Characteristics and genesis of shear zone-related gold mineralization ...

Ore Geology Reviews 34 (2008) 445–470

Contents lists available at ScienceDirect

Ore Geology Reviews j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / o r e g e o r ev

Characteristics and genesis of shear zone-related gold mineralization in Egypt: A case study from the Um El Tuyor mine, south Eastern Desert Basem A. Zoheir Department of Geology, Benha Faculty of Science, 13518 Benha, Egypt

A R T I C L E

I N F O

Article history: Received 29 March 2007 Accepted 8 May 2008 Available online 8 June 2008 Keywords: Shear zone Orogenic gold Um El Tuyor Eastern Desert Allaqi Egypt

A B S T R A C T Although gold production from orogenic deposits in the Arabian–Nubian Shield is currently relatively minor, extensive alluvial and lode fields were exploited by the ancient Egyptians along the western side of the Red Sea in Upper Egypt and northern Sudan. In the Eastern Desert of Egypt, numerous but small gold deposits are generally related to auriferous quartz veins commonly associated with brittle–ductile shear zones cutting the Neoproterozoic crystalline basement rocks. The Um El Tuyor gold deposit, in the extreme south of the Eastern Desert, consists of a series of quartz and quartz–carbonate veins, lenses and veinlets, along foliation-concordant and foliation-discordant fault segments. The orebodies are mainly quartz–carbonate (quartz+Fe-dolomite/ankerite+muscovite/sericite+ sulfides ±gold/electrum) and laminated quartz veins (quartz+sericite/chlorite±graphite±free gold). Gold appears as microscopic blebs within arsenopyrite and pyrite or along fractures and grain boundaries in quartz veins and finely disseminated grains in mixtures of chlorite±sericite and carbonate in the wallrock. The bulk vein system is controlled by a NNW- to NW-trending brittle–ductile shear zone, cutting sequences of pelitic metasediments (garnet–biotite schist with intercalations of metamudstone and metagreywacke). Pervasive sericitization, carbonatization and chloritization overprint the metamorphic assemblages in rocks adjacent to the mineralized quartz veins. Temporal and lateral evolution of the ore fluid composition through interaction with the wallrock is inferred from the concentric alteration pattern. Stages of increasing hydrothermal alteration are identified as initial, intermediate and advanced. Hydrolysis reactions by a near acid fluid, whose pH was buffered by the wallrock mineralogy, prevailed in the initial stage. The transitional stage involved intense carbonatization, sulfidation and redox reactions along with hydrolysis. During the advanced stage, intense sericitization consumed K+, released H+, and lowered the solution pH. Sulfidation continued, and unbuffered conditions were locally attained under high fluid/rock ratios. This collectively indicates that the ore fluids evolved progressively towards lower temperatures and sulfur fugacity with time. Clustered and intragranular trail-bound aqueous-carbonic [H2O–NaCl–CO2(±CH4 ±N2)] inclusions are common in cores of the less deformed quartz crystals, whereas carbonic [CO2 ±CH4 ±N2] and aqueous [low salinity H2O–NaCl] inclusions occur along intergranular and transgranular trails. The aqueous-carbonic and aqueous inclusions are common in the quartz–carbonate veins, whereas the carbonic inclusions are by far the dominant — or virtually the sole fluid inclusions — in most parts of the laminated quartz and quartz–carbonate veins. Clathrate melting temperatures indicate low fluid salinities (3 to 8 wt.% NaCl equiv.). Bulk densities and salinities of the aqueouscarbonic inclusions are high in the quartz–carbonate veins compared to those in the laminated quartz veins, further constraining interpretation of the field observations suggesting that the laminated quartz veins were formed through a fault–valve system and cyclic opening and annealing of reactivated quartz–carbonate reefs. Destabilization of gold– bisulfide complexes and lowering of gold solubility through interplay of fluid mixing (±unmixing), cooling, changes in pH and fO2 along with fluid–wallrock interaction brought about gold deposition. In the quartz–carbonate veins, in which the highest Th total (336 °C) indicates a pressure range of 1.7 to 2.1 kbar using isochores for highest and lowest bulk density aqueous-carbonic inclusions. In the laminated quartz veins, gold deposition likely took place due to fluid mixing and fluid–wallrock interaction at ≤325 °C at pressures of 1.6 to 1.3 kbar. Gold deposition due to fluid–wallrock interaction (sulfidation) does not conflict with intermittent fluid unmixing or mixing, as the latter does not appear to have extended into the wallrock. Pressure estimates indicate depths of 6 to 8 km for quartz veining and gold deposition in the area, compatible with crustal conditions of greenschist metamorphism and brittle–ductile transition. © 2008 Elsevier B.V. All rights reserved.

1. Introduction

E-mail address: [email protected]. 0169-1368/$ – see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.oregeorev.2008.05.003

Gold was mined from more than a hundred locations in the Eastern Desert of Egypt during the Dynastic, Ptolemaic, Roman, Byzantine,

446

B.A. Zoheir / Ore Geology Reviews 34 (2008) 445–470

Coptic and Islamic periods; there is no available data on the quantity of gold extracted. Gold production was episodic during this long time, but continued into the mid-20th Century. Between 1902 and 1958, gold output was about 7 t (Kochin and Bassyunni, 1968), including 2.4 t from the El Sid mine between 1944 and 1958 (Neubauer, 1962 and references therein). Genesis, stages of development and relationships with magmatic intrusions are subjects that remain controversial for many Egyptian gold deposits. This fact has attracted — and continues to attract — numerous authors to better understand their genesis and controls on their distribution. Hume (1937) assumed that gold mineralization in the Eastern Desert was linked to the hydrothermal activity that accompanied emplacement of development of Proterozoic diorite intrusions. Amin (1955) and El Shazly (1957) suggested that gold mineralization in the area may be polyphase and connected it to Late Proterozoic–Early Paleozoic calc-alkaline granites. Kochin and Bassyuni (1968) classified Egyptian gold deposits, based on their mode of occurrence and nature of mineralization, into dyke-, vein- and placertypes. Sabet et al. (1976) and Sabet and Bordonosov (1984) suggested a three-fold classification for gold deposits in the Eastern Desert of Egypt, including gold–sulfide formation, gold–ferruginous quartzite formation and gold–quartz formation. They indicate that gold ore formation took place during four time spans (pre-orogenic, syn-lateorogenic, Riphean–Lower Paleozoic, and Mesozoic–Cenozoic epochs). Most gold occurrences in the Eastern Desert are related to the syn-late

orogenic and Riphean–Lower Paleozoic activation periods. They are characteristically of vein-type and are closely associated with small tonalite–granodiorite–plagiogranite massifs and stocks and commonly display strong structural control. Botros (1991, 1995) suggested that gold mineralization in Egypt is of different ages and is genetically associated with volcanic cycles that repeatedly occurred in the Nubian Shield from the Precambrian to the Tertiary. An updated classification, also taking the tectonic setting into consideration, was introduced by Botros (2004), in which three broad categories of deposits were suggested: stratabound deposits (gold-bearing Algoma-type BIFs, tuffaceous sediments and massive sulfides), non-stratabound deposits (vein-type mineralization and disseminated gold hydrothermal alteration zones), and placer deposits (gold-bearing modern alluvial/beach placers and lithified conglomerates). Most of the important gold deposits belong to the vein-type group of Botros (2004) and are collectively interpreted as products of hydrothermal activity induced either by metamorphism or through cooling of Lower Palaeozoic magmatism and/or Early Cambrian subduction-related calc-alkaline intrusions (El Ramly et al., 1970; Garson and Shalaby, 1976; Pohl, 1988; El Gaby et al., 1988). Almond et al. (1984) suggested that gold deposition in the northern Red Sea Hills was related to a shearing episode that post-dated emplacement of all batholithic intrusions, but may have been coeval with regional cooling. Hussein (1990) argued that most of these hydrothermal vein deposits are epithermal rather than mesothermal. Although some information on the geological and

Fig. 1. Geological map of the central part of the Allaqi-Heiani Suture (modified after Zoheir, 2004; Zoheir and Klemm, 2007). Inset showing the location of this part of the suture in Egypt.


447

Fig. 2. Geologic map of the Um El Tuyor gold mine area (modified after original mapping by Zoheir, 2004).

structural context of most gold occurrences in the Eastern Desert is available, no adequate data exist to specify the age of any of them. Although the source of the hydrothermal ore fluids may have varied from one occurrence to another, some authors favor either a metamorphic origin (e.g., Hassaan and El Mezayen, 1995) or a combined metamorphic–magmatic origin (e.g., Harraz, 2000; Klemm et al., 2001; Botros, 2002, 2004) for these fluids. Leaching and remobilization from a buried source are suggested by Hilmy and Osman (1989), Takla et al. (1990), and Harraz and El Dahhar (1993). P–T conditions at or below greenschist facies conditions favored development of brittle–ductile and brittle structures in which gold–quartz veins were preferentially emplaced (El Gaby et al., 1988). Many authors have recognized structural controls on gold mineralization in the Eastern Desert, e.g., Sabet and Bondonsosov (1984), Harraz and Ashmawy (1994), and Loizenbauer and Neumayr (1996). At the Fawakhir mine area (26°16′N, 33°38′E), Loizenbauer et al. (2002) concluded that a genetic relationship existed between gold mineralization and transpressional tectonics of Late Pan-African age. In the Wadi Allaqi district, El Kazzaz (1996) and Klemm et al.

(2001) suggested a genetic and spatial link between gold mineralization and shear zones that crosscut the ophiolitic sequences, island arc assemblages and late-orogenic granitoids. Kusky and Ramadan (2002)

Fig. 3. Densely veined wallrocks from zones adjacent to the main gold-bearing veins. The quartz veinlets are disposed along foliation planes.

448


assumed that gold–quartz veins of the Wadi Allaqi region are generally structurally controlled and are genetically related to altered ultramafic rocks associated with imbricate thrust faults, or to shear zones truncating the thrust slices. El Shimi (1996) emphasized the spatial and genetic relationship between gold mineralization in the Wadi Allaqi district, which includes the Um El Tuyor deposit, and granitoids intruding the metamorphed ophiolitic-island arc rocks. He suggested that the auriferous shear zones in the area truncate large synclines commonly associated with huge ophiolitic blocks. The Um El Tuyor gold deposit is one of fifteen gold occurrences in the Allaqi district. The mine area is located at lat. 22°18′15″N, long. 34°38′00″E, ca. 175 km west of the Red Sea coast. Sadek et al. (1995, 2005) suggested that the Um El Tuyor gold mineralization is associated with NNW-trending alteration shear zones cutting highly sheared metavolcanic rocks. Ahmed et al. (2001) reported that the mineralization is associated with disseminated pyrite-rich rocks and NW-trending smoky quartz veins, conformable with the schistosity of the host metasedimentary rocks. Two grab samples from the mine area gave 1.8 and 4.2 g/t Au (Ahmed et al., 2001). Ramadan et al. (2005) suggested that the auriferous alteration shear zone at the Um El Tuyor mine was developed coeval with tight overturned anticlines with NW–SE axes and plunge towards the NW. Although completely abandoned at present, the mine area contains numerous traces (stone huts, tailings and dumps) of significant past mining activities. The deposit was intensively exploited at the beginning of the 20th Century and likely abandoned in 1925 (Hume, 1937 and references therein). Quartz veins were worked through several vertical and inclined shafts and stopes. Numerous open pits into milky quartz veins are spread over the mine area and surroundings. Two internal reports by the Egyptian Geological Survey and Mining Authority (EGSMA, 1912, 1913), describing the mining work and general conditions during that time, resulted from early English inspections of the Um El Tuyor mine (‘Um Tiour’ in these

reports). The reports state that the mineralized veins have been stoped out through six vertical or steeply inclined shafts varying in depth from 36 to 88 m. In 1912, the manager of the mine reported the average gold grade in the mineralized veins as ca. 23 to 25 g/t (EGSMA, 1912). 2. Geological setting The basement complex cropping out in the Um El Tuyor area is part of a curvilinear ophiolitic belt associated with the ca. 750 to 720 Ma Allaqi-Heiani belt in the southernmost part of the Egyptian Eastern Desert (Kröner et al., 1987; Stern et al., 1989; Abdelsalam and Stern, 1996). The mine area is mainly underlain by mafic/ultramafic ophiolitic sequences, island arc metavolcanic ± volcaniclastic rocks, successions of back arc metasediments, syn-orogenic granitoids and subordinate post-orogenic granite (Fig. 1). This rock pile experienced a multi-stage deformational history and related metamorphic events terminated with local metasomatism. The structural evolution of the area included four main phases of deformation. An early period of crustal shortening (Dm) involved transportation and overriding of huge ophiolitic sheets from north onto sequences of island arc rocks is manifested by major thrust faults, imbricate ophiolitic thrust slices and recumbent fold structures. Regional folds and pervasive foliation cleavage represent a NE–SW compressional regime (D2) superimposed on the thrust fabrics. The related NW–SE foliation (S2) is the most abundant structural element in the area. A third deformation phase (D3) is indicated by the presence of NNW-trending major folds and left-lateral faults superimposed on the older structural fabrics. This stage was accompanied or likely terminated by an episode of transcurrent shearing yielding slip reactivation of the pre-existing NW-trending faults and formation of discrete shear zones, accommodating gold mineralization in the mine area. This increment was likely the last ductile deformation event in the area. Subsequently, a

Fig. 4. Microscopic and scanning electron images illustrating features of the quartz–carbonate veins. (A) Muscovite and sericite-rich zone bounds an inner carbonatized, silicified zone in a quartz–carbonate vein (crossed polars, width of view 3.03 mm). (B) Large muscovite flakes embedded in a carbonate-rich matrix and bounded by rims, mainly of sericite. (C) Patchy sericite and carbonate with irregular outlines. (D) Dispersed sericite flakes forming networks in a quartz–carbonate vein.


weak, mainly brittle shear strain (D4) is indicated by the intersecting fault and joint trends traversing the post-orogenic rocks. Textural relationships indicate that the peak metamorphic conditions were attained during development of a NW–SE schistosity and related hornblende, garnet and staurolite (±sillimanite). Pressure–temperature estimates based on the garnet–biotite geothermometry of Bhattacharya et al. (1992) and plagioclase–biotite–muscovite–garnet geobarometer of Powell and Holland (1988) indicated a temperature range of 530 to 560 °C and pressures between 5.3 and 6.2 kbar for the peak metamorphic conditions (Zoheir, 2004). Gold mineralization relates to a system of quartz veins along a NNW- to NW-trending shear zone, cutting through successions of garnet–biotite schist with subordinate intercalations of metasiltstone and metagreywacke (Fig. 2). Field and microscopic observations suggest the presence of several types of quartz veins and veinlets in the mine area; only two of these are gold-bearing — the quartz– carbonate and laminated quartz veins. Other barren milky quartz veins, mostly follow NNW, NE and NNE directions, and clearly postdate the mineralized veins. Notably, the laminated quartz veins are restricted to the northern part of the mine area. Compared with the quartz–carbonate veins, their sulfide content is much higher. In both these vein types, gold is irregularly distributed, but is usually associated with sulfides. Next to the gold-bearing quartz veins, an assemblage of quartz, ‘sericite’, carbonate, albite and subordinate graphite almost completely replaces the primary mineralogy of the metasediments. Structural controls on the mineralization are observed at different scales. The occurrence of the main oreshoots

449

within a NNW-trending map-scale shear zone, and their tabular geometry represent the prime control on gold distribution. Field relationships and internal deformation features within the veins suggest formation during different stages of shearing. Dacite sheets, commonly as sills concordant with the main foliation (S2) — less commonly as dyke-like bodies, are abundant in the mine area. They are clearly older than the shear zone and related quartz vein system. Adjacent to the mineralized quartz veins, the host rocks and dacite sills are bleached and brecciated. Stretched asymmetric biotite and feldspar prophyroblasts and rootless microfolds indicate dextral transpression kinematics prevailed during the formation of this shear zone. Besides the shear zone, the mine area is densely dissected by NW–SE and NE–SW trending faults (Fig. 1). 3. Mineralization style Geometric and textural features including replacement, ribbon and serrate structures and intermingled quartz veinlets and altered wallrocks, suggest multiple stages of quartz veining and ore deposition in the mine area. Complexity is also inferred from the overprinting relationships between milky quartz–carbonate, laminated grey-quartz and massive milky quartz veins. The quartz–carbonate and laminated quartz veins are gold-bearing, whereas the massive milky clear quartz veins are barren. The laminated quartz veins are confined to the intensely sericitized rocks in the northern part of the mine area, replacing the quartz–carbonate veins. Both vein types are characterized by pinch-and-swell and bifurcation geometries and high

Table 1 Trace element concentrations (ppm) of the quartz–carbonate veins Sample

230

265

269

250

267

249

Detect. limit

Au Ag As Ba Sr Co Cr Rb Cs Hf Hg Mo Ni Sb Sc Sn Ta Th U W Zn Pb V Y Cu Mn Bi Cd Br Na K S Fe Mg Ratios Fe/Mg K/Na Au/Ag As/Au

15.48 3.48 3609 370 213 23 22 106 4 3 2 – 5 – 6.7 199 – 3.3 – – 54 18 56 – 13 345 – 0,8 0.43 2567 8753 6270 26,040 5950 0.24 4.38 3.41 4.45 233

16.44 3.63 7746 130 204 44 b5 62 2 2 1 2 6 – 3.5 138 0.5 2 – – 24 27 28 – 16 257 – 0.5 0.39 1899 11,598 4890 34,280 6863 0.14 4.99 6.11 4.53 471

12.75 2.99 4231 334 218 6 19 99 4 3 – – 5.5 – 6.3 141 – 3 – – 37 10 48 – 14 287 – 0.3 0.46 3276 9090 6600 30,782 4675 0.21 6.58 2.78 4.26 332

11.93 2.82 3251 166 222 17 21 87 2 – – – 3 – 4.1 103 – 2.8 – – 35 9 43 4 12 208 – – 0.34 2838 8011 7200 22,309 5611 0.32 3.98 2.82 4.23 273

9.82 1.84 4026 189 199 3 7 117 – – – – 4 – 3.1 127 – 2.5 – – 41 13 32 4 10 343 – – 0.45 2682 9796 7650 22,464 5179 0.34 4.34 3.65 5.34 410

14.53 3.22 3518 298 246 7 11 111 5 3 – – 5.3 – 5.9 97 – 3.1 – – 46 21 49 5 13 339 – 0.5 0.41 2530 8634 6800 23,397 5179 0.29 4.52 3.41 4,51 242

0.002 0.3 0.5 1 1 1 2 15 1 1 1 1 1 0.1 0.1 0.01% 0.5 0.2 0.5 1 1 3 2 1 1 1 2 0.3 0.5 0.01% 0.01% 0.01% 0.01% 0.01%

– below detection limit. Analyses were made by the ACTLABS Group, Canada, using the INAA and ICP-MS techniques (Code 1H).

450


concentrations of sulfides (up to 15 vol.%). Both quartz vein types occur along foliation-concordant and foliation-discordant shear segments; vary significantly in thickness over short distances along strike or up and down dip. These veins have greater thicknesses where occurring along the central shear planes and where the host rock foliation deflects into the shear zone. Boudinage of the quartz–carbonate veins is ubiquitous along the shear zone, forming the typical linear fabric in the host rocks. The host rock fabrics adjacent to the vein margins display progressive changes in orientation, such that they coincide with the vein margins or intersect the veins at low (15 to 20°) angles. Both quartz–carbonate and laminated quartz veins are composed of central shear (fault–fill) shoots and lensoidal bodies mostly associated with high-angle reverse faults and fracture arrays along the shear zone. Generally, these veins strike in a NNW- to NW-direction and dip steeply towards SW. The laminated quartz veins dip almost subvertically. Brittle deformation fabrics are dominated by S–C structures, asymmetric partitioned quartz ribbons, tension gashes, inclusion trails, and microfractures. Evidence for crystal-plasticity is derived from the undulating extinction, recrystallization, sub-grain development and irregular serrate boundaries between quartz crystals. The characteristics of the different types of quartz veins encountered in the mine area are discussed in more detail below. 3.1. Quartz–carbonate veins The quartz–carbonate veins occur as single tabular bodies, vein arrays, irregular swellings and pinches (generally 75 to 120 cm thick). They show anastomosing and undulating geometries, both down dip and along strike. Intermittent quartz (±carbonate) and host rock stringers are common in the marginal parts of the main veins (Fig. 3). In places, these veins cut the shear planes, but are commonly buckled around it. Vein mineralogy consists mainly of quartz (ca. 55 vol.%), carbonate (ca. 22 vol.%), muscovite/‘sericite’ (ca. 17 to 11 vol.%), and 6 to

12 vol.% sulfides (dominant pyrite and arsenopyrite, with subordinate sphalerite, pyrrhotite, chalcopyrite and galena). Carbonates are dominated by ferrodolomite [Ca(Fe,Mg)(CO3)2 solid solution]. Carbonate-rich domains are typically massive with irregular outlines. Muscovite is abundant along the vein margins, whereas ‘sericite’ commonly intermingles with carbonate or disseminated in a network pattern within the body of the vein (Fig. 4). Coarse muscovite flakes display compositional zoning typifying paragonitic substitution; Na/(Na + K + Ca) ranges from 0.15 to 0.39. They have variable (Fe + Mn + Mg) values (0.08 to 0.63 atoms per formula unit; a.p.f.u.). Albite occurs as subhedral to anhedral crystals associated with quartz and is commonly embedded in a carbonate-rich matrix. Compositionally, albite is homogenous and pure with less than 3 mol% anorthite. Large albite crystals have irregular outlines against the surrounding carbonate and quartz. Fine-grained albite also occurs at the interface between veins and the adjacent wallrocks. Albite crystals disseminated in and adjacent to the auriferous quartz–carbonate veins show similar compositions. Trace element data for the quartz–carbonate veins are given in Table 1. Silver values are lower than those of Au in all samples. There is a positive correlation between Au and Ag contents, with Au/Ag ratios ranging between 4.23 and 5.34, suggesting a cogenetic relationship of Au and Ag (e.g., Shelton et al., 1988). The highest Au and Ag values (16.44 and 3.63 ppm, respectively) occur in deep-seated hematitized parts of the main quartz–carbonate veins, adjacent to intensively ‘sericitized’ wallrock. Arsenic contents range from 3251 to 7746 ppm, and the As/Au ratios exceed 200 in all analyzed samples. Base-metal values are generally low; they nevertheless support the field and petrographic observations that chalcopyrite is more abundant than sphalerite and galena — Pb has a median value of 16 ppm. The maximum Pb content lies near 27 ppm and occurs in the sample contains the highest Cd value (0.8 ppm). Values of Co are erratic, ranging from 3 to 44 ppm, and Cr values are mostly in the range 7 to

Fig. 5. Features of the laminated quartz veins. (A) Close-up view of the laminated quartz veins showing alternated milky quartz bands and fine laminations rich in wallrock material. (B) Sketch drawing showing the relationship between the quartz–carbonate and laminated quartz veins. (C) Chlorite–sericite–sulfide ± graphite-rich lamination separating active and passive bands of markedly different grain sizes. Notice the spatial concentration of recrystallized fine quartz adjacent to the laminations depicts invalidation through shearing (crossed polars, width of view 3.03 mm). (D) Sinistral shearing along the lamination deformed the coarse grained quartz crystals (crossed polars, width of photo 5.93 mm).


22 ppm. In contrast to Cu or Pb, Zn (maximum 54 ppm) shows a poor correlation with Au. Although reflecting the mafic provenance of the host rock schists, Ni, Co, V, and Cr are associated with high concentrations of Au and As. Elevated Ba, Sr, Sn concentrations may be explained by the enrichment in carbonate and Sr release during feldspar alteration. 3.2. Laminated quartz veins The laminated quartz veins are referred to as ‘laminated’ due to their appearance, in which abundant deep grey or black-colored planar zones are separated by milky quartz (Fig. 5A). These veins occur as boudin-like bodies along the shear zone, commonly less than 50 cm wide. In the northern part of the mine area, laminated quartz replaces parts of a quartz–carbonate vein (Fig. 5B). These veins are largely composed of bluish grey to white quartz, pyrite, arsenopyrite, galena, ±chalcopyrite and free gold. They contain deep grey or black-colored narrow planar

451

elongate slivers of country rock (chlorite–‘sericite’–sulfides ± graphite) separated by milky quartz domains. Visible gold particles occur mainly in the dark bluish quartz domains. Pyrite and arsenopyrite crystals cross the main fabric of the wallrock or slivers of the wallrock within quartz veins, and occasionally show incipient to conspicuous development of quartz-dominated pressure shadows. In other places, these sulfides are variably fragmented. Large elongated clasts of the wallrocks are commonly enclosed in the veins, generally close to their contacts. The extremities of these clasts display continuity with vein laminations. In the vicinity of vein margins, the laminations extend parallel to the vein walls and display continuity with mesoscopic fault/fracture planes in the host rock schist. These fault planes are locally parallel to the vein margins or intersect them at acute angles (up to ca. 25°). Study of the vein margins indicated that many laminations are initiated where the host rock protruded into the veins along the fault planes. The laminations become progressively thinner away from the host rock protrusions and commonly acquire a stylolitic character. This change in

Fig. 6. Textures of the auriferous quartz veins from the Um El gold deposit. (A) Incipient recrystallization defined by small polygonal quartz grains and subgrains along grain boundaries and vein margins (crossed polars, width of photo 5.93 mm). (B) Shear planes filled with carbonate + clay minerals (crossed polars, width of photo 5.93 mm). (C) Intensively recrystallized quartz porphyroclasts with serrated grain boundaries and recrystallized bulges along grain boundaries and cracks (crossed polars, width of photo 2 mm). (D) Vein quartz crystals displaying undulose extinction and conjugate fracture sets (crossed polars, width of photo 0.95 mm). (E) and (F) Elongate, curved quartz ribbons with irregular boundaries and variable intergrain bulging recrystallization (crossed polars, width of photos 4 mm and 2 mm, respectively). The textural temperatures are roughly estimated at 350 °C in sense of Stipp et al. (2002).

452


character, from laminar to stylolitic, commonly corresponds to the disappearance of carbonate and sulfides in the lamination mineralogy. A difference in grain size across the individual lamination (Fig. 5C) suggests that these laminations represent boundaries/discontinuities of grain size reduction (generally a few mm wide) in the vein material. Formation of these laminations is attributed to a progressive shear strain-associated movement on the fault planes. Evidence for shearing along the laminations is indicated by development of recrystallized quartz grains with deformation bands, incipient subgrains and asymmetric orientation of c-axes at high angles to the lamination planes (Fig. 5D). As a result, strain-enhanced dissolution of quartz took place and residual graphite was enriched in the quartz veins (e.g., Bell and Cuff, 1989). The fine-grained bands are composed of recrystallized polygonal grains corresponding to shear bands in which the fault movement was accommodated. These fine-grained bands are herein referred to as

‘active’ bands (see Fig. 5C), in regard to their behavior during initiation of lamination (e.g., Davis and Hippertt, 1998). In contrast, the ‘passive’ bands correspond to the relatively non-deformed domains between the shear bands. Formation of polygonal quartz grains, abundant in the active bands, is attributed to plastic deformation and recrystallization under temperatures (ca. 250 °C, c.f. Hickey and Bell, 1996). The lamination mineralogy includes Fe-dolomite, sericite, sulfide minerals, ±chlorite, ±calcite, ±graphite. Very fine gold grains are confined to the quartz sub-grain boundaries, preferentially in zones of relatively coarse grained quartz adjacent to the shear planes. 3.3. Barren quartz veins Barren massive, milky quartz veins and pockets are generally common in the southern part of the mine area. These veins are ∼15 cm up to 2 m wide and follow NNW, NE and NNE directions. No direct

Fig. 7. Backscattered electron images illustrating ore minerals in the Um El Tuyor deposit. (A) and (B) Textural features indicating coeval crystallization of pyrite and arsenopyrite in the quartz–carbonate veins. Note that arsenopyrite occurs adjacent to the As-bearing (lighter) zone in pyrite. (C) Fine, oscillatory growth zoning of arsenopyrite disseminated in the quartz–carbonate veins. (D) Gold inclusions oriented along growth zones in a wedge-shaped arsenopyrite. (E) Subhedral gold grain on surface of a subhedral arsenopyrite crystal in a laminated quartz vein. (F) Gold wires and grain accumulations (with high fineness) in goethite grains (pseudomorphic after arsenopyrite or pyrite).


453

the southern part of the mine area. It has cavities and vugs either empty or filled by kaolinite. Other milky quartz veins, relatively thin and limited in extension, are common in the southwestern part of the mine area. 3.4. Quartz textures

Fig. 8. Sulfur activity–temperature projection of the stability field of arsenopyrite (Barton, 1969), with atom. wt.% As arsenopyrite-buffered curves from Kretschmar and Scott (1976). The hatched area depicts the temperature range and f S2 conditions of precipitation for the auriferous sulfides. As = arsenic, aspy =arsenopyrite, l = liquid, lö = löllingite, bn = bornite, po = pyrrhotite, py = pyrite.

relationship between the auriferous and barren quartz veins was observed, however, some of the barren quartz veins are hosted by fault planes that dissect the gold-bearing veins on surface. Most of the barren veins are tension gashes and fracture-fillings along N–S and conjugate NE–SW and NW–SE trending fault systems. A 2 m-wide milky quartz vein extends in a N–S direction for a distance of 100 m in

The gold-bearing quartz veins display variable degrees of shearing and dynamic recrystallization (sensu Drury et al., 1985). In the large quartz crystals, the undulose extinction is moderate to strong and mortar textures and deformation lamellae are localized or widespread. These criteria indicate that ductile deformation and dynamic recrystallization were variable in intensity during the evolution of the veins (e.g., Poirier and Guillope, 1979). In the quartz–carbonate veins, quartz porphyroblasts are dissected by conjugate sets of shear bands and large irregular subgrains (Fig. 6). Small bulges are developed along the older grain boundaries, indicating recrystallization by local boundary bulging (Vollbrecht et al., 1999). In the laminated quartz veins, quartz crystals show different but coexisting plastic and brittle deformation, manifested by the seriate irregular grain boundaries, sub-grain development, undulose extinction, microcracks and deformation lamellae. The large quartz crystals are commonly coarsely crushed (granulated) and surrounded by a matrix of smaller grains, or are completely mylonitized. The non-recrystallized cores of the large quartz grains are sutured and exhibit smooth undulose extinction instead of sharp deformation bands. With increasing strain, the ‘original’ large grains become flatter and more elongate, where abundant trails of fluid inclusions mark the healed fractures.

Fig. 9. Simplified hydrothermal alteration map of the Um El Tuyor gold deposit (modified after original mapping by Zoheir, 2004).

454


Table 2 Schematic geographic distribution of the hydrothermal alteration types in the Um El Tuyor mine

Table 2Schematic geographic distribution of the hydrothermal alteration types in the Um El Tuyor mine

Unstable = phase undergoes partial or complete alteration (minor alteration products given in parentheses). Stable = phases not affected by alteration. Abbreviations: bio = biotite, kaol = kaolinite, ser = sericite, musc = muscovite, chl = chlorite, epi = epidote, qz = quartz, rut = rutile, dolo = dolomite, ank = ankerite, py = pyrite, aspy = arsenopyrite, and ccpy = chalcopyrite.

Unstable = phase undergoes partial or complete alteration (minor alteration products given in parentheses). Stable = phases not affected by alteration.Abbreviations: bio = biotite, kaol = kaolinite, ser = sericite, musc = muscovite, chl = chlorite, epi = epidote, qz = quartz, rut = 3.5. Ore mineralogy Ore microscopy combined with the microprobe data revealed premineralization opaque minerals in the wallrock, including cobaltite, magnetite and hematite, and two assemblages of sulfide minerals in the mineralized quartz veins and wallrock, which are interpreted as distinct generations. The early generation includes arsenopyrite, pyrite, chalcopyrite, pyrrhotite, sphalerite, and gold. This assemblage occurs commonly in the quartz–carbonate veins as disseminations or clustered in quartz veinlets traversing the carbonate aggregates. A later generation, dominated by pyrite, arsenopyrite and subordinate galena and chalcopyrite, is common in the laminated quartz veins as disseminated, fine prismatic crystals. Marcasite, rare diginite and unidentified hydroxyl arsenate are disseminated in the altered wallrocks and post-date gold mineralization. Arsenopyrite is the most abundant ore mineral (N50 vol.% of sulfides), commonly occurring as aggregates of euhedral to subhedral crystals. Early arsenopyrite occurs as individual or clustered rhombic grains disseminated in the quartz–carbonate veins, commonly in the carbonate-rich domains. It is associated with euhedral As-bearing pyrite, with serrate or interlocking outlines (Fig. 7A, B) and is generally zoned and inclusion-rich (Fig. 7C). The inclusions, commonly close to the grain margins, are mostly chalcopyrite, pyrrhotite, sphalerite and gold and/or electrum; silver inclusions are rare. In arsenopyrite adjacent to pyrite, gold inclusions are located close to the grain boundaries. A schematic paragenetic sequence of the ore and gangue minerals (Fig. 16) includes the three main generations: (a) pre-ore stage: garnet, staurolite, biotite, ilmenite, magnetite, hematite and cobaltite; (b) ore stage: involved an early sub-stage during which arsenopyrite, pyrite, chalcopyrite, pyrrhotite, sphalerite, and gold were deposited and a late sub-stage during which pyrite, arsenopyrite and subordinate galena, chalcopyrite and gold were deposited. Quartz and quartz–Fe–Mg–carbonate veins and disseminations were also developed during this stage, in addition to chlorite and sericite; and

(c) post-ore stage: marcasite, diginite, rutile, goethite and less common unidentified hydroxyl arsenate. Ore mineral chemistry was determined on a Zeiss 950 DSM electron microprobe at the Institute of General and Applied Geology, University of Munich. The applied accelerating voltage was 20 kV, and specimen current was 25 nA. A fixed time of 20 s was used and each analysis was based on counts from two or three spots on a single grain. The whole data set and additional analytical parameters are available in Zoheir (2004) or on demand from the author. Variation in the As content within a single arsenopyrite crystal (b1 at.%) is typically less than variations between different crystals in the same sample (up to 3 at.%). Arsenopyrite from the late assemblage occurs as small (b150 μm) grains, optically homogenous and devoid of inclusions, common in the chlorite–‘sericite’ ± graphitebearing laminations within the laminated quartz veins. Gold occurs as ≤ 50 μm inclusions in arsenopyrite (Fig. 7D), ≤ 10 μm blebs at pyrite– arsenopyrite contacts (Fig. 7E), ≤ 70 μm dispersed grains in the quartz–carbonate-rich domains and ≤ 5 μm stringers along rhythmic zones goethite (Fig. 7F) or along microfissures. Microprobe data indicate that Au correlates positively with As/S, and inversely with at.% Fe in both pyrite and arsenopyrite. This relationship may denote substitution of Fe by Au, for which the mechanism: [2As[Fe] ↔ (Au,Sb) + Fe], was proposed by Johan et al. (1989), where As[Fe] is As in Fe sites. Fleet and Mumin (1997) suggested that invisible Au in arsenopyrite represents Au removed from ore fluids by chemisorption at As-rich, Fe-deficient surface sites and which was then incorporated as a metastable solid solution. This implies that invisible structurally bound Au is present in arsenopyrite and As-bearing pyrite from the Um El Tuyor deposit. Assuming equilibrium conditions, the composition of arsenopyrite is a function of fS2 and temperature. It is possible to estimate the formation conditions of arsenopyrite based on the Fe–As–S mineral assemblage and the As at.% content of arsenopyrite (Barton, 1969; Kretschmar and Scott, 1976; Scott,1983; Sharp et al.,1985). In order to detect any zonation within arsenopyrite, points from grain cores and rims were routinely analyzed. In


455

4. Hydrothermal alteration

some samples, arsenopyrite grains are compositionally more variable than others. The conspicuous presence of pores in the marginal zones of some grains indicates subsequent chemical overprinting. This textural evidence may reflect varying degrees of retention of equilibrium; these grains were therefore avoided for geothermometry. Arsenopyrite of the early assemblage (early mineralization), associated with pyrite and pyrrhotite, shows a range of 30.83 to 31.14 at.% As. This assemblage implies a temperature range of 353 to 372 °C (Fig. 8) using the calibration of Kretschmar and Scott (1976). Pyrite-associated arsenopyrite from the late assemblage has a range of 29.3 to 30.41 at.% As, assuming temperatures of 312 to 325 °C. Sulfur fugacity (−f S2) in the earlier mineralization phase was around 10− 8.5 bar, whereas the late assemblage indicate f S2 ∼10− 10 bar.

The gold-bearing quartz veins are surrounded by a roughly 80 m-wide halo of bleached rocks with disseminated arsenopyrite, pyrite and carbonate; sulfidation is limited to a narrow zone, 10 m wide, adjacent to the shear zone. Around the veins which are parallel to the foliation, this halo is symmetric, whereas the alteration halo is indented where veins cutting the foliation. This observation suggests that fluid percolation was guided by the host rock foliation. The alteration types include ‘sericitisation’, carbonatisation, sulfidation, chloritisation, and silicification. A sample map (Fig. 9) shows the distribution of hydrothermal alteration in the mine area. Three main alteration zones are distinguished, including (a) an inner quartz–carbonate–‘sericite’/ muscovite–sulfide ± graphite ± albite-rich zone, (b) an intermediate locally pervasive ‘sericite’–chlorite–carbonate–quartz–sulfide ± biotite zone, and (c) an outer zone of chlorite–‘sericite’ alteration. Boundaries among these zones are gradational over the cm- to dm-scales. Generally, the degree of quartz veining, gold and base metal enrichment increases from the outer towards the inner alteration zones. The inner zone displays the highest gold grade, especially adjacent to the laminated quartz veins. The spatial distribution of the hydrothermal phases and mineralogical changes throughout the three alteration types is summarized in Table 2. Alteration mineralogy, style, and intensity of the three alteration types are discussed in the following sections. Microanalytical data on hydrothermal silicate minerals was determined with a JEOL JSM-6310 electron microprobe at the Institute of Geology and Mineralogy, Graz University, Austria. Analytical conditions were 15 kV and 30 nA, using a defocused electron beam with a diameter of 5 to 10 μm.

Fig. 11. (A) Composition of chlorites from the intermediate alteration zone (classification of Hey, 1954). (B) Frequency histogram showing ranges of the estimated temperatures of chlorites (using the calibration of Cathelineau, 1988) from the distal (dotted blocks) and intermediate (dark blocks) alteration zone in the Um El Tuyor mine area.

Fig. 12. Scannng electron microscope images showing; (A) Large muscovite flake with carbonate minerals along the cleavage planes and dispersed sericite flakes near margin. (B) Zoned albite, carbonate and sericite rimming relict chlorite.

Fig. 10. Backscattered electron image showing biotite replaced by sericite, chlorite and carbonate from the distal (outer) alteration zone.

456


4.1. Outer zone (chlorite–sericite ± carbonate alteration)

4.2. Intermediate zone (quartz–chlorite–‘sericite’–carbonate alteration)

The outer alteration zone is characterized by preferential replacement of the primary garnet–biotite ± staurolite ± hornblende mineralogy by an association of chlorite, ‘sericite’ and less abundant carbonate and sulfides. Replacement of biotite plates by fine-grained aggregates of ‘sericite’, chlorite and carbonate is common in most samples (Fig. 10). Generally, the chlorite-rich domains contain very fine acicular shreds of magnetite ± hematite. Sericitic alteration overprints the biotite + chlorite (± minor rutile) alteration. The chlorite–‘sericite’ alteration is traversed by few b2 cm-wide veinlets of vuggy quartz, enclosing thin chlorite selvages. Chlorite from this zone shows variable AlIV (1.67 to 1.98) but rather constant Fe/(Fe + Mg) ratios (0.43 to 0.47). Replacement of the metamorphic assemblage to form a ‘sericite’–chlorite–carbonate–quartz rock was likely by way of interaction with a CO2 + K+-rich fluid that may evolve to a H2O + Na+ fluid.

The intermediate alteration zone is marked by an assemblage of chlorite, ‘sericite’, ferrodolomite, quartz, arsenopyrite and pyrite. Chlorite (±Fe-oxides) gives this zone with a greenish to brick-reddish color. This alteration type is typically developed in the sheared rocks, and is traversed by a dense set of 3 to 12 cm-wide barren quartz–hematite and quartz ± ‘sericite’ veins. Relics of fine-grained biotite are preserved. Chlorite forms colorless to olive green aggregates with rosette morphology, filling the interstices between the porphyroblasts. Microprobe data show that chlorite from the intermediate alteration zone have Fe- and Mg-chlorite compositions (Fig. 11A). Chlorite thermometry (Cathelineau, 1988) indicates temperatures of 278 to 346 °C for this zone. Temperatures based on chlorite from the outer alteration zone range from 221 to 261 °C (Fig. 11B). ‘Sericite’ occurs as fine-grained flakes, commonly associated with chlorite. Chemical compositional variations are controlled by phengitic substitution of Si for Al (Fe(VI), Mg(VI) + Si(IV) ⬄Al(VI)+ Al(IV)).

Table 3 Compositional variations of the least-altered and altered rocks from the different alteration zones Least-altered host rocks (initial precursor)

Distal alteration zone

Sample

258

291

243

261

293

264

299

252

Intermediate alteration zone 269

271

265

266

Proximal alteration zone 301

9M

SiO2 TiO2 Al2O3 Fe2O3⁎ MnO MgO CaO Na2O K2O P2O5 LOI SO2 Total

53.96 0.69 9.78 8.86 0.17 13.81 2.52 0.73 4.06 0.53 4.75 0.05 99.91

54.02 0.73 10.13 7.87 0.21 12.37 3.23 0.98 3.99 0.39 4.33 0.10 98.35

51.81 0.57 9.56 9.06 0.29 13.35 4.13 0.93 3.88 0.42 4.82 0.08 98.91

52.37 0.58 9.93 10.97 0.20 11.58 3.42 0.87 3.37 0.54 4.67 0.07 98.57

50.51 0.63 9.14 10.05 0.36 13.44 4.82 0.65 4.47 0.38 5.03 0.22 99.70

51.34 0.77 9.23 8.94 0.23 12.91 5.11 0.94 4.98 0.49 4.87 0.21 100.1

51.87 0.64 8.79 9.52 0.26 13.08 5.17 0.43 4.64 0.27 5.13 0.35 100.2

53.93 0.78 11.62 8.28 0.31 8.94 2.98 0.86 5.95 0.17 4.98 0.43 99.2

56.26 0.64 11.29 7.68 0.28 7.42 3.85 1.23 5.73 0,44 5.01 0.51 100.4

55.35 0.80 12.32 8.53 0.26 7.31 3.22 1.09 5.77 0.39 4.69 0.49 100.2

56.44 0.66 11.14 8.66 0.30 7.32 3.95 1.51 5.43 0.42 4.20 0.39 100.4

55.14 0.49 11.62 8.16 0.28 7.15 3.35 1.97 6.05 0.33 5.03 0.55 100.1

57.81 0.51 11.5 9.03 0.31 6.24 2.83 2.14 5.18 0.29 4.27 0.47 100.6

56.12 0.69 12.00 8.18 0.24 6.42 3.71 1.83 5.06 0.38 5.02 0.68 100.3

Ppm Au Ag As Ba Br Co Cr Ni Rb Sb Sc Sn Sr Th U Zn Pb V Y Cu Mn

0.03 – 301 63 – 19 83 16 36 – 0.9 13 71 b 0.2 – 24 9 26 2 11 413

– – 296 60 16 66 15 34 – – 30 73 b0.2 – 22 8 32 3 10 361

0.02 – 243 51 – 15 67 13 29 – 0.73 11 57 b 0.2 – 19 7 21 2 12 334

0.05 – 389 76 – 16 58 19 41 – 0.66 71 102 b 0.2 – 28 8 53 5 9 393

0.62 0.21 557 196 0.58 13 51 15 b 15 b 0.1 1.09 143 365 b 0.2 b 0.5 53 8 44 4 6 533

0.65 0.09 683 153 0.48 12 68 6 b15 b0.1 0.8 165 203 b0.2 b0.5 77 8 29 1 4 709

0.58 0.16 574 222 0.59 13 46 24 b 15 b 0.1 0.41 139 255 b 0.2 b 0.5 28 13 91 6 10 293

1.14 0.21 1237 219 0.59 23 33 37 51 b 0.1 4.77 133 237 2.34 b 0.5 26 0 123 5 0 1237

1.33 0.25 1522 232 0.48 12 49 21 43 b 0.1 7.50 109 222 3.89 b 0.5 31 12 77 6 10 352

1.42 0.16 1759 290 0.36 17 44 28 56 b 0.1 10.9 153 162 7.3 2.26 38 9 95 8 9 463

1.14 0.09 1426 271 0.31 13 49 21 48 b 0.1 7.84 80 118 5.13 1.31 22 11 74 5 11 405

1.72 0.26 2016 271.2 0.43 14 48 26 59 b 0.1 11.17 126 169 5.74 1.74 30 13 97 7 11 352

2.13 0.43 2454 199 0.5 8 41 25 64 b 0.1 13 91 152 3.6 0.8 13 17 104 7 12 163

1.61 0.28 2182 245 0.72 16 44 42 57 0.2 8.94 119 196 3.93 – 25 10 108 6 8 234

REEs La Ce Nd Sm Eu Tb Yb Lu

23.1 48 19 4.6 0.9 0.6 0.7 0.13

20.7 42 17 4.8 0.8 0.57 0.64 0.09

19.7 39 15 5.1 0.7 0.5 0.6 0.11

22.5 46 18 4.7 0.8 0.6 0.9 0.08

3.1 34 16 2.6 0.7 0.7 b 0.1 0.13

2.5 23 14 1.4 0.5 0.66 b0.1 0.15

18.9 42 18 2.9 0.7 0.61 b 0.1 0.21

16.8 44 18 3.6 0.8 b 0.5 2.11 0.19

18.3 45 16 4.3 0.8 b 0.5 2.01 0.15

27.9 55 22 5.3 0.9 b 0.5 2.8 0.12

23.7 46 14 4.9 0.8 b 0.5 1.93 0.13

27.6 54 21 5.3 0.9 b 0.5 2.78 0.15

26.3 50 18 4.8 0.9 b 0.5 2.7 0.18

23.6 42 19 4.7 0.9 b 0.5 2.33 0.18

Density g/cm3

2.71

2.74

2.75

2.81

3.02

2.97

2.97

2.87

2.84

2.89

2.88

2.86

2.87

2.87

Major and trace element analyses have been carried out by using an ICP-MS technique in the Activation Laboratories, ACTLABS Group, Canada (Code 4B).


457

This is confirmed by the negative correlation between (Si +Mg+ Fe2+) versus (AlIV +AlVI). In individual flakes, cores show higher celadonite content relative to the margins. This observation, together with the enrichment in AlVI and K in grain margins is attributed to some degrees of paired substitution of (Mg,Fe) for Al, and Si for Al. Carbonate minerals partially or completely replace chlorite and biotite. They are mainly ironbearing dolomite (ferroan dolomite), Ca(Fe,Mg,Mn)(CO3)2 solid solution, with a rhodochrosite component not exceeding 2 mol%. 4.3. Inner zone (quartz–‘sericite’–carbonate–sulfides ± albite ± graphite alteration) Although not extensive, the quartz–carbonate–‘sericite’/muscovite– sulfide± graphite ± albite assemblage is the most conspicuous hydrothermal alteration in the mine area. The thickness of this alteration zone is proportional to the width of the quartz veins, rarely exceeding a few meters. Dense quartz veining is distinct within this zone. Within and adjacent to the quartz veins, carbonate intermingles with fine-grained quartz and sericite. Carbonates are mainly Fe-dolomite and ankerite; calcite is a minor component. Higher concentrations of sulfide minerals are common in zones where carbonate minerals are substantial constituents of the veins and wallrock. Backscattered electron images show a spatial relationship between sulfides and zones rich in carbonate having high siderite component, suggesting that aFe and fS2 of the ore fluids controlled carbonate composition. ‘Sericite’ occurs as randomly to weakly-oriented fine-grained flakes and aggregates associated with carbonate (Fig. 12A). Na2O contents in ‘sericite’ are higher compared to that of muscovite, whereas contents of TiO2 and MgO are higher in muscovite. ‘Sericite’ from the inner alteration zone is rich in Na+ and poor in K+ and phengite content (atomic Fe + Mg a.p.f.u.) compared to ‘sericite’ from the intermediate alteration zone. Albite is confined to the contact between quartz veins and the quartz–‘sericite’–carbonate-rich wallrock (Fig. 12B). It occurs as fine-grained, subhedral crystals in and adjacent to the quartz–carbonate veins. The large crystals show compositional zoning, Na-rich rims and K-rich cores. Graphite occurs as tiny laths commonly associated with ‘sericite’ in zones enriched in sulfides. 5. Hydrothermal alteration geochemistry Direct observations indicated higher K2O, Fe2O3⁎ and lower MgO contents in the hydrothermally altered wallrocks compared to the leastaltered rocks (Table 3). Loss on ignition (LOI) is considerably higher in samples displaying intense alteration. In the altered and unaltered rocks, the Rb/Sr, Rb/Ba, and K/Rb ratios are low and almost identical. However, the Ba/La ratios are generally higher in the intensely altered rocks relative to those in the unaltered rocks (Table 3). A single alteration trend is obtained by plotting TiO2 vs. Al2O3 contents in the altered samples and average of the least- or unaltered precursor (far away from the alteration zone where the country rocks look clearly unbleached). This single trend is considered as indicative of an initially homogeneous precursor (Fig. 13). The ‘sericitized’ rocks show an apparent mass loss, whereas a net mass gain is revealed for samples of the propylitized (chloritized/carbonatized) rocks compared to the least or unaltered precursor. 5.1. Mass balance calculations The isocon method of Grant (1986) is used in this context to evaluate whether the compositional changes involved significant changes in concentrations of components as well as in mass and volume. Al, Ti, P, Y, Sc and Tb are used in the present study to monitor mass transfer in the isocon calculations (e.g., Selverstone et al., 1991; Leitch and Lentz, 1994; Kolb et al., 2000). For each alteration type, convergence among the concentrations of these elements in the leastor unaltered rocks and in the hydrothermally altered wallrocks defines a best fit isocon.

Fig. 13. Binary TiO2 vs. Al2O3 plot of the altered samples and average of the least-altered precursor. The correlation coefficient (r) is calculated using all altered samples, with significance value of b0.01 in all cases.

Chemical comparisons were made among the respective medians of the unaltered and altered rocks in the mine area. The use of the statistical median, instead of the absolute concentration in a specific sample of the altered and least-altered rocks, minimizes error introduced by heterogeneity which might be expected in the host metasedimentary rocks. In many cases, the median values are similar to the chemical analyses corresponding to the most representative samples of each group, selected on basis of petrographic observations. Mass changes in each of the alteration types, with respect to the least-altered precursor, are summarized in Table 4 and shown in the isocon plots (Fig. 14). Estimation of the mass factor may be doubtful if the metasedimentary rocks vary compositionally with lateral and vertical extension. However, the convergence between compositions of each of the four representative samples with the median of these four samples yielded isocons having mass factors very close to 1, which reveals a very similar composition of the chosen precursor samples. Based on the latter observation, the median of the four representative samples of the leastor unaltered host rocks has been found to be reliable as representing the precursor for comparison with the altered wallrocks. Based on the mass balance calculations (according to equations of Gresens, 1967 and Grant, 1986), the calculated volume factors (fV) are very close to unity (b1.06), indicating that there was no significant volume or mass change. Slight loss of SiO2, Na2O, Al2O3 and P2O5, and addition of CaO, MgO, Fe2O3⁎, K2O, SO2 and LOI characterize the major chemical changes in the outer alteration zone. Samples from the intermediate alteration zone exhibit slight addition of SiO2, K2O, CaO and SO2, and loss of MgO. Rocks from the inner alteration zone underwent enrichment in SiO2, Na2O, K2O and SO2, and significant depletion in MgO (Fig. 15). The three alteration zones are generally enriched in a range of trace elements (S, Au, As, Ba, Ni, Rb, Sn, Pb, V, Zn, Cu and Sr). However, it is noted that the light rare earth elements (i.e., La and Ce) are depleted in the outer zone, whereas, the intermediate and inner zones show variable increase in their concentrations. The addition of V, Ni and less commonly Cr in most of the altered samples reflects a source enriched in these elements. With regard to the major elements, the progressive changes indicated by mass balance calculations reveal a systematic increase in SiO2, K2O, Fe2O3⁎, LOI, and S and less distinctly in Na2O. Mg shows systematic depletion from the outer to inner alteration zone, whereas, Ca is mobile and shows variable patterns in the alteration zones. Al2O3, is commonly increases with increase of K2O, suggesting progressive sericitization. 5.2. Alteration model Changes in host rock composition, mineral microanalyses and modal mineralogy are used to decipher the type and extent of fluid– wallrock interactions during alteration. Based on microscopic and field relationships, the paragenetic sequence of mineral phases associated with the investigated deposit is presented in Fig. 16. Applying the local equilibrium concept to the investigated hydrothermal alteration zones

458

Table 4 Mass changes, grams added or removed per 100 g in the major elements and per 1000 g in trace elements during hydrothermal alteration in the Um El Tuyor deposit Distal (propylitized) rocks

% gains and losses

Median

293

264

299

293

264

299

av.

252

269

271

265

252

269

271

265

av.

266

301

9M

266

301

9M

av.

53.65 0.64 9.94 9.04 0.20 12.98 3.36 0.91 3.97 0.48 4.77 0.07 100.2

50.51 0.63 9.14 10.05 0.36 13.44 4.82 0.65 4.47 0.38 5.03 0.22 99.7

51.34 0.77 9.23 8.94 0.23 12.91 5.11 0.94 4.98 0.49 4.87 0.21 100

51.87 0.64 8.79 9.52 0.26 13.08 5.17 0.43 4.64 0.27 5.13 0.35 100.2

− 2.82 − 0.01 − 0.74 1.07 0.16 0.54 1.49 − 0.26 0.53 − 0.10 0.31 0.15

−2.72 0.12 −0.79 −0.17 0.02 −0.17 1.71 0.02 0.97 0.01 0.08 0.14

− 2.28 − 0.01 − 1.24 0.39 0.05 − 0.03 1.77 − 0.48 0.62 − 0.22 0.33 0.27

−2.60 0.03 −0.92 0.43 0.08 0.11 1.65 −0.24 0.71 −0.10 0.24 0.19

53.93 0.78 11.62 8.28 0.31 8.94 2.98 0.86 5.95 0.17 4.98 0.43 99.2

56.26 0.64 11.29 7.68 0.28 7.42 3.85 1.23 5.73 0.44 5.01 0.51 100.3

55.35 0.80 12.32 8.53 0.26 7.31 3.22 1.09 5.77 0.39 4.69 0.49 100.2

56.44 0.66 11.14 8.66 0.30 7.32 3.95 1.51 5.43 0.42 4.20 0.39 100.4

0.88 0.14 1.81 −0.67 0.11 −3.94 −0.35 −0.04 2.04 −0.31 0.28 0.36

2.81 0.00 1.39 −1.34 0.08 −5.53 0.50 0.32 1.78 −0.04 0.28 0.44

2.57 0.17 2.57 −0.38 0.06 −5.55 −0.09 0.20 1.89 −0.08 0.01 0.43

3.85 0.04 1.40 − 0.22 0.10 − 5.52 0.66 0.63 1.56 − 0.06 − 0.47 0.33

2.53 0.09 1.79 −0.65 0.09 −5.14 0.18 0.28 1.82 −0.12 0.02 0.39

56.12 0.69 12.00 8.18 0.24 6.42 3.71 1.83 5.06 0.38 5.02 0.68 100.3

55.14 0.49 11.62 8.16 0.28 7.15 3.35 1.97 6.05 0.33 5.03 0.55 100.1

57.81 0.51 11.55 9.03 0.31 6.24 2.83 2.14 5.18 0.29 4.27 0.47 100.6

4.99 0.08 2.59 −0.49 0.05 −6.27 0.52 1.00 1.32 −0.08 0.50 0.64

3.97 −0.13 2.19 −0.51 0.09 −5.51 0.14 1.14 2.35 −0.13 0.50 0.50

6.76 −0.10 2.12 0.40 0.12 −6.46 −0.40 1.33 1.44 −0.17 −0.29 0.42

5.24 −0.05 2.30 −0.20 0.09 −6.08 0.09 1.16 1.70 −0.13 0.23 0.52

g/1000 kg S 353 Au 0.03 Ag 0.00 As 298 Ba 62 Co 16 Cr 67 Ni 16 Rb 35 Sn 22 Sr 72 Zn 23 Pb 8 V 29 Y 2 Cu 11 Mn 377 La 21.58 Ce 44.13 Nd 17.52 Sm 4.73 Eu 0.83 Tb 0.58 Yb 0.67 Lu 0.10

1110 0.62 0.21 557 196 13 51 15 b 15 143 365 53 8 44 4 6 533 3.06 34.27 16.03 2.58 0.68 0.70 b 0.1 0.13

1050 0.65 0.09 683 153 12 68 6 b15 165 203 77 8 29 1 4 709 2.50 23.00 14.00 1.37 0.49 0.66 b0.1 0.15

1733 0.58 0.16 574 222 13 46 24 b 15 139 255 28 13 91 6 10 293 18.98 41.69 17.83 2.89 0.74 0.61 b 0.1 0.21

764 0.59 0.21 263 135 −3 − 15 −1 – 123 295 30 0 16 2 −5 160 − 18.51 − 9.65 − 1.39 − 2.13 − 0.14 – – 0.03

688 0.61 0.09 379 90 −4 1 −9 – 142 129 53 0 0 −1 −6 326 −19.11 −21.31 −3.63 −3.37 −0.34 – – 0.05

1363 0.54 0.16 270 158 −4 − 20 8 – 116 180 5 5 61 4 0 − 87 − 2.79 − 2.84 0.15 − 1.86 − 0.09 – – 0.11

939 0.58 0.15 304 128 −4 −12 −1 – 127 202 29 1 26 1 −4 133 −13.47 −11.27 −1.62 −2.45 −0.19 – – 0.06

2157 1.14 0.21 1237 219 23 33 37 51 133 237 26 0 123 5 0 1237 16.81 44.39 18.36 3.60 0.82 b0.5 2.11 0.19

2553 1.33 0.25 1522 232 12 49 21 43 109 222 31 12 77 6 10 352 18.27 45.20 15.94 4.25 0.83 b0.5 2.01 0.15

2453 1.42 0.16 1759 290 17 44 28 56 153 162 38 9 95 8 9 463 27.90 55.00 22.00 5.30 0.90 b0.5 2.80 0.12

1953 1.14 0.09 1426 271 13 49 21 48 80 118 22 11 74 5 11 405 23.70 46.12 13.61 4.90 0.81 b 0.5 1.93 0.13

1828 1.13 0.21 953 160 7 −33 22 17 113 167 3 −8 96 3 −10 874 −4.59 0.75 1.05 −1.09 0.00 – 1.46 0.09

2209 1.31 0.25 1229 171 −4 −17 5 8 88 151 8 4 48 3 −1 −24 −3.26 1.23 −1.52 −0.46 0.01 – 1.35 0.05

2138 1.41 0.17 1488 233 1 −22 12 22 134 92 15 1 67 5 −1 94 6.75 11.73 4.83 0.66 0.09 – 2.17 0.02

1637 1.13 0.10 1154 215 −3 − 16 6 14 60 49 −1 3 47 3 0 36 2.56 2.85 − 3.65 0.27 0.00 – 1.29 0.03

1953 1.25 0.18 1206 195 0 −22 11 15 99 115 6 0 65 3 −3 245 0.37 4.14 0.17 −0.15 0.03 – 1.57 0.05

3399 2 0 2016 271 14 48 26 59 126 169 30 13 97 7 11 352 27.6 53.6 20.6 5.3 0.9 b0.5 2.8 0.2

2744 2 0 2454 199 8 41 25 64 91 152 13 17 104 7 12 163 26.4 50.0 18.0 4.8 0.9 b 0.5 2.7 0.2

2336 2 0 2182 245 16 44 42 57 119 196 25 10 108 6 8 234 23.6 41.7 19.2 4.7 0.9 b0.5 2.3 0.2

3199 2 0 1808 222 −2 −17 12 27 110 104 8 6 73 5 1 −9 7.26 11.85 3.99 0.84 0.15 – 2.24 0.06

2503 2 0 2255 146 −8 −24 10 32 73 86 −10 10 80 5 2 −207 5.83 7.90 1.21 0.27 0.11 – 2.14 0.09

2095 2 0 1989 195 0 −21 28 25 103 134 3 2 84 4 −2 −131 3.11 −0.41 2.66 0.20 0.12 – 1.77 0.09

2599 2 0 2017 188 −3 −20 17 28 96 108 0 6 79 5 0 −116 5.40 6.45 2.62 0.44 0.13 0.00 2.05 0.08

g/100 g SiO2 TiO2 Al2O3 Fe2O3⁎ MnO MgO CaO Na2O K2O P2O5 LOI SO2 Total

Intermediate (sericitized) rocks

% gains and losses

Inner (sericitized– albitized) rocks

% gains and losses


Least-altered


459

minerals. The result is carbonate- and sulfide-bearing, quartz–sericite (±chlorite) zone next to mineralized veins. The transitional stage is characterized by carbonatization and sulfidation of the ferromagnesian silicates (reaction 6). Time relationship between sulfidation and carbonatization is not clear, as sulfides and carbonate-rich domains intermingle and overlap with one another. However, the occasional presence of sulfide-rich veinlets crosscutting the carbonate domains suggests that carbonatization was likely earlier. The spatial association of chlorite and Fe-sulfides with hydrothermal quartz is consistent with chloritization (reaction 7). The change from chlorite or chlorite+calcite (products of reaction 3) to sericite+ ankerite is considered as a result of reactions involving addition of CO2 and K+ and release of H+ (reactions 8 and 9). The presence of biotite relics in the intermediate alteration zone implies incomplete transformation (reaction 10). Sulfidation is seen through the conversion of biotite to sericite in the presence of H2S (reaction 11). This is evident by development of

Fig. 14. Isocon diagram comparing median composition of samples representing the least-altered rocks and median composition of samples from the (A) distal, (B) intermediate and (C) proximal alteration zones. The dashed line (isocon) is drawn for the largest number of potentially immobile elements. The best constraints on mass changes are given by TiO2, Al2O3, P2O5, Ni and Ce. The inset histogram shows the distribution of mass factors calculated for each element.

was aimed at understanding the evolution of the ore fluid during progressive alteration. Based on published investigations (Meyer and Hemley, 1967; Mueller and Saxena, 1977; Kishida and Kerrich, 1987; Groves and Foster, 1991; Barton and Skinner, 1979), a group of simplified chemical reactions are assumed to have been responsible for the main mineral transformations induced by hydrothermal alteration in the mine area (Table 5). The incipient mineralogical transformations are typified by near-complete conversion of the Fe– Mg-silicates into chlorite and muscovite by means of hydrolysis, probably according to one or more of the reactions 1, 2, 3 and 4. Formation of the hydrothermal biotite implies fluid–rock interaction in presence of CO2 (reaction 5). The iron released by such reactions might have been fixed by the formation of ferrodolomite and sulfide

Fig. 15. Profile-histograms of mass changes (grams added or removed per 100 g country rock) in major element oxides during alteration based on median of the least-altered samples and averages of mass changes in the hydrothermally altered samples from the three alteration types in the investigated deposit. (A) Distal zone, (B) Intermediate zone and (C) Proximal alteration zone.

460


thin envelopes of sericite chlorite adjacent to the sulfide-rich/ carbonatized zones, and bordering the quartz–pyrite veinlets. It might also have taken place by oxidation of chlorite or magnetite by H2S-bearing solutions (reaction 12). Alternatively, if HS− was the main sulfur species in the fluid instead of H2S, the pyrite-forming reaction might directly involve acid-base equilibria (reactions 13 and 14). The latter are redox reactions (S2− is oxidized to S−) and, therefore, tend to reduce the interacting fluids particularly where the fluid/rock ratio is low, thus providing favorable conditions for gold deposition (Phillips and Groves, 1983). However, the abundance of CO2 in the fluid, − inferred from the abundant carbonate suggests that H2CO3, CO2− 3 or HCO3 rather than H2S, HS− and S equilibria controlled the pH. In such case, Phillips and Groves (1983) considered that pyritization does not directly affect gold solubility via pH changes, but may operate through redox reactions. The result is a bleached ankerite+sericite+quartz-rich zone with discrete gold blebs. A zone with these characteristics has been identified in the vicinity of the auriferous quartz veins in the mine area. A sericite–ankerite–albite–quartz assemblage is typical for the advanced alteration stage in the studied deposit. Sericitic alteration is indicated by the complete replacement of chlorite and biotite by sericite and quartz. This replacement can be explained by a reaction involving a release of H+ (reaction 15). If Al3+ cations were present in the fluid, the reaction can balanced as in reaction 16. Locally, cation-exchange reactions between K-bearing mineral species (sericite) and circulating fluids promoted deposition of albite, favored by relatively high aNa+/aK+ in the fluid at higher temperature (Audétat et al., 1998). A chemical reaction for albitization of a K-bearing mineral species can be expressed in reaction 17. Circulation of cations released from the last reaction would facilitate the development of additional sericite at the expense of chlorite relics (reaction 15). More or less contemporaneous with the aforementioned reactions is reduction of magnetite to form sulfides by (reaction 18 or 19). However, the presence of CO2-fluid in the hydrothermal system might have prevented such reaction from proceeding to convert all the magnetite to pyrite. Alternatively, reduction of CO2 on the magnetite surface might have initiated graphite precipitation. Only one molecule of graphite can be derived from four molecules of magnetite (reaction 20). This may provide an explanation for the scarcity of graphite in the wallrock. Further, it explains the presence of graphite mainly in the hematized domains. Finally, advanced argillic alteration produced kaolinite from albite, sericite and/or detrital K-feldspar by hydrolysis under acid low temperature conditions.

Changes in the solution pH through fluid–rock buffer reactions occur in response to the exchange of ions between fluids and the primary minerals of the wallrock. According to the aforementioned alteration reactions and stages, the present model assumes that interaction between the host rocks and a weakly acidic solution resulted in an overall increase in pH as the rock neutralizes the solution. This pH change has likely occurred when the aqueous H+ was consumed and K+ was released, reaction 21 (e.g., Giggenbach, 1987, 1988). Precipitation of muscovite/sericite consumed K+ and silica, and concentrated H+ until muscovite saturation (reaction 22). As a result, the solution pH has likely decreased during the transitional and late (advanced) stages. 6. Fluid inclusions Detailed petrographic study of some 9 representative samples (5 quartz–carbonate and 4 laminated quartz veins), allowed the relative chronology of fluid inclusions to be established using the criteria of Roedder (1984). Microthermometric measurements were obtained on a Linkam THM 600 heating/freezing stage, fitted with a thermal control unit TMS-93 and equipped with a binocular Olympus microscope at the Institute of Applied Mineralogy and Geochemistry, TU Munich. Additional measurements were made using a similar stage at the Institute of Earth Sciences, Tübingen. The stage enables measurements within the range of −196 and 600 °C. Freezing and heating runs were, respectively, undertaken using liquid nitrogen and a thermal resistor. Calibration of the stage was carried out by using standard natural and synthetic inclusions. To ensure accurate observation of melting behavior, slow heating rates (0.2 °C/min) have been used, whilst very fast freezing rates were used to avoid meta-stability. For all final melting temperatures of CO2, precision is ±0.1 °C. These studies aimed at assessing the nature and evolution of the mineralizing fluids and to consider the physicochemical parameters which controlled gold deposition in regard to the regional metamorphic and magmatic framework of the country rocks in the study area. The succession of fluids has been studied by examining relationships between fluid inclusions, their host mineral, geometry of the host microstructures and location of ore minerals. Under conditions of intermediate temperatures, brittle and ductile deformation, the rates of dynamic recrystallization may exceed rates of recovery, destroying the primary depositional features and producing abundant secondary and pseudosecondary fluid inclusions along

Fig. 16. Simplified paragenetic sequence of the hydrothermal mineral species associated with the Um El Tuyor gold deposit.

ðFeþþ ; MgÞ2 Al9 ðSi; AlÞ4 O20 ðO; OHÞ4 staurolite þ 2KðFe; Mg; TiÞ3 AlSi3 O10 ðOHÞ2 biotite þ 2SiO2 fluid þ6Hþ þ Kþ þ 4O2− Y ðFe; Mg; TiÞ5 Al2 Si3 O10 ðOHÞ8 chlorite þ 3KAl3 Si3 O10 ðOHÞ2 muscovite þ TiO2 rutile þ 3ðFe; MgÞ2þ ðaq:Þ .....(1) 2KðFe; Mg; TiÞ3 AlSi3 O10 ðOHÞ2 biotite þ 4Hþ fluid þ12O2− Y ðFe; Mg; TiÞ5 Al2 Si3 O10 ðOHÞ8 chlorite þ 3SiO2 quartz þ 6TiO2 rutile þ ðFe; MgÞ2þ ðaq:Þ þ2Kþ .....(2) 2Ca2 ðMg; Fe; AlÞ5 Si8 O22 ðOHÞ2 hornblende þ 6H2 Ofluid þ4CO2 þ 3O2 Y 2ðFe; MgÞ5 Al2 Si3 O10 ðOHÞ8 chlorite þ 4CaCO3 calcite þ 10SiO2 quartz .....(3) ðFe; MgÞ3 Al2 ðSiO4 Þ3 garnet þ 2KðFe; Mg; TiÞ3 AlSi3 O10 ðOHÞ2 biotite þ Hþ fluid þ2O2− Y ðFe; Mg; TiÞ5 Al2 Si3 O10 ðOHÞ8 chlorite þ 3KAl3 Si3 O10 ðOHÞ2 muscovite þ 3SiO2 quartz .....(4) 2Ca2 ðMg; Fe; TiÞ4 Al2 Si7 O22 ðOHÞ2 hornblende þ 2ðFe; MgÞ3 Al2 ðSiO4 Þ3 garnet þ8CO2 þ 2H2 Ofluid þ4Kþ þ 7O2 Y 2KðFe; Mg; TiÞ3 AlSi3 O10 ðOHÞ2 biotite þ 2KAl3 Si3 O10 ðOHÞ2 sericite þ 4CaMgðCO3 Þ2 dolomite þ 2TiO2 rutile þ 4Fe2 O3 hematite þ 8SiO2 quartz .....(5) 4CaAl2 Si2 O8 plagioclase þ Ca2 ðMg; FeÞ4 Al2 Si7 O22 ðOHÞ2 hornblende þ 4CO2 fluid þ3H2 OY 2Ca2 ðFe; MgÞAl2 Si3 O12 ðOHÞepidote þ 2CaðFe; MgÞðCO3 Þ2 Fe−dolomite þ Fe2 O3 hematite þ 9SiO2 quartz þ 6AlðOHÞ2þ ðaq:Þ þ2O2 .....(6) 2KðMg; FeÞ3 ðAlSi3 O10 ÞðOH; FÞ2 biotite þ 4Hþ þ 2H2 S− fluid Y ðMg; FeÞ5 AlðAlSi3 O10 ÞðOHÞ8 chlorite þ 3SiO2 quartz þ FeS2 þ 2Kþ pyrite þ 4F− aq: …..(7) 3ðFe; MgÞ5 Al2 Si3 O10 ðOHÞ8 chlorite þ 15CO2 fluid þ2Kþ þ 8H2 OY 2KAl3 Si3 O10 ðOHÞ2 sericite þ 15CaðMg; FeÞðCO3 Þ2 ankerite þ 3SiO2 quartz þ 32Hþ ðaq:Þ .....(8) 3ðFe; Mg; TiÞ5 Al2 Si3 O10 ðOHÞ8 chlorite þ 15CaCO3 calcite þ15CO2 þ 2Kþ fluid þ14:5O2 Y 2KAl3 Si3 O10 ðOHÞ2 sericite þ 15CaðMg; FeÞðCO3 Þ2 ankerite þ 15TiO2 rutile þ 3SiO2 quartz þ 8H2 Ofluid þ4Hþ .....(9) KðFeMgÞ3 AlSi3 O10 ðOHÞ2 biotite þ ðMg; FeÞ5 Al2 Si3 O10 ðHOÞ8 chlorite þ 8CaCO3 calcite þ8CO2 Y 2KAl3 Si3 O10 ðOHÞ2 sericite þ 8CaðFe; MgÞðCO3 Þ2 ankerite þ 3SiO2 quartz þ4H2 O.....(10) 3KðFe; Mg; TiÞ3 AlSi3 O10 ðOHÞ2 biotite þ 18H2 S þ 10O2 fluid Y KAl3 Si3 O10 ðOHÞ2 sericite þ 9FeS2 pyrite þ 3SiO2 quartz þ 9TiO2 rutile þ9Mg2þ þ 2Kþ ðaq:Þ þ20H2 O.....(11) ðFe; Mg; TiÞ5 Al2 Si3 O10 ðOHÞ8 chlorite þ 10H2 S þ 8O2 fluid Y 5ðFe; MgÞS2 pyrite þ 3SiO2 quartz þ 5TiO2 rutile þ10H2 O þ 2AlðOHÞ2þ ðaq:Þ .....(12) 3KðFe; Mg; TiÞ3 AlSi3 O10 ðOHÞ2 biotite þ18HS− þ 5:5O2 fluid Y KAl3 Si3 O10 ðOHÞ2 sericite þ 9FeS2 pyrite þ 3SiO2 quartz þ 9TiO2 rutile þ11H2 O þ 9Mg2þ ðaq:Þ þKþ .....(13) ðFe; MgÞ5 Al2 Si3 O10 ðOHÞ8 chlorite þ 10HS− fluid þ3:5O2 Y 5FeS2 pyrite þ 3SiO2 quartz þ 5TiO2 rutile þ 9H2 Oðaq:Þ þ5Mg2þ .....(14) 3ðMg; FeÞ5 Al2 Si3 O10 ðOHÞ8 chlorite þ 2Kþ þ 2H2 Ofluid Y 2KAl3 Si3 O10 ðOHÞ2 sericite þ 3SiO2 quartz þ 15ðFe; MgÞþþ þ26H2 O þ 3Hþ ðaq:Þ .....(15) 3þ Mg5 Al2 Si3 O10 ðOHÞ8 chlorite þKþ þ Al fluid þ6Hþ f KAl3 Si3 O10 ðOH; FÞ2 sericite þ 5Mg2þ ðaq:Þ þ6H2 O.....(16) 2KAl3 Si3 O10 ðOHÞ2 sericite þ Fe3 O4 magnetite þ6H2 S þ 4SiO2 fluid þ2Naþ þ 2O2− Y 2NaAlSi3 O8 albite þ 2Al2 Si2 O5 ðOHÞ4 kaolinite þ 3FeS2 pyrite þ 2Kþ ðaq:Þ þ4H2 O.....(17) 2Fe3 O4 þ 3H2 SY3FeS2 þ 2H2 O þ 0:5O2 .....(18) Fe3 O4 magnetite þ2Hþ þ 6HS− fluid Y 3FeS2 pyrite þ 4H2 Ofluid .....(19) 4Fe3 O4 magnetite þ CO2 fluid Y 6Fe2 O3 hematite þ Cgraphite .....(20) 2KðFe; Mg; TiÞ3 AlSi3 O10 ðOHÞ2 biotite þ 4Hþ fluid þ2O2− Y ðFe; Mg; TiÞ5 Al2 Si3 O10 ðOHÞ8 chlorite þ 3SiO2 quartz þ TiO2 rutile þðFe; MgÞ2þ þ2Kþ .....(21) 3ðFe; Mg; TiÞ5 Al2 Si3 O10 ðOHÞ8 biotite þ 15CaCO3 calcite þ15CO2 þ 2Kþ fluid þ14:5O2 Y 2KAl3 Si3 O10 ðOHÞ2 sericite þ 15CaðMg; FeÞðCO3 Þ2 ankerite þ 15TiO2 rutile þ 3SiO2 quartz þ 8H2 Ofluid þ4Hþ .....(22)

Table 5 Simplified chemical reactions likely responsible for the hydrothermal alteration in the Um El Tuyor mine


461

inter- and transgranular fractures. Early entrapped inclusions are, however, preserved in the least-deformed domains at distance from planes of strain accommodation (e.g., Gebre-Mariam et al., 1991; Peters, 1993). 6.1. Inclusion types and modes of occurrence Deformation and recrystallization processes were intensive in places and limited in others, as indicated by moderate to strong undulose extinction, local mortar texture and deformation lamellae in the gold-bearing quartz veins. As a result, a strict application of the concept of primary and secondary inclusions (sensu Roedder, 1984) is not possible. Deformation might have caused post-entrapment modifications, such as necking down, partial or total leakage, selective leaching or diffusion of some components. To minimize the effect of post-entrapment modifications, inclusions suspected to have been affected by such modifications were clearly avoided. Instead, cores of the host quartz in quartz–carbonate veins show less deformed domains contain isolated and/or clusters of randomly and three-dimensionally distributed inclusions. Inclusions having these characteristics were selectively chosen for the microthermometric study. The investigated fluid inclusions are compositionally grouped into type I carbonic [CO2(± CH4 ± N2)-rich], type II aqueous [low salinity H2O–NaCl] and type III aqueous-carbonic [H2O–NaCl–CO2 (± CH4 ± N2)]. The aqueous-carbonic and aqueous inclusions are commonly observed in the quartz–carbonate veins, whereas the carbonic inclusions are by far the dominant or virtually the sole fluid inclusions in most parts of the laminated quartz and quartz–carbonate veins. Tiny elongate and oval inclusions along planar arrays, crosscutting populations of the carbonic and aqueous-carbonic inclusions, are considered late and likely have no relation to mineralization. The abundance of these inclusions is directly related to proximity to micro-fissured domains and sealed cracks. In this approach and due to their very small size (∼ 2 μm), these inclusions have not been further considered. The carbonic inclusions vary from negative crystals, through subrounded or oval to irregular, and in diameter from b3 to 12 μm. They are commonly two-phase (liquid–vapor) inclusions at room temperature, but occasionally mono-phase (liquid or vapor) inclusions are observed. The two-phase inclusions mostly have liquid N vapor. In some cases, these inclusions display thick dark boundaries, maybe suggesting the presence of a thin H2O film rimming their inner walls. The large carbonic inclusions occur as isolated singles or in clusters of randomly aligned inclusions (Fig. 17A). Also, they occur as planar intergranular arrays commonly confined to boundaries of the recrystallized vein quartz. Data for the isolated and clustered carbonic inclusions are considered more reliable and were therefore used for further processing in the present study. The aqueous (H2O–NaCl) inclusions can be easily identified by their transparency and low relief boundaries. Morphologically, they show variations from negative-crystal shapes to subhedral (especially the smaller inclusions), oval, spherical, tubular or irregular forms, and vary in size from 2 to 10 μm. At room temperature, two-phase inclusions (liquid and vapor) of this type show a fill range of 80 to 90% (Fig. 17B). Minimum trapping temperatures of some of these inclusions were difficult to assess because they invariably decrepitated prior to total homogenization. The aqueous-carbonic inclusions are rather common in the quartz– carbonate veins compared to the laminated veins. These inclusions exhibit a variety of polygonal forms and range in size from ∼5 to 17 μm. They are two- or three-phase inclusions (vapor CO2 ±liquid CO2 +aqueous liquid). In the three-phase inclusions, two immiscible liquids (liquid CO2 +aqueous liquid) and a small CO2 vapor bubble are commonly observed. The degree of fill (volumetric proportion of the aqueous phase relative to the total inclusion volume: DF=VH2O/Vt) in a single population is significantly variable (∼0.30 up to 0.85). In some cases, the carbonic inclusions tend to

462


occur as sets of diffuse planar arrays, restricted to the interiors of quartz grains or transecting their borders especially in domains rich in the aqueous-carbonic inclusions.

Based on the distribution and characteristics of the individual and grouped fluid inclusion planes at room temperature, there is no evidence for any pristine primary fluid inclusions in the laminated

Fig. 17. Simplified schematic sketches showing the distribution of the three types of fluid inclusions (early) and late transgranular inclusion trails in the auriferous quartz–carbonate (A, B) and laminated quartz (C) veins. Numbers indicate the total homogenization temperatures of the nearby inclusions in °C into liquid or vapor.


quartz veins. However, some of these inclusions show features analogous to pseudosecondary inclusions elsewhere (Roedder, 1984, Shepherd et al., 1985). The carbonic inclusions are by far the most predominant in these veins. They are mostly small (b6 μm), but some attain a diameter of ∼8 μm. They are oval to round in shape, but the larger inclusions are commonly irregular. Many of these fluid inclusions are aligned sub-parallel to the shear planes in the wallrocks. These inclusions are usually two-phase (liquid CO2 N vapor CO2) and mono-phase (liquid CO2 or vapor CO2) at room temperature. The mono-phase inclusions display two phases through cooling down to 10 °C or below. These mono-phase inclusions occur as planar arrays or sharp intergranular trails commonly confined to the recrystallized domains in quartz veins (Fig. 17C). Although they cannot be considered as early or primary, these inclusions may maintain ambient fluids entrapped when crystal growth took place during shearing. Therefore, relevant information about conditions of entrapment, and sequence of deformation events could be derived from them, especially if spatial and genetic relationships between gold mineralization and the shear zone are taken into account. 6.2. Microthermometry Data obtained from the microthermometric measurements and volume fraction estimates of the different phases are processed using the program packages CLATHRATES (Bakker, 1997) and FLUIDS (Bakker, 1999), in order to transform melting and homogenization temperatures and optical volume fraction estimates into bulk compositions and densities. Salinities were calculated after Bodnar (1993), from Tm ice, for aqueous inclusions and after Bakker (1997), from Tm clath, for aqueouscarbonic inclusions. Densities of H2O and CO2 are calculated from the equation of state by Zhang and Frantz (1987) and Holloway (1981), respectively. For the isochore calculations, the equations of state from Bowers and Helgeson (1983) and Bakker (1999) have been applied. Total freezing of the CO2-bearing inclusions typically required cooling to −110 °C, whereas the aqueous inclusions were commonly completely frozen at −47 °C. Homogenization of CO2 in the carbonic inclusions occurred mainly into the liquid phase. CO2 final melting

463

temperatures (Tm CO2) are slightly variable, from zero to ∼ 3 °C below the typical triple point of CO2 (−56.6 °C), assuming the presence of other dissolved gasses in addition to CO2 in the carbonic and aqueouscarbonic inclusions, maximum of 14 mol.% CH4 or 20 mol.% N2 (Burruss, 1981; Heyen et al., 1982; Thiéry et al., 1994). The total homogenization temperatures (Th total) of some large inclusions could not be obtained due to decrepitation prior to homogenization, generally above 340 °C. Th total does not appear to vary as a function of inclusion size, and therefore it is assumed that the local composition and density of the vein fluid were not differentially influenced by any physio-chemical effects dependent on the surface area (c.f. Binu-Lal et al., 2003). The physicochemical characteristics of the carbonic, aqueous and aqueous-carbonic inclusions are summarized in Table 6. Almost 45% of the studied inclusions are carbonic (CO2 ± CH4 ± N2), ranging from 1 to 10 μm in diameter. At room temperature, one third of these carbonic inclusions are mono-phase inclusions (LCO2), whereas the remainder are two-phase, liquid surrounding a bubble of variable size (20 to 90 vol.%). These inclusions occur as isolated, clustered and intragranular and intergranular trail-bound. Their variable modes of occurrence, in both quartz–carbonate and laminated quartz veins, suggest the relative timing of their formation. The isolated and clustered inclusions are considered earlier than the trailbound inclusions (e.g., Touret, 1981). In rare cases, isolated inclusions with oval or negative crystal shapes are associated with the aqueouscarbonic inclusions. The carbonic inclusions in the quartz–carbonate veins showed Tm CO2 varying from −56.6 to −58.4 °C and Th CO2 in the range of 3.7 to 18.3 °C. The carbonic inclusions in the laminated veins showed a range of Tm CO2 from −56.6 to −60.5 °C and Th CO2 from 13.6 to 26.8 °C. In one laminated quartz vein sample, several inclusions exhibited homogenization to the vapor phase, mostly between 20.4 and 28.9 °C. In both types of the quartz veins, the intergranular trailbound inclusions showed generally higher Th CO2 compared to the individual and clustered inclusions. Nearly 30% of the inclusion population consists of aqueous inclusions (H2O–NaCl, L + V) with diameters from 3 to 8 μm. They occur as solitary or preferentially along intragranular trails and are sporadically associated with aqueous-carbonic inclusions. The isolated inclusions

Table 6 Summary of the characteristics and microthermometry of the observed fluid inclusion types in the Um El Tuyor deposit Type 1: CO2 ± CH4 ± N2 inclusions

Type 2: H2O–NaCl inclusions

Type 3: H2O–NaCl–CO2 (± CH4 ± N2) inclusions

In the quartz–carbonate veins

Abundance: ∼44% of the bulk inclusions population Shape: oval, oblate Size: 2–9 μm Tm CO2 = −56.6 to −58.4 °C Th CO2 = 3.7 to 18.3 °C (into liquid) Molar volume = 48.62–55.59 cm3/mol Bulk density = 0.79–0.91 g/cm3

Abundance: ∼ 20% of the bulk inclusions population Shape: equant, elongate, irregular Size: 3–12 μm Vol.% H2O vapour = 5–20 Tm ice = −0.8 to − 4.7 °C Th total = 301.9 to 328.3 °C Salinity = 1.4–7.5 wt.% NaCl eq. Molar volume = 20.78–25.4 cm3/mol Bulk density = 0.88–0.75 g/cm3

In the laminated quartz veins

Abundance: ∼78% of the bulk inclusions population Shape: elongate oval, negative crystals, stretched irregular Size: 2–9 μm Tm CO2 = −56.6 to −60.5 °C Th CO2 = 13.6 to 26.8 °C (into liquid) Th CO2 = 20.4 to 28.9 °C (into vapor) Molar volume = 52.8–223.2 cm3/mol Bulk density = 0.20–0.83 g/cm3

Abundance: ∼ 24% of the bulk inclusions population Shape: sub-rounded, irregular Size: 2–10 μm Vol.% H2O vapour = 2–10 Tm ice = −0.1 to −2.9 °C Th total = 143.6 to 189.2 °C Salinity = 0.2–4.8 wt.% NaCl eq. Molar volume = 19.4–20.4 cm3/mol Bulk density = 0.93–0.91 g/cm3

Abundance: ∼36% of the bulk inclusions population Shape: sub-rounded, irregular Size: 4–15 μm DF = 0.20–0.80 Tm CO2 = − 56.6 to − 58.1 °C Th CO2 = 7.6–23.8 °C Tm Clath = 5.9–8.4 °C Th total = 321.4–336.1 °C Salinity = 3.2–7.6 wt.% NaCl eq. XCO2 = 0.05–0.20 mol% Molar volume = 20.84–41.24 cm3/mol Bulk density = 0.79–1.01 g/cm3 Abundance: ∼2% of the bulk inclusions population Shape: oval, wedged, irregular Size: 6–10 μm DF = 0.30–0.40 Tm CO2 = − 56.6 to − 57.3 °C Th CO2 = 18.4–26.6 °C Tm Clath = 7.7–9.3 °C Th total = 297.9–323.5 °C Salinity = 1.4–4.5 wt.% NaCl eq. XCO2 = 0.09–0.18 mol% Molar volume = 31.9–34.4 cm3/mol Bulk density = 0.81–0.86 g/cm3

More details about the equations of state used in fluid inclusions composition modelling are given in the text.

464


are oblate, oval or irregular in shape. Generally, the aqueous inclusions in the investigated samples have less variable phase proportions and lower final homogenization temperatures in comparison with the associated aqueous-carbonic inclusions. In the quartz–carbonate veins, the aqueous inclusions associated with the aqueous-carbonic inclusions in a single trail showed ice melting values (Tm ice) of −0.8 to −4.7 °C and total homogenization temperatures from 301.9 to 328.3 °C. The aqueous inclusions in the laminated quartz veins are more dilute compared to those in the quartz–carbonate veins; Tm ice and Th span the ranges −0.1 to −2.9 °C and 143 to 189 °C, respectively. Applying equations of state (Potter et al.,1978; Zhang and Frantz,1987), these Tm ice and Th ranges correspond to salinities between ∼1.2 and 5.1 wt.% NaCl equiv., and bulk fluid densities of 0.91 and 1.02 g/cm3. The aqueous-carbonic inclusions (H2O–CO2 ± CH4 ± N2) showed two, and occasionally three, phases at room temperature (liquid H2O, vapor CO2 ± liquid CO2). In these inclusions, the aqueous phase (of variable vol.%) surrounds a CO2 bubble, or forms a thin film along the walls of the cavity. Within a single quartz grain, the CO2 always homogenizes into the liquid phase, at a wide range of temperature (Th CO2 from 3.4 to 26.6 °C). No relationship between Th CO2 and the volumetric proportions of the carbonic phase was observed. Total homogenization (Th total) of the carbonic and aqueous phases occurred between 321 and 336 °C, commonly but not consistently into liquid. In rare cases, aqueous-carbonic, vapor-dominant inclusions, in which H2O is inferred only from the appearance of clathrate upon heating, total homogenization occurred into the vapor phase. Clathrate melting temperatures (Tm clath) range from 4.2° to 9.3 °C. The few observed aqueous-carbonic inclusions in the laminated quartz veins show less variable liquid to vapor ratios, smaller sizes, and commonly lower total homogenization temperatures compared to the quartz–carbonate veins. 6.3. Interpretation of the fluid inclusion data For the quartz–carbonate veins, the variably weak to intense deformation (e.g., undulose extinction) may account for the fluid inclusion diversity and variable degrees of filling observed in the aqueous-carbonic inclusions. Necking-down and decrepitation features are, however, uncommon. In addition, the coexistence of aqueous and carbonic inclusions along the same intragranular trails, and the variable liquid/vapor ratios of the aqueous-carbonic inclusions, point to possible fluid immiscibility (e.g., Ramboz et al., 1982). However, removal of H2O from aqueous-carbonic inclusions through grain boundary migration by crystal-plastic deformation could also account for the association of aqueous-carbonic and carbonic inclusions in the same inclusion trail (e.g., Hollister, 1990; Johnson and Hollister, 1995). The similar densities of CO2 in the carbonic and aqueous-carbonic inclusions along the same trail argue against postentrapment leakage (e.g., Yao and Robb, 2000). Plotting the total homogenization temperatures versus salinities of the aqueouscarbonic inclusions in the quartz–carbonate veins shows a positive correlation, suggesting a preferential fractionation of salt into the aqueous phase (e.g. Frantz et al., 1992; aqueous-rich inclusions with higher salinities than water-poor inclusions; Fig. 18A). A positive correlation between XCO2 and total homogenization temperatures of the aqueous-carbonic inclusions may, on the other hand, stand for grain boundary migration as water will be lost to the host quartz so that the residual fluid will be enriched in both salt and CO2 (Fig. 18B; e.g., Hollister, 1990; Audétat and Günther, 1999). A complex fluid evolution history including early partial unmixing and subsequent chemical re-equilibrium during crystal-plastic deformation is therefore, the appropriate interpretation of these mixed relationships. 6.4. Pressure–temperature conditions Textural relationships indicate that aqueous-carbonic inclusions in isolation or along intragranular trails are likely the early fluid inclusions

Fig. 18. (A) Th total versus salinities of the aqueous-carbonic fluid inclusions in the quartz–carbonate veins. (B) Th total versus XCO2 of the aqueous-carbonic inclusions in the quartz–carbonate veins.

in both quartz–carbonate and laminated quartz veins. In this case, the maximum homogenization temperature of these inclusions is considered as the entrapment temperature. For the quartz–carbonate veins, maximum Th total corresponds to a pressure range of 1.7 to 2.1 kbar (Fig. 19). If plotted in the same space, isochores for the aqueous-carbonic inclusions in the laminated quartz veins are less steep if compared with analogous inclusions from the quartz–carbonate veins. Using the isochores for these inclusions and applying the formation temperature of the late stage sulfides (325 °C), a pressure range of 1.6 to 1.3 kbar is defined (Fig. 19). 7. Discussion 7.1. Controls on mineralization and timing relative to the regional tectonothermal evolution According to Kerrich and Cassidy (1994), Goldfarb et al. (2001) and Groves et al. (2003), orogenic lode gold deposits are typically formed during late stages of the deformational-metamorphic–magmatic history of the evolving orogen, syn-kinematic with at least one main penetrative deformation stage of the host rocks. This conclusion is based on the widespread observations that the deposits structurally cut across fabrics formed during orogenesis, and most of the deposits are largely unaffected by early orogenic deformation. In the Um El Tuyor deposit, gold occurs as free-milling inclusions and refractory Au in arsenopyrite and pyrite. A similar bimodal distribution of gold has been described from many orogenic gold deposits (e.g., the Ashanti belt of Ghana; Oberthür et al., 1994). Accommodation within a shear zone, tabular vein geometries and host rock foliation anisotropy chiefly control the localization of the studied deposit. Within and close to the shear zone, S2 foliation is superimposed by crenulations and related pervasive cleavages (S3; N18–22°W/21–34°SW). Deflection of S2 into the shear zone parallel to S3 and development of oblique striation along the reverse fault planes assumes that development of the shear zone was contemporaneous with the NE–SW compressional stress regime prevailed during the D3 deformation phase (Fig. 2). Boudinage of foliation-parallel veins within the shear zone further


465

Fig. 19. Isochores for the highest and lowest density aqueous-carbonic inclusions from gold-bearing quartz veins from the Um El Tuyor deposit. The ‘a’ and ‘b’ letters represent the arsenopyrite temperature ranges of the early sulfides (common in quartz–carbonate veins) and late sulfides (common in the laminated quartz veins), respectively. The shaded box represents P–T of peak metamorphic conditions of the pelitic metasediments (Zoheir, 2004). Geothermal gradients of 25 and 50 °C/km (dotted lines) are given and the scale on the right side refers to depth in km at rock density of 2.6 g/cm3 for lithostatic overburden.

confirms a predominant shortening roughly normal to the S2 foliation planes during shearing. Other fabrics, such as intersection lineation and slickenfibers on cleavage planes, generally plunging steeply towards south, formed during cleavage-parallel shearing. Bends and dilation jogs formed by deflection of the host rock foliation into the shear zone have likely become favorable sites for quartz precipitation. Replacement of the metamorphic mineralogy by hydrothermal alteration assemblages within and adjacent to the shear zone indicates that the mineralizing event occurred after peak metamorphism. The peak metamorphic event under amphibolite facies conditions along the Allaqi-Heiani belt is dated at 630 to 550 Ma (Abd El Naby et al., 2000). In the central part of the Allaqi-Heiani belt, the peak metamorphic conditions were attained during D2 (SW–NE compressional regime). This deformation increment was terminated by a NW-trending transpression, leading to the formation of abundant kmscale sinistral NW–SE strike-slip faults. These faults are common in the central Eastern Desert of Egypt, known as the Wadi Hodein–Wadi Kharit shear system (Stern, 1994; Greiling et al., 1994) and in Arabia as the Najd Fault System (NFS; Stoesser and Camp, 1985). The NFS imparted deepseated regional fabrics that controlled the locus of later episodes of deformation (e.g., Agar, 1987). In the Wadi Allaqi region, gold mineralization is chiefly controlled by a system of discrete NNW-trending brittle– ductile shear zones (e.g., El Shimi,1996; El Kazzaz,1996), formed through a transcurrent shearing event during D3 post-dating the late Proterozoic NFS (∼640–530 Ma). Accordingly, D3 and related mineralization should have taken place after 640 to 530 Ma. According to Stern et al. (1989), the magmatic and metamorphic events in the region ceased at 510±40 Ma, late in the Neoproterozoic. Botros (2004) described sixteen gold deposits and occurrences in the Eastern Desert of Egypt in which gold mineralization is clearly spatially and genetically related to brittle–ductile shear zones, mostly developed during late deformational stages among the evolution history of the Eastern Desert. Loizenbauer et al. (2002) concluded that Late Pan-African transpressional tectonics controlled gold mineralization in the Fawakhir mine area. The isochron model for Pb isotope data of galena associated with gold in quartz veins from the Fawakhir mine indicates an age of 480 Ma (Delevaux et al., 1967). Helmy et al. (2004) suggested that gold mineralization at the El Sukari mine area, took place during a late stage extension and strike-slip tectonics (∼580 Ma), due to a combined strike-slip and vertical motion along flower type shear zones. For a similar deposit at the Betam mine area, 20 km west of

the Um El Tuyor mine, Zoheir (2008) suggested that gold was transported as bisulphide complexes by low salinity aqueous-carbonic fluids and precipitated because of variations in pH and fO2 through pressure fluctuation and CO2 effervescence as the ore fluids infiltrated the shear zone. 7.2. Hydrothermal alteration characteristics of the Egyptian vein-type gold deposits Wallrock alteration around the hydrothermal gold-bearing quartz veins of the Eastern Desert shows distinct mineralogical changes suggestive of metasomatic alteration as a result of percolation of external hydrothermal fluids (Osman and Dardir, 1989; Harraz, 1991; Harraz and El Dahhar, 1994). The metasomatic styles were strongly influenced by variations in the CO2 and alkali contents of the mineralizing fluid (Harraz et al., 1992; Harraz and El Dahhar, 1994). Precipitation and nature of the hydrothermal alteration minerals are dependent on the host rock composition. Botros (1993) reviewed the common types of hydrothermal alteration in zones bounding the mineralized veins in the Egyptian gold deposits and suggested that wallrock alteration shows regularity with respect to the host rock mineralogy. In the acid rocks, the most prevalent types are sericitization, argillic alteration, silicification, whereas chloritization, carbonatization, sericitization, pyritization and propylitization prevail in the intermediate and basic rocks. Several authors concluded that these zones are of economic interest as they contain reasonably high gold values (e.g., Osman and Dardir, 1989; Azzaz, 1987; Takla et al., 1995; Botros, 1991). For the Atud gold deposit, hosted by a gabbro–diorite complex, Harraz (1999) described an alteration halo around the mineralized quartz veins composed of three zones, namely a) chlorite and calcite; b) albite and ankerite; and c) albite, muscovite and kaolinite. He suggested that these alteration zones have been formed through addition of CO2, H2O, K and lesser Na. In the Barramyia gold mine, where numerous auriferous quartz veins cutting through highly altered tuffogene metasediments, ultramafites, and listwaenite, Osman and Dardir (1989) described a complex mixture of alteration processes, including sericitization, silcification, sulfidation, carbonatization, ankeritization and chloritization. They reported the highest gold grade in the altered rocks (1 to 5 ppm Au) in the carbonaceous wallrocks adjacent to the mineralized veins. At Hutit mine, hydrothermal

466


alteration halo enveloping the Au-bearing quartz veins grades up to 35 ppm Au, made by mainly of silicified, carbonatized serpentinite (Takla et al., 1995). A narrow alteration halo bounds the main mineralized quartz vein at the Um Tenedba mine, replacing the original granitic/gabbroic material. Sericitization and sulfidation are the main alteration types, with 6 ppm Au at maximum (Takla et al., 1995). In the Dungash gold mine area, Khalil et al. (2003) and Helba et al. (2001) described four alteration zones, mainly mixtures of chloritization, sericitization, carbonatization and silicification overprinting the original volcani-sedimentary host rock mineralogy, in which disseminated gold occur. They suggested that foliation, fracturing and shearing of the wall rocks are the major factors controlling the alterations and mineralization. As for the Um El Tuyor deposit, the proximal alteration zone grades up to 2.13 ppm Au, in zones tickly seamed with hydrothermal quartz, carbonate, sericite and sulfide minerals. This may serve as evidence for the significant rule of wallrock alteration in depositing gold, likely through pH or redox state reactions, in several examples of the vein-type gold deposits of Egypt. 7.3. Fluid evolution and gold deposition Considering that the Au-bearing quartz veins are hosted by deformed rocks and abundant evidence for ductile deformation, it is doubtful that the aqueous-carbonic inclusions have not undergone post-entrapment modifications and/or re-equilibration. However, a

Fig. 20. Stereographic plots of structural data of mineralized shear fractures at the mine area. (A) Equal area projection of fault planes and related slip striations. (B) Strain directions, graphically determined according to the P (pressure) and T (tension) dihedral method of Angelier and Gogeul (1979), and the paleostress tensor is comparatively calculated using the inversion algorithm developed by Etchecopar et al. (1981).

Fig. 21. Log fO2–T diagram showing the redox state of the ore fluids with regard to with the quartz–fayalite–magnetite (Q–F–M) and hematite–magnetite (H–M) buffers. T–fO2 ranges of the studied deposit along with main gold deposit types (diagram adapted from Huston and Large, 1989).

vertical trend in the Th versus salinity bivariant plot, typical evidence for necking down and leakage under heating (Shepherd et al., 1985) has not been recognized. Unequivocal petrographic evidences, such as variable deformation of the investigated quartz veins and the small size of the fluid inclusions brought much difficulty in associating fluid inclusions with the fluids responsible for gold mineralization. However, no later event with higher pressure–temperature conditions, which could have overprinted the studied inclusions, was identified. This observation allowed to suggest that data derived from these fluid inclusion assemblages are directly related to the gold mineralizing event, and represents the physicochemical conditions of the gold deposition. In the quartz–carbonate veins, Fe–As ± Cu sulfide association formed at temperatures of 353 to 372 °C, but gold started to precipitate when immiscibility conditions prevailed at 336 °C and 1.7 to 2.1 kbar (Fig. 20). On the other hand, temperatures estimated for the later assemblage (312 to 325 °C) correspond to 1.3 to 1.6 kbar. The pressure values estimated for quartz–carbonate and laminated veins imply formation at 6 to 8 km under lithostatic conditions. This wide pressure range may reflect fluid pressure fluctuations within the shear zone through fissure re-opening and refilling (e.g., Mullis et al., 1994; Robert and Kelly, 1987). It is assumed that the early quartz–carbonate reefs were reactivated through a late strike-slip to oblique-slip environment by fault–valve behavior at lower greenschist facies conditions, led to the formation of the laminated quartz veins. Since solubility of the metal species in aqueous hydrothermal solutions is significantly affected by fO2, the redox state of the ore fluid is considered critical to the transport and deposition of metals. Considering a low salinity aqueous-carbonic ore, evaluation of the C–O–H system is appropriate to address fO2 by using the program COHFLUID (Huizenga, 1995). Calculations are based on the presence of graphite in the deposit and assuming ideal mixing of the compounds. The fO2 of the mineralizing fluids calculated for a temperature range of 325 to 372 °C and pressures of 1.3 to 2.1 kbar varies from 10− 27 to 10− 32 bar. These values, under the given P–T conditions, plot slightly above the quartz–fayalite–magnetite (QFM) buffer and below the Ni–NiO buffer (Fig. 21). Accordingly, a relatively reduced nature of H2O–CO2 gold-bearing fluids is demonstrated (e.g., Vigneresse, 2007). Such reducing conditions, together with the textural evidence for deposition of the iron sulfides concomitant with gold (at least


partially), suggests that HS−-bearing or neutral HS complexes Au(HS)−2 or HAu(HS)02 were the predominant agents in the transport of gold in the aqueous-carbonic fluid (Hayashi and Ohmoto, 1991; Seward, 1991). Similar fluid evolution histories are described for other gold deposits in the Eastern Desert of Egypt. Loizenbauer and Neumayr (1996) described pseudosecondary liquid-rich, vapor-rich and mixed low salinity CO2–H2O fluid inclusions in the auriferous quartz veins from the El Sid gold deposit (25°59′15″N, 33°36′40″E), allowing them to propose formation at pressures of 1.8 to 2.3 kbar. Harraz (2002) suggested that conditions of gold deposition in Atud gold deposit (25°01′19″N, 34°24′6″E) are similar to many mesothermal vein systems and likely occurred because of a sharp decrease in sulfur activity through unmixing due to an abrupt drop in fluid pressure from 272 to 160 MPa. He further concluded that low salinity (2.8 to 8.2 wt.% NaCl equiv.) H2O–CO2-rich inclusions and their homogenization temperatures assume a metamorphic source, likely originated through devolatization of metasediments and serpentinites at depth and which migrated upwards through the structural conduit offered by dilation structures. The ore fluids obtained their metal contents by leaching the country rocks through fluid–rock interaction. Similar assemblages of aqueous, aqueous-carbonic and carbonic fluid inclusions have been described from gold-bearing quartz veins at the El Sukari gold deposit (24°56′49″N, 34°42′53″E), south Eastern Desert (Helmy et al., 2004). The authors assumed that early circulating CO2-rich fluids caused alteration along the shear zones, and that these fluids were mixed with meteoric water during exhumation. They suggested a long-term cyclic crack-seal mechanism through pressure fluctuation (between 210 and 1,890 bar) along with weakening by alteration. The authors further suggested that similar gold deposits within the Eastern Desert of Egypt, located at the contacts between granitoids and mafic–ultramafic rocks and along shear zones, result from analogous formation processes. 8. Conclusions Criteria including the disseminated mineralization style, alteration mineral assemblages, spatial and genetic relationships with a brittle– ductile shear zone, and occurrence within a metamorphic terrane are characteristic of orogenic gold deposits (Groves et al., 1998). In view of the deformation history of the mine area and surroundings, formation of

467

the shear zone is linked to a late transpression deformation that postdates peak metamorphism. Furthermore, the location where the shear zone traverses the tectono-stratigraphic contact between ophiolitic and pelitic metasedimentary rocks suggests that contrasting rock competency and rheology may have imparted dilational sites for the circulating ore fluids. The abundance of dacite sills in the mine area and close to the mineralization zone suggests a possible contribution of magmatic source at depth at least in the early mineralization stages, characteristic of the orogenic-type gold deposits. Replacement of the metamorphic assemblage by hydrothermal mineral phases points to a post-peak metamorphic hydrothermal event. The zoned alteration pattern around the auriferous vein system (Fig. 22) is attributed to temporal and lateral evolution in fluid composition through interaction with wallrock. The formation of sericite is the most characteristic alteration process, from the incipient to the advanced stages. The generally high K2O, SO2, As, Ba, Sr, Sn and Au contents in zones adjacent to the quartz veins, compared to unaltered host rocks, are considered indicative of their addition during alteration. SiO2 contents are either invariable or elevated, suggesting that quartz veining is derived from direct silica precipitation. CO2 addition is inferred from the higher contents of carbonate minerals in wallrock. Stages of increasing hydrothermal alteration are identified as initial, intermediate and advanced. Considering that both reactants and products are observed in the initial stage, it is suggested that hydrolysis reactions operated in presence of a near acid fluid, whose pH was buffered by the wallrock mineralogy. The transitional stage involved hydrolysis reactions along with intense carbonatization, sulfidation and redox reactions. The original mineralogy is almost completely replaced by hydrated minerals (muscovite and chlorite) and carbonate. The advanced stage was likely a phase of intense sericitization, consumed K+, released H+, and raised pH of the solution. Cationexchange reactions were limited to the time when favored aNa+/aK+ and temperature conditions promoted deposition of albite. Sulfidation continued operating throughout this stage, and unbuffered conditions were locally attained under high fluid/rock ratios. This collectively indicates that the ore fluids progressively evolved towards lower temperatures and sulfur fugacity with time. Decrease of sulfur fugacity, through sulfide precipitation and/or H2S loss, may have favored gold deposition through destabilization of gold disulfide complexes such as Au(HS)2 (Seward, 1984; Drummond and Ohmoto, 1985; Bowers, 1991).

Fig. 22. Sketch cross section of the Um El Tuyor gold deposit, proposed on basis of the various field and petrographic observations (not to scale).

468


An interplay of mixing (and local unmixing) of heterogamous aqueous and carbonic fluids and changes in the physicochemical conditions via a fault–valve system triggered brittle–ductile transition within the shear zone (e.g., Walsh et al., 1988). Sulfidation of the host rocks also contributed to destabilization of the gold–di-sulfide complexes, thereby leading to gold precipitation in the adjacent wallrock (e.g., Phillips and Groves, 1983; Groves and Phillips, 1987; Groves et al., 1998). Gold deposition due to sulfidation doe not conflict with intermittent fluid unmixing, as the latter does not appear to have extended into the wallrock (e.g., Guha et al., 1991). Pressure estimates based on isochores of the aqueous-carbonic inclusions indicate depths of 6 to 8 km for quartz veining and gold deposition in the area, which is compatible with crustal conditions of greenschist metamorphism and brittle–ductile transition. Formation of the quartz–carbonate veins was early in the mineralization process, whereas the laminated quartz veins were formed through a fault–valve system and cyclic opening and annealing of reactivated reefs under greenschist facies conditions (e.g., Kolb et al., 2005). Although the origin of the fluids remains speculative, certain trends nevertheless appear relatively clear. If extrapolated to higher temperatures–pressures, isochores for the highest bulk density aqueous-carbonic inclusions will pass through the box representing the peak metamorphic conditions on the P–T diagram (Fig. 19). This observation suggests that the low salinity aqueous-carbonic ore fluid may have been generated during the metamorphic dehydration/ decarbonation of the ophiolite–pelitic sedimentary rocks; contributions from a magmatic source at depth cannot, however, be formally excluded. Acknowledgements This study builds on the Ph.D. thesis of the author with additional fluid inclusion work at the University of Tübingen in 2006. Thorough supervision and sincere guidance in both contextual and scientific contents of the thesis were offered by Profs. D. Klemm and R. Marschik, Univ. of Munich. I have been especially fortunate to have them as advisors and mentors. My colleagues at the Benha University Department of Geology (Profs. Wetait, Mehanna) and Prof. Harraz are thanked for help in the first field trip. For technical support, a standing ovation goes, in no particular order, to Prof. Walter (Karlsruhe), Dr. Reinhard Kaindl (Austria), Dr. Albert Gilg and Andreas Murr (Munich) for their help with the microprobe and microthermometric measurements and interpretations. I thank the German Academic Exchange Service (DAAD) for immense support during my stay as visiting fellow in Munich. Two anonymous journal reviewers and Chief Editor Dr. N. Cook and Dr. A. Gilg are thanked for their time spent revising and improving this manuscript. References Abd El Naby, H.W., Frisch, W., Hegner, E., 2000. Evolution of the Pan-African Wadi Haimur metamorphic sole, Eastern Desert, Egypt. Journal of Metamorphic Petrology 18, 639–651. Abdelsalam, M.G., Stern, R.J., 1996. Sutures and shear zones in the Arabian–Nubian Shield. Journal of African Earth Sciences 23, 289–310. Agar, R.A., 1987. The Najd fault system revisited: a two-way strike-slip orogen in the Saudi Arabian shield. Journal of Structural Geology 9, 41–48. Ahmed, A.M., Said, M.M., El Baghdady, M.M., Abdel Wahab, G., 2001. Mineral potential of the eastern part of Wadi Allaqi, South Eastern Desert, Egypt. Annals of the Geological Survey of Egypt XXIV, 451–462. Almond, D.C., Ahmed, F., Shaddad, M.Z., 1984. Setting of gold mineralization in the northern Red Sea hills of Sudan. Economic Geology 79, 389–392. Amin, M.S., 1955. Geological features of some mineral deposits in Egypt. Bulletin of the Institute of Desert, Egypt 5, 208–239. Angelier, J., Goguel, J., 1979. Sur une méthode simple de détermination des contraintes pour une population de failles. Comptes Rendus Hebdomadaires des Séances de l'Academie des Sciences 288, 307–310. Audetat, A., Gunther, D., 1999. Mobility and H2O loss from fluid inclusions in natural quartz crystals. Contributions to Mineralogy and Petrology 137, 1–14. Audétat, A., Günther, D., Heinrich, C.A., 1998. Formation of a magmatic-hydrothermal ore deposit: insights with LA-ICP-MS analyses of fluid inclusions. Science 279, 2091–2094.

Azzaz, A.S., 1987. Distribution of gold and associated elements in the wallrocks of the Sukkari gold deposit, Central Eastern Desert, Egypt. Arabic Journal for Science and Engineering 12 (3), 243–250. Bakker, R.J., 1997. Clathrates: computer programs to calculate fluid inclusion V–X properties using clathrate melting temperatures. Computer Geosciences 23, 1–18. Bakker, R.J., 1999. Optimal interpretation of microthermometrical data from fluid inclusions: thermodynamic modelling and computer programming. Habilitation thesis, Ruprecht-Karls-University, Heidelberg, Germany, 50 pp. Barton Jr., P.B., 1969. Thermochemical study of the system Fe–As–S. Geochimica et Cosmochimica Acta, 33, 841–857. Barton Jr., P.B., Skinner, B.J., 1979. Sulfide mineral stabilities, In: Barnes, H.L. (Ed.), Geochemistry of Hydrothermal Ore Deposits, 2nd edition. Wiley Interscience, New York, pp. 278–403. Bell, T.H., Cuff, C.,1989. Dissolution, solution transfer, dissolution versus fluid flow and volume loss during deformation/metamorphism. Journal of Metamorphic Geology 7, 425–447. Bhattacharya, A., Mohanty, L., Maji, A., Sen, S.K., Raith, M., 1992. Non-ideal mixing in the phlogopite–annite binary: constraints from experimental data on Mg–Fe partitioning and a reformulation of the biotite–garnet thermometer. Contributions to Mineralogy and Petrology 111, 87–93. Binu-Lal, S.S., Wilbert Kehelpannalab, K.V., Satish-Kumara, M., Wadaa, H., 2003. Multistage graphite precipitation through protracted fluid flow in sheared meta-granitoid, Digana, Sri Lanka: evidence from stable isotopes. Chemical Geology 197, 253–270. Bodnar, R.J., 1993. Revised equation and table for determining the freezing point depression of H2O–NaCl solutions. Geochimica et Cosmochimica Acta 57, 683–684. Botros, N.S., 1991. Geological and geochemical studies on some gold occurrences in the north Eastern Desert, Egypt. Ph.D. thesis, Zagazig Univ., Zagazig, Egypt, 146 pp. Botros, N.S., 1993. New prospects for gold mineralization in Egypt. Annals of the Geological Survey of Egypt 19, 47–56. Botros, N.S., 1995. Genesis of gold mineralization in the North Eastern Desert, Egypt. Annals of the Geological Survey of Egypt 20, 381–409. Botros, N.S., 2002. Metallogeny of gold in relation to the evolution of the Nubian Shield in Egypt. Ore Geology Reviews 19, 137–164. Botros, N.S., 2004. A new classification of the gold deposits of Egypt. Ore Geology Reviews 25, 1–37. Bowers, T.S., 1991. The deposition of gold and other metals: pressure induced fluid immiscibility and associated stable isotope signatures. Geochimica et Cosmochimica Acta 55, 2427–2434. Bowers, T.S., Helgeson, H.C., 1983. Calculation of thermodynamic and geochemical consequences of non ideal mixing in the system H2O–CO2–NaCl on phase relations in geologic systems: equation of state for H2O–CO2–NaCl fluids at high pressures and temperatures. Geochimica et Cosmochimica Acta 47, 1247–1275. Burruss, R.C., 1981. Analysis of phase equilibria in C–O–H–S fluid inclusions. In: Hollister, L.S., Crawford, M.L. (Eds.), Short Course in Fluid Inclusions — Applications to Petrology. Mineralogical Association of Canada Short Course Handbook, vol. 6, pp. 39–74. Cathelineau, M., 1988. Cation site occupancy in chlorites and illites as a function of temperature. Clay Minerals 23, 471–485. Davis, B.K., Hippertt, J.F.M., 1998. Relationships between gold concentration and structure in quartz veins from the Hodgkinson Province, northeastern Australia. Mineralium Deposita 33, 391–405. Delevaux, M.H., Doe, B.R., Brown, G.F., 1967. Preliminary lead isotope investigations of brine from the Red Sea, Galena from the Kingdom of Saudi Arabia, and galena from United Arab Republic (Egypt). Earth and Planetary Science Letters 3, 139–144. Drummond, S.E., Ohmoto, H., 1985. Chemical evolution and mineral deposition in boiling hydrothermal systems. Economic Geology 80, 126–147. Drury, M.R., Humphreys, F.J., White, S.H., 1985. Large strain deformation studies using polycrystalline magnesium as a rock analogue. Part II: dynamic recrystallization mechanisms at high temperatures. Physics of the Earth and Planetary Interiors 40, 208–222. EGSMA, 1912. Um Tiour Mine. Inspection report. Lease No. 13. Internal report, Serial Nr. 7050. Department of Publication and Documentation, Cairo, 6 pp. EGSMA, 1913. Um El Tiour Mine. Inspection report. Lease No. 12. Internal report, Serial Nr. 7066. Department of Publication and Documentation, Cairo, 7 pp. EI Gaby, S., List, F.K., Tehrani, R., 1988. Geology, evolution and metallogenesis of the PanAfrican Belt in Egypt. In: EI-Gaby, S., Greiling, R.O. (Eds.), The Pan-African Belt of Northeast Africa and Adjacent Areas. Friedrich Vieweg und Sohn, Braunschweig Wiesbaden, pp. 17–68. El Kazzaz, Y.A., 1996. Shear zone-hosted gold mineralization in South Eastern Desert, Egypt. Proceedings of the Geological Survey of Egypt Centennial Conference, pp. 185–204. El Ramly, M.F., Ivaanov, S.S., Kochin, G.C., 1970. The occurrence of gold in the Eastern Desert of Egypt. Studies on some mineral deposits of Egypt. Part I, Section A, Metallic minerals. Geological Survey of Egypt 21, 1–22. El Shazly, E.M., 1957. Classification of Egyptian mineral deposits. Egyptian Journal of Geology 1, 1–20. El Shimi, K.A., 1996. Geology, structure and exploration of gold mineralization in Wadi Allaqi area (SW, Eastern Desert, Egypt). Ph.D. Thesis, Ain Shams University, Egypt, 326 pp. Etchecopar, A., Vasseur, G., Daignieres, M., 1981. An inverse problem in microtectonics for the determination of stress tensors from fault striation analysis. Journal of Structural Geology 3, 51–65. Fleet, M.E., Mumin, A.H., 1997. Gold-bearing arsenian pyrite and marcasite and arsenopyrite from Carlin Trend gold deposits and laboratory synthesis. American Mineralogist 82, 182–193. Frantz, J.D., Popp, R.K., Hoering, T.C., 1992. The compositional limits of fluid immiscibility in the system H2O–NaCl–CO2 as determined with the use of synthetic fluid inclusions in conjunction with mass spectrometry. Chemical Geology 98, 237–255.

B.A. Zoheir / Ore Geology Reviews 34 (2008) 445–470 Garson, M.S., Shalaby, I., 1976. Precambrian–lower Palaeozoic plate tectonics and metallogenesis in the Red Sea region. Geologic Association Canada, Special Paper 14, 573–596. Gebre-Mariam, M., Groves, D.I., Ho, S.E., McNaughton, N.J., Veamcombe, J.R., 1991. The Archaean lode-gold deposit at Racetrack, near Kalgoorlie, Western Australia: a transitional mesothermal–epithennal hydrothermal system. In: Pagel, M., Leroy, J.L. (Eds.), Source, Transport and Deposition of Metals. Balkema, Rotterdam, pp. 661–664. Giggenbach, W.F., 1987. Redox processes governing the chemistry of fumarolic gas discharges from White Island, New Zealand. Applied Geochemistry 2, 143–161. Giggenbach, W.F., 1988. Geothermal solute equilibria: derivation of Na–K–Mg–Ca geoindicators. Geochimica et Cosmochimica Acta 52, 2749–2765. Goldfarb, R.J., Groves, D.I., Gardoll, S., 2001. Orogenic gold and geologic time: a global synthesis. Ore Geology Reviews 18, 1–75. Grant, J.A., 1986. The isocon diagram — a simple solution to Gresen's equations for metasomatic alteration. Economic Geology 81, 1976–1982. Greiling, R.O., Abdeen, M.M., Dardir, A.A., El Akhal, H., El Ramly, M.F., Kamal El Din, G.M., Osman, A.F., Rashwan, A.A., Rice, A.H.N., Sadek, M.F., 1994. A structural synthesis of the Proterozoic Arabian–Nubian Shield in Egypt. Geologische Rundschau 83, 484–501. Gresens, R.L., 1967. Composition–volume relations of metasomatism. Chemical Geology 2, 47–65. Groves, D.I., Phillips, G.N., 1987. The genesis and tectonic control on Archaean gold deposits of the western Australian shield — a metamorphic replacement model. Ore Geology Reviews 2, 287–322. Groves, D.I., Foster, R.P., 1991. Archean lode-gold deposits. In: Foster, R.P. (Ed.), Gold Metallogeny and Exploration. Blackie, Glasgow, pp. 63–103. Groves, D.I., Goldfarb, R.J., Gebre-Mariam, M., Hagemann, S.G., Robert, F., 1998. Orogenic gold deposits: a proposed classification in the context of their crustal distribution and relationship to other gold deposit types. Ore Geology Reviews 13, 7–27. Groves, D.I., Goldfarb, R.J., Robert, F., Hart, C.J., 2003. Gold deposits in metamorphic belts; overview of current understanding, outstanding problems, future research, and exploration significance. Economic Geology 98, 1–29. Guha, J., Lu, H.Z., Dubé, B., 1991. Fluid characteristics of vein and altered wallrock in Archaean mesothermal gold deposits. Economic Geology 86, 667–684. Harraz, H.Z., 1991. Lithochemical prospecting and genesis of gold deposit in El Sukari gold mine, Eastern Desert, Egypt. PhD thesis, Faculty Science, Tanta University, Egypt, 494 pp. Harraz, H.Z., 1999. Wall rock alteration, Atud gold mine, Eastern Desert, Egypt: processes and P–T–XCO2 conditions of metasomatism. Journal of African Earth Sciences 28, 527–551. Harraz, H.Z., 2000. A genetic model for a mesothermal Au deposit: evidence from fluid inclusions and stable isotopic studies at El-Sid Gold Mine, Eastern Desert. Journal of African Earth Sciences 30, 267–282. Harraz, H.Z., 2002. Fluid inclusions in the mesothermal gold deposit at Atud mine, Eastern Desert, Egypt. Journal of African Earth Sciences 35, 347–363. Harraz, H.Z., Ashmawy, M.H., 1994. Structural and lithochemical constraints on the localization of gold deposits at the El Sid-Fawakhir gold mine area, Eastern Desert, Egypt. Egyptian Journal of Geology 38, 629–648. Harraz, H.Z., EI Dahhar, M.A., 1993. Nature and composition of gold-forming fluids at Umm Rus area, Eastern Desert, Egypt: evidence from fluid inclusions in vein materials. Journal of African Earth Sciences 16, 341–353. Harraz, H.Z., El Dahhar, M.A., 1994. Fluid–wall rock interaction and its implication in gold mineralization at the Umm Rus gold mine area, Eastern Desert, Egypt. Egyptian Journal of Geology 38, 713–747. Harraz, H.Z., Hassanen, M.A., El Dahhar, M.A., 1992. Fluid inclusions and stable isotopic studies at El-Sid gold mine, Eastern Desert, Egypt. Egyptian Journal of Geology 36 (1–2), 243–332. Hassaan, M.M., El Mezayen, A.M., 1995. Genesis of gold mineralization in the Eastern Desert, Egypt. Al-Azhar Bulletin of Science 6, 921–939. Hollister, L.S., 1990. Enrichment of CO2 in fluid inclusions in quartz by removal of H2O during crystal-plastic deformation. Journal of Structural Geology 12, 895–901. Hussein, A.A., 1990. Mineral deposits. In: Said, R. (Ed.), The Geology of Egypt. Balkema, Rotterdam, pp. 511–566. Hayashi, K.-I., Ohmoto, H., 1991. Solubility of gold in NaCl- and H2S-bearing aqueous solutions at 250–350 °C. Geochimica et Cosmochimica Acta 55, 2111–2126. Helba, H.A., Khalil, K.I., Abdou, N.M., 2001. Alteration patterns related to hydrothermal gold mineralization in meta-andesites at Dungash area, Eastern Desert, Egypt. Resource Geology, 51, 19–30. Helmy, H.M., Kaindl, R., Fritz, H., Loizenbauer, J., 2004. The Sukari Gold Mine, Eastern Desert—Egypt: structural setting, mineralogy and fluid inclusion study. Mineralium Deposita 39, 495–511. Hey, M.H., 1954. A new review of the chlorites. Mineralogical Magazine 30, 277–292. Heyen, G., Ramboz, C., Dubessy, J., 1982. Simulation des équilibres des phases dans le systè me CO2–CH4 en dessous de 50 °C et de 100 bar. Application aux inclusions fluides. C R Acadamie des Sciences Paris 294, 203–206. Hickey, K.A., Bell, T.H., 1996. Syn-deformational grain growth: matrix coarsening during foliation development and regional metamorphism rather than by static annealing. European Journal of Mineralogy 8, 1351–1373. Hilmy, M.E., Osman, A., 1989. Remobilization of gold from a chalcopyrite–pyrite mineralization Hamash gold mine, Southeastern Desert, Egypt. Mineralium Deposita 24, 244–249. Holloway, J.R., 1981. Composition and volumes of supercritical fluids in the Earth crust. In: Hollister, L.S., Crawford, M.L. (Eds.), Fluid Inclusions: Applications to Petrology. Mineralogical Association of Canada, Short Course Handbook, vol. 6, pp. 13–38. Huizenga, J.M., 1995. Fluid evolution in shear zones from the late-Archean Harare– Shamva–Bindura greenstone belt NE Zimbabwe. Thermodynamic calculations of the C–O–H system applied to fluid inclusions. Ph.D. Thesis, Vrije Universiteit, Amsterdam, 146 pp.

469

Hume, W.F., 1937. Geology of Egypt, V.II, part III. Geological Survey of Egypt. Huston, D.L., Large, R.R., 1989. A chemical model for the concentration of gold in volcanogenic massive sulfide deposits. Ore Geology Reviews 4, 171–200. Johan, Z., Marcoux, E., Bonnemaison, M., 1989. Arsénopyrite aurifère: mode de substitution de Au dans le structure de FeAsS. CR Academy of Science Paris 308 (II), 185–191. Johnson, E.L., Hollister, L.S., 1995. Syn-deformational fluid trapping in quartz: determining the pressure–temperature conditions of deformation from fluid inclusions and the formation of pure CO2 fluid inclusions during grain-boundary migration. Journal of Metamorphic Geology 13, 239–249. Kerrich, R., Cassidy, K.F., 1994. Temporal relationships of lode gold mineralization to accretion, magmatism, metamorphism and deformation — Archean to present: a review. Ore Geology Reviews 9, 263–310. Khalil, K.I., Helba, H.A., Mucke, A., 2003. Genesis of the gold mineralization at the Dungash gold mine area, Eastern Desert, Egypt: a mineralogical–microchemical study. Journal of African Earth Sciences 37, 111–122. Kishida, A., Kerrich, R., 1987. Hydrothermal alteration zoning and gold concentration at the Kerr Addison Archean lode gold deposit, Kirkland Lake, Ontario. Economic Geology 82, 649–690. Klemm, D.D., Klemm, R., Murr, A., 2001. Gold of the Pharaohs — 6000 years of gold mining in Egypt and Nubia. Journal of African Earth Sciences 33, 643–659. Kochin, G.G., Bassyunni, F.A., 1968. The mineral resources of United and Arab Republic. Report on Generalization of Geological Data on Mineral Resources of the UAR Carried Out Under Contract 1247 (1966–1968). Part 1: Metallic Minerals: Internal Report 18/68, Egyptian Geological Survey and Mining Authorities (EGSMA). Kolb, J., Kisters, A.F.M., Hoerne, S., Meyer, F.M., 2000. The origin of fluids and nature of fluid–rock interaction in mid crustal auriferous mylonites of the Renco mine, Southern Zimbabwe. Mineralium Deposita 35, 109–125. Kolb, J., Rogers, A., Meyer, F.M., 2005. Relative timing of deformation and two-stage gold mineralization at the Hutti Mine, Dharwar Craton, India. Mineralium Deposita 40, 156–174. Kretschmar, U., Scott, S.D., 1976. Phase relations involving arsenopyrite in the system Fe–As–S and their application. Canadian Mineralogist 14, 364–386. Kröner, A., Greiling, R., Reischmann, T., Hussein, I.M., Stern, R.J., Durr, S., Kruger, J., Zimmer, M., 1987. Pan-African crustal evolution in the Nubian segment of Northeast Africa. In: Kroner, A. (Ed.), Proterozoic Lithospheric Evolution. AGU Geodynamics Series, vol. 17, pp. 237–257. Kusky, T.M., Ramadan, T.M., 2002. Structural controls on Neoproterozoic mineralization in the South Eastern Desert, Egypt: an integrated field, Landsat TM, and SIR-C/X SAR approach. Journal of African Earth Sciences 35, 107–121. Leitch, C.H.B., Lentz, D.R., 1994. The Gresens approach to mass balance constraints of alteration systems: methods, pitfalls, examples. In: Lentz, D.R. (Ed.), Alteration and Alteration Processes Associated With Ore-Forming Systems. Geological Association of Canada, Short Course Notes, pp. 161–192. Loizenbauer, J., Neumayr, P., 1996. Structural controls on the formation of the Fawakhir gold mine, El Sid-Eastern Desert, Egypt: tectonic and fluid inclusion evidence. Proceedings of the Geological Survey of Egypt Centennial Conference, pp. 477–488. Loizenbauer, J., Neumayr, P., Fritz, H., 2002. Genesis of the Fawakhir gold deposit in the Eastern Desert of Egypt: the result of late Pan-African transpressional tectonics. 19th Colloquium of African Geology, El Jadida, Morocco, pp. 120–121. Abstract volume. Meyer, C., Hemley, J.J., 1967. Wallrock alteration. In: Barnes, H.L. (Ed.), Geochemistry of Hydrothermal Ore Deposits. Holt, Rinehart and Winston, New York, pp. 166–235. Mueller, R.F., Saxena, S.K., 1977. Chemical Petrology, with Application to the Terrestrial Planets and Meteorites. Springer-Verlag, New York. 396 pp. Mullis, J., Dubessy, J., Poty, B., O'Neil, J., 1994. Fluid regimes during late stages of a continental collision: physical, chemical and stable isotope measurements of fluid inclusions in fissure quartz from a geotraverse through the Central Alps, Switzerland. Geochimica et Cosmochimica Acta 58, 2239–2267. Neubauer, W.H., 1962. Geologie der Goldlagerstaitte von El Sid in Oberaegypten mit einem Beitrag zur Geologie der zentralen Arabischen Wuste. Geologische Jahrbuch 80, 117–160. Oberthür, T., Vetter, U., Schmidt Mumm, A., Weiser, T., Amanor, J.A., Gyapong, W.A., Kumi, R., Blenkinsop, T.G., 1994. The Ashanti gold mine at Obuasi in Ghana: mineralogical, geochemical, stable isotope and fluid inclusion studies on the metallogenesis of the deposit. Geologische Jahrbuch 100, 31–129. Osman, A.M., Dardir, A.A., 1989. On the mineralogy and geochemistry of some goldbearing quartz veins in the central Eastern Desert of Egypt and their altered wallrocks. Annals of the Geological Survey of Egypt 21, 17–25. Peters, S.G., 1993. Formation of oreshoots in mesothermal gold–quartz vein deposits: examples from Queensland, Australia. Ore Geology Reviews 8, 277–301. Phillips, G.N., Groves, D.I., 1983. The nature of Archaean gold bearing fluids as deduced from gold deposits from Western Australia. Journal of the Geological Society of Australia 30, 25–40. Pohl, W., 1988. Precambrian metallogeny of NE-Africa. In: El-Gaby, S., Greiling, R.O. (Eds.), The Pan-African Belt of NE Africa and Adjacent areas. Friedrich Vieweg und Sohn, Brauschweig - Wiesbaden, pp. 319–341. Poirier, J.P., Guillope, M., 1979. Deformation induced recrystallization of minerals. Bulletin of Mineralogy 102, 67–74. Potter II, R.W., Clynne, M.A., Brown, D.L., 1978. Freezing point depression of aqueous sodium chloride solutions. Economic Geology 73, 284–285. Powell, R., Holland, T.J.B., 1988. An internally consistent dataset with uncertainties and correlations; 3, Applications to geobarometry, worked examples and a computer program. Journal of Metamorphic Geology 6, 173–204. Ramadan, T.M., Sadek, M.F., Abu El Leil, I., Salem, S.M., 2005. Um El Touyur El Fouqani gold mineralization, South Eastern desert, Egypt. Annals of the Geological Survey of Egypt XXVIII, 263–281.

470


Ramboz, C., Pichavant, M., Weisbrod, A., 1982. Fluid immiscibility in natural processes: use and misuse of fluid inclusion data. II. Interpretation of fluid inclusion data in terms of immiscibility. Chemical Geology 37, 29–48. Robert, F., Kelly, W.C., 1987. Ore-forming fluids in Archean goldbearing quartz veins at the Sigma Mine, Abitibi greenstone belt, Quebec, Canada. Economic Geology 82, 1464–1482. Roedder, E., 1984. Fluid inclusions. Reviews in Mineralogy 12 644 pp. Sabet, A.H., Bondonosov, V.P., 1984. The gold ore formations in the Eastern Desert of Egypt. Annals of the Geological Survey of Egypt 14, 35–42. Sabet, A.H., Tscgoev, V.B., Bordonosov, V.P., Babourin, L.M., Zalata, A.A., Francis, M.H., 1976. On gold mineralization in the Eastern Desert of Egypt. Annals of the Geological Survey of Egypt 6, 201–212. Sadek, M.F., Youssef, M.M., Abdel Gawad, G.M., Salem, S.M., Hashash, M., 1995. Geology of Gabal Adraq Area, South Eastern Desert of Egypt. Internal Report 21/95. EGSMA, Cairo. 58 pp. Sadek, M.F., Ramadan, T.M., Abu, El Leil, Salem, S.M., 2005. The basement rocks of Gabal Al Adraq area, South Eastern Desert, Egypt: geology, geochemistry and tectonic setting. Annals of the Geological Survey of Egypt XXVIII, 37–57. Scott, S.D., 1983. Chemical behaviour of sphalerite and arsenopyrite in hydrothermal and metamorphic environments. Mineralogical Magazine 47, 427–435. Selverstone, J., Morteani, G., Staude, J.M., 1991. Fluid channeling during ductile shearing: transformation of granodiorite into aluminous schist in the Tauern Window, Eastern Alps. Journal of Metamorphic Geology 9, 419–431. Seward, T.M., 1984. The transport and deposition of gold in hydrothermal systems. In: Foster, R.P. (Ed.), Gold'82. The Geology, Geochemistry and Genesis of Gold Deposits. Balkema, Rotterdam, pp. 165–182. Seward, T.M., 1991. The hydrothermal geochemistry of gold. In: Foster, R.P. (Ed.), Gold Metallogeny and Exploration. Blackie and Son, Glasgow, pp. 37–62. Sharp, Z.D., Essene, E.J., Kelly, W.C., 1985. A re-examination of the arsenopyrite geothermometer: pressure considerations and applications to natural assemblages. Canadian Mineralogist 23, 517–534. Shelton, K.L., So, C.S., Chang, J.S., 1988. Gold rich mesothermal vein deposits of the Republic of Korea: geochemical studies of the Jungwon gold area. Economic Geology 83, 1221–1237. Shepherd, T.J., Rankin, A.H., Alderton, D.H., 1985. A Practical Guide to Fluid Inclusion Studies. Blackie and Sons, Glasgow. 239 pp. Stern, R.J., 1994. Arc assembly and continental collision in the Neoproterozoic East African Orogen: implications for the consolidation of Gondwanaland. Annual Review of Earth and Planetary Sciences 22, 319–351. Stern, R.J., Kröner, A., Manton, W.I., Reischmann, T., Mansour, M., Hussein, I.M., 1989. Geochronology of the late Precambrian Hamisana shear zone, Red Sea Hills, Sudan and Egypt. Journal of the Geological Society (London) 146, 1017–1030. Stipp, M., Stünitz, H., Heilbronner, R., Schmid, S.M., 2002. Dynamic recrystallization of quartz: correlation between natural and experimental conditions. In: De Meer, S.,

Drury, M.R., De Bresser, J.H.P., Pennock, G.M. (Eds.), Deformation Mechanisms, Rheology and Tectonics: Current Status and Future Perspectives. Geological Society Special Publication, vol. 200, pp. 171–190. Stoesser, D.B., Camp, V.E., 1985. Pan-African microplate accretion of the Arabian shield. Geological Society of America, Bulletin 96, 817–826. Takla, M.A., El Dougdoug, A.A., Rasmay, A.H., Gad, A.A., El Tabbal, H.K., 1990. Origin of Um Eleiga gold mineralization, south Eastern Desert, Egypt. Egyptian Mineralogist 2, 3–20. Takla, M.A., El Dougdoug, A.A., Gad, M.A., Rasmay, A.H., El Tabbal, H.K., 1995. Goldbearing quartz veins in mafic and ultramafic rocks, Hutite and Um Tenedba, south Eastern Desert, Egypt. Annals of the Geological Survey of Egypt 20, 411–432. Thiéry, R., Van den Kerkhof, A.M., Dubessy, J., 1994. VX properties of CH4–CO2 and CO2– N2 fluid inclusions: modelling for T b 31 °C and P b 400 bars. European Journal of Mineralogy 6, 753–771. Touret, J.L.R., 1981. Fluid inclusions in high grade metamorphic rocks. In: Hollister, L.S., Crawford, M.L. (Eds.), Short Course in Fluid Inclusions: Applications to Petrology. Mineralogical Association of Canada Short Course Notes, vol. 6, pp. 182–207. Vigneresse, J.L., 2007. The role of discontinuous magma inputs in felsic magma and ore generation. Ore Geology Reviews 30, 181–216. Vollbrecht, A., Stipp, M., Olesen, N.O., 1999. Crystallographic orientation of microcracks in quartz and inferred deformation processes: a study on genesis from the German Continental Deep Drilling Project (KTB). Tectonophysics 303, 279–297. Walsh, J.F., Kesler, S.E., Duff, D., Cloke, P.L., 1988. Fluid inclusion geochemistry of high grade, vein hosted gold ore at the Parmour Mine, Porcupine Camp, Ontario. Economic Geology 83, 1347–1367. Yao, Y., Robb, L.J., 2000. Gold mineralization in Paleoproterozoic granitoids at Obuasi, Ashanti region, Ghana: ore geology, geochemistry and fluid characteristics. South African Journal of Geology 103, 255–278. Zhang, Y., Frantz, J.D., 1987. Determination of the homogenization temperatures and densities of supercritical fluids in the system NaCl–KCl–CaCl2–H2O using synthetic fluid inclusions. Chemical Geology 64, 335–350. Zoheir, B.A., 2004. Gold mineralization in the Um El Tuyor area, South Eastern Desert, Egypt: geologic context, characteristics and genesis. Ph.D. Thesis, LudwigMaximilians-Universität, München, Germany, 159 pp. Zoheir, B.A., 2008. Structural controls, temperature–pressure conditions and fluid evolution of orogenic gold mineralisation in Egypt: a case study from the Betam gold mine, south Eastern Desert. Mineralium Deposita. Mineralium Deposita 43, 79–95. Zoheir, B.A., Klemm, D.D., 2007. The tectono-metamorphic evolution of the central part of the Neoproterozoic Allaqi-Heiani suture, south Eastern Desert of Egypt. Gondwana Research 12, 289–304.