Abstract. The Jabali deposit (3.8 Mt at 16% Zn, 2% Pb and 132 g/t Ag) is hosted by dolomitized platform carbon- ates of Kimmeridgian age at the southwestern ...
Mineral. Deposita 29, 44 -56 (1994)
~neralium
Deposita ©
Springer- Verlag 1994
Jabali, a Zn-Pb-(Ag) carbonate-hosted deposit associated with Late Jurassic rifting in Yemen I. AI Ganad', P. Lagny-, J.L. Lescuyer/, C Ramboz-, and J.e. Touray' Mineral Exploration Board , Ministry of Mineral Resources, P.O. Box 297, Sana'a, Republic of Yemen BRGM, Departement Exploration, BP 6009, F-45060 Orleans Cedex 02, France 3 CRSCM-CNRS et GdR 69, La rue de la Ferollerie, F-45071 Orleans Cedex 02, France 4 GdR 69 et URA 1366 du CNRS , Ecole superieure de l'Energie et des Materiaux, Universite d'Orleans, BP 6759, F-45067 Orleans Cedex 02, France 1
2
Received: 3 April 1992 / Accepted: 23 November 1993
Abstract. The Jabali deposit (3.8 Mt at 16% Zn, 2% Pb and 132 g/t Ag) is hosted by dolomitized platform carbonates of Kimmeridgian age at the southwestern edge of the oil-producing Wadi al Jawf rift basin in northern Yemen. Paleogeographical reconstructions demonstrate that tensional synsedimentary tectonic acti vity from the Late Jurassic to Early Cretaceous was responsible for the thick accumulation of argillaceous and evaporitic sediments in the subsident rift basin , the unstable margin of which was the site of rapid facies changes , local disconformities and periods of emergence, as well as of dolomitization along the WNW- and NNW-striking boundary fault system. In the Jabali area, the upper part of the Jurassic sequence underwent two stages of dolomitization before emergence and deep karstic erosion. Solution cavities and depressions in the eroded surface were filled by dolomite sand and black pyritic mudstone prior to a last marine transgression of limited extent. Subsequent ore deposition and associated late dolomitization sealed the network of solution cavities, impregnating the dolomite sands and the host dolomites. Sphalerite I and wurtzite, followed by silverbearing zoned sphalerite II associated with galena, crystallized from a cyclic influx of low-temperature (75-100 °C) saline solutions. Lead isotope geochemistry indicates that the lead, zinc and silver probably originated from an Early Proterozoic basement. The dissolved metals were likely derived from the basal aquifer (detrital material of basement origin) of the evaporite-bearing sequence filling the Wadi al Jawf trough. Migrating metalliferous brines from the basin to the uplifted Jabali area, where ore deposition was favoured by a reducing environment, were probably channelled by the boundary fault system during the last stages of synsedimentary tectonic activity.
The Jabali Pb-Zn-Ag deposit lies 70 km northeast of Sana'a, capital of the Republic of Yemen, at lat. 15°36'44" N, long . 44°46'10" W (Fig. 1). The mineralization is exposed on a small plateau (150 m wide by 200 m long) at an altitude of 1800 m in a mountainous area to
the southwest of the Mesozoic Wadi al Jawf basin, the site of large oil discoverie s in the early 1980s. Jabali was the largest silver mine of the Islamic world between the 1st and 3rd centuries A.H. (7th-9th centuries A.D.), with the old mine workings covering an area of about 10 ha where cavities were filled with soft, locall y silver-rich carbonate ore (Christmann et al. 1983). The deposit was rediscovered in 1980 during an exploration programme carried out by BRGM on behalf of the Yemeni Ministry of Oil and Mineral Resources, and subsequent drilling has revealed 3.8 Mt of oxide ore grading 16% Zn, 2% Pb and 132 g/t Ag (Michel et al. 1988). New exploration data since the first compilation concerning the discovery (Christmann et al. 1983) has justified a more advanced scientific study of the deposit to ascertain the depositional conditions of the primary sulphide mineralization that was locally preserved from oxidation. The main results of this study (AI Ganad 1991) are presented here.
Geological setting The Jabali deposit lies within the topm ost beds of the Jurassic Amran Formation (defined by Lamare 1930), which was deposited during a regional marine transgression that reached its peak in the Late Oxfordian-Early Kimmeridgi an when the SomaliaEthiopian- Yemen Sea and the Arabian Sea became joined (Le Nind re et al. 1990). The nearby subsident Late Jurassic Wadi al Ja wf basin, which developed as a half graben (Maycox 1989) initiated by synsedimentary tensional tectonism along a pre-existing northw eststriking transcurrent fault system in the Late Proter ozoic basement (Christmann et al. 1983), was filled with a thick ( > 6000 m) sedime ntary succession of dominantly argillaceous and evaporit ic facies overlain by detrital format ions of the Late Jurassic-Early Cretaceous Transition Beds and the Cretace ous Tawilah Formation (Lamare 1930). The Amran Formation, cropping out along the southwestern edge of the Wadi al Jawfbasin, contains numerous Pb-Zn-Ba occurrences associated with local dolom itization controlled by WNW- to NNWstriking faults. These occurrences develop over a strike length of some 60 km, but are mainly concentrated in the Jabali -M ajnah area where, in addit ion to the Jabali depo sit located at the top of the carbonate sequence, non-economic Pb-Ba occurrences are hosted by the lower units of the Amran Formation.
45 The J urassic succession of the area (defined by Christmann et al. 1983) has recently been studied in detail and pa rtly revised by Al Ga na d (199 1) who shows the 300-m-thick sequence at Jab ali to co mprise eight unit s which, from the base upwa rds, a re (Fig. 2):
c=J
Recent .lIuvl.1 and eoh,n I.dlmenn
M I Jor oilfield
~ Ml jo r IIUII
~ Qu.t.rn.ry volc.mc fl.lds
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Other pre -Qu.ternary form.tlons l l •• Jur •• s,c ••It dl.pl'S
Bound.ry of the li te JurliliC W.d, .1 Jaw' b.ltn
ca rbon.tes
Jut' I SIC
Precambr ian basement
Fig. 1. Geo logica l sketch map of the southwestern Arab ian peninsula showing the locat ion of the Jabali dep osit
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late O xfordl.n - El r ly Kim meridgian
E
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Fig. 2. Synt hetic litho stra tigraph ic sectio n of the Ja bali area showing Jurassic Unit s I to 8 (see text) and local facies variations (no horizontal scale)
Unit 1 ( 10 m ) . Sand stone and co nglomerate tran sgressive over the Lat e Pro terozoic basemen t (schist and quart zite intru ded by gran ite and pegmatite). Unit 2 ( 25 m ) . Gypsiferou s mud stone over lain by dolom itized calca renite int erb edd ed with marl and nodul ar limestone. Unit 3 (50 m ). Micritic and biomi crit ic limeston e of Ca llovia n age co ntaining nodular co ncretions and laye rs of grey chert. The presence of Characea at the top of th is un it indicat es a limited period of emerge nce. Unit 4 ( 15 m ) . Micritic limeston e and yellowis h, finely bedd ed dolom ite of lago onal/l acust rine facies . Unit 5 ( 40 m ) . Bryozoan calcarenite of Late Ox fordian - Ea rly Kim merid gian age, overlain by co ra l-bea ring oo litic and onco litic limesto ne. T his unit is partially dolomitized, with a local disconformit y at its top . Unit 6 (80 m ) . Greenish gypsifero us mudstone grading to interbedded lam inated micriti c ammo nite-bea ring limestone and marl with (a) lenses of calcareo us sa nds to ne co ntai ning plant remain s and (b) bedd ed d olom itized biocalcarenite. Th e ammon ite fauna indica te a Kimmeridgian age. Unit 7 (80 m ). Massive bioclastic and biomicritic limesto ne, locally oo litic and conta ining coral bioh erms. Th e unit is almos t co mplet ely dolom itized and its top has been a ffected by significa nt kar stic erosio n. Unit 8 (0- 30 m ) . Black mud stone with dolomitic intercalat ion s, gra ding northwards to micrit ic limeston e co nta ining Lat e Kimmerid gian - Tith on ian ammo nites. The black mud stone is a lagoonal facies containing sma ll gypsum cryst als and lignite debri s. A slight disco nformity (local hardground o r kar stic ero sion ) is found at the boundar y bet ween Unit 8 and the ove rlying Lat e Jurassic-Early Cre taceo us T ransition Beds composed of var iega ted gypsum-bea ring mud ston e with lenses of arkosic san dsto ne. Th e Tran sition Beds in the Jab ali area do not exceed 20 m in thickness, whereas in the Wad i al Jawf basin they have been intersected over a thickness of more tha n 2000 m by petroleum dr illing (Maycox 1989). Th e fluvio-d eltaic sa ndsto ne of the Tawilah Format ion is not exposed, and was prob abl y not deposited in the Jab ali a rea. Num erou s tra chytic dykes, sills and laccolith s were empl aced at th e so uthwestern edge of the Wad i al Ja wf basin du ring Early Miocene times. Dated a t 22 Ma by the K/A r method (AI G anad 1991), th ese intrus ions were contemporaneous with the reactivation of the regional northwest-stri king boundar y fault system th at ga ve rise to the present basin morphology (marked notably by collapses and mass slides of the limestone cliffs), a nd so with th e rifting that preceded the opening of the Red Sea. The lat era l facies varia tions of the Jurassic succession, both to the so uthwest an d to the northeast towards the Wad i al Jawf basin, a re indi cated in Fig. 2. Th ey show that th e sedimenta tion was relatively hom ogeneous up to th e end of O xford ian times (U nit 5) with no significa nt cha nges in either facies or th ickness. From the Kimmerid gian on, a period of instab ility was related to the init iat ion of the Wad i al Jawfbasin, the de velopment of which peaked during the Tithonian , a per iod of widespread salt deposition, before being filled by Early Cretaceou s detrital facies (M aycox 1989). Th is period of insta bility is reflected in th e uplifted Jabali-Majnah are a by notable lateral facies and th ickn ess varia tions (revealed du ring I : 100 000scale regional mapping, par ticularl y in Unit 7), a nd by severa l local eme rgence surfaces (Chris tma nn et al. 1983). Th e upli fted Jab ali- Majn ah area (F ig. 3) is cha racterized by a dolomitization cont rolled by the bounda ry faults of the basin. It is located along a bar rier zo ne con taining ooli tic facies an d coral patches th at is bounded to the northeast by slope facies composed of slumped marl an d black micrit ic limeston e, includ ing partially dolom itized shelf-limesto ne olisto liths. To the so uthwest, a subsident inner platform (Unit 7 and part of Un it 6) consists of th ick massive micritic limestone with cor als, lam ellibr an chs, gastropod s and
46
Fig. 5. Irregular eroded top of dolomitized Unit 7 overlain by transgressive facies of Unit 8, locally well mineralized (weathered black pyritic mudstone and dolomite sand) and filling in small depressions and cavities exploited by the ancients (see also Fig. 9)
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late Jur,ssic to Earlv Cretaceous baSin faCies (rmcr mc , argillaceous and Iv.poril ic sediments) Ba,,..r faCies (oo ht lc and reef hmes tone of Umts 5 and 7 1 ~ M iddle to late JurassIc carbonate platform
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~"·:'::. a1 Faull conuolled do lom it izat io n
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Synse dimentary f.ull
~ OIlSlollths ~ Slope breCCia
from the underlying dolomitized Unit 7, which filled a karstic depression (interpreted as a doline ; AI Ganad 1991)and plugged the underlying solution cavities. Karst erosion continued at the stillemergent edges of the depres sion over which Microcodium-bearing breccia (mudstone fragments in a dolomitic matrix) was deposited. A temporary return to marine conditions, transgressive from the northeast over the edges of the lagoonal paleodepression, is marked by sedimentation of micritic limestone and calcarenite.
I ...,.. " I Slumps
CD
M ineral occurrence'
~ Thickness (In metres) of the Jurlsllc sequence
Fig. 3. Location of the Pb-Zn occurrences and dolomitized zones along the southwestern margin of the Late Jurassic Wadi al Jawf basin oncolites. During the uplift and emergence of this area following the dolomitization of Unit 7, sedimentation continued both on the inner carbonate platform and along the fault-controlled slope margin of the basin. An accurate reconstruction of the paleomorphology of the Jabali site (Fig. 4), which was possible due to detailed I: 1000-scale mapping and drill-hole data, shows that Unit 8 transgressed over the eroded top of Unit 7 (Fig. 5). This transgression began with deposition of black pyritic mudstone, containing dolomite sand reworked
The Jabali deposit The Jabali deposit is hosted mainly by the dolomitized limestone of Unit 7 which forms an elongate NW -SE halo of about 2 km x 1 km (Fig. 6). The irregularly mineralized volume within the dolomitized host rock , is a 500 x 500 m block, the thickness of which ranges from 20 to 100 m. The western half of this block , which has been explored by drilling , was partially preserved from erosion due to the protection of a massive limestone slab (Jabal Barik) that broke away from and slid down the southwestern cliff during the Tertiary and covered Unit 8 and the Transition Beds (Christmann et al. 1983). The ore is almost completely oxidized , except locally where primary sulphide mineralization has been preserved by an impervious cover of Unit 8 black mudstone.
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Dolomi te sand (base of Unl ' 8)
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M lcro co dlum ·b ear lOg br eccia
~ M ICrit iC limes tone and calcareni te With local bioherms
Loc al emergence surface of Unit 8
Fig. 4. Paleogeographic reconstruction of the uplifted Jabali zone at the end of Unit 8 deposition
47
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Fig. 6. Geological map of the Jabali deposit
Table 1 gives the major- and trace-element analysis result s for composite samples of the oxidized (71 kg) and sulphide (17.6 kg) ores taken from drill holes. The Zn , Pb and Ag grades are of the same order as the average grades of the ore deposit, but among the trace elements the presence of significant quantities of cadmium and germanium should be emphasized. The oxidized mineralization consists mainly of smithsonite, hydrozincite, cerussite, and iron and manganese oxides and hydroxides; anglesite is rare. The trace elements (Cd, Ag, Cu) released during oxidation of the sulphides are found in the form of greenockite, argentite, native silver and covellite. Zinc silicates (hemimorphite and willemite) occur locally at the contact zone between the oxidized ore and black pyritic mudstone that has been transformed into silico-ferruginous gossan; they have been observed only in outcrop and correspond to a late oxidation phase that occurred, after the downslip of Jabal Barik , under tropical weathering conditions of probable Pliocene-Quaternary age (as interpreted by Lagny and Feraud 1983). Pervasive oxidation of the deposit therefore predated the downslip of Jabal Barik and the subsequent intrusion of Miocene trachyte sills and dykes , which also cut the rock slide. Thus, the oxidation may have begun during the Cretaceous, developed mainly during the Paleogene (a period characterized by extensive lateritic weathering), and continued up to the present. Most of the mineralization is concentrated in the dolomitized and karstified Unit 7 (Fig. 7) where it (a) impregnates the dolomite sand filling the solution cavities, (b) is disseminated in the host dolostone, and (c) lines a fracture network (Fig. 8a, b, d). The orebodies are found mainly at intersections between bedding joints and the solution-fracture network, with maximum
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Impregnetlon 01 kerst IllIing M,n e rehze d frecture net wor k Scettered ore Preserved sulph ide ore
Fig. 7. Schematic section of the Jabali deposit (see Fig. 6 for location)
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50 m
48 Table I. Chemical composition of oxidized and sulphide ores from Jabali Major elements (%)
Si0 2
AI20 3
FC203
Ti0 2
MnO
K 20
Na20
MgO
CaO
S
Oxidized ore Sulphide ore
0.60 0.80
0.19 0.26
4.15 3.80
0.02 0.02
0.61 0.38
0.03 0.05
0.05 0.04
10.00 10.80
17.70 21.00
0.34 8.90
Trace elements (ppm)
Cd
Ag
Ge
As
Cu
Sb
Ni
Mo
Bi
Sn
Oxidized ore Sulphide ore
659 585
170 130
67 107
620 410
60 60
10
39 34
28 27
280 230
61 60
Zn
Pb
18.50 15.00
2.50 2.60
Major elements, Ag, Cu, and Ge were analyzed by atomic absorption; other minor elements by ICP
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Fig. 9. Detail of the mineralized Unit 8 facies filling solution cavities at the top of Unit 7
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Host r oc k dalasI one (U m t 7 )
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Dedol omlllzed host rock M onerallzed (Zn and Pb c ar b o n at e s) do tormte sands
Smlthson lte after sphalerote Ir o n OXIdes and hydroxIdes ( f o r m e r lion su lptude s} La l e e areue
Fig. 8a-d. Typical aspects of the oxidized mineralization observed on drill cores. a, b, d Mineralized dolom ite sand filling small cavities or cementing solution breccia (Unit 7). c Barren do lomite sand (base of Unit 8) containing reworked fragments of dolostone (Unit 7) and threads of slumped black mudstone. Note the preservation of primary ore textures : pseudomorphs of millimetre-sized crysta ls of smithso nite after spha lerite (a); early reworking of ore fragments (d)
concentrations on the edges of the paleodepression at the top of Unit 7. Paleokarstic cavities severa l metres in length generally show a typical zoning with a Zn-Pb impregnated footwall (former sulphid es tran sformed into carbonates), massive smit hsonite and iron oxides at the base of the cavity (0.1-2.0 m thick), impregnated dolomite sand and breccia with altered black mudstone fragmen ts at the top of the cavity, and vertical fissures extending upwards in the hanging wall. Th e yellowish to reddish internal sediments show typical well-preserved sedimentary features such as slumps and ear ly reworking of mineralization, and the original macroscopic and microscopic sulphide fabric is generally recognizable in the oxidized ore (Fig. 8a, d). Ore grades in the karst cavities commonly exceed 20% Zn over thicknesses of several metres (e.g. drill ho le J4B with 10 mat 28.2% Zn, 1.6% Pb and 160 glt Ag) and locally exceed 40% Zn over a thickness of 1 m. Lenticular layers of do lomite sand, intercalated at the base of the black mudstone overlying the Unit 7 erosion surface, are also locally well mineralized (Fig. 8c, 9). The black mudstone facies itself displays a high geochemical zinc content (as much as 2500 ppm) related to
49 Table 2. Sequence of events associated with emplacem ent of the Jabali ore deposit Unit 7 1st period of dolomitizati on
Emergence and erosion of Unit 7, transgression of Unit 8
Units 7 and 8 2nd period of dolomitizat ion , mineralization
01 mosaic dolomite
Dissolut ion, disaggregation and part ial dedolomitization of Unit 7 do lomitic rocks
Fr acturing and deposition of 0 3 baroque ferroan dolomite 04 baroque dolomit e
Sphalerite I (and former wurtzite), pyrite, marca site
02 ferroan dolomite enclosing fine disseminated pyrite
Sedimentation of dolomite sand in karst cavities and at the base of tran sgressive Unit 8
Recrystallization of dolomite sand Dolomit ization of micrite and calcarenit e of Unit 8
Zoned sphalerite II, galena, pyrite, marcasite
Early Kimmeridgian to Tith onian
a dissemination of small sphalerite crystals, and the dolomitized micrite at the top of Unit 8 locally contains a well-developed fracture mineralization. Subordinate sulphide mineralization has also been encountered by deep drilling at several levels in the Jurassic below Unit 7 (fractures and impregnations) and even in the basement (dolomite and sphalerite veinlets). It has been possible, through laboratory studies, to specify the relationships between the different phases of dolomitization and sulphide deposition, to determine the trace -element content of the sphalerite, and to obtain new inform ation (fluid-inclusion, organic-matter and lead- isotope data) constraining the nature and origin of the mineralizing fluids. Dolom itization
Dolomitization was studied in thin sections stained with alizarine red-S and potassium ferricyanide, in order to determine the relationships between the different kinds of carbonate (calcite, ferroan dolomite and non-ferroan dolomite). A detailed study of the different generations of carbonate performed by cathodoluminescence confirmed the thin-section observations and showed a growth zonation in the latest stage of dolomitization. Detailed results are given in AI Ganad (1991). Two major periods of dolomitization can be distin guished, each comprising two successive stages (Table 2): the first period (01 and 02 dolomite), interpreted as early diagenetic, occurred before the transgression of Unit 8 and the emplacement of the mineralization. The second period was contemporaneous with the sulphide depo sition (0 3 dolomite) or occurred slightl y later (0 4 dolomite ). Microprobe analyses of seven samples representative of the 01 - 04 generations show the composition of the dol omite to be near stoichiometric or to contain a slight excess of calcium (the molar CaC0 3 fraction reaching a maximum of 53%). These results also confirm the thinsection ob servations where stai ning makes it easy to distingui sh the non-ferroan 01 and 04 dolomite from the ferroan 02 and 03 dolomite: Oland 04 contain less than 1% FeC0 3 and less than 0.8% MnC0 3 , whereas 02 and 03 contain between 1 and 5% FeC0 3 and up to 1.5% MnC0 3 ·
+
Early Cretaceous (?)
First period of dolomitization (DJ and D2 dolomite). The first period of dolomitization was widespread, affecting large volumes of the carbonate series. It developed below the erosion surface of Unit 7 and formed the host rock of the major part of the lead- zinc mineralization. The 01 dolomite is fairly discrete compared with the 0 2 dolomite. It corresponds to a first transformation of the limestone into a grey, inclusion-rich, mosaic dolosparite preserving remn ants of bioclastic and intraclastic allochems. The crystal size is very variable (a few to several hundred microns), depending on the initial texture of the replaced carbonate facies. The 0 2 dolomite developed either (1) at the expense of 01 , retaining its mosaic texture but with a slight incre ase in average grain size (of the order of several tens of microns) and almost totally replacing the original limestone, or (2) by direct transform ation of the limestone into an assemblage of euhedral to subhedral zoned dolomite crystals. Unlike 01 , 02 is a ferroan dolomite (i.e. it is stained by ferricyanide) that also contains fine disseminated pyrite, and is easily recognizable in the field by its tan colour due to recent oxidation. Weathering and disaggregation of the dolomitized limestone during emergence and erosion of Unit 7 provided the dolostone fragments and dolomite sand that were resedimented at the base of Unit 8 and in the underlying solution cavities (Fig. 9). Iron-oxide and hydroxide pigmentation marking the growth zones of the crystals show that the reworked crystals and fragment s of 02 dolomite were oxidized, probably during a phase of dedolomitization preceding redolomitization during diagenesis of Unit 8. The first period of dolomitization, which clearl y predate s local erosion and karst ification of the barrier facies of Unit 7, can be interpreted according to a mixing-zone model (Bad iozamani 1973; Folk and Land 1975). Despite relevant objections, expressed in particular by Machel and Mountjo y (1986) and Hardie (1987), this seems to be the onl y existing model consistent with the geological data: moreover, Hardie (1987) admits th at 'mixing-zone waters remain potential dolomitizing waters' (op. cit., p. 171). Second period of dolomitiza tion (D3 and D4 dolomite). The second period of dolomitization, which presumably began during earl y diagenesis of Unit 8, comprises two new successive phases of dolomite crystallization associated with the fracture and microcavities filling sulphides.
50
The 03 and 04 dolomite crystals, with a size ranging from several hundred microns to a few millimetres, are baroque (Radke and Mathis 1980) and cut the earlier dolomitized facies and the dolomite sand. The 03 dolomite is ferroan and is rimmed by non-ferroan 04 dolomite crystals. The recrystallization of the dolomite sand (both in karst cavities and at the base of Unit 8) and the dolomitization of the micritic calcarenite at the top of Unit 8 occurred during this second period . Post dolomitization. The last generation of carbonate prior to sulphide oxidation and late dedolomitization consists of white calcisparite. Late dedolomitization is associated with oxidation of the ore deposit, and also developed in the dolomitized units along faults, such as the Jabal Salab fault (see Figs. 6 and 7), during Tertiary reactivation. Circulation of fresh water with a high Ca/Mg ratio was certainly involved in destabilizing the dolomite with respect to calcite. Sulphide m ineralization
The sulphide paragenesis consists of sphalerite, galena, pyrite and marcasite. Sphaler ite is predominant and displays various fabrics, depending on its position in the deposit: 1. In the dolomitized facies of Unit 7, it occurs as disseminations of fine subhedral crystals along the intergranular solution joints of 02 ferroan dolomite (Fig. 10); these disseminations are mainly observed in the immediate vicinity of karst cavities and mineralized fractures.
2. In the recrystallized dolomite sand (solution cavities and base of Unit 8), it occurs as euhedral millimetric crystals that are generally zoned (Fig. 11). 3. In the fractures and microcavities, it occurs as ribbons of zoned crystals encrusting the walls (Fig. 12). The 03 ferroan baroque dolomite associated with the sulphides may also contain micro-inclusions of sphalerite in its growth zones, whereas the 04 baroque dolomite generally seals the sulphide deposition. Microscopic observation reveals two types of sphalerite. The first (sphalerite I) is dark coloured and is characterized by a radial microtexture reminiscent of a pseudomorph after wurtzite (Fig. 13), remnants of which were revealed through X-ray diffractometry. The second (sphalerite II) is more abundant and occurs as zoned euhedral to subhedral crystals varying between honey-coloured and brownish-red (Figs. 11 to 13). The sphalerite II is closely associated with galena, with the presence of euhedral and in places skeletal inclusions of this mineral (Fig. 14) indicating syncrystallization of the two sulphides . Sphalerite II is clearly later since it caps sphalerite I (see Fig. 13). Microprobe analysis of the two crystal types shows that: 1. Sphalerite I has an iron content that varies between 2 and 8 wt %, with most of the values being between 5 and 8 wt %. All trace elements, except Cd and Cu, are below detection level; in particular, Mn is below 600 ppm.
Fig. 10. Disseminated sphalerite (dots) along intergranular solution joints of D2 dolomite (dolomiti zed host rock, Unit 7)
2. Sphalerite II has a variable iron content between 1 and 8 wt %, the darkest-red areas being the most iron-rich; Mn was not detected. The trace metals present above the detection limit (600 ppm) are concentrated mainly in the honey-coloured zones and include Cd (max. 7300 ppm), Cu (max. 5600 ppm), Ag (max. 3900 ppm) and Ge (max. 3000 ppm); Hg (max. 2300 ppm) and As (max. 1500 ppm) have been detected locally. The galena associated with sphalerite II does not contain significant quantities of silver, but bismuth has been detected locally (max. 1000 ppm). Zoned sphalerite II is therefore the major silver-bearing mineral, with 15% of the measured crystals revealing Ag-contents higher than 1000 ppm . This was confirmed by atomic absorption analyses of sphalerite II-rich sulphide ores containing variable amounts of galena. Statistical interpretation of the results shows a good correlation between Zn and Ag, whereas Pb and Ag are independent.
Fluid-inclusion studies
Only fluids in the honey-colored and dark-red zones of sphalerite II were studied, the fluid inclusions in sphalerite I being hardly visible due to the high refraction index of the host mineral and those in baroque 03 and 04 dolomite being rare despite the fact that saddle dolomite is known to be a prime candidate for performing fluidinclusion studies (Radke and Mathis 1980). The inclusions in sphalerite II, generally flattened and of variable shape (Fig. 15), are approximately 5 to 40 urn long and contain more than 95% liquid . They could not be posit ively identified as primary for lack of definite criteria; however this is our preferred interpretation due to the fact that most of them do not underline healed fractures . Microthermometric measurements , using a Chaix-Meca stage properly calibrated for temperatures between - 56.6° and 307 °C, and using both reference fluid inclusions
51
Fig. 11. Euhedral sphalerite II crysta ls in recrystallized dolomite sand (D3 dolomitization). Reflected light
Fig. 13. Zoned sphalerite II capping sphalerite I in recrystallized dolomite sand. Transmitted light
Fig. 12. Detail of a veinlet developed in dolomite sand (left), coated by zoned subhedral sphalerite II (Sp) associated with bladed marcasite (Ma ), and cemented by ferroan D3 dolomite (D 3). Reflected light. Py pyrite part ially altered to goethite
Fig. 14. Syncrystallization of zoned sphalerite II (Sp) and skeletal galena (Ga) in recrystallized dolomite sand. Reflected light
and solids as standards (Poty et al. 1976; Touray 1989), were carried out on about 50 fluid inclusions. The precision of the measurements is estimated to be ± 0.2 °C for cryometric data and - ± 1 °C in the high T-range. Eutectic temperatures were estimated around - 55 °C for all fluid inclusions, which is compatible with the presence of brines in the H 20-NaCI-CaCI 2 system (cf. Crawford 1981). Due to the lack of data concerning the Na jCa ratio, the results were interpreted in a simplified way with reference to the NaCl-H 20 system. It is worth noting that some fluid inclusions in sphalerite II contain, besides liquid water, an additional transparent liquid meniscus of probable organic origin. This fluid emitted no fluorescence under the light of a mercury lamp, but did degrade under the laser beam.
Both the histograms of ice-melt temperatures (Tm) and homogenization temperatures to the liquid (Th) are bimodal, with two distinct groups of values in the intervals between - 6 to - 11°C and - 15 to - 20 °C and between 60- 85 °C and 85-110 °C respectively (Fig. 16). The existence of two trapped brine populations is illustrated in Fig . 17, which shows a plot of the homogenization temperatures (Th) against salinities (in terms of wt % equivalent NaCl, interpreted from the Tms: Potter et al. 1978). Figure 17 shows that the most-saline fluid inclusions displa y the highest homogenization temperatures (mean of 98 °C, against 74 °C for the least- saline solutions). Careful examination of the data revealed that most fluid inclusions in the honey-coloured sphalerite were of medium salinity, whereas those in the dark-red Fe-rich parts of the
52 +
130
Er ror bar
120 110
100
~ s: 90 f-
80 70
60 50+--.-~~-~-.-~~-~-.-~~-~-.-~~---,--J 8
9
10
II
12
13
14
15
16
17
18
19
20
21
22
23
24
Sa l i ni t y (eq . wI. % Noel)
Fig. 17. Salinity (interpreted from the Tms after Potter et al. 1978) versus homog enization tempe rature of fluid inclusions in sphalerite II
20
Fig. IS. Pr imary fluid inclusions in zoned sphalerite II
I-
lJ)
(\)
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12
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6
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Fig. 18. Distribution of reflectan ce values measured on organic matter in a Unit 7 sample . Similar bimod al histograms are characteristic of Unit 7 samples a nd of sample s at the Unit 6-Unit 7 contact
-
-
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z
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Fig. 16a, b. Histograms of a melting temperatures of ice (Tm) and b homogenization temperatures to liquid (Th) of fluid inclusions in sphalerite II
crystals mostly exhibited higher salinities and Th. Such a distribution is an indirect argument in favour of the primary nature of the fluid inclusions in sphalerite II.
Organic-matter reflectance Reflectance measurements on polished-rock sections of seven drill-hole samples from Units 6 and 7 were made at U.R.P .O . (University of Orleans; F . Defarge-Lagoun, analyst) using a reflected-light microscope equipped with a reflectometer, a photomultiplier and a potentiometric recording device. The reflectance value of each sample is the mean value of some 100 measurements made preferentially on vitrinite grains. All the studied samples yielded a first group of low reflectance val ues. Reflectances of 0.13 and 0.16 were obtained on two Unit 6 samples; values of 0.30 were recorded for two samples located at the contact between Units 6 and 7; intermediate values between 0.19 and 0.24 were obtained for two Unit 7 samples. Additionally, vitrinite particles from Unit 7 samples and from samples at the Unit 6-Unit 7 contact yielded higher values in the range 0.77-1.08 (Fig . 18). We propose that the low reflectance data characterize the in situ organic matter and that they point to its overall low maturity, compatible with a peat to peat /lignite grade (Kubler 1984; Robert 1985). This interpretation fits well with the low diagenetic grade of the sedimentary pile at Jabali proposed by Gauthier et al. (1985) based on organic
53
,
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(Stace y et al. 1980; Stace y and Hedge 1984). At Jabali, therefore, it can be considered that the lead and the associated metals were derived either directly, or via the basal detrital fill of the Wadi al Jawf basin, from a basement of Earl y Proterozoic age. The exact location of this basement is unknown at pre sent due to lack of a sufficient number of radiometric datings in Yemen.
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Discussion and conclusions
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19.5
Fig. 19. Lead isotope composition of the Jabali deposit (25 analyses: 9 galenas, and 16 oxidized ores) and of three oth er Pb -Zn prospects located south of the Wadi al Jawfbasin (see Fig. 3): Haylan (Stacey et al. 1980), Majn ah (2 galenas) and Barr an (2 oxidized ores). N Nujum Pb-Zn prospect in Om an (Stacey et al. 1980) A composition field of galenas from the Arabian-Nubian Shield (Stacey et al. 1980) B isotop ic compo sition of the Pb-Zn Miocene deposits along the Red Sea. Evolution curves from Doe and Zartman (1979)
geochemistry of Unit 8 samples. Higher reflectance values measured around 0.90 are typically characteristic of samples located in the oil window (Robert 1985). Based on the characteristic Th-range of fluid inclusions in sphalerite II, we also propose that part of the organic matter from the Unit 6-Unit 7 contact and from Unit 7 was matured by the mineralizing solutions, yielding the characteristic reflectance data in the range 0.77-1.08. Finally, Gauthier et al. (1985) were able to show that the thermal influence of the Tertiary volcanic dykes on organic matter was not visible beyond 3-5 m from the contact.
Lead isotope geochemistry Isotopic analysis of 9 galenas and 16 cerussite-rich oxide ores from Jabali show s great homogeneity in the values, with the initial isotopic signature being preserved in spite of the oxidation and regardless of the unit from which they were collected (respectively Units 3, 7 and 8). The results range as follows: 206Pb/ 204Pb : 18.85-18.95 20 7PbF o4Pb : 15.66-15.72 20 8PbF o4Pb: 39.71-39.92 Lead isotope compositions of two galen as from the Majnah prospect and two ceru ssite-rich facies from the Barran prospect (see Fig. 3 for location) show respectively lower and higher 206Pb / 204Pb ratios, indicating different sources for the lead. On standard correlation diagrams however (Fig. 19), all the values lie close to the evolution curve of Upper Continental Crust lead (Doe and Zartman 1979). They are much more radiogenic than mo st values obtained from the Late Proterozoic of the Arabian Shield , and are close to those found at Haylan (east of Jabali) and at Nujum (in Oman), which have been considered to be indicative of the contribution of an Early Proterozoic crustal component
Like many carbonate-hosted lead-zinc deposit s, the Jabali deposit was controlled by karstic processes related to a surface of emergence (the 'below unconformity' type of Callahan 1964, 1967). The relationships between the mineralization and the host rock demonstrate that the sulphide emplacement postdated two successive phases of dolomitization immediately preceding the local emergence of the carbonate shelf. The second and major phase is characterized by a ferroan dolomite that forms the host rock of the orebody, a situation that has been described in many deposits of this type (Lagny and Rouvier 1976). The ferroan dolomite was reworked into the dolomite sand filling the karstic network and doline depressions of the erosion surface, and the major part of the mineralization was sub sequently emplaced as impregnations of the dolomite sand. This succession of events (dolomitization, emergence, karst solution and sedimentation, mineralization) and the setting of the various ore facies are similar to those of the Early Jurassic Treves depo sit of the Causse s basin margin (southern French Massif Central; Macquar and Lagny 1981; Lagny et al. 1981). At a regional scale, the location of the Jabali ore deposit corresponds to a particular paleogeographic and paleostructural setting - the Jabali-Majnah block - the uplift of which was initiated during Late Jurassic times along the unstable margin of the carbonate platform at its boundary with a subsiding evaporitic basin. This situation resulted from rifting-type tensional tectonic activity characterized by rapid lateral and thickness changes, synsedimentary movements in the transition zone between shelf and basin (slumping and slope breccia) , and local emergence. The site of ore depos ition within this regional paleostructural environment appears to have been controlled by the intersection of a WNW-striking boundary fault system and NNW-striking fractures. Such an environment is typical of numerous carbonatehosted lead-zinc deposits and districts such as: Treves (previously mentioned); the eastern Alpine Triassic Province with Salafossa and Raibl in Italy , the B1eiberg district in Austria , and Mezica in Slovenia (Brigo et al. 1977; Bechstadt 1979; Klau and Mostler 1986); several Jurassic and Cretaceous deposits of North Africa, in particular those of the Touissit-Bou Beker-EI Abed district in the Algeria -Morocco border area (Samson 1973; Touahri 1983; Wadjinny 1989); most of the Early Cambrian deposits in Iglesiente (Sardinia, Ital y) (Boni 1985; Bechstadt et al. 1988; Courjault-Rade and Gandin 1988); and the Irish Earl y Carboniferous Province (Deeny 1987; McArdle 1990). The geotectonic setting of these depo sits is quite different from that of the Mississippi Valley-type (MVT) deposits of the USA, as emphasized by Maucher and
54
Schneider (1967) and recently by Sangster (1983, 1990), which are located at the margins of stable intracratonic basins (MVT stricto-sensu). In spite of numerous common features between the two types of deposit, this fundamental difference clearly characterizes rift-related lead-zinc deposits (Sawkins 1990),whether these rifts are mature, reaching the oceanization stage as in the Alps, or aborted as is the case for Yemen and North Africa. Brine circulation in the rift settings is likely to have occurred during synsedimentary tectonic activity (as previously discussed). The age of ore emplacement at Jabali cannot be determined precisely , although ore deposition took place after erosion and karstification following the second stage of dolomitization, and possibly slightly later than the sedimentation of Unit 8 (Late Jurassic-Early Cretaceous?). The observation of reworked mineralized fragments in the ore (see Fig. 9) could be indicative of sulphide deposition prior to lithification of the karst infill. Whatever the exact age of ore deposition, it probably occurred within about 15 Ma following the sedimentation of the host carbonates. This would imply a relatively short time span compared with that deduced for some MVT ore deposits, as for example in the Appalachians and in the midcontinental region of the USA (Hearn et al. 1987; Sangster and Symon s 1991). A definitive determination of the age of the mineralization could provide constraints for developing a genetic model (Sangster 1983). However, discussions about ore genesis do not really concern the timing of ore deposition only. According to the models proposed by Anderson and MacQueen (1982), Cathles and Smith (1983), Sverjensky (1989), the discussion ofthe genesis of the Jabali deposit will be based on the three following points: • The basinal sources of the mineralizing fluids and base metals. • The nature of the main aquifer that drained and possibly modified these solutions. • The mechanisms of metal transport and deposition (mixing models, sulphate reduction models, reduced sulphur models). At Jabali, the lead-isotope composition of the ore indicates an initial metal source from an old Proterozoic crust, probably reworked before being incorporated at the base of the sedimentary pile in the Wadi al Jawf basin. On the other hand, the microthermometric characteristics of the fluid inclusions, typical of Na-Ca-CI brines, plot in the same field as those obtained for many carbonate-hosted Pb-Zn deposits (Roedder 1976; Hanor 1979; Touray 1989). They show a basinal character and temperatures higher than the diagenetic grade of the carbonate wallrock . These features, according to geological data, suggest that the solutions could only have migrated from the Wadi al Jawf basin. The main channelling aquifer was probably the detrital Unit 1, and the boundary fault system could have played a significant role in transferring the brines from the deeper parts of the basin to the uplifted Jabali-Majnah zone. Assessment of the exact role played by the various mechanisms involved in the transportation and deposition of base metals is, as often, difficult. The mineralogical and geochemical studies have revealed variations in the
physico-chemical conditions during the crystalli zation of the zoned sphalerite II; these are compatible with successive, possibly periodic, influxes of distinct mineralizing solutions originating from the deep parts of the basin (cf. Anderson and MacQueen 1982; Cathles and Smith 1983), as evidenced by: • The presence of low-temperature wurtzite, which suggests a very reducing environment and a pH close to neutral (Scott and Barnes 1972). Conversely, the pre sence of marcasite is indicative of an acid pH and an oxidizing environment (Arne et al. 1991). • The successive crystallization of two generations of sphalerite, the later one being associated with lead . Also, the sph alerite II is richer in Cd , Cu and Ag, and contains Ge, which suggests that these trace metals were transported with lead and zinc during the second stage of ore deposition. • The presence of two populations of fluid inclusions in sphalerite II, with contrasted salinities and homogenization temperatures (F ig. 17). • The strong variations in the iron content of sphalerite II, which probably resulted from variations in oxygen fugacity, temperature, pH , or total sulphur in solution (Barton and Skinner 1979). Oilfield brines have become the most favoured ore-forming solutions for MVT deposits (Billings et al. 1969; White 1974; Sverjensky 1984). However, the concentrations of Pb and Zn in oilfield waters are usually well below the minimum values of about 10 mg /l required for the formation of MVT ore deposits (Barnes 1969, 1979; Barnes and Czamanske 1967). Lead- and zinc-ri ch oilfield waters have been reported from a large area in the Central Mississippi Salt Dome Basin (Ca rpenter et al. 1974; Kharaka et al. 1987) where high metal concentrations are onl y pre sent in waters with extremely low concentrations of HzS. Geochemical modelling (Kharaka et al. 1987) show s that mixing of metal-rich/S-deficient water with S-rich /metal-deficient water is a likely process for the precipitation of ores from brines . Such a mixing model could account for a casual precipitation of wurtzite as observed at Jabali; this ZnS polymorph may form when sulphur-deficient or zinc-rich solutions are involved (Scott and Barnes 1972). At Jabali, however, models involving mixing of various proportions of two fluids from two reservoirs a re hardly compatible with the existence of a gap between the two distinct populations of fluid inclusion in sphalerite II (Figs. 16, 17). If both populations are primary, as suggested by the correlation of each population with zones of specific colour, and mixing was the driving force for ore precipitation, one would expect fluid inclusions with intermediate characteristics. At best, a rough linear trend is observable within population I, suggesting dilution and a temperature drop after limited mixing. It is to be noted that simple brine dilution is not a valuable sulphide precipitation mech an ism, except at pH values well below conditions believed to be geologically reasonable (Rowan and Leach 1989). Following classical discussions on the phy sico-chemical conditions of formation of Mississippi Valley type deposits, the T-X properties of the mineralizing solutions at Jabali can
55
hardly be reconciled with present knowledge of appropriate base-metal solubilities (cf. Anderson and MacQueen 1982) since (1)experimental data show that, below 150°C, chlorides are inappropriate ligands for transporting Zn and Pb in sufficient amounts to form a deposit (Barrett and Anderson 1982), and (2)metal solubilities in solutions containing appreciable amounts of reduced sulphur are low (Anderson 1975). Indeed, the only reported oilfield brines with high Pb- and Zn-contents in the Central Mississippi Valley have extremely low concentrations of H 2S (Carpenter et al. 1974; Kharaka et al. 1987). Besides, Pb-complexing by bisulfides is ruled out (Giordiano and Barnes 1981). To account for the transportation of sufficient amounts of Pb and Zn below 150°C, some ligands other than chlorides or bisulfides are required. Organic complexes are frequently invoked in this T-range (see Drummonds and Palmer 1986). At Jabali, some inclusions in sphalerite indeed contain, besides water, an additional liquid of probable organic origin. Acknowledgement s. We should like to thank Mr . Ali Jaber Alawi, Deputy Minister at the Ministry of Oil and Minerals (Republic of Yemen), for his support of our work . We should also like to express our gratitude to V. Perthuisot and J.R. Disnar of the Univers ity of Orleans (France) and to E. Marcoux of BRGM (France) for their invaluable contributions in the fields of lithostratigraphy, organic matter and lead isotopes respectively. L. Fontbote and a second reviewer are acknowledged for their improvements to the manuscript which was translated from the French by Sir Patrick Skipwith, Bt.
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