Multi-stage emerald formation during Pan-African regional metamorphism: The Zabara, Sikait, Umm Kabo deposits, South Eastern desert of Egypt

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Journal of African Earth Sciences 50 (2008) 168–187 www.elsevier.com/locate/jafrearsci

Multi-stage emerald formation during Pan-African regional metamorphism: The Zabara, Sikait, Umm Kabo deposits, South Eastern desert of Egypt G. Grundmann a, G. Morteani a

b,*

Lehrstuhl fu¨r Ingenieurgeologie, Technische Universita¨t Mu¨nchen, Arcisstr. 21, 80333 Mu¨nchen, Germany b Gmain Nr. 1, 84424 Isen, Germany Received 10 May 2006; received in revised form 5 March 2007; accepted 13 September 2007 Available online 23 October 2007

Abstract The genesis of gem-quality deep green emeralds of Zabara, Sikait and Umm Kabo (South Eastern Desert, Egypt) is to date a controversial topic. The emerald-bearing biotite schists and quartz lenses are interpreted alternatively as a product of (i) thrust-fault-shear zone – controlled large scale alkali-metasomatism driven by post-magmatic fluid flow or of (ii) a large scale interaction between syntectonic pegmatitic magma or hydrothermal fluids with pre-existing basic to ultrabasic rocks, or of (iii) a syn- to post-tectonic regional metamorphism and small scale blackwall metasomatism. Detailed microstructural and chemical analyses of the Egyptian emeralds and their host rocks show that three generations of beryl can be distinguished: a colourless pegmatitic beryl; a pale green Cr-poor beryl crystallized from pegmatite-related hydrothermal fluids; and a deep green Cr- and Mg-rich emerald. The crystallization of the Cr- and Mg-rich emerald was controlled by the very local availability of Cr, Mg and Be-rich metamorphic fluids during the Pan-African tectonothermal event. Emerald-rich quartz lenses demonstrate that those fluids locally did mobilize quartz, too. The pale green emeralds found within the pegmatites in association with colourless beryl are the product of a mobilization of colourless pegmatitic beryl and/or phenakite by late pegmatitic fluids slightly enriched in Cr by an interaction with the Cr-rich country rocks. The late pegmatitic fluids are typically Na-rich as is demonstrated by the pervasive albitization of the pegmatites. The complex interplay of magmatic and regional metamorphic events during the genesis of the Egyptian emeralds/beryls makes it impossible through stable oxygen isotope data to relate their genesis to the one or the other event. Ó 2007 Published by Elsevier Ltd. Keywords: Beryl; Egypt; Emerald; Metamorphism; Pegmatite; Pan-African

1. Introduction The emeralds of the South Eastern Desert of Egypt (SEDE) have been exploited by the world’s first emerald mines (Harrell, 2004). They deserve a special interest because the excellent outcrops of a fully arid area give the opportunity to investigate in detail the much debated genesis of the so called ‘‘schist-type’’ emerald mineralizations. *

Corresponding author. E-mail addresses: [email protected] (G. Grundmann),gmorteani@ gmx.de (G. Morteani). 1464-343X/$ - see front matter Ó 2007 Published by Elsevier Ltd. doi:10.1016/j.jafrearsci.2007.09.009

MacAlister (1900) and Hume (1934) recognized that the entire rock succession of emerald-bearing ‘‘schists and granitoid gneisses’’ of the SEDE has been subjected to a metamorphism on a regional scale and that the emeralds are therefore of metamorphic origin. On the contrary, Fersman (1929) concluded by analogy with his theory on the genesis of the emerald deposits in the Urals, that the emerald mineralization in the SEDE is produced by desilicification of pegmatites due to the mutual interaction of a pegmatite-derived fluid phase with a melanocratic, ultrabasic rock complex. Sinkankas (1981) accepted the genetic model of Fersman (1929) and classified the Egyptian emerald occurrences as ‘‘schist-type exometamorphic’’ and

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Fig. 1. (a) Geological sketch map giving the position of the study area near the northern end of the Nubian shield; (b) simplified geological sketch map from Kro¨ner et al. (1987) of the area surrounding the emerald deposits of Zabara, Sikait and Umm Kabo along the Nugrus thrust; (c) profile showing the position of the emerald deposits at the Nugrus thrust, which is the basis of the imbricated volcano-sedimentary sequence (VSS) found on top of the gneisses and foliated granitoids of the Hafafit uplift.

unrelated to any regional tectono-metamorphic event. Grundmann and Morteani (1993a,b) emphasized, in contrast the regional-metamorphic genesis of the emeralds in the schist-type deposits of Egypt. Further information on geology, tectonics, geochemistry, fluid inclusions, oxygen stable isotopes, gemmology and archaeology of the emerald deposits and their host rocks can be found in Hassan and El-Shatoury (1976), Soliman (1986), Grundmann and Morteani (1993a,b), Mohamed and Hassanen (1997), Hassan (1998), Abdalla and Mohamed (1999), Harrell (2004) and Surour (1995). The term ‘‘schist-type’’ refers to the biotite schists which are the main host of emerald mineralizations and the main part of more or less foliated metasomatic ‘‘blackwall zones’’ developed in a tectonic melange between ultrabasic serpentinite bodies and alkali-alumina-rich pegmatitic/granitic rocks such as meta-pelites, meta-volcanic rocks, metagranites, meta-aplites, meta-pegmatites and meta-greisen (Grundmann and Morteani, 1989). The blackwall zones consist, under greenschist- to epidote–amphibolite-facies metamorphic conditions, of a sequence of: talc schist, +/  actinolite schist, +/ chlorite schist, biotite/phlogopite schist (Phillips and Hess, 1936; Curtis and Brown, 1969, 1971; Brady, 1977; Grundmann, 1983; Grundmann and Morteani, 1989; Grundmann and Giuliani, 2002). Only very few emerald schist-type deposits have been investigated from a microstructural point of view (among others: Grundmann, 1983; Grundmann and Morteani,

1989; Laurs et al., 1996; Zwaan, 2006). The present work complements the information existing on the genesis of the Egyptian emeralds with a new microstructural study and new geochemical data, and gives a complete picture of the poly-metamorphic emerald genesis in the context of an ophiolitic melange zone generated and overprinted by the multistage Pan-African tectono-thermal event. For clarity and to avoid a discussion of ‘‘what is an emerald’’ in the following, the macroscopically colourless beryl retains this name, whereas beryl that shows the faintest green colour up to deep green colour will be called emerald (Schwarz and Schmetzer, 2002). 2. History of emerald mining in the SEDE The historical importance of the here studied emerald mineralizations (Giuliani et al., 2000) justifies to start with a short introduction into the history of emerald mining in the SEDE. Until the emeralds of Colombia were found by the Spaniards, the legendary Egyptian ‘‘Cleopatra’s’’ mines were probably the most important source for emeralds for Europe and the Near East and Middle East (Sinkankas, 1981; Jennings et al., 1993; Harrell, 2004). The mines are located in the SEDE about 30–40 km south of the Red Sea harbour of Marsa el-‘Alam. The size of the emeralds in ancient jewels rarely, if ever, exceeds 3 cm in length. The weight of the cloudy and fractured ‘‘Egyptian emeralds’’ is mostly less then 2 g. In this context it has to

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be considered that many of the beads and amulets supposed to be made from emerald are fakes made from much more easily available green feldspar (Lucas and Harris, 1962). The roman writers Strabo (Geographica XVI: IV; p. 20; XVII: I, p. 45) and Plinius the Elder (Historia Naturalis XXXVII: pp. 16–18) are the first to mention, from Egypt, ‘‘Smaragdus Mons’’. Therefore it has to be supposed that the emerald occurrences of the SEDE were known and mined for the first time in the Graeco-Roman period between 30 BC and 312 AD. No information exists about emerald mining in early Egyptian times. A very consistent amount of ruins of Graeco-Roman affinity close to the mines in the Wadi Nugrus at Sikait, Wadi Gemal, Gebel Zabara and Umm Kabo in an otherwise desert area confirm such a Graeco-Roman emerald mining in southern Egypt. The best preserved ruin is a temple excavated in a talc body near Sikait. It shows columns of the Doric type and inscriptions in Greek language of late roman age (R. Klemm, pers. comm.). From Graeco-Roman times till early Arabic times the emerald mining was dormant. According to Maqrizi (in Murray, 1925), the mines were definitely closed in 1342. Recent attempts to reactivate emerald mining did not succeed due to the poor quality of the emeralds, and very expensive logistics (Lucas and Harris, 1962).

3. Geological setting The emerald deposits of Zabara, Sikait and Umm Kabo are found in the Red Sea Hills of the SEDE, near the northern end of the Nubian shield along the Nugrus thrust, on the eastern border of the gneisses and foliated granitoids of the Hafafit uplift (Fig. 1a–c). The gneisses and foliated granitoids forming the core of the Hafafit uplift are summarized in the following as meta-granitoids. The country rock of the emeralds is a volcano-sedimentary sequence (VSS) that is a tectonic melange of serpentinite, talc schist, biotite schist, chlorite schist, graphitic schist, graphitic quartzite, mica schist, amphibole schist, amphibole gneiss and amphibolite. The different terranes now forming the VSS converged and became amalgamated between 780 and 680 Ma, with over-thrusting of the metagranitoid cupolas of the Hafafit uplift in consequence of the collision of east and west Gondwana (Kro¨ner et al., 1987). It is open to debate whether the meta-granitoids of the Hafafit uplift are part of the African plate (El-Ramly and Akaad, 1960; El-Ramly, 1972) or whether those rocks belong to the Nubian shield and were metamorphosed at about 710–680 Ma during the onset of the Pan-African tectono-thermal event (Kro¨ner et al., 1987, 1994). The ophiolites inter-layered in the VSS document the rifting of Rodinia and the formation of the Mozambique Ocean at about 900–750 Ma. The core complexes, like the Hafafit uplift, were produced by extensional detachment systems at 600–560 Ma (Blasband et al., 2000). 40Ar/39Ar determi-

nation of syntectonic muscovites within shear zones gave Neo-Proterozoic Pan-African ages of 595.9 ± 0.5 and 588.2 ± 0.3 Ma (Fritz et al., 1996). The gold-baryte mineralization that is found predominantly in the VSS formed around 650 Ma, whereas the more important shear zone hosted gold mineralization formed at about 550 Ma (Elsamani et al., 2001). According to Kro¨ner et al. (1987), four main rock associations make up the bulk of the Hafafit uplift. These are: Association 1: quartzites and quartzitic schists deposited on a continental shelf. Association 2: tholeiitic as well as calc-alkaline volcanic and volcano-clastic assemblages with island arc characteristics. Association 3: dismembered ophiolitic complexes and melanges that occur as large nappes, and Association 4: large masses of syntectonic foliated granitoids of I type and post-tectonic non foliated A-type granites. Kro¨ner et al. (1987) reported, for the syntectonic granites, ages between 850 and 660 Ma and for the undeformed ones, ages around 670 Ma. The Sikait S-type leucogranite with the associated pegmatitic veins (Mohamed and Hassanen, 1997) belongs to association 4. The volcano-sedimentary sequence (VSS) of the study area can be considered as a tectonic melange of associations 1–3. The granitoids of association 4 (Hafafit gneisses and Hafafit foliated granitoids, Fig. 1b) form large metagranitoid cupolas in the study area. The primary magmatic contacts between the metagranitoid cupolas of the Hafafit uplift and the VSS are lost due to strong overprinting by the Nugrus thrust. The volcano-sedimentary sequence found on top of the meta-granitoid cupolas consists in the study area of meta-sediments (garnet–mica schist, carbonate-mica schist, graphite quartzite, tourmaline–chlorite–mica schist) and metabasites (garnet–amphibole–biotite schist, amphibole–biotite schist with lawsonite(?) pseudomorphs, carbonate–albite–biotite schist, amphibole gneiss, garnet amphibolite). The roof zone of the Hafafit meta-granitoid cupolas is marked at the contact to the VSS by boudinaged and foliated meta-pegmatite, meta-greisen, quartz veins and meta-aplite veins intermingled with rocks of the VSS. The main emerald mineralization is bound to quartz-veined biotite schists of the blackwall zones, that are found at the contact of elongated lenslike talc bodies and meta-pegmatites on top of the meta-granitoid cupolas. The most intense emerald mining activity is localized therefore along the roof zone of the meta-granitoid cupolas. The blackwall zones are produced by a metasomatic reaction driven by contact and regional metamorphism between the ultrabasic and the surrounding quartzo-feldspathic rocks (Phillips and Hess, 1936; Curtis and Brown, 1969; Grundmann and Morteani, 1989).

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4. Studied localities

4.2. Zabara

4.1. Sikait

At Zabara the old diggings follow very closely the contact between the meta-granitoids (HMG) of the Hafafit uplift and the overlying VSS The latter consists of a tectonic melange of meta-basic volcanic rocks and garnetiferous meta-sediments (Fig. 5). The host rock of the emerald is biotite schist, that belongs together with carbonate, chlorite, actinolite and talc schist to the blackwall envelope of serpentinite bodies. Subordinately a rich mineralization of deep green emerald can be found in small quartz lenses embedded in biotite schist. Thick quartz veins and bodies which are not embedded in biotite schists of a backwall zone are typically beryl/ emerald-barren. Lens-like colourless beryl and/or aquamarine mineralizations are found scattered along the foliation of the gneisses of the Hafafit uplift near the roof zone. Along the roof of the Hafafit meta-granitoid cupola remarkable polymict fault breccias can be found, which consist of fragmented talc bodies cemented by a network of biotite schists and quartz veins.

In the Gebel Sikait area dismembered and boudinaged antigorite serpentinite bodies with their blackwall envelope form at least three rather persistent horizons within the VSS (MacAlister, 1900; Hume, 1934; Basta and Zaki, 1961; Grundmann and Morteani, 1993a). Remarkable local enrichment of green emerald can be found: (a) within biotite schist of the blackwall zone at the contact between boudinaged meta-pegmatite, meta-aplite and meta-greisen veinlets of up to 10 cm in thickness (Fig. 2), or (b) within strongly folded quartz layers and lenses of up to 10 cm in thickness found in the biotite schist of the blackwall zones (Grundmann and Morteani, 1993b) (Fig. 3). Fig. 4 shows the clustering of emeralds along the border between a meta-pegmatite and the blackwall biotite schist. The strongly foliated meta-pegmatite shows aggregates of fine-grained recrystallized feldspar and elongated aggregates of fine-grained recrystallized quartz. The quartz aggregates represent the distorted equivalents of the original granophyric quartz texture. In spite of a rather pervasive strong deformation and metamorphic overprint, the meta-pegmatite veins display locally relics of their typical graphic quartz–feldspar texture (granophyric and/or myrmekitic texture) and subhedral colourless beryl and/or very pale green subhedral to euhedral emerald crystals. In almost all cases, beryls and emeralds are macroscopically ruptured and in some cases bent.

4.3. Umm Kabo At Umm Kabo the distribution of the emerald mineralization and consequently of the diggings over an area of about 4 km2 outlines the sub-horizontal contact between the meta-granitoids of the Nugrus uplift and the talc schist of the VSS with their emerald- bearing blackwall zones. The NW/SE orientation of the diggings follows the NW/ SE strike of elongated talc bodies (Grundmann and Morte-

Fig. 2. Schematic sketch of a specimen of a fractured and boudinaged lens of a meta-pegmatoid vein (#6905) set in biotite schist from Sikait. The emerald porphyroblasts (black) grew post-tectonically with respect to the intrusion and boudinage of the granophyric pegmatite on the surface and in the pressure shadows of the boudins. Picture redrawn from Grundmann and Morteani (1993a,b).

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Fig. 3. Schematic sketch of emerald-bearing strongly folded quartz lenses embedded in a biotite schist from Sikait. The emeralds (black) are found within or on the border of the quartz lenses. Picture redrawn from Grundmann and Morteani (1993a,b).

Fig. 4. Microphotograph of the contact between biotite schist and meta-pegmatoid vein (#6905). The emeralds crystallized along the contact between the biotite schist and the granophyric meta-pegmatite. Plane polarized light.

ani, 1993a, p. 36). Rich deep green emerald mineralization can only be found in biotite schist in contact with small quartz lenses, as is the case in Zabara and Sikait. Metapegmatite lenses are locally rich in colourless beryl, very light green emerald of a maximum length of 10 cm and in fragmented fine-grained tourmaline. Typical and very frequent in this area are sub-horizontal meta-greisen veins with a quartz-core zone and a border zone consisting of white mica books up to 20 cm in length

and rare feldspar (Grundmann and Morteani, 1993a, p. 38). The contact between the mica border zone of the meta-greisen veins and the talc schist is marked by a zone of fuchsite-bearing biotite schist up to 30-cm thickness. It is remarkable that in and around such meta-greisen veins no emerald mineralization could be found. In all three deposits the emerald mineralization is rather strictly bound to a specific type of rock, but also in favourable rock sequence the mineralization is very localized and

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Fig. 5. NE flank of the Gebel Zabara showing the alignment of the adits along the tectonized sub-horizontal roof zone at the contact between the VSS (top) and the foliated granitic gneiss (MHG) of the Hafafit uplift (bottom).

distributed very heterogeneously. The local enrichment gives to the mineralization its nugget-like character. In all three old mining areas deep green, inclusion- and fracture-free gem-grade emeralds are extremely rare. 5. Chemical, whole rock and mineral composition 5.1. Analytical methods Major and selected trace elements of 40 rock samples were analyzed by a fully automated Siemens XRF using Li-tetraborate glass disks, at the Technische Universita¨t Mu¨nchen, Germany. Electron-microprobe analyses were obtained at the Institut fu¨r Geowissenschaften, University of Kiel, Germany, using a Cameca ‘‘Camebax Microbeam’’ electronmicroprobe with four wavelength-dispersive spectrometers and an online PaP-correction program (Pouchou and Pichoir, 1984) using fluor-apatite (Ca, P, F, Cl), wollastonite (Ca.Si), synthetic corundum (Al), synthetic periclase (Mg), Fe-metal (Fe), synthetic SrCuSi4O10 (Sr), pyrite (S) for major elements and different glass standards for trace elements. Operating conditions were 15 kV accelerating voltage and 15 mA beam current. 5.2. Whole rock major and trace element analyses Major and trace element content and the CIPW normative composition of selected meta-granitoids of the Nugrus uplift and meta-sediments, meta-basites, and blackwall rocks of the VSS are given in Table 1.

The CIPW normative composition of the gneisses of the Hafafit uplift is rhyolitic to trachydacitic. More basic compositions are missing. The amphibolites of the VSS show an andesitic to basaltic composition (Fig. 6). The beryllium contents in all rocks of the Zabara, Sikait and Umm Kabo areas are above the Be abundance of the earths crust, of 2 ppm (Emsley, 1989) (Table 1). Soliman (1986) already noted that the emerald mines are located within a geochemical province of beryllium. He refers the elevated Be content of granite, gneiss, schist, quartz veins and pegmatite dykes to pneumatolytic hydrothermal processes associated with the emplacement of stanniferous Stype granites. In the study area the highest mean Be content can be found in the tourmaline–chlorite–mica schist (43.9 ppm), biotite–actinolite schist (35.6 ppm), meta-aplitic veins and meta-pegmatites (38 ppm) and biotite schist (13.4 ppm) (Table 2). By contrast the meta-basites and amphibolites, garnet–mica schist, chlorite schist and talc schist, have low Be (<6.4 ppm) contents (Table 2). The mean Cr content decreases from the talc schist (3079 ppm) to the biotite–actinolite schist (1700 ppm) to the biotite schist (607 ppm) and to the tourmaline schist (500 ppm). All other rock types have less than 500 ppm Cr. The low Be content of the emerald-bearing biotite schist is explained by the fact that particular care was taken to eliminate, by hand picking from the sample, all material including emeralds. The determined mean Be content of 13.4 ppm corresponds closely to the mean Be content in biotite schist of 9.6 ppm, as given by Grundmann and Morteani (1989) and to the maximum possible Be uptake of 20 ppm in dark mica (Grew, 2002). The trace element

78.80 0.10 11.50 1.07 0.01 0.75 0.07 0.14 3.98 d.l. 1.40

97.82

d.l 354 9.8 d.l. 24 d.l. d.l. 47 d.l. 61 28 d.l. 203 d.l. 60 31 d.l. 25 d.l. 12 36 188 80 243

1401.8

SiO2 (%) TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI

Total

As (ppm) Ba Be Co Cr Cs Cu Ga Mo Nb Ni Pb Rb Sc Sn Sr Ta Th U V W Y Zn Zr

Total

2686.3

d.l. 1493 3.3 20 59 d.l. 107 16 14 10 75 d.l. 34 33 d.l. 439 d.l. d.l. d.l. 117 33 31 44 158

99.26

65.3 0.78 16.20 5.82 0.08 1.96 1.07 3.50 2.11 0.14 2.30

1093.5

d.l. 139 3.5 29 432 d.l. 11 8 d.l. d.l. 133 d.l. 45 d.l. d.l. 142 d.l. d.l. d.l. 111 d.l. d.l. 40 d.l.

101.89

46.70 0.27 12.20 6.11 0.11 4.87 14.10 2.14 1.49 d.l. 13.90

1164

d.l. d.l. 4.5 50 56 d.l. 14 23 d.l. 9 63 d.l. 16 39 d.l. 34 d.l. d.l. d.l. 276 32 23 194 335

99.44

27.90 1.77 22.10 12.10 0.08 23.70 0.64 d.l. 0.13 0.42 10.60

ZAB 6917

ZAB 6916

ZAB 6909

ZAB 6915

Phlogopite– talc schist

Muscovite– Garnet– Quartz– quartzite mica albite– schist chlorite schist

2468.6

d.l. 578 9.6 34 540 d.l. 85 11 d.l. 9 265 d.l. 8 45 d.l. 432 d.l. d.l. d.l. 174 d.l. 36 83 159

98.89

44.80 0.98 22.50 9.00 0.09 10.00 1.01 4.64 0.59 0.18 5.10

ZAB 6919 74.00 0.14 13.90 1.10 0.03 0.43 1.13 6.80 0.72 d.l. 0.90

ZAB 6930 74.90 0.22 13.30 1.93 0.03 0.17 0.91 3.92 2.48 d.l. 1.20

ZAB 6931 57.00 0.39 15.90 6.65 0.11 6.51 6.25 4.77 0.26 d.l. 1.20

ZAB 6932 50.90 1.39 14.90 9.24 0.13 5.60 7.16 5.05 0.87 0.27 2.80

d.l. 58 10.1 d.l. 26 d.l. d.l. 35 15 42 26 d.l. 29 d.l. d.l. 70 d.l. d.l. d.l. 9 d.l. 69 12 127

d.l. 98 6.7 d.l. 20 d.l. 13 30 9 54 30 d.l. 49 75 d.l. 75 d.l. 21 d.l. 11 d.l. 109 32 294

d.l. 152 16.2 d.l. 23 d.l. 23 33 9 54 29 d.l. 229 d.l. 125 56 d.l. 31 d.l. 10 261 83 93 301

d.l. 44 3.5 38 81 d.l. 434 14 d.l. d.l. 65 d.l. 14 d.l. d.l. 204 d.l. d.l. d.l. 188 d.l. d.l. 37 22

d.l. 505 6 34 131 42 9 18 10 8 98 d.l. 46 d.l. d.l. 780 d.l. d.l. d.l. 197 d.l. 17 99 107 3504.1

d.l. 62 7.1 67 1544 d.l. 22 8 13 9 1391 d.l. 14 d.l. d.l. 174 d.l. d.l. d.l. 82 d.l. 15 56 40

98.48

52.40 0.33 6.10 7.23 0.31 16.70 11.20 1.75 0.28 0.08 2.10

SIK SIK 6868 A 6868 B

101.69 99.15 99.06 99.04 98.31

76.80 0.05 12.70 0.39 0.01 0.14 0.20 6.42 0.48 d.l. 4.50

ZAB 6929

5167.1 528.1 926.7 1528.2 1144.5 2107

136 101 7.1 100 1809 34 26 8 32 d.l. 1575 d.l. 157 d.l. d.l. 745 d.l. d.l. d.l. <8 d.l. 37 400 d.l.

99.00

30.10 0.07 2.39 7.82 0.29 27.59 6.24 d.l. 2.00 d.l. 22.50

ZAB 6920

881

d.l. 48 6 14 77 d.l. 20 d.l. 34 d.l. 190 d.l. 8 d.l. d.l. 205 d.l. d.l. d.l. 47 d.l. 32 45 155

100.00

78.10 0.38 4.64 2.93 0.06 5.45 0.94 2.17 0.04 0.09 5.20

SIK 6870

d.l. 874 5.3 d.l. 19 d.l. 17 28 23 9 25 d.l. 40 d.l. d.l. 642 d.l. d.l. d.l. 19 30 13 82 241

98.66

70.70 0.28 15.40 2.11 0.04 0.76 1.46 6.13 1.16 0.12 0.50

SIK 6874 36.60 1.08 22.30 9.03 0.06 16.70 0.68 0.77 6.38 0.28 3.80

SIK 6882

d.l. 58 8.4 70 212 d.l. 14 19 d.l. 16 151 d.l. 12 42 d.l. 20 d.l. d.l. d.l. 215 d.l. 40 156 322

d.l. 1056 63.3 45 869 42 d.l. 11 d.l. 12 479 d.l. 291 36 d.l. 145 d.l. d.l. d.l. 200 d.l. 24 92 177

100.03 97.68

28.00 1.26 22.50 9.77 0.16 27.00 0.09 d.l. 0.15 d.l. 11.10

SIK 6881

1843.1 2067.3 1355.4 3542.3

d.l. 746 7.1 d.l. 17 d.l. 9 24 12 8 24 d.l. 33 d.l. d.l. 851 d.l. d.l. d.l. 31 d.l. 21 60 418

98.91

67.40 0.47 17.00 2.25 0.04 1.09 1.81 6.96 0.98 0.21 0.70

SIK 6873 64.20 0.77 15.30 6.80 0.20 1.93 3.28 4.00 1.47 0.10 1.10

SIK 6892

d.l. 347 4.9 18 60 d.l. 61 18 12 12 43 d.l. 42 d.l. d.l. 270 d.l. d.l. d.l. 110 d.l. 32 81 117 3428.4 1227.9

d.l. 781 19.4 47 514 97 d.l. 50 d.l. 66 541 d.l. 808 33 43 82 d.l. d.l. d.l. 107 d.l. 24 124 92

98.23 99.15

46.10 0.49 14.50 8.28 0.10 16.30 0.51 1.63 8.37 0.05 1.90

SIK 6890

607

d.l. 75 10 d.l. 24 d.l. 43 21 28 14 48 d.l. d.l. d.l. d.l. 298 d.l. d.l. d.l. d.l. d.l. 19 d.l. 27

99.04

72.50 0.07 15.40 0.61 0.01 1.49 0.48 7.93 0.05 d.l. 0.50

SIK 6897

d.l. 649 10 d.l. 26 d.l. 22 19 d.l. d.l. 42 d.l. 126 d.l. d.l. 220 d.l. d.l. d.l. 64 d.l. 40 57 171

100.10

61.00 0.56 13.10 4.68 0.24 2.58 6.40 3.41 2.99 0.14 5.00

SIK 6899

5147.2 1446

d.l. 665 13.2 60 1471 73 d.l. 48 11 63 1068 d.l. 955 36 44 37 d.l. d.l. d.l. 141 d.l. 22 440 d.l.

97.33

40.70 0.76 14.30 11.00 0.16 18.70 0.43 0.43 9.75 d.l. 1.10

SIK 6898

Albite– Biotite–Meta- Meta- Meta- Meta- Amphi- Carbonate– Tremolite– Meta- Meta- Chlorite Turmaline– Biotite Garnet— Meta- Biotite Carbonate– chlorite– talc granite granite granite diorite bolite aktinolite talk granite granite schist biotite schist mica granite scshist biotite talc schist schist quarzite schist schist gneis schist

1579.1

d.l. 580 17.1 11 53 d.l. 107 17 20 9 117 d.l. 123 d.l. d.l. 197 d.l. d.l. d.l. 85 d.l. 41 40 162

99.61

68.70 0.67 13.70 3.98 0.08 3.00 0.77 5.39 1.84 0.08 1.40

SIK 6900 67.40 0.01 18.00 0.41 0.01 0.61 1.75 7.86 0.27 d.l. 0.50

73.50 0.02 15.10 0.50 0.05 0.11 0.67 4.72 4.11 d.l 0.60

58.50 0.28 17.90 2.81 0.05 6.32 0.84 7.62 2.75 0.20 0.70

52.20 0.06 2.81 6.15 0.25 20.00 11.50 0.58 1.33 d.l. 2.20

UMK UMK UMK 6851 6854 6855

84 48 106 d.l. 19 d.l. d.l. 57 d.l. 1993 259 d.l. 29 d.l. d.l. 759 218 604 164 d.l. 85 183 d.l. 8320

d.l. 1735 3 d.l. 33 d.l. d.l. 23 d.l. 11 29 37 60 d.l. d.l. 318 d.l. d.l. d.l. d.l. d.l. 8 40 22 5947.9 12928 2319

d.l. 45 2.9 87 4348 d.l. 40 d.l. 23 d.l. 1208 d.l. 50 d.l. d.l. 21 d.l. d.l. d.l. 26 d.l. d.l. 97 d.l.

d.l. 51 38.5 74 1856 d.l. d.l. d.l. d.l. d.l. 1502 d.l. 80 d.l. d.l. 104 d.l. d.l. d.l. 31 d.l. 29 85 d.l. 2156.3 3850.5

d.l. 184 14.3 24 517 32 d.l. 55 16 42 454 d.l. 276 d.l. d.l. 230 d.l. d.l. d.l. 50 d.l. 39 120 103

96.13 96.82 99.38 97.97 97.08

57.80 0.05 1.44 6.09 0.04 25.70 0.11 0.21 0.69 d.l. 4.00

SIK SIK 6904 6905

2455.1

d.l. 366 48.1 d.l. 92 d.l. d.l. 173 19 113 303 33 700 30 257 91 d.l. d.l. d,l, 85 d.l. 9 136 d.l.

97.78

50.30 0.32 26.00 2.78 0.04 3.98 0.27 2.44 8.25 d.l. 3.40

d.l. 373 5.1 13 140 d.l. 37 19 d.l. 14 59 d.l. 74 30 d.l. 100 d.l. d.l. d.l. 95 d.l. 43 75 184

99.17

66.90 0.68 15.30 5.01 0.11 1.71 1.66 5.12 1.86 0.12 0.70

4689.9 1261.1

d.l. 291 11.9 65 1030 81 d.l. 72 d.l. 84 1137 d.l. 1161 d.l. 63 18 d.l. 21 d.l. 59 d.l. 20 576 d.l.

96.21

41.60 0.34 14.00 8.40 0.13 20.70 0.05 0.19 10.20 d.l. 0.60

1644.3

d.l. 483 3.3 13 55 d.l. 115 16 12 10 76 d.l. 15 d.l. d.l. 438 d.l. d.l. d.l. 90 34 47 22 215

99.32

68.30 0.77 15.10 4.10 0.13 1.88 1.27 6.50 0.29 0.08 0.90

UMK 6862

4262.5

d.l. 2855 24.5 25 131 d.l. 12 34 37 <8 65 d.l. 276 33 45 212 d.l. d.l. d.l. 217 32 48 51 165

98.14

42.30 1.17 25.50 6.97 0.08 7.59 0.92 1.97 6.97 0.17 4.50

UMK 6864

Biotite Garnet– Garnet– Turmaline– schist mica mica mica schist schist schist

UMK UMK UMK 6856 A 6856 B 6861

Biotite– Talc Meta- Meta- Biotite Actino- Metaalbite– schists greisen granite schist lite pegmgraphite schist atite gneis

Table 1 Main and trace element composition of meta-granitoids (HMG) of the Hafafit uplift and meta-sediments, meta-basites, and metasomatic rock types of the VSS from the emerald deposits of Sikait (SIK), Zabara (ZAB) and Umm Kabo (UMK)

174 G. Grundmann, G. Morteani / Journal of African Earth Sciences 50 (2008) 168–187

G. Grundmann, G. Morteani / Journal of African Earth Sciences 50 (2008) 168–187

Fig. 6. SiO2 vs. Na2O + K2O diagram showing the CIPW normative composition of the meta-granitoids of the Nugrus uplift and of the amphibolites of the VSS.

behaviour of the studied blackwall rocks correlates well with that of the blackwall zones of other schist-type emerald deposits (Grundmann and Morteani, 1989). The tourmaline–chlorite–mica schist has both a high Be (mean 43.9 ppm) as well as a high chromium (mean 500 ppm) content and can be considered as an exhalite dominated meta-sediment (Table 2) (Slack et al., 1984). Such a genesis is supported by its very high Ba content, too. Ba is a typical element in gangue minerals of the hydrothermal gold mineralization of the area (Elsamani et al., 2001). The Cr and Ni contents in the tourmaline– mica schist can be explained by the close association of ultrabasic rocks and tourmalinites which facilitated a Cr and Ni leaching from the ultrabasic rocks by the hydrothermal fluids of the tourmalinites. The highest Sn content is displayed by the muscovite meta-greisen vein (#6856a) (up to 257 ppm) of Umm Kabo. This sample also has an elevated Nb content of 113 ppm. The meta-pegmatoid vein (#6905), shown as a sketch in Fig. 2, has a Sn content below the detection limit, but very high Nb (1993 ppm), Ta (218 ppm), Th (606 ppm), U (164 ppm), W (85 ppm), Y (193 ppm) and Zr (8320 ppm) contents.

175

The Sn content of the granitoid gneisses forming the core of the Hafafit uplift at Umm Kabu (# 6851) and Sikait (# 6873, 6874, 6897) is below the detection limit. At Zabara (#6909, 6929, 6930, 6931) the Sn content in the granitoid gneisses ranges from the detection limit up to 125 ppm. Soliman (1986) gives a mean Sn content of 88 ppm for the granitoids of the Zabara area. The tungsten content of the granitoid gneisses of the Hafafit uplift at Zabara is up to 261 ppm. At Sikait the W content is up to 30 ppm and at Umm Kabo it is below the detection limit. The Nb contents in granitoid gneisses at Zabara are between 42–61, at Sikait between 8–14 ppm, and at Umm Kabu the content is 11 ppm. High Ti, Sn and W contents in the granitoid gneisses of the Hafafit uplift shows that their granitic protolith was a highly differentiated one. Due to the strong tectonic overprint a direct connection of the meta granitoid veins with the granitoid gneisses of the Hafafit uplift cannot be demonstrated, but from a geochemical point of view their genetic relation to the magmatic event that produced the highly differentiated Hafafit meta-granitoids is very probable. 5.3. Chemical composition of minerals Representative electron-microprobe point analyses of minerals are given in Table 3. The amphiboles of the study area range in composition between actinolitic (# 6880, 6884) in the actinolite–talc schist, and pargasitic (# 6809) in the garnet–amphibole schist (# 6809). A pargasitic amphibole composition is typical of amphibolite-facies basic rocks. The garnets in the garnet–mica schist are almandines. In the biotite schist, the dark mica has a phlogopitic (# 6884, 6878, 6876) composition with Cr contents of up to 0.52 wt.%, whereas the biotite from the garnet mica schist forming the adjacent country rock shows a Cr content of only 0.04 wt.% (# 6809a). The emeralds show a Cr-enriched rim with Cr contents of up to 0.72 wt.% (# 6876). In all cases the iron content is higher than that of Cr. It can be up to 1.01 wt.%. Remarkable is the rather elevated Mg content of the green emeralds of up to 3.2 wt.% (# 6866) and a Na2O content of up to 0.04 wt.%. High Mg content indicates, according to Schwarz and Giuliani

Table 2 Mean Be and Cr contents of the main studied rock types with sample numbers Rock type

Zabara

Sikeit

Umm Kabo

Range Be (ppm)

Mean Be (ppm)

Mean Cr (ppm)

Metabasites, amphibolites Garnet–mica schists Granitoid gneisses Meta-greisens, -aplites, -pegmatoids Tourmaline mica schists Biotite–actinolite schists Biotite schists Chlorite schists Talc schists

6916, 6932 6915 6930, 6929 6931, 6909 – – 6917 6919 6920

6868a, 6870 6899, 6892 6874, 6873 6897, 6905 6882 6868b 6900, 6898, 6890 6881 6904

– 6861, 6862 6851 6856a 6864 6855 6854, 6856b – –

3.5–6.0 3.3–10.9 2.4–10.1 9.8–106 24.5–63.3 32.6–38.5 4.5–19.4 8.4–9.6 2.9–7.1

4.8 5.5 6.4 38 43.9 35.6 13.4 9 5

180 68 23 36.4 500 1700 607 376 3079

For further information see text.

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Table 3 Representative electron-microprobe mineral analyses Mineral

act

Locality

act

amph amph plag

grt

phl

phl

phl

chl

brl core

brl rim

brl core

Sikeit Sikeit Sikeit Sikeit Sikeit

Sikeit

Sikeit Sikeit Sikeit Sikeit Sikeit

Sikeit Sikeit

Sample no.

6880

6884

6809

6809

6809

6809

6884

6878

6876

6884

6878

6879

SiO2 TiO2 Al2O3 Cr2O3 FeO(tot) MnO MgO CaO Na2O K2O

55.61 0.07 2.13 0.05 6.22 0.28 20.6 12.3 0.72 0.04

51.52 0.23 4.64 0.04 8.23 0.26 18.27 12.09 1.63 0.41

42.60 0.19 15.99 0.08 13.54 0.49 10.74 10.91 2.87 0.37

42.38 0.20 16.15 0.26 13.08 0.45 10.81 10.96 2.82 0.42

69.59 d.l. 21.94 d.l. 0.73 0.01 0.19 1.82 6.61 0.06

38.25 0.02 21.22 d.l. 30.33 1.44 2.30 8.29 0.01 0.02

41.6 0.51 11.94 0.52 9.05 0.15 21.59 n.d. 0.29 10.87

42.5 0.58 12.16 0.09 9.18 0.06 20.93 0.09 0.08 10.58

42.5 0.45 12.13 0.18 8.04 0.09 22.32 0.03 0.37 10.45

28.9 0.05 15.25 0.87 16.82 0.13 20.88 0.16 0.05 0.61

67.35 0.02 14.55 0.04 0.56 d.l. 2.57 0.02 1.26 0.04

67.12 0.01 14.68 0.15 0.66 0.01 2.39 0.02 1.20 0.02

brl rim

brl core

brl rim

brl core

brl rim

Sikeit U. Kabo

U. Kabo

U. Kabo

U. Kabo

6876

6876

6866

6866

6866

6866

67.17 d.l. 14.28 0.15 0.54 0.03 2.51 0.04 1.29 0.04

66.99 0.01 14.04 0.40 0.72 0.04 2.58 0.04 1.33 0.04

68.06 d.l. 15.16 0.08 0.39 d.l. 2.35 0.01 1.16 0.04

67.07 0.01 14.74 0.15 0.52 d.l. 2.38 0.03 1.28 0.02

66.25 0.03 13.09 0.12 1.01 d.l. 3.2 0.02 1.49 0.01

67.23 d.l. 14.29 0.23 0.46 0.01 2.77 0.03 1.23 0.03

Total

98.02 97.32 97.78 97.53 100.95 101.88 96.52 96.25 96.56 83.72 86.41

86.26 86.05

86.19 87.25

86.2

85.22

86.28

Oxygens Si Ti Al Cr Fe2+ Mn Mg Ca Na K

46 15.49 0.01 0.69 0.01 1.44 0.07 8.55 3.67 0.39 0.01

8 2.97 – 1.11 – 0.03 – 0.01 0.08 0.55 –

12 3.00 – 1.96 – 1.99 0.10 0.27 0.70 – –

22 5.99 0.06 2.03 0.06 1.09 0.02 4.63 – 0.08 1.99

18 7.30 – 1.84 0 0.05 – 0.41 – 0.26 0.01

18 7.29 – 1.88 0.01 0.06 – 0.39 – 0.25 –

18 7.31 – 1.83 0.01 0.05 – 0.41 0.01 0.27 0.01

18 7.30 – 1.80 0.03 0.06 – 0.42 – 0.28 0.01

18 7.29 – 1.91 0.01 0.03 – 0.37 – 0.24 –

18 7.29 – 1.88 0.01 0.05 – 0.38 – 0.27 –

18 7.32 0 1.70 0.01 0.09 0 0.53 0 0.32 0

18 7.30 0 1.83 0.02 0.04 0 0.45 0 0.26 0

Total

30.33 30.95 31.52 31.52 4.75

8.02

15.95 15.79 15.88 19.95 9.87

9.88

9.9

9.9

9.85

9.88

9.97

9.9

46 14.74 0.05 1.57 0.01 1.97 0.06 7.79 3.71 0.9 0.15

46 12.53 0.04 5.55 0.02 3.33 0.12 4.71 3.44 1.64 0.14

46 12.49 0.04 5.61 0.06 3.22 0.11 4.75 3.46 1.61 0.16

22 6.10 0.06 2.06 0.01 1.10 0.01 4.48 0.01 0.02 1.94

22 6.05 0.05 2.04 0.02 0.96 0.01 4.74 0.01 0.10 1.90

(2001), genesis of emerald in a metamorphic environment. Na is an essential partner to magnesium, as they together constitute a coupled substitution for Al. Na also plays a key role in the hydrothermal remobilization of pegmatitic beryl (Markl and Schumacher, 1997). 6. Microstructural features In the following paragraphs, the microstructures of emeralds and of their country rocks from the Zabara, Sikait and Umm Kabo districts are examined to determine whether the emeralds formed during a pegmatitic or rather during a regional-metamorphic event. In the case of pegmatitic crystallization, the emeralds and cogenetic minerals should show features indicating static growth, with unoriented inclusions of magmatic minerals. In the case of regional-metamorphic origin, the emerald and cogenetic phases should show ductile and brittle deformation and oriented inclusions, indicating medium pressure and medium temperature conditions corresponding to greenschist- to amphibolite-facies conditions. 6.1. Beryl/Emerald As already mentioned by Grundmann and Morteani (1993) three beryl/emerald generations (types 1–3) can be distinguished:

28 6.14 0.01 3.82 0.15 2.99 0.02 6.61 0.03 0.02 0.16

(a) The first and oldest generation (type 1) is a typically colourless beryl which occurs as individual crystals or as a core surrounded by up to two rims of lightto deep green emerald. The colourless beryl core contains randomly distributed inclusions of biotite, white mica, quartz, plagioclase, tourmaline, apatite, ilmenorutile, columbite, tantalite, zircon, allanite, phenakite and thorite, and often shows marked cracks (Fig. 7) produced by brittle deformation. The randomly distributed and randomly oriented inclusions demonstrate a growth under static conditions. (b) The second type (type 2) can be observed as light green individual homogeneous beryl crystals, or as a rim overgrowing a colourless beryl, or as the inner light green rim between a colourless core and a deep green outer rim. This second type of emerald contains inclusions of mainly biotite, white mica, quartz and albite, but also of tourmaline, apatite, rutile, ilmenorutile, columbite, tantalite, zircon, allanite, phenakite and thorite. The inclusions are often oriented parallel to the prism faces of the emerald crystal (Fig. 8). Tensional cracks found in the colourless beryl (type 1) of the inner core may propagate into the pale green inner rim (type 2), but are missing in the deep green emerald forming the outer rim (type 3) (Fig. 7). The distribution of the inclusions demonstrates also for this second type of emerald a growth under static conditions.

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Fig. 7. Photomicrograph of euhedral emerald porphyroblasts (type 1) from a strongly deformed quartz lens (compare Fig. 3). The colourless core shows randomly oriented mica inclusions. The core and the pale green rim are characterized by tensional cracks that predominantly do not find a continuation in the outer rim formed by deep green emerald (type 3). Crossed nicols, polarized light.

Fig. 8. Photomicrograph of a subhedral emerald porphyroblast in mica schist. The colourless beryl of the core (type 1) shows many randomly oriented mica and quartz inclusions. The grain boundary of the euhedral light green emerald (type 2) is marked by inclusions that are oriented parallel to the prism faces of the emerald. Emerald of type 3 forms a very irregular outer rim masking the euhedral grain boundary of emerald type 2. Crossed nicols, polarized light.

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Fig. 9. Photomicrograph of a zoned deep green emerald porphyroblast of type 3 in biotite schist. The inner part of the porphyroblast (type 3) shows sigmoidal orientated mica inclusions. The outer part of the emerald is almost inclusion-free. The foliation and crenulation cleavage outlined by mica inclusion trails inside the core have a different orientation as compared to the main foliation present in the rim and in the surrounding matrix, indicating syn-tectonic growth of the core and post-tectonic growth of the rim. Crossed nicols, polarized light.

(c) The third type (type 3) is a green to very deep green emerald. It can be observed as single crystals without any zoning, or as an outer rim surrounding a core of colourless beryl of type 1, or as a rim surrounding pale green emerald of type 2. Emerald type 3 shows sigmoidal or strongly oriented inclusion trails indicating a syn- to post-tectonic growth (Fig. 9). Typical inclusions in emerald of type 3 are biotite, white mica, quartz, albite, tremolite, actinolite, talc, calcite, but only rarely tourmaline, apatite, rutile, ilmenorutile, columbite, tantalite, zircon, allanite, phenakite and thorite. In the case of the emerald crystallization of type 3 surrounding beryl of type 1 and emerald of type 2, the tensional cracks found in the core (type 1) and inner rim (type 2) may be healed by emerald type 3 (Fig. 7). Green emerald of type 3 healing the cracks in colourless beryl of type 1 and emerald of type 2 gives them a better green colouration.

tinue into the outer rim. The irregular outer rim is inclusion-poor. Undeformed biotite and chlorite partly replacing the outer rim of amphibole are the product of a late stage retrograde biotitization and chloritization of the amphibole. 6.3. Garnet Garnet porphyroblasts such as those in garnet mica schist from Zabara display in the subhedral core a sigmoidal texture outlined by fine-grained quartz and ilmenite inclusions (Fig. 11, see insert). Tensional cracks in the core cross-cutting the inclusion trails may be healed by quartz, mica and chlorite and do not continue in the inner and outer rims (Fig. 11). The inner and outer rims are generally inclusion-poor (Fig. 12). The subhedral to anhedral outer rim of the garnet replaces minerals of the main foliation (Fig. 11). During retrograde metamorphism garnet is partly replaced by undeformed biotite and chlorite.

6.2. Amphibole 6.4. Tourmaline Fig. 10 shows a zoned amphibole porphyroblast in an amphibolite from Zabara, surrounded by a matrix of fine-grained plagioclase. The core of the actinolitic amphibole shows tensional cracks outlined by dusty-looking minute inclusions. The tensional cracks cut the cleavage, that is marked in the core by opaque inclusions, and do not con-

Zoned tourmaline porphyroblasts typically show a dark inner core and a multiple sequence of inclusion-poor and inclusion-rich rims. The foliation and crenulation cleavage outlined by graphite inclusion trails inside the core have a different orientation as compared to the main foliation pres-

G. Grundmann, G. Morteani / Journal of African Earth Sciences 50 (2008) 168–187

179

Fig. 10. Photomicrograph of a zoned amphibole porphyroblast. The inclusion-rich core shows sub-vertical tensional cracks outlined by a dark pigmentation. The rim is inclusion-poor and almost fracture-free. Plane polarized light.

Fig. 11. Photomicrograph of a garnet porphyroblast showing a core with tensional cracks cross-cutting the inclusion trails of fine-grained quartz, an inclusion- poor inner rim, and an anhedral outer rim. Insert schematically emphasizes the different zones and the tensional cracks crosscutting the foliation in the core of the garnet. Plane polarized light.

ent in the rims and in the surrounding matrix, indicating syntectonic growth (Fig. 13). Inclusion trails in the outer rim indicate post-tectonic growth with respect to the main

foliation of the schist. Frequent fractures and cracks can be healed by a late generation of tourmaline and additionally by quartz, white mica, biotite and chlorite.

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Fig. 12. Microphotograph of two garnet porphyroblasts in garnet–mica schist, showing dark inclusion-rich cores outlined by sigmoidal inclusion trails of fine-grained quartz. The margin of the subhedral inner rim is outlined in the lower garnet crystal by minute inclusions. The outer rim is almost inclusionfree. Plane polarized light.

Fig. 13. Photomicrograph of a zoned tourmaline porphyroblast in tourmaline–chlorite–mica schist. The foliation and crenulation cleavage outlined by graphite inclusion trails inside the core have a different grain size and orientation as compared to the main foliation present in the inclusion-poor rim and in the surrounding matrix. Plane polarized light.

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181

Fig. 14. Photomicrograph of a pseudomorph after a probable euhedral zoned lawsonite porphyroblast. The dark inclusion trails consist of mainly graphite. Notice the relic multiple zoning. Plane polarized light.

Fig. 15. Microphotograph of large elongated quartz grains with undulose extinction surrounded by fine-grained recrystallized quartz grains in an emeraldbearing quartz lens. This type of microstructure in quartz indicates strain-induced grain boundary migration under greenschist- to amphibolite-facies conditions. Crossed nicols, polarized light.

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6.5. Tremolite-talc-plagioclase pseudomorphs At Sikait (# 6870) the graphite-bearing amphibole schist from the south western slope of Gebel Sikait contains pseudomorphs after euhedral rhombic porphyroblasts (Fig. 14). The shape of the pseudomorphs recalls that of andalusite or lawsonite. The pseudomorphs consist of a fine-grained association of tremolite, talc and plagioclase and display a distinct zoning emphasized by inclusion trails given by very fine-grained graphite. The trails outline a relic crenulation cleavage that does not correspond to the main foliation found outside the pseudomorphs. Healed sub-parallel tensional cracks cross-cutting the core zone of the pseudomorphs indicate a brittle deformation that predates the break-down of the porphyroblasts. 6.6. Quartz Large elongate quartz grains as typically found in the meta-pegmatites and quartz-veins show undulose extinction and low angle subgrain boundaries, grain size reduction and fine-grained recystallized quartz (Fig. 15). This type of microstructure in quartz indicates strain-induced grain boundary migration under greenschist- to epidote– amphibolite-facies conditions (Kruhl, 1996). Minerals that are more refractory to the deformation than quartz (e.g. tourmaline, garnet, beryl and feldspar) display a corona of elongated deformed old quartz grains and undeformed recrystallized new quartz and mica aggregates. In the

meta-pegmatites, tensional cracks in the elongated large quartz grains are perpendicular to the main foliation and are outlined by dusty-looking minute fluid inclusions. 6.7. Feldspar In strongly deformed meta-pegmatites and meta-granitoids coarse-grained albite shows multiple deformation twinning with undulose extinction and fine-grained recrystallization along ruptures due to a post-magmatic ductile deformation, recrystallization and albitization. In less intensely deformed rocks feldspar is fractured and fragmented whereas quartz flows plastically, recrystallizing later to fine-grained aggregates. The original magmatic granophyric intergrowth of quartz and feldspar in the meta-pegmatites is modified by brittle deformation in feldspar, contrasted by ductile flow and fine-grained recrystallization in quartz (Fig. 16). Cracks in feldspar are often filled and healed by deformed and or undeformed quartz (Fig. 16). 7. Discussion The internal and external microstructures of beryl/emerald, amphibole, tourmaline, garnet, and pseudomorphs after probable lawsonite formed during the Pan-African sequence of crystallization and deformation events of the Zabara, Sikait and Umm Kabo districts are summarized in Fig. 17.

Fig. 16. Microphotograph of a deformed meta-pegmatite showing a granophyric texture. Notice brittle deformation and extensional cracks in coarsegrained feldspar, contrasted in quartz by ductile flow. Fine-grained recrystallized quartz follows diagonal fractures. Crossed nicols, polarized light.

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183

Fig. 17. Diagram summarizing the microstructures produced by the sequence of Pan-African crystallization and deformation events in emerald, amphibole, tourmaline, garnet and probable lawsonite from the host rocks of the Egyptian emerald deposits.

The microstructures show that the studied rocks were affected by at least two tectono-thermal events (M1 and M2) including two pervasive ductile (D1 and D3) deformations and one pervasive brittle (D2) deformation (Fig. 17). The two tectono-thermal events are separated by a period of granitoid magmatism that delivered the protolith of the meta-granitoid cupolas of the Hafafi uplift. A distinct retrograde metamorphism producing among others biotitization and chloritization of garnets and amphiboles marks the end of the tectono-thermal events and of related fluid circulation. This last event is not shown in Fig. 17. The given sequence of events deduced from microstructural data finds its equivalent in the succession of two regionalmetamorphic events and three phases of deformation, as deduced from a detailed textural fold pattern analysis by Hassan (1998). The best microstructural documentation of the first PanAfrican regional metamorphic event (M1) is the inclusionrich core and inclusion-poor inner rim of the almandinerich garnet, tourmaline and amphibole porphyroblasts of the garnet–mica schist, amphibole schist and amphibolites. The crystallization of probable lawsonite in the graphitebearing garnet–amphibole schist at Zabara and Sikait can be referred to this event, too (Fig. 17). A first foliation, formed under increasing grade at blueschist- to greenschist-facies conditions of the M1 event, is documented in the cores of the porphyroblasts by the very fine-grained inclusion trails that are typical for foliated and folded sedimentary rocks. The M1 event can be correlated with the age of the onset of the multistage Pan-African tectono-thermal event.

Rb/Sr-ages between 700 and 677 Ma from granitic gneisses of the Sikait S-type leucogranite indicate a period of strong thermal activity and magmatism (Kro¨ner et al., 1994). Such an age corresponds in the Arabian–Nubian shield, according to Windley (1984), to the middle PanAfrican period. The middle Pan-African event was characterized by the emplacement of ophiolitic complexes, greywacke sedimentation and calc-alkaline andesitic to basaltic volcanism in an island-arc setting (Windley, 1984), now forming the VSS. The S-type leucogranites intruded into an already metamorphosed VSS sequence of rocks. The emplacement of the pegmatites and aplites and the crystallization of the colourless and the light green emeralds had to be related to this granitoid magmatism. Fine-grained zircon, epidote, thorite, and columbite/tantalite in the Be-rich meta-pegmatite veins produce in the whole rock analyses high contents of thorium (680 ppm), uranium (170 ppm), zirconium (6100 ppm), niobium (1990 ppm) and tantalum (220 ppm), and classify the protolith of the meta-pegmatites as highly differentiated. The light green emeralds have been produced by a remobilization of Be from pre-existing pegmatitic beryl and/or phenakite driven by a metasomatic reaction between synto post-magmatic hydrothermal fluids of pegmatitic origin and Cr-rich ultrabasic country rocks. The remobilization of primary magmatic beryl by late- to post-magmatic fluids is emphasized by Markl and Schumacher (1997) and Giuliani et al. (1997a,b). The randomly oriented mica inclusions in the colourless or light green euhedral to anhedral beryl of the meta-pegmatites confirm their static growth in a pegmatitic environment.

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A definition of the metamorphic grade of M1 is difficult. The presence of pseudomorphs after possible lawsonite and the recrystallization of magmatic amphibole/pyroxene to actinolite with opacitic inclusions, suggest that the metamorphic M1 event was at the beginning of blueschist-facies conditions. The formation of the core complexes such as the Hafafit uplift, by extensional detachment systems, is dated by Blasband et al. (2000) at 600–560 Ma. Such ages correspond to the upper Pan-African event (Windley, 1984). Age determinations of 600–565 Ma (Fullagar, 1980; Stern and Hedge, 1985) that are supposed to give the age of the intrusion of the granites now forming the meta-granitoid cupolas and connected meta-pegmatites, give probably the upper Pan-African age of the extensional detachment and metamorphism. According to Soliman (1986), Mohamed and Hassanen (1997), Hassan (1998) and Abdalla and Mohamed (1999) the granitoid rocks of the study area are emplaced by a ‘‘syntectonic intrusion of leucogranite and pegmatite along major ductile shear zones’’. But the absence of blocky subgrain pattern of high quartz (‘‘chess board’’ subgrain pattern), and the absence of sillimanite that, according to Kruhl (1996), are typical for such high-temperature granitoid intrusive bodies put some doubt on their ‘‘syntectonic intrusion’’. Such high temperature intrusions produce necessarily a high-temperature contact aureole characterized by the presence of sillimanite (Kruhl and Vernon, 2005). In such a high temperature aureole, beryl would also react with K-feldspar or albite to form chrysoberyl (Franz and Morteani, 1984). Chrysoberyl could not be found in the beryl-bearing rocks of the study area. This leads to the conclusion that, around the metagranitoid cupola of the Hafafit uplift, such a high-temperature contact aureole typical for syntectonic granites, is missing. The strongly foliated granitoids and boudinaged meta-pegmatites show, on the contrary, large elongate aggregates of fine-grained recrystallized quartz under the microscope, that represent the distorted equivalents of coarse-grained magmatic low quartz, as they are typical constituents of granitoid rocks overprinted under greenschist- to amphibolite-facies conditions during the middle, but much more likely the upper Pan-African event. An own Rb/Sr chronological determination on biotite from an emerald type 3-bearing biotite schist gives an age of 591 ± 5.94 Ma for the crystallization of biotite and cogenetic green emerald of type 3. This age puts the second tectono-thermal event (M2) into the time span of 500 to 600 Ma given by Windley (1984) for the upper Pan-African event. The P–T conditions given, for example, by Surour (1995) for the amphibolites of the VSS in the Zabara and Sikait area, can be referred to this M2 event, that concluded the metamorphic history in the study area. Surour (1995) estimated a metamorphic temperature range for the amphibolite-facies of between 485 and 571 °C and a pressure range of 6.8–7.7 kb. From our own electron-

microprobe analyses of biotite surrounding garnet in amphibolite of the VSS, the garnet–biotite thermometer calculated according to Perchuk et al. (2000) gives, at 7 kb, a temperature of 415 °C. This temperature, that is somewhat lower in comparison to the minimum temperature determined by Surour (1995), very likely gives the temperature of the retrograde overprint. A tschermakitic composition of amphibole in the amphibolites confirms the medium pressure characteristic of the metamorphism (Surour, 1995). In the microstructures the M2-event is documented by the outer rim of the emeralds, tourmalines and garnets. From a deformational point of view, the late Pan-African metamorphic event (M2) encompasses a pervasive brittle (D2) deformation followed by an extensional ductile deformation (D3) (Fig. 17). The deformation D2 was characterized by a low temperature brittle deformation and shearing of the middle Pan-African porphyroblasts. The deformation (D3) is documented by extensional cracks in, for example, almandine-rich garnet, amphibole, tourmaline and beryl/emerald. The extensional cracks do not find a continuation in the outermost post-tectonically grown and quite often very irregular outermost rim of the given minerals. Macroscopic and microscopic structures related to this event are pencil cleavage, crenulation cleavage and/or extensional boudinage, as described by Hassan (1998). A final retrograde metamorphism (diaphthoresis) is documented in the host rocks of the emerald by the replacement of pre-existing minerals by fine-grained albitic plagioclase, white mica, dark mica, chlorite, actinolite and talc. The typically deep green emerald of type 3 grew during the regional tectono-metamorphic event, M2. It occurs as strongly zoned individual subhedral porphyroblasts or as rims around pre-existing colourless or pale green beryl. The green colour of the emerald was produced by chromium that was in situ available in the metamorphic fluid phase and by a replacement of chromium-rich minerals, such as biotite, white mica (Cr-phengite) and amphibole. The Cr uptake by replacement of pre-existing minerals is documented by the green Cr-rich halos that are found in emeralds around partly replaced mica and amphibole inclusions. The colourless pegmatitic beryl (type 1) and the emerald (type 2) suffered an intense brittle deformation and shearing followed by the crystallization of emerald type 3, during M2. The deformation and coeval thermal event caused a partial decrepitation of primary and secondary fluid inclusions in beryl (type 1), emerald (type 2) and quartz. In order to characterize the fluid inclusions, a total of 20 doubly polished wafers of loose emerald crystals were studied with a Linkham TH 600 heating–freezing stage. Three main types of fluid inclusions were found: CH4-rich, CO2-rich, aqueous-rich and subordinately supposedly nitrogen-bearing inclusions. Our data are not presented here because they are not sufficient to reconstruct the complicated multi-stage emerald formation as revealed from

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the microstructural analysis. The cores and the inner rims of the emerald porphyroblasts often show a dense distribution of the well known thin tubular inclusions common to beryl, and numerous irregularly shaped decrepitated inclusions. The outer rims are remarkably poor in fluid inclusions. The core/rim boundary outlined by the density of fluid inclusions often shows an irregular, smooth, rounded shape compared to the euhedral crystal faces. Numerous fractures cross-cutting the core and/or inner and outer rim obviously caused a pervasive necking down and decripitation of the fluid inclusions. The wide range and scatter of composition of our data and the data given by Abdalla and Mohamed (1999) suggest a strong fluctuation in fluid pressure, with an opening and resealing of fractures during multistage syn- to post-tectonic formation of the emerald crystals. In this case only a combination of microthermometry and Raman probe analysis would probably clarify the complicated history of necking down, decripitation and fluid-mixing. Differences between the lithostatic pressure and the fluid pressure have been found by Nwe and Grundmann (1990) in emeralds which crystallized during regional metamorphism in a shear zone of the Alpine Tauern Window. It is therefore important to realize that in the case of multi-stage magmatic and metamorphic events, the terms ‘‘primary, secondary and pseudosecondary’’, as used by Abdalla and Mohamed (1999), have to be used with caution. Own stable oxygen isotope data on emerald from the Zabara (sample #12800, d18O = 10.5; 10.45) Sikait (sample #6878, d18O = 10.3; sample #6884 = d18O = 9.9) and Umm Kabo (sample # 6866, d18O = 10.7) deposits range in the same interval as those given by Giuliani et al. (1998a,b). In view of the multistage magmatic and metamorphic crystallization of beryl/emerald, intercalated by deformational events and driven by fluids of changing composition, oxygen isotope data cannot give very reliable genetic information. The present data can be in fact related to magmatic as well as metamorphic fluids (Giuliani et al., 1998a,b). A better criterion to distinguish between magmatic and regional metamorphic beryl/emerald seems to be the Mg content. Beryl/emerald crystallized in a regional metamorphic event, like the present one, seems to be significantly higher in Mg then those crystallized in a magmatic/ pegmatitic environment. The very limited transport distance of Cr and Be in a metamorphic environment is demonstrated by the very local crystallization of emerald at the contact between Be-rich (up to 106 ppm Be) beryl-bearing meta-pegmatoids (# 6905) and their Cr-rich biotite schist envelope and ultrabasic rocks. In the present example the maximum distance of Cr and Be transport can be estimated to be about 50 cm. Also, the emerald-bearing small quartz-veins and lenses found close to the Be-rich meta-pegmatites are the result of such short-range mobilization. These small emeraldbearing quartz veins and lenses (maximum 10 cm in thickness) have to be distinguished strictly from the numerous and often large quartz veins and lenses of up to 5 m in

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thickness, that form a regional net and which are completely beryl- and emerald-barren. The very limited transport distance of the key elements Be and Cr in a metamorphic environment impedes the growth of large emerald crystals. The contrast in size between emeralds and pegmatitic beryls is striking – they differ by orders of magnitude (Grundmann and Morteani, 1989). Beryl forms bigger crystals than any other mineral species – up to 200 metric tons. Aquamarines, for example, may reach 100 kg and retain their gem quality, whereas emeralds rarely, if ever, exceed 1 kg. The mean weight of individual emeralds of Zabara, Sikait and Umm Kabo is very likely less than 1 g. The rather limited transport distance of Be and Cr and the importance of local regional-metamorphic reactions freeing those elements during emerald genesis is demonstrated by the mass balance calculation of Grundmann and Morteani (1989). The preferred emerald crystallization in the biotite schists is the consequence of the low Be storage capacity of the biotite schists (mean Be content = 13,5 ppm) in comparison to that of the Be-rich metasomatically replaced rocks (mean Be content = 38 ppm) (Grundmann and Morteani, 1989; Grew, 2002).The consequent Be surplus in the biotite schists is stored (together with Cr) in beryl/emerald. An emerald genesis, that includes the formation of the remarkably thick blackwall zones, driven only by the heat and fluids given off by the crystallization of the Be-rich pegmatites and greisen veins is, at least in the study area, unlikely in view of their very small size and heat capacity. 8. Conclusion For the emerald deposits of Southern Egypt, a syn- to post-tectonic crystallization of three beryl/emerald generations during a magmatic, a post-magmatic hydrothermal and a regional-metamorphic event is supported by the detailed microstructural and chemical analysis of beryl/ emerald, garnet, amphibole, mica, tourmaline, quartz, chlorite and probable lawsonite. Precursors of the emeralds are the colourless to light green beryl and/or phenakite of pegmatitic origin that has been partly replaced during the formation of blackwall metasomatic reaction zones found between Be-rich meta-pegmatites and Cr-rich ultrabasic rocks. All evidence points to an emerald-producing process that was controlled in detail by the very local availability of Be ions, Cr ions and metamorphic fluids in the context of the late Pan-African tectono-thermal event. The importance of the very local environment is underlined by the fact that emerald mineralizations are extremely rare (nugget-type) compared to the huge amount of Be-rich meta-granites, meta-pegmatites, metaaplites, meta-greisen and quartz veins amalgamated with ultrabasic rocks in the volcano-sedimentary series of the SEDE. In the present study no persuading evidence can be found for an emerald genesis related only to:

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(a) long distance infiltration of Be- and Cr-rich fluids (Soliman, 1986); (b) thrust-fault-shear zone – controlled genesis unrelated to pegmatites and granites (Giuliani et al., 1998a,b; Schwarz and Giuliani, 2001); (c) non-pegmatitic beryllium occurrence (Barton and Young, 2002); (d) infiltration of Be-rich solutions supplied by syntectonically emplaced leucogranites (Abdalla and Mohamed, 1999). All these models describe only particular aspects of the complicated process that is needed in the case of the emerald deposits in Egypt, but also in the case of all schist-type emerald deposits of the world, to bring together Be- and Cr-rich rocks and to create the geological conditions necessary for the formation of valuable deep green emeralds of gem quality. As long as a robust model of the genesis of the individual emerald occurrences does not exist, it is suggested to use the non-genetic descriptive classification of Schwarz et al. (2002). These authors distinguished: (a) ‘‘Pegmatites without schist seams’’ (type Gwantu, Nigeria); (b) ‘‘Pegmatites and greisens with a schist seam’’ (type Malychevo, Russia); (c) ‘‘Schists without pegmatites’’ (type Tauern, Austria; Swat, Pakistan; Goia´s, Brazil); (d) ‘‘Black shales with veins and breccias’’ (type Chivor, Muzo, Colombia). The studied emerald deposits in the South Eastern Desert of Egypt would fit best in category b. Acknowledgements We thank Dr. G. Giuliani for the thoughtful review and for stable oxygen isotope determination on four of our emerald samples from Egypt. We thank R. Klemm (Munich) for the preliminary archaeological dating of the ruins existing in the area of emerald mining. Dr. El Gawli (Cairo) is gratefully acknowledged for financing and care of organizing the field work in 1992/93. We thank A.A. Rashwan (Cairo) for help with the field work and discussions about the emerald genesis, D. Ackermand (Kiel) helped with the microprobe analyses. We thank M. Satir (Tu¨bingen) for the Rb/Sr age determination on biotite. C. Preinfalk (Munich) helped much with editing and artwork. References Abdalla, H.M., Mohamed, F.H., 1999. Mineralogical and geochemical investigation of emerald and beryl mineralisation, Pan-African Belt of Egypt: genetic and exploration aspects. J. African Earth Sci. 28 (3), 581–598.

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