortening and halokinesis (Fig. 1c; e.g., Charriere et al., 2009; 107 Michard et al., 2011). The famous Cenozoic alkaline-carbonatite magmatic episode (~ 108 45‒35 Ma) that occurred at NW Central High-Atlas Tamazight area (Fig. 1b) coincides 6 109 with the African‒European collisional event (e.g., Bouabdli et al., 1988; Bernard-Griffiths 110 et al., 1991; Bouabdellah et al., 2010). 111 In addition, the core of the Central High-Atlas anticlinal ridges is formed by an 112 important Jurassic–Cretaceous alkaline to transitional magmatism between ~ 165 Ma 113 and ~ 125 Ma (Fig. 1b‒c; e.g., Hailwood and Mitchell 1971; Westphal et al., 1979; 114 Beraâouz and Bonin, 1993; Rahimi et al. 1997; Armando, 1999; Haddoumi et al., 2010; 115 Michard et al., 2011, 2013; Bensalah et al., 2013). This magmatism occurs as 116 numerous dykes, sills and laccoliths, either intrusive along faults and narrow anticlinal 117 ridges or intercalated in the thick Mesozoic sedimentary series (Fig. 1b‒c; e.g., 118 Armando, 1999; Teixell et al., 2003). Basaltic lava flows within wide synclines are also 119 reported (e.g., Guezal et al., 2011; Bensalah et al., 2013; Michard et al., 2013). The 120 magma emplacement was mainly controlled by faults coeval with halokinetic tectonic 121 and moderate syn-schistous deformation, might link to the widespread exhumation 122 resulting from the orogen during the Upper Jurassic‒Early Cretaceous (e.g., Laville and 123 Pique, 1992; Frizon de Lamotte et al., 2008, 2009; Michard et al., 2011; Saura et al., 124 2014; Torres-Lopez et al., 2016; Essaifi and Zayane, 2018). Metamorphic aureole 125 around the intrusions is almost absent at the contact with the country sedimentary 126 rocks, which are mostly faulted (e.g., Fig. 1b‒c). The tectonic boundaries are mostly 127 vertical/sub-vertical faults marked by brecciated formation composed of fragments of 128 igneous intrusive and carbonate rocks. However, some rare normal contacts between 129 the intrusions and the sedimentary sequence are observed with chilled marginal 130 magmatic rocks with carbonates recrystallization that occurred at about 300‒500 °C 131 (Lhachmi, 1992, Armando, 1999; Lhachmi et al., 2001; Essaifi and Zayane, 2018). The 7 132 intrusive complexes include almost all terms of alkaline magma differentiation series 133 (e.g., Armando, 1999; Lhachmi et al., 2001; Zayane et al., 2002; Essaifi and Zayane, 134 2018). The intrusions are made up to 60‒70% of troctolites, gabbros, and 135 monzogabbros, while the remaining volume is represented by differentiated rocks 136 (mostly monzonites and syenites). Depside the fragmentation of the intrusive bodies 137 into several blocks which hampers to infer more information about the primary 138 relationships, the differentiated unit seems to be intrusive into the mafic unit (e.g., 139 Armando, 1999; Lhachmi et al., 2001; Essaifi and Zayane, 2018). The evolution of the 140 magma sequence from mafic to felsic members would be mainly linked to fractional 141 crystallization, whereas crustal contamination during the magma ascent cannot be 142 excluded (e.g., Armando, 1999; Lhachmi et al., 2001, Zayane et al., 2002; Essaifi and 143 Zayane, 2018). The magmatism is interpreted to be formed in a transpressional tectonic 144 setting where the Central High-Atlas experienced a significant uplift and orogenesis 145 during the Jurassic (e.g., Laville and Pique, 1992), or in an asthenosphere mantle 146 upwelling context beneath the orogen with partial melting of enriched upper mantle 147 components without significant crustal extension (e.g., Frizon de Lamotte et al., 2009; 148 Michard et al., 2013; Essaifi and Zayane, 2018). From our field observations, the major 149 apatite deposits of the Moroccan Central High-Atlas —enclosing Anemzi, Tirrhist150 Inougzane, Tassent-Tarstaft, Tarstaft, and Ait Daoud-Toumliline deposits, Figs. 1b, c— 151 are always spatially associated to these Jurassic-Cretaceous alkaline intrusions. The 152 apatite ores form mmto several cm-thick veins without any preferential trending and 153 the current mining is done only in an artisanal way and the only economic opportunities 154 are for mineral collectors. Unfortunately, the limited extension of the exposed apatite 8 155 veins hides more information about the lateral extensions, but the scraping of the 156 apatite ores generated by the mineral collectors could vary between 2 and 40 m long. 157 These apatite-producing localities are isolated and less abundant. The contacts 158 between the apatite veins and their host magmatic rocks are mostly sharp without any 159 clear zonation and the mineralogical compositions of the vein-type apatite ores and its 160 host rocks from different localities are slightly variable, whereas the main components 161 are always alkali feldspars, pyroxene, amphibole, quartz, prehnite, calcite, magnetite, 162 epidote, and titanite (detail below). 163 3. Sampling and analytical methods 164 The five Moroccan Central High-Atlas apatite deposits cited above (Fig. 1) were 165 investigated. For each deposit, we sampled apatite mineralized veins and their direct 166 country magmatic rocks (see Table 1). Apatite crystals have been separated by hand, 167 included in epoxy mounts and polished. These apatites, the gangue-forming minerals, 168 and the country rocks were also studied in polished thin-sections. For petrographical 169 observations, we used standard microscope Leica DM2700P in the Geo-AnalyticalLab 170 of Geology and Sustainable Mining Department, Mohammed VI Polytechnic University 171 (UM6P, Benguerir, Morocco). 172 Seven representative samples of the magmatic rocks hosting the apatite 173 mineralization veins were selected for whole-rock geochemical analyses of major 174 elements and some trace elements performed at the Instituto Andaluz de Ciencias de la 175 Tierra (IACT-CSIC, Granada, Spain). The samples include Anemzi AP45 and AP54, 176 Tirrhist-Inouzane AP10, Ait Daoud-Toumliline AP08 and AP09, Tassent AP27 and 9 177 Tasraft AP61 (Table 1). Rock powders were obtained using a jaw crusher and an agate 178 ring mill. Loss on ignition (L.O.I) was determined by drying the samples at 900°C, and 179 ranges from 0.07 to 5.86 wt.%. (Cl, Cr, Co, Ni, Cu, Zn, Ga, Zr, Sr, Y, Nb) The analyses 180 were performed using a sequential spectrometer Bruker AXS S4 Pioneer equipped with 181 three analyzers (LiF200, OVO-55, PET). For X-ray fluorescence (XRF) major element, 182 Cr, Ni, Zr, Sr and Y analyses, rock powders (c. 1 g) were weighed with di-lithium 183 tetraborate flux, and then the mixture was fused at 1000 ºC for 15 min. Within run 184 precision (% RSD), measured by repeated analyses of USGS BHVO-2 and AGV-2 185 reference materials as external standards was better than 1.5% for all the elements 186 except P (2.7%) (Varas-Reus et al., 2018). The other minor elements and SO3 were 187 analysed in pressed powder pellets in an independent calibration. The analyses of 188 certified reference materials (BIR-1) have been used as external standard and the 189 accuracies are <5% for major elements and <11% for minor elements. 190 In situ analyzes of major and minor elements for minerals were carried out using 191 the Cameca SX100 Electron Microprobe at the Geosciences Montpellier (Montpellier, 192 France). Analyses were obtained using a beam size of 30–40 µm for apatite and 5 µm 193 for the minerals of gangue and country rocks, an accelerating voltage of 15 kV and a 194 probe current of 20 nA. The counting time was fixed at 20 s for major elements and 40 s 195 for F, Cl, and S. Combination of natural —including fluorapatite (F, P), chlorapatite (Cl), 196 orthoclase (K), albite (Al, Na), wollastonite (Si, Ca), forsterite (Mg, Fe), baryte (S), 197 rhodochrosite (Mn)— and synthetic (TiO2 and Cr2O3) standards are used for the 198 calibrations. The apatite and the associated pyroxene and amphibole were analyzed for 199 trace elements using the Laser Ablation Coupled Plasma-Mass Spectrometry (LA–ICP- 16 329 6. Mineralogy 330 6.1. Apatite 331 6.1.1. Major element composition 332 Table 3 shows representative major element analyses of the studied apatites. 333 The two major constituents —i.e., CaO and P2O5— show rather little variations: CaO 334 and P2O5 vary between 53 and 57 wt.% and between 39 and 44 wt.%, respectively. The 335 F content of the Anemzi apatite associated with calcite, feldspars, and magnetite matrix 336 (white gangue; see above) is quite homogenous with about 2 wt.% and low Cl content 337 (0.67–0.90 wt.%). In contrast, the Anemzi apatites associated with dark gangue 338 (amphibole matrix), are rather heterogeneous: their F contents decrease from core to 339 rim (from ~ 1.9 to ~ 0.8 wt.%) while Cl increases (from ~ 0.6 to ~ 2.5 wt.%) (Table 3). All 340 analyzed Tirrhist-Inouzane apatites have relatively homogeneous F and Cl contents (~ 341 1.9–2.4 and ~ 0.34–0.47 wt.%, respectively). Similarly, the apatites from Tassent and 342 Tasraft (fine gems) have F concentrations between 1.7 and 2.8 wt.% and low to 343 moderate Cl contents (0.5–1.47 wt.%). On the other hand, distinct apatite compositions 344 have been yielded by amphibole-bearing apatites from both Tasraft and Ait Daoud345 Toumliline: the first ones show compositions intermediate between those of Cl-rich 346 apatites and those of Cl-F apatites (F contents vary between 0.47 and 1.47 wt.% and Cl 347 between 3.44 and 1.5 wt.%); the second ones are Cl-rich apatites with high Cl (up to 3.3 348 wt.%) and low F (~ 0.6–1 wt.%) contents. If we considered all apatites analyses, they 349 show a good linear trend and progressive compositional variations from the F-rich to Cl350 rich apatites as shown by Fig 8a. 17 351 In all studied apatites, the element substitutions are always very low, and they 352 vary between the different apatite deposits and even between the crystals of the same 353 deposit (see Table 3 and Figs. 8b, c, d). The major substituent element is Si (0.1 < SiO2 354 wt.% < 0.8). Na2O contents can reach 0.55 wt.%, whereas FeOtot and MgO never 355 exceed 0.11 and 0.08 wt.%, respectively. The S contents are particularly variable: SO2 356 varies between 0 and 0.57 wt.%. In detail, the amphibole-bearing apatites from Ait 357 Daoud-Toumiline and Tasraft show quite similar compositions with relatively high Na2O 358 and MgO contents (up to 0.57 wt.% and 0.9 wt.%, respectively), FeOtot contents 359 between 0.04 and 0.11 wt.% and low concentrations of SiO2 (< 0.4 wt.%) and SO2 (~ 0– 360 0.3 wt.%). Similar FeOtot and SiO2 abundances (0.04–0.10 wt.% and 0.18–0.40 wt.%, 361 respectively) have been measured in Tassent apatite but with generally low Na2O 362 (0.06–0.29 wt.%) and SO2 (< 0.04 wt.%) contents and no MgO. The apatites from 363 Tirrhist-Inouzane and Tasraft (fine gems) have MgO below detection limit and they 364 generally contain similar low Na2O (≤ 0.08 wt.%) and FeOtot (≤ 0.04 wt.%) 365 concentrations, whereas SiO2 (up to 0.76 wt.%) and SO2 (0.30–0.57 wt.%) abundances 366 are relatively high. The Anemzi apatites from both white and dark gangues have 367 comparable Na2O (0–0.15 wt.%), MgO (≤ 0.01 wt.%) and FeO tot (0.02–0.09 wt.%) 368 concentrations. However, these two types of apatite are distinguished by their SiO2 and 369 SO2 contents: the apatites from the dark gangue have SiO2 (0.12–0.32 wt.%) and SO2 370 (0–0.04 wt.%) significantly lower than those of the white gangue apatites (0.62–0.70 371 and 0.01–0.42 wt.%, respectively). 372 In addition, the rim–core chemical variations (e.g., SiO2, FeOtot, Na2O, Cl, and F) 373 in some apatite crystals of veins are shown in Figs. 8e–h and Table 3. Despite their 18 374 large size, almost all fluorapatites —e.g., Tirrhist-Inouzane and Anemzi (white 375 gangue)— have no chemical zoning. In contrast, the Anemzi amphibole-bearing 376 apatites (dark gangue) display F-rich cores and Cl-rich rims and the SO2 contents 377 decrease from core to rim while Na2O contents increase (Table 3). 378 To further constrain the relationships between the apatites and their magmatic 379 host rocks, the apatites in Tirrhist-Inouzane syenites have also been analyzed (see 380 Table 3). These igneous apatite grains have similar compositions compared to the 381 Tirrhist-Inouzane, Anemzi (white gangue), Tassent and Tasraft (fine gems) apatite 382 crystals in veins in terms of F (2–2.3 wt.%), Cl (0.5–0.8 wt.%), Na2O (0.1–0.19 wt.%), 383 MgO (< 0.01) and FeOtot (0.04–0.09 wt.%), and they are clearly distinguished from both 384 Ait Daoud-Toumliline and Tasraft amphibole-bearing apatite gems in veins as shown by 385 Figs. 8a–c and Table 3. On the other hand, SiO2 and SO2 abundances in the syenite 386 apatites are similar to those found in the gemmy apatites in veins from Tassent and 387 Anemzi (dark gangue) (SO2 <0.04 wt.%, SiO2 ~ 0.35 wt.%) (Fig. 8d). 388 6.1.2. Trace element composition 389 The trace element abundances of studied apatites given in Table 3 show that 390 these minerals can incorporate a wide range of trace elements and often in large 391 quantities. The most abundant trace elements in the Central High-Atlas gem-quality 392 apatites in veins are REE (mainly La, Ce and Nd), Sr, Y, V, Mn, Th and U, whereas Sc, 393 Co, Rb, Zr, Nb, Ba and Pb contents are very low (< 6 ppm) (Fig. 9 and Table 3). The 394 highest REE concentrations are observed in the Ait Daoud-Toumliline apatites (ΣREE: 395 18190‒22587 ppm), with variable abundances in the Anemzi (white gangue) and 19 396 Tasraft amphibole-bearing apatites (ΣREE: 3846‒17432 ppm), moderate contents in 397 Tassent apatites (ΣREE: 13140‒13554 ppm) and low to moderate values (ΣREE: 398 4823‒8495 ppm) in Tirrhist-Inouzane, Anemzi (dark gangue) and Tasraft (fine gems) 399 apatites. Figs. 10 and 11 show, respectively, their REE patterns and multi-element 400 spider diagrams normalized to chondrite abundances from McDonough and Sun (1995). 401 In all apatites, a pronounced negative Eu-anomaly can be observed (0.11 < Eu/Eu* < 402 0.63; Fig. 10). Almost all analyses show consistent enrichment in light REE relative to 403 heavy REE with (Ce/Yb)n ratios ranging from 12 to 78. 404 The V contents of apatites vary greatly between the different deposits. They are 405 very low (between 3 and 4 ppm) in the Tassent apatites and slightly higher (between 18 406 and 34 ppm) in the Tirrhist-Inouzane and high in Ait Daoud-Toumliline (55–77 ppm) 407 (Fig. 9a). We also find these high values but with variability in both Tasraft apatite types: 408 3‒128 ppm in the amphibole-bearing apatites and 96‒146 ppm in the fine apatite gems. 409 In Anemzi, the V contents of the apatites from the dark gangue are low but relatively 410 constant (9–14 ppm), while these from the white gangue are surprisingly variable and 411 they can even reach high values (0.17‒128 ppm) as in Tasraft. The Mn contents are 412 very variable in the Anemzi apatites from the white gangue (56–397 ppm), while they 413 are relatively stable in those from the dark gangue (274–312 ppm) (Fig. 9b). These 414 latter values are relatively close to that found in the Tassent apatites (222‒241 ppm 415 Mn). The two types of apatite from Tasraft have Mn contents significantly different: 63– 416 96 ppm for the fine gem-crystals and 138–250 ppm for the Amp-bearing crystals. The 417 Mn contents of the apatites from Tirrhist-Inouzane (140–148 ppm) and Ait Daoud418 Toumliline (116–131 ppm) are homogeneous and quite similar. Strontium 20 419 concentrations are moderate to high in all studied apatites (Fig. 9c): 407–686 ppm in 420 Anemzi white gangue, 673–792 ppm in Anemzi dark gangue, 763–781 ppm in Tirrhist421 Inouzane, 720–730 ppm in Tassent, 382–764 ppm in Tasraft and 386–483 ppm in Ait 422 Daoud-Toumliline. Yttrium concentration ranges from 230 ppm to 4118 ppm (Fig. 9e): 423 1567–1672 ppm in Tassent, 1454–1692 ppm in Ait Daoud-Toumliline, 230–4118 ppm in 424 Anemzi (white gangue), 425–947 ppm in Anemzi (dark gangue), 358–696 ppm in 425 Tirrhist-Inouzane and 240–979 ppm in Tasraft. Thorium concentrations are moderate to 426 high and usually ranges from 76 ppm to 530 ppm (Fig. 9f): 90–227 ppm in Anemzi white 427 gangue, 79–112 ppm in Anemzi dark gangue, 140–182 ppm in Tirrhist-Inouzane, 76–83 428 ppm in Tassent, 63–253 ppm in Tasraft and 183–253 ppm in Ait Daoud-Toumliline. 429 Uranium concentrations vary generally from 1.5 to 26 ppm (Fig. 9g). Rather similar high 430 U values (16–26 ppm) are found in the apatites form the Anemzi (white gangue), 431 Tirrhist-Inouzane and Tasraft (fine gems). The lowest U contents (1.5–7 ppm) are found 432 in Tassent and mostly in amphibole-bearing apatites from Anemzi (dark gangue) and in 433 these from Ait Daoud-Toumliline and Tasraft. 434 In addition, if we considered all apatites analyses from all studied veins, they 435 show a rough positive correlation between total REE and Na2O from the less to more 436 Cl-rich apatites as shown by Fig 12. Although some compositional variations can be 437 observed among the different apatites in veins (Figs. 9, 10, 11, 12, Table 3), their REE 438 and multi-trace element patterns overlap and are globally parallel with those of the 439 magmatic host rock apatites investigated here (Figs. 10‒11). 21 440 6.2. The gangue minerals 441 6.2.1. Clinopyroxene 442 Clinopyroxene is exclusively observed in the F-rich apatites veins (e.g., Figs. 3, 4 443 and 6). The crystals are euhedral, up to 3 cm in size and dark to light green in color, 444 sometimes with well-marked zoning corresponding to an alternation of dark and light 445 green domains. Dark and light green grains also occur in almost all magmatic rocks 446 hosting the apatite veins: they appear as euhedral to subhedral grains (0.5 and 2 mm, 447 Figs. 3–4, 6). Major and trace element abundances of clinopyroxenes are given in 448 Tables 4–5. The two major constituents —i.e., SiO2 and CaO— show little variations; 449 the representative values are ~ 49–54 wt.% SiO2 and 20–26 wt.% CaO. In detail, all 450 analyzed dark green clinopyroxene zones and crystals from both apatite-bearing veins 451 and country alkaline rocks are similar and have hedenbergite to Fe-rich augite 452 compositions (Wo44–49, En12–26, Fs27–43, see Fig. 13a and Table 4; e.g., Morimoto et al., 453 1988) with low Mg# [Mg/(Mg + Fe2+) ~ 0.32–0.66], high total FeO content (~ 16–23 wt. 454 %), and moderate to high Na2O values (up to ~ 3 wt. %) and MnO (up to 1.4 wt. %). 455 However, the light green zones and crystals, are mainly diopside in compositions (Wo46– 456 51, En27–41, Fs8–25) with high Mg# ~ (up to 0.95), and low total FeO content (~ 5–14 wt. 457 %), MnO (~ 0.1–0.4 wt. %) and Na2O values (mostly < 2 wt. %). All these clinopyroxene 458 types have similar variable compositional range, with, Al2O3 (~ 0.1–2.6 wt. %), TiO2 459 (mostly < 1 wt. %) and Cr2O3 (0–0.07 wt. %) contents. 460 Fig. 13b shows REE patterns normalized to chondrite values (McDonough and 461 Sun, 1995) of clinopyroxenes from Tasraft and Tassent F-rich apatite veins, and 22 462 Tirrhist-Inouzane host syenites (Table 5). All clinopyroxene REE concentrations are 10– 463 160 times enriched relative to chondrite abundances (ƩREE ~ 80–147 ppm). They have 464 parallel REE patterns [with slight high values in light REE in clinopyroxene from Tasraft 465 apatite veins with (Lan/Ybn) up to 1.3] and show strongly developed negative Eu 466 anomaly (Eu/Eu* ~ 0.19–0.37). General high depletion for Cu, Rb, Ba, Nb, Ta, Pb, Th, 467 and U (< 0.07 ppm) is present for all analyzed clinopyroxenes. They have generally ~ 468 3–12 ppm Sc, 9–23 ppm Co, ~ 17–80 ppm Sr, ~ 26–98 ppm Zn, and ~ 23–166 ppm V. 469 6.2.2. Amphibole 470 Amphibole occurs as main gangue mineral in Cl-rich apatites veins (e.g., Ait471 Daoud-Toumliline; Fig. 5) and as inclusions in almost all the Cl-rich apatites (e.g., Ait 472 Daoud-Toumliline, Tasraft and Anemzi; Figs. 2, 5‒6). It forms euhedral and acicular 473 green to brown crystals up to 2 cm in size. Amphibole also occurs as euhedral to 474 subhedral grains (~ 0.2–1 mm in size) in magmatic host rocks with colors similar to that 475 of amphibole from apatite veins (e.g., Figs. 4–5). Table 4 shows representative major 476 element compositions of the amphiboles. All amphiboles from apatite veins and their 477 host rocks are calcic with high Ca content (~ 1.67–2.10 atoms per formula unit = 478 a.p.f.u), Si (~ 6.87–7.78 a.p.f.u) and Mg (Mg# ~ 0.60 to 0.90) with moderate Fe (mostly 479 ~ 0.54–1.20 a.p.f.u), low to moderate (Na+K) values (0.18–0.90 a.p.f.u) and low Al 480 concentrations (~ 0.1–1 a.p.f.u), characteristics of magnesiohornblende-actinolite and 481 edinite association (Fig. 13c; e.g., Leake et al., 1997). The amphibole trace element 482 abundances have been analyzed for Anemzi apatite veins (dark gangue) and are given 483 in Table 5. The REE abundance in amphiboles range from 46 to 101 ppm. In general, 23 484 their chondrite-normalized REE patterns are remarkably constant (Fig. 13d): they all 485 display a gentle shape, flat pattern (Lan/Ybn ~ 0.7–1) with a pronounced negative Eu 486 anomaly (Eu/Eu* ~ 0.4–0.5) similar to those of clinopyroxenes (Fig. 13c). They also 487 have 32–650 ppm Cr and 364–405 ppm V, and Sc, Co, Zn, Sr, Y, Zr, and Nb 488 abundances are less than 68 ppm, with low values (< 2 ppm) for Cu, Rb, Ba, Ta, Pb, 489 Th, and U (Table 5). 490 6.2.3. Feldspars 491 Feldspars are commonly present in all studied vein-type apatite ores. Table 6 492 displays representative feldspar analyses. In all apatite ores and magmatic host rocks 493 (Figs. 2–6), albite is the most abundant feldspar: the crystals are euhedral to subhedral 494 and of centimeter size (up to 5 cm). Their compositions are similar, and they correspond 495 almost pure to albite (Ab ~ 90 to 100). K-feldspar is also present but less abundant than 496 albite: the crystals are euhedral to subhedral in the veins, subhedral to anhedral in 497 magmatic host rocks and with sizes varying between 0.1 and 1 cm. Their compositions 498 are between Or94 and Or99. Anorthoclase compositions (Ab ~ 70–75, An ~ 1–3, Or ~ 499 23–29) are exclusively found in the magmatic host rocks (see Table 6). 500 6.2.4. Other minerals 501 Magnetite is found in all apatite veins: the crystals are euhedral to subhedral with 502 up to 1 cm in size. This oxide is also observed in magmatic host rock as subhedral to 503 anhedral grains up to 2 mm in size (e.g., Fig. 3). Representative compositions are 504 reported in Table 7. Analyses of magnetite indicate moderate to high FeOtot contents of 24 505 about ~ 91–95 wt.% in the apatite veins. In the magmatic host rock, the magnetite 506 composition varies from magnetite (FeOtot ~ 90–95 wt.%) to Ti-magnetite [FeOtot ~ 87 507 wt.% and TiO2 ~ 3 wt.%], sometimes with ilmenite exsolutions (TiO2 ~ 45–54 wt.% and 508 FeOtot ~ 46‒50 wt.%) (Table 7). 509 Titanite is also present in both apatite veins and host rocks and the analyses are 510 reported in Table 7. It is associated with apatite veins and forms greenish yellow 511 euhedral crystals up to 2 cm in size; in magmatic host rocks, it occurs as dark brown to 512 colorless euhedral to subhedral crystals (0.5–2 mm). Titanite from veins has relatively 513 homogeneous contents in SiO2, TiO2, and CaO (~ 29.5, ~ 41.5 and 29 wt.% 514 respectively), whereas in the host rocks the titanite compositions are more variable: 20– 515 32 wt.% SiO2, 29–54 wt.% TiO2 and 19–29 wt.% CaO. On the other hand, rutile is 516 observed and analyzed exclusively in the magmatic host rock, forming small euhedral to 517 suhedral dark-red grains (up to 0.2 mm in size) and its analyses are also given in Table 518 7. 519 In the apatite mineralized veins, epidote (pistachite) occurs as green to greenish 520 yellow euhedral crystals up to 3 cm in size, whereas in the alkaline host rocks it appears 521 as yellow subhedral to anhedral grains of small size (0.02–0.5 cm). As shown in Table 522 8, the epidote compositions are similar in veins and their host rocks: 36‒38 SiO2 wt.%, 523 18.50‒22.70 Al2O3 wt.%, 21‒25 CaO wt% wt.%, with slight variation FeOtot content (~ 524 10–18 wt.%). 525 Prehnite is observed in the Tassent and Ait Daoud-Toumliline apatite mineralized 526 veins. It occurs as radiating fibrous aggregates and is commonly translucent with 25 527 colorless to white greenish color (up to 4 cm in size, e.g., Figs. 4‒5). Prehnite can occur 528 also in the magmatic host rocks where it appears both as veinlets and as microcavity 529 filling products. Analyses of this calcic mineral (~ 28 wt.% CaO) are shown in Table 8 530 and are homogenous and similar in both mineralized veins and host alkaline rocks. 531 7. Discussion and conclusions 532 According to our field observations, and petrological, mineralogical, and 533 geochemical study, all Moroccan Central High-Atlas vein-type apatite ores (mostly gem534 quality) are hosted by differentiated igneous intrusions ranging from syenite‒quartz 535 monzonite (dominant) to monzodiorite‒gabbrodiorite (Fig. 7 and Table 2). The spatial 536 association with Mesozoic (Jurassic-Cretaceous) alkaline/transitional rocks is typical of 537 phosphates derived from silicate alkaline‒carbonatite complexes (e.g., Simandl and 538 Paradis, 2018 and references therein). However, carbonatites have not been found in 539 the studied area, although the ore bodies are not so far to the well-known Tamazeght 540 Eocene carbonatite complex (see Fig. 1b). 541 Our new data lead us to distinguish two types of apatite ores: (i) high gem-quality F-rich 542 apatites associated with clinopyroxene (diopside to hedenbergite-augite) and (ii) gemmy 543 Cl-rich apatites mainly formed in an amphibole matrix (hornblende‒edenite to actinolite) 544 (Figs. 2‒6, 8, Table 3). Both ore bodies are also associated with feldspars (albitic 545 plagioclase dominant and orthoclase), quartz, magnetite, titanite, epidote, calcite, and 546 prehnite. All studied apatites of both deposit types have similar REE and multi-trace 547 element patterns which suggest that they could have a relatively close origin (Figs. 9, 548 10, 11, Table 3). A comparison between the Central High-Atlas apatites and those of 32 682 from the early Late Cretaceous of southeastern Morocco: An overview. J. Afr. 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