The Jabali nonsulfide Zn–Pb–Ag deposit, western Yemen Ore

Transcription

Ore Geology Reviews 61 (2014) 248–267
Contents lists available at ScienceDirect
Ore Geology Reviews
journal homepage: www.elsevier.com/locate/oregeorev
The Jabali nonsulfide Zn–Pb–Ag deposit, western Yemen
N. Mondillo a,⁎, M. Boni a, G. Balassone a, M. Joachimski b, A. Mormone c
a
b
c
Dipartimento di Scienze della Terra, dell'Ambiente e delle Risorse, Università degli Studi di Napoli Federico II, Via Mezzocannone 8, 80134 Napoli, Italy
GeoZentrum Nordbayern, University of Erlangen-Nuremberg, Schlossgarten 5, 91054 Erlangen, Germany
INGV Osservatorio Vesuviano, Via Diocleziano 328, 80124 Napoli, Italy
a r t i c l e
i n f o
Article history:
Received 19 August 2013
Received in revised form 30 January 2014
Accepted 5 February 2014
Available online 14 February 2014
Keywords:
Yemen
Nonsulfide deposit
Pb–Zn–Ag
Mineralogy
Stable isotopes
a b s t r a c t
The Jabali Zn–Pb–Ag deposit is located about 110 km east of Sana'a, the capital of Yemen, along the western
border of the Marib-Al-Jawf/Sab'atayn basin. The economic mineralization at Jabali is a nonsulfide deposit,
consisting of 8.7 million tons at an average grade of 9.2% zinc, derived from the oxidation of primary sulfides.
The rock hosting both primary and secondary ores is a strongly dolomitized carbonate platform limestone of
the Jurassic Shuqra Formation (Amran Group). The primary sulfides consist of sphalerite, galena and pyrite/
marcasite. Smithsonite is the most abundant economic mineral in the secondary deposit, and is associated
with minor hydrozincite, hemimorphite, acanthite and greenockite. Smithsonite occurs as two main generations:
smithsonite 1, which replaces both host dolomite and sphalerite, and smithsonite 2, occurring as concretions and
vein fillings in the host rock. At the boundary between smithsonite 1 and host dolomite, the latter is widely
replaced by broad, irregular bands of Zn-bearing dolomite, where Zn has substituted for Mg. The secondary
mineralization evolved through different stages: 1) alteration of original sulfides (sphalerite, pyrite and galena),
and release of metals in acid solutions; 2) alteration of dolomite host rock and formation of Zn-bearing dolomite;
3) partial dissolution of dolomite by metal-carrying acid fluids and replacement of dolomite and Zn-bearing
dolomite by a first smithsonite phase (smithsonite 1). To this stage also belong the direct replacement of
sphalerite and galena by secondary minerals (smithsonite and cerussite); 4) precipitation of a later smithsonite
phase (smithsonite 2) in veins and cavities, together with Ag- and Cd-sulfides.
The δ18O composition of Jabali smithsonite is generally lower than in other known supergene smithsonites,
whereas the carbon isotope composition is in the same range of the negative δ13C values recorded in most supergene nonsulfide ores. Considering that the groundwaters and paleo-groundwaters in this area of Yemen have
negative δ18O values, it can be assumed that the Jabali smithsonite precipitated in different stages from a combination of fluids, possibly consisting of local groundwaters variably mixed with low-temperature hydrothermal
waters. The carbon isotope composition is interpreted as a result of mixing between carbon from host rock
carbonates and soil/atmospheric CO2.
The most favorable setting for the development of the Jabali secondary deposit could be placed in the early
Miocene (~17 Ma), when supergene weathering was favored by major uplift and exhumation resulting from
the main phase of Red Sea extension. Low-temperature hydrothermal fluids may have also circulated at the
same time, through the magmatically-induced geothermal activity in the area.
© 2014 Elsevier B.V. All rights reserved.
1. Introduction
The Jabali zinc deposit is located at an altitude of 1800 m above sea
level in a mountainous desert terrain along the western border of the
Marib-Al-Jawf/Sab'atayn basin, about 110 km east of Sana'a, the capital
of Yemen. It represents the most significant base metal deposit of
Yemen (Grist, 2006; Watts, Griffis, and McOuat, 1993; Yemen Geological Survey and Mineral Resources Board, 1994, 2009). The artisanal
mine workings in the area are thought to be over 2500 years old. The
old workings for silver and lead were extended over an area of about
⁎ Corresponding author. Tel.: +39 081 2535063.
E-mail address: nicola.mondillo@unina.it (N. Mondillo).
http://dx.doi.org/10.1016/j.oregeorev.2014.02.003
0169-1368/© 2014 Elsevier B.V. All rights reserved.
10 h, tracing cavities filled by relatively soft oxidized ore, locally rich
in silver. The ore was processed on site. Still existing waste dumps
contain about 120,000 t, at average grades of 24% Zn, 3.5% Pb, and
160 ppm Ag. The artisanal metallurgical process was not effective,
since slags still contain 23% Zn, 6.5% Pb, and 40 ppm Ag (SRK
Consulting, 2005).
The Jabali site was re-discovered by the Bureau de Recherches
Géologiques et Minières (BRGM) and the Yemen Geological Survey
and Mineral Resources Board (YGSMRB) in 1980. Between 1981 and
1986, an exploration and evaluation program, based on 57 drill holes,
reported an accessible-by-open-pit resource of 3.0 Mt, at 15.2% Zn,
and an amenable-by-underground mining tonnage of 1.24 Mt at 13%
Zn. The exploration was mainly aimed to nonsulfide ores. During and
Author's personal copy
N. Mondillo et al. / Ore Geology Reviews 61 (2014) 248–267
after the period of exploration at Jabali, BRGM and YGSMRB produced
several reports, a few scientific papers on the characterization and
genesis of the deposit (Al Ganad et al., 1994; Christmann et al., 1989)
and a PhD thesis (Al Ganad, 1991). No further scientific research papers
were published in more recent years, with the exception of a preliminary mineralogical study by Mondillo et al. (2011).
In June 1993, Watts, Griffis and McOuat (WGM) issued a prefeasibility study indicating reserves of 3.6 Mt at 16.4% Zn. In 1996,
Minorco and Ansan Wikfs were granted a license over the deposit. In
1998, ZincOx Resources – whose main interest was in nonsulfides –
entered into a joint venture with Minorco and Ansan Wikfs, thus
becoming the manager and operator for the development of the deposit.
In 2004–2005, the exploitation and development rights of the Jabali zinc
deposit were owned by Jabal Salab Company (Yemen) Ltd., a company
in which ZincOx previously had a 52% interest, with the balance (48%)
held by Ansan Wikfs Investments Ltd. In that period, ZincOx, together
with SRK Consulting, concluded the feasibility study of the deposit,
and reported resources consisting of 12.6 Mt at 8.9% Zn, 1.2% Pb, and
68 ppm Ag. In March 2013, due to the realignment of the ZincOx
strategy towards recycling instead of exploitation of primary natural
resources (ZincOx press release, 2013), ZincOx sold its interest in the
Jabali Project to Ansan Wikfs, which actually owns the deposit.
The feasibility study on the Jabali deposit considered a number of
processing routes, owing to its rather complex mineralogical association. The metallurgical problems encountered over the years are essentially a consequence of the incomplete knowledge of the mineralogy
and genesis of the Jabali nonsulfide orebody.
In the present study, after a brief review on the primary sulfide
mineralization, we focus on the mineralogical, geochemical and isotopic
characterization of the nonsulfides and their relationships to the host
rock, with the aim of identifying the supergene oxidation processes.
2. Regional geology
The geology of Yemen is characterized by (i) Precambrian basement
transected by a failed Jurassic rift system forming during the break-up of
the Gondwana supercontinent; (ii) Jurassic pre-, syn-, and post-rift
carbonate and siliciclastic sedimentary sequences; and (iii), Tertiary to
recent sedimentary and magmatic rocks associated with the opening
of the Gulf of Aden-Red Sea rift (Fig. 1) (Menzies et al., 1994).
249
The basement rocks in Yemen are considered as part of the
Proterozoic Arabian–Nubian Shield, which covers northeast Africa and
the Arabian Peninsula (Whitehouse et al., 2001). They consist of
metavolcanic and metasedimentary suites, deposited in arc environments, deformed and intruded by post-tectonic granites and granodiorites during the Pan-African orogeny (Whitehouse et al., 1998). The
southern Arabian Peninsula underwent post-orogenic extension
accompanied by magmatism. This was associated with uplift and
erosion of the basement rocks, followed by the formation of several
basins related to major wrench-fault systems (e.g. Najd fault-system;
Ellis et al., 1996). The basins are filled with Paleozoic to early Mesozoic
siliciclastic sequences, deposited in marine epicontinental to deltaic
environments, locally containing evaporites (Beydoun, 1997).
From Triassic to middle Jurassic, Yemen was part of the Afro-Arabian
plate of western Gondwanaland. The break-up of Gondwana initiated
the extensional tectonics in Yemen and the formation of normal faults
oriented along tectonic trends inherited from the Precambrian
wrench-fault systems (Ahlbrandt, 2002; Bosence, 1997 and references
therein). The Mesozoic extensional tectonic regime resulted in the
formation of the Siham-Ad-Dali', Sab'atayn, Say'un–Masilah, Balhaf,
and Jiza'–Qamar basins (Fig. 1) (As-Saruri et al., 2010). The Sab'atayn
and the Say'un–Masilah basins are the only hydrocarbons producing
basins in Yemen; the Jabali mineral concentrations are located at the
western border of the Sab'atayn basin.
Pre-rift sedimentation within the basins (Toarcian–Bathonian) is
initially represented by the continental fluviatile red bed sediments of
the Kuhlan Formation. The red beds grade upward into shallowmarine sediments, representing the early transgressive phase of the
Middle Jurassic sea (Bathonian–Callovian) (Beydoun et al., 1998). The
most widespread Jurassic marine deposits correspond to the Amran
Group, which consists of pre-rift sediments, as well as of syn- and
post-rift sequences (Csato et al., 2001). Beydoun et al. (1998), in a
review on the stratigraphy of Yemen, established the “official” subdivision of the Amran Group. They distinguished: (i) the Shuqra Formation
(Callovian–Oxfordian); (ii) the Madbi Formation (Kimmeridgian–
Tithonian); (iii) the Sab'atayn Formation; and (iv) the Naifa Formation
(upper Tithonian–Berriasian). These formations have been further
subdivided into several members, based on the heteropic facies of the
sediments (Beydoun et al., 1998). After Csato et al. (2001) the Shuqra
Formation was deposited in a pre-rift regime, whereas the Madbi and
Fig. 1. Geological map of Yemen, showing the location of Jabali (Yemen Geological Survey and Mineral Resources Board, 2009, modified). Hydrocarbon producing Mesozoic
basins: 1 =Sab'atayn basin; 2 = Say'un–Masilah basin.
250
Sab'atayn Formations have the characteristics of syn-rift successions.
The Naifa Formation is interpreted as a post-rift sedimentary sequence.
In the following descriptions of the sedimentary sequences, we will
keep the original terminology of the authors quoted in the literature
(Al Ganad et al., 1994; As-Saruri et al., 2010; Beydoun et al., 1998;
Youssef, 1998).
The Shuqra Formation, which is the main host of the Jabali ores, is
characterized by the following lithological units: 1) basal intertidal
limestones; 2) marls and thinly bedded carbonaceous biomicrites;
3) foraminiferal biomicrite with chert nodules; 4) dolomitic marls interbedded with dolostones, and oolitic–oncolitic limestones interbedded
with fossiliferous limestones and bioclastic sandstones; 5) coral–algal
stromatolitic limestones and black foraminiferal biomicrites (Youssef,
1998).
The Madbi Formation consists of bituminous marine shales, deltaic
sandstones, debris flow breccias, well-bedded limestones, and
turbidites (As-Saruri et al., 2010). The Sab'atayn Formation consists
mainly of evaporites (Beydoun et al., 1998). In the Naifa Formation
dolostones, dolomitic shales, and fine-grained limestones occur, locally
yielding a rich ammonite fauna (Beydoun et al., 1998; Menzies et al.,
1994).
In the Sab'atayn and Say'un–Masilah basins, the siliciclastic and
carbonate members of the Madbi Formation represent the main
onshore petroleum reservoirs of Yemen (Ahlbrandt, 2002; As-Saruri
et al., 2010). The oil generation may have started in the Late Cretaceous
(Csato et al., 2001).
Cretaceous sediments, ranging in age from Berriasian to
Maastrichtian (Tawilah Group), completed the infilling of the Jurassic
graben (Beydoun et al., 1998). In western Yemen, the main Cretaceous
lithofacies consist of medium- to very coarse-grained, cross-bedded
sandstones, interrupted by numerous paleosols. The total thickness of
the Tawilah Group can vary from a few hundred meters in the Sana'a
area to a few thousand meters in the Sab'atayn basin (Beydoun et al.,
1998).
A second important rifting phase was associated with the opening of
the Red Sea and the Gulf of Aden during Oligocene and Miocene. At that
time, widespread volcanism followed by uplift and denudation affected
the whole of western Yemen, whereas the eastern part of the country
was characterized by continuous sedimentation and complete absence
of magmatic activity (Menzies et al., 1994). The volcanism was associated
with the upwelling of the Afar plume beneath the Africa-Arabian plate
(Bosworth et al., 2005). The Oligocene igneous rocks, known as the
Yemen volcanic group or Yemen trap series (Moseley, 1969), consisting
of more than 3 km-thick flood basalts to massive ignimbrites, erupted
between ~31 Ma and ~29 Ma, and between ~29 Ma and 26 Ma (Baker
et al., 1996; Coulié et al., 2003). A stratigraphic gap (~ 26 to ~ 19 Ma),
marked by an unconformity at the top of the deformed Yemen volcanic
group, indicates that the main rifting phase started in late Oligocene or
early Miocene (Ukstins et al., 2002). Early Miocene magmatic rocks
consist of small plutonic bodies (age ~22–21 Ma), and abundant mafic
and felsic dykes (age ~ 25–16 Ma), both emplaced along the Red Sea
margin (Zumbo et al., 1995). Pliocene to Quaternary volcanism is also
fairly widespread in the areas NW of Sana'a–Amran, Dhamar–Rada,
Marib–Sirwah, Balhalf–Bir Ali, and Shuqra (Menzies et al., 1994).
Fission track ages and length data for apatite specimens from
Pan-African basement rocks of Yemen indicate a significant exhumation
phase at ~ 17–16 Ma, associated with the main continental rifting
(Menzies et al., 1992). The fossil beaches along the Gulf of Aden record
a continuous tectonic uplift during most of the Plio-Quaternary
(Brannan et al., 1997).
About twenty zinc–lead occurrences have been reported in Yemen,
most of them located along the margin of the rifts, or in rift-affected
blocks. Lead and zinc deposits are hosted in Jurassic to Paleocene carbonate rocks within the Sab'atayn basin, and form two main clusters
in the Jabali and Tabaq areas (Fig. 1) (Yemen Geological Survey and
Mineral Resources Board, 1994, 2009).
The Jabali area includes the Jabali deposit and several other mineral
occurrences, which are characterized by maximum grades of 16.5% Zn
and 6% Pb (Yemen Geological Survey and Mineral Resources Board,
1994, 2009).
The Tabaq area, where nine small Pb–Zn occurrences (with some
barite) have been identified, is located in southern Yemen, approximately 360 km east-northeast of Aden and 500 km east of the Jabali
mine site, in the same rift system hosting the Jabali deposit (Fig. 1).
The ore deposit is characterized here by maximum values of 12% Zn
and 3.8% Pb.
Another base-metal mineralized district is located in the Mukalla
area, in the southwestern part of the Say'un–Masilah basin (Fig. 1).
The ore concentration is fault-controlled, and consists of barite and
galena in veins. A secondary mineralization containing willemite,
smithsonite, cerussite, descloizite, calcite, pyrolusite and celestine,
with anomalous grades of Ag, Cd, Ga, Ge, and Mo was also identified
in this area (Mattash, 2008; Mattash et al., 2005).
3. Jabali geology and ores — previous studies
3.1. Local geology
The Jabali deposit (15°37′ N latitude, 44°46′ W longitude) covers an
area of about 2 km2, at an altitude between 1850 and 1950 m.s.l. (Fig. 2).
The deposit is located on a small plateau on the eastern flank of a
NW–SE-elongated mountainous area that is a segment of the western
boundary of the Sab'atayn basin. In the southeast the plateau is
Fig. 2. Geological map of the Jabali mining site with the location of analyzed drill holes, and
the future open pit area (from SRK Consulting, 2005, modified). Description of the units in
text and Fig. 3.
dissected by the Wadi Khaynar valley, which cuts the mineralized successions exposed along its flanks.
The deposit area is dominated by the sedimentary rocks of the
Amran Group, hosts of the deposit (Al Ganad, 1991; Al Ganad et al.,
1994; Christmann et al., 1983, 1989). The maximum thickness of
the Amran Group at the mine site is 300 m. The Shuqra Formation
(Callovian–late Kimmeridgian) is locally subdivided in seven units
(members), which from the base of the succession are (Al Ganad
et al., 1994) (Figs. 2, 3):
Unit 1: sandstones and conglomerates, transgressive on the
Proterozoic basement (10 m);
Unit 2: gypsiferous mudstones and dolomitized calcarenites, interbedded with marls and nodular limestones (25 m);
Unit 3: biomicritic limestones (Callovian) with nodular concretions
and chert layers (50 m);
Unit 4: micritic limestones and laminated dolomites (15 m);
Unit 5: partly dolomitized bryozoan calcarenites (late Oxfordian–
early Kimmeridgian), overlain by coral-bearing oolitic and oncolitic
limestones. A local disconformity occurs at the top of the unit (40 m);
Unit 6: gypsiferous mudstones, followed by ammonite-bearing
limestones interbedded with marls and calcareous sandstones
(Kimmeridgian) (80 m);
Unit 7: partly dolomitized massive bioclastic and biomicritic limestones, locally oolitic limestone with coral bioherms (Kimmeridgian).
This unit is exposed at the top of the Jabali plateau, and is strongly
affected by karstic erosion (80 m).
The Madbi Formation (maximum thickness of 30 m) is locally called
Unit 8 and consists of black mudstones and argillites with gypsum
Fig. 3. Stratigraphy of the Jabali area with the units established by Al Ganad et al. (1994)
(modified), correlated with the corresponding formations of the “standard” Amran
Group of Beydoun et al. (1998). Shuqra Formation — Unit 1: sandstone and conglomerate,
transgressive on the Late Proterozoic basement; Unit 2: gypsiferous mudstone overlain
by dolomitized calcarenite, marl and nodular limestone; Unit 3: micritic-biomicritic
limestone (Callovian), with nodular concretions and chert layers; Unit 4: micritic limestone and lagoonal/lacustrine dolomite; Unit 5: partly dolomitized bryozoan calcarenite
(Late Oxfordian–Early Kimmeridgian), overlain by coral-bearing oolitic limestone;
Unit 6: gypsiferous mudstone grading into micritic limestone (Kimmeridgian) and marl;
Unit 7: massive bioclastic–biomicritic limestone, locally oolitic with coral bioherms
(Kimmeridgian). Madbi Formation — Unit 8: black mudstone and argillite with gypsum
and dolomite intercalations, grading laterally into micritic ammonite-bearing limestone
(Late Kimmeridgian–Tithonian). Sab'atayn Formation — Unit 9: biomicrite with oncolites
and bio-oocalcarenite (Late Jurassic).
251
crystals and dolomite intercalations. It grades laterally into ammonitebearing limestones.
The Sab'atayn Formation (locally known as Unit 9) consists of
biomicrites with oncolites and bio-oocalcarenites; intercalations of
gypsum lenses and arkosic sandstones also occur.
In the Jabali area, early Miocene trachytic sills and dykes (dated at 22
Ma) cut the sedimentary rocks (Al Ganad et al., 1994). Several Tertiary
alkaline granite bodies occur at Jabal as Saad, 15 km to the west of the
mine site. Less than a kilometer south of Jabal as Saad, a thick Pliocene
to Holocene travertine deposit indicates that thermal springs were
active in this area, in analogy to travertine deposits in the Sirwah area,
associated to the Marib volcanic field (Weiss et al., 2009).
The structural setting at Jabali is dominated by extensional rift
tectonics and the rocks were affected mostly by brittle deformation.
The sedimentary succession is almost horizontal at the periphery of
the basin, and only small areas are characterized by west-slightly
dipping strata. The most prominent normal faults strike 120 to 140°
(Fig. 2). This fault set includes the main Jabal Salab fault that borders
the shoulder of the plateau below the Jabal Salab peak. Another set of
normal faults has 65° to 80° trends. A further major system is developed
at 0° to 5°, while a more subtle fracture trend ranges between 25° and
40° (SRK Consulting, 2005). All fault planes have dip angles between
60° and 80°.
3.2. Mineralization
The Jabali ores (Fig. 4A) are hosted in the higher part of the Shuqra
Formation. The ore is almost completely oxidized (Fig. 4B, C), even
though primary sulfides have been locally preserved in the interval
underlying the Madbi Formation (Unit 8) (Fig. 4D). The nonsulfide ore
at Jabali is massive, semi-massive and disseminated, and consists of
vuggy to highly porous, brown-orange to white nonsulfide mineral
aggregates (Fig. 4E, F). A porous cellular boxwork structure, accompanied by numerous cavities coated with secondary zinc minerals is also
widespread.
The ore bodies are only partly exposed, since at least half of the
mineralized lithologies occur in the subsurface below Jabal Barrik
(SRK Consulting, 2005). The mineralization is structurally and lithologically controlled. This is reflected in the stratiform-to-tabular architecture of the ore bodies, and in the mineral enrichments occurring along
vertical fractures, faults and at the intersection of the above structures.
Specifically, the mineralization is developed along the three main structural trends recognized in the field. At the intersection of faults, big
vertical ore bodies occur, called “chimneys” by local exploration geologists. The stratiform bodies occur in three different zones: a laterally
extensive upper zone, and more sporadic lower and middle zones.
These bodies are generally flat and manto-like but, at the base of the
Jabal Salab massif, along the NW–SE fault they dip towards NE with
angles greater than 30° (SRK Consulting, 2005).
The more extensive studies on the Jabali primary mineralization
were performed by Al Ganad et al. (1994). According to these authors
the primary sulfide association consists of sphalerite (predominant),
galena, and pyrite/marcasite. Sphalerite occurs as two distinct generations: a first dark-colored, and a second represented by zoned euhedral
to subhedral honey-colored or brownish-red crystals. The second
generation is the most abundant; it is not pure, but contains iron, silver,
cadmium, copper, germanium and mercury.
The ore deposit at Jabali is hosted in the topmost horizons of the
Shuqra Formation (Unit 7), specifically when this interval is dolomitized
(Al Ganad et al., 1994). Two dolomitization stages have been distinguished: an early district-scale diagenetic (dolomite phases D1, D2)
and a second one, more local, associated with sulfide deposition (dolomite phases D3, D4). The dolomite phases are strongly fabric destructive
and have obliterated both depositional and diagenetic features of the
precursor limestone. In detail, D3 and D4 dolomites are characterized
by a saddle structure (Radke and Mathis, 1980), with crystal sizes
252
Fig. 4. A) View of Wadi Jabali, looking E: old mine workings. B) JS MON 3. Smithsonite in outcrop, with a vuggy-highly porous texture. C) JS MON 28. Hydrozincite coating smithsonite and
host rock. D) J125-2. Partly oxidized ore, with remnants of sphalerite and galena. E) J125-15. Massive smithsonite, replacing Zn-dolomite, and smithsonite crusts in cavities. F) J125-3.
Gypsum veins, cutting both dolomite and smithsonite. Dol = dolomite; Gn = galena; Gp = gypsum; Hz = hydrozincite; Sm = smithsonite; Sp = sphalerite; Zn-Dol = Zn-bearing
dolomite.
comprised from a few hundred microns to few millimeters. The saddle
crystals occur within cavities and fill fractures crosscutting the previous
dolomite generations. A calcite phase is associated with saddle
dolomite, and mainly occurs as late cement completely filling the
porosity.
Fluid inclusions in sphalerite (Al Ganad et al., 1994) show a bimodal
grouping (60–85 °C and 85–110 °C) of homogenization temperatures
to the liquid, mirrored by salinities of 10–14 eq. wt.% NaCl and
19–23 eq. wt.% NaCl, respectively. These data are compatible with a
basinal fluid origin: salinities and temperatures are similar to many
other carbonate-hosted zinc–lead Mississippi Valley-type deposits
(Roedder, 1976). Lead isotope ratios of galena and cerussite are in the
range of: 18.85 and 18.95 206Pb/204Pb, 15.66 and 15.72 207Pb/204Pb,
and 39.71 and 39.92 208Pb/204Pb (Al Ganad et al., 1994). After these
authors the lead isotope compositions point to an original metal source
in the early Proterozoic crustal basement of the Sab'atayn basin.
Based on these observations, Al Ganad et al. (1994) stated that
the primary mineralization was deposited in karst cavities related
to the emersion surface at the top of the Shuqra Formation (Unit
7), from fluids migrating from the Sab'atayn basin during Mesozoic
rifting. They also proposed that sulfide ores were emplaced slightly
after sedimentation of the Madbi Formation (Unit 8) (Late Jurassic–
Cretaceous). The black shales of the Madbi Formation acted as an
impermeable barrier to fluid migration. Al Ganad et al. (1994) also
suggested that oxidation of the primary deposit may have begun
during Cretaceous, developed during Paleogene and continued up
to the present.
C. Allen (unpublished data, 2000) speculated on a different genetic concept, hypothesizing a hybrid Carbonate Replacement Deposit–Mississippi Valley-type character for the primary Jabali
mineralization.
Reynolds and Large (2010), in their revision on the Tethyan zinc–
lead metallogeny in Europe, northern Africa and Asia, have considered
Jabali as a Mississippi Valley-type deposit formed in a carbonate
platform during the Sab'atayn basin evolution.
4. Materials and methods
The mineralogical study was carried out on 40 sections (each one ca.
1 m in length) of drillcores from the Jabali drill holes J109, J125 and J138
(Table 1) (Mondillo et al., 2011). The sampled core intervals are on the
drill hole logs shown in Fig. 5. The selected holes were drilled by ZincOx
Resources in 2004 in the Jabal Barrik area, near the center of the planned
pit (Fig. 2). Several mineralized outcrop samples (49) (Table 2; Fig. 2)
were analyzed as well.
Core and outcrop samples have been studied using petrographic and
cathodoluminescence (CL) microscopy. Cathodoluminescence measurements were done on carbon-coated, polished thin sections using a
hot cathode cathodoluminescence microscope HC1-LM (cf., Neuser
et al., 1995), at TU Bergakademie Freiberg, Germany. The system was
operated at 14 kV accelerating voltage and a current of 0.2 mA (current
density of about 10 μA/mm2). Luminescence images were captured
on-line during cathodoluminescence operations using a peltier cooled
digital video camera (OLYMPUS DP72). Cathodoluminescence spectra
in the wavelength range 380 to 1000 nm were recorded with an
Acton Research SP-2356 digital triple-grating spectrograph with a
Princeton Spec-10 CCD detector that was attached to the cathodoluminescence microscope by a silica-glass fiber guide. Cathodoluminescence spectra were measured under standardized conditions
(wavelength calibration by a Hg-halogen lamp, spot width 30 μm,
measuring time 2 s).
Scanning electron microscopy (SEM) examination was carried out
using a Jeol JSM 5310 instrument at the University of Napoli Federico
II (CISAG). Element mapping in backscattered electron mode (BSE)
and qualitative and quantitative analyses by energy-dispersive spectrometry (EDS) were obtained with the INCA X-stream pulse processor
Table 1
Jabali drillcore samples. Drillcore location from Fig. 2.
Drillcore n.
From (m)
To (m)
ZincOx assay n.
Sample n.
J109
57.3
58.3
59.3
60.3
61.65
50.78
51.78
53.1
54.73
55.73
57.92
59.45
60.97
62
64
65
66
67
68
69
74.5
75.5
76.5
77.5
78.5
79.5
84.5
85.5
86.5
87.5
88.5
89.5
90.5
68
69
70
71
72
73
74
58.3
59.3
60.3
61.65
62.7
51.78
53.1
54.73
55.73
56.73
59.45
60.97
62
64
65
66
67
68
69
70
75.5
76.5
77.5
78.5
79.5
80.5
85.5
86.5
87.5
88.5
89.5
90.5
91.5
69
70
71
72
73
74
75
1
2
3
4
5
1
2
3
4
5
6
7
8
9
10
11
12
13
14
15
19
20
21
22
23
24
29
30
31
32
33
34
35
4
5
6
7
8
9
10
J109-1
J109-2
J109-3
J109-4
J109-5
J125-1
J125-2
J125-3
J125-4
J125-5
J125-6
J125-7
J125-8
J125-9
J125-10
J125-11
J125-12
J125-13
J125-14
J125-15
J125-19
J125-20
J125-21
J125-22
J125-23
J125-24
J125-29
J125-30
J125-31
J125-32
J125-33
J125-34
J125-35
J138-4
J138-5
J138-6
J138-7
J138-8
J138-9
J138-10
J125
J138
253
Fig. 5. Stratigraphic logs of selected drill holes. The sampled intervals are indicated.
and the 4.08 version Inca software (Oxford Instruments detector),
interfaced with the JEOL JSM 5310. The following reference standards
were used: albite (Si, Al, Na), orthoclase (K), wollastonite (Ca), diopside
(Mg), almandine (Fe), rutile (Ti), barite (Ba), strontianite (Sr), Cr2O3
(Cr), rhodonite (Mn), sulfur (pyrite), sphalerite (Zn), galena (Pb), fluorite (F), apatite (P), sylvite (Cl), Smithsonian phosphates (La, Ce, Nd, Sm,
Y), pure vanadium (V) and Corning glass (Th and U). Analytical errors
are 1% relative for major elements and 3% relative for minor elements.
X-ray diffraction analysis was carried out on all samples. The core
sections and the outcrop samples have been crushed to 1 mm, milled
and homogenized. X-ray diffraction analyses have been carried out
with a Philips PW 3020 automated diffractometer at the University of
Heidelberg, with CuKα radiation, 40 kV and 30 mA, 10 s per step and
a step scan of 0.02° 2θ. The data were collected from 3 to 110° 2θ.
X-ray diffraction quantitative phase analysis of the core samples
(Mondillo et al., 2011) was performed on the X-ray diffraction patterns
using the Rietveld method (Bish and Howard, 1988; Bish and Post,
1993; Hill, 1991; Rietveld, 1969). X-ray powder diffraction data were
analyzed using the GSAS package (General Structure Analysis System,
Larson and Von Dreele, 2000) and its graphical interface EXPGUI
(Toby, 2001). In this paper we report only the mean values of the data
previously published by Mondillo et al. (2011), calculated for each
mineralized interval occurring in the drillcores. The mean values
were calculated weighting the data relatively to the length of the samples, and to the total length of the mineralized intervals, using the formula:
Mx ¼ ð∑I Wxi Li Þ=∑I Li
where Mx is the mean amount of a mineral x in a specific mineralized
interval I, Wxi is the amount of mineral x in the sample i, and Li is
the length of the core sample i; the sums ∑I have been obtained considering all the samples i of the interval I.
X-ray diffraction qualitative analyses on outcrop samples were
performed with the PANalytical B.V. software HighScore version 3.0e
(Osservatorio Vesuviano-INGV) and JCPDS PDF-2 database. For quantitative analysis of X-ray diffraction patterns we used the HighScore
254
Table 2
Jabali outcrop samples. Sampling area n. from Fig. 2.
Sampling
area n.
Sample n.
1A
JS MON 1
JS MON 2
Sample description
Yellowish-brown concretionary smithsonite
Weathered red-brown dolomite with sulfide
remnants
JS MON 3
Smithsonite ore encrusted by hydrozincite
JS MON 4
Brownish dolomite, partly dedolomitized by calcite
JS MON 5
Porous smithsonite encrusted by hydrozincite
JS MAR 1
Massive smithsonite with galena remnants
JS MAR 2
Massive smithsonite with macrocrystalline galena
remnants
1B
JS MON 6
Reddish-brown massive ore
JS MON 7
Brown massive ore
JS MON 8
Gossan; (hydr)oxide black-red crusts
JS MON 9
Reddish-brown massive ore; few smithsonite
crusts in voids
JS MON 10
Reddish-brown massive ore; few smithsonite
crusts in voids
JS MON 11
Clay sample from gossan
JS MON 12
Gossan sample
JS MON 13
Smithsonite–dolomite transition
JS MON 36
Dedolomitized dolomite
JS MAR 3
Massive smithsonite with hydrozincite coatings
2
JS MON 14
Smithsonite breccia encrusted by hydrozincite and
hemimorphite
JS MON 15
Smithsonite breccia encrusted by hydrozincite and
hemimorphite
JS MON 16
Gossan and smithsonite ore
JS MON 17
Gossan sample
JS MAR 5
Rusty dolomite with galena spots
JS MAR 6
Smithsonite breccia with hydrozincite encrustations
3
JS MON 18
Gossan sample
JS MON 19
Silicified gossan sample with calcite spots
JS MON 20
Smithsonite–hydrozincite massive ore
4
JS MON 21
JS MON 22A Smithsonite–hydrozincite massive ore
JS MON 22B Smithsonite–hydrozincite massive ore
JS MON 23
Massive dolomite
JS MON 24
Smithsonite–hydrozincite ore
JS MON 25
Gossan sample
JS MON 26
JS MON 27
JS MON 28
5
JS MON 29
Smithsonite–hydrozincite massive ore;
hemimorphite spots
JS MON 30
Gossan sample with hydrozincite
JS MON 31
Gossan sample
6
JS MON 32_1 Smithsonite–hydrozincite massive ore;
hemimorphite spots
JS MON 32_2 Smithsonite–hydrozincite massive ore;
hemimorphite spots
JS MON 33
Massive dolomite, partly dedolomitized,
with galena veins
JS MON 34
Trachyte dyke
JS MON 35
Trachyte dyke
7
JS MAR 7
Gossan sample
JS MAR 8
Massive smithsonite with hydrozincite spots
8
JS MAR 4
Massive smithsonite with hydrozincite spots
Jabal Salab area JS MON 37
Barite vein
JS MON 38
Limestone with bioherms of Unit 7
JS MON 39
Gossan sample
Plus software with Rietveld structural models based on the American
Mineralogical Crystal Structure Database (AMCSD).
Whole rock chemical analyses of major and minor elements for the
core samples were carried out by OMAC Laboratories Ltd. (Co. Galway,
Ireland). Diamond drillcore was cut and the entire half-core samples homogenized and powdered to obtain 30 g aliquots for chemical analysis.
After aqua regia digestion, the samples were analyzed by multi-element
inductively-coupled plasma mass spectrometry (ICP-MS). Samples with
Zn contents N 9% have been also analyzed by atomic absorption
spectrometry (AAS), with an excellent agreement between the two
data sets (SRK Consulting, 2005). As for the previously mentioned Xray diffraction analyses, we here only report the mean values of chemical assay data from Mondillo et al. (2011), calculated for each mineralized interval recognized in the drillcores. The mean values were
calculated weighting the data relatively to the length of the samples
and to the total length of the mineralized intervals, using the formula:
Ex ¼ ð∑I Axi Li Þ=∑I Li
where Ex is the mean amount of element x in a specific mineralized
interval I, Axi is the amount of element x in the sample i, and Li is the
length of the core sample i; the sums ∑I have been carried out considering all the samples i of the interval I.
Whole rock chemical analyses of the outcrop samples were performed by ACME Laboratories (Vancouver) to determine the major
and minor elements, using aqua regia digestion and multi-element
inductively-coupled plasma energy and mass spectrometry (ICP-ES/
ICP-MS). Loss on ignition was not evaluated.
Stable carbon and oxygen isotope analyses were carried out at the
University of Erlangen-Nürnberg (Germany). Carbonate powders and
picked minerals were reacted with phosphoric acid at 70 °C using a
GasBench II connected to a Thermo Finnigan Five Plus mass spectrometer. All values are reported in per mil relative to V-PDB by assigning a
δ13C value of +1.95‰ and a δ18O value of −2.20‰ to NBS19. Reproducibility was checked by replicate analysis of laboratory standards and
was better than ± 0.07‰ (1σ) for both carbon and oxygen isotope
analyses. Oxygen isotope values of dolomite and smithsonite were
corrected using the phosphoric acid fractionation factors given by Kim
et al. (2007), Rosenbaum and Sheppard (1986) and Gilg et al. (2008).
Sulfur isotope analyses of a few sulfide and sulfate samples were
carried out at Actlabs (Ancaster, Ontario, Canada). Single crystals and
mineral fragments for analysis were collected with a dental drill from
core and outcrop samples. Pure BaSO4 and sulfide samples were
combusted to SO2 gas under ~ 10−3 tor vacuum. The SO2 was transferred directly from the vacuum line to the ion source of a VG 602
Isotope Ratio Mass Spectrometer (Ueda and Krouse, 1986). Quantitative
combustion to SO2 was achieved by mixing 5 mg of sample with 100 mg
of a V2O5/SiO2 mixture (1:1). The reaction was carried out at 950 °C for
7 min in a quartz glass tube. Pure copper turnings were used as a
catalyst to ensure conversion of SO3 to SO2. Internal Lab Standards
(SeaWaterBaSO4 and FisherBaSO4) were run at the beginning and at end
of each set of samples and were used to normalize the data as well as
to correct for instrumental drift. All results are reported in the permil
(‰) notation relative to the international VCDT standard. Precision
and reproducibility (1 std.dev.) based on replicate laboratory standard
analyses (n = 10) was typically better than ± 0.2‰ (1σ) (www.
actalabs.com).
5. Results
5.1. Mineralogy of the nonsulfide orebody
The mineralogical results obtained with X-ray diffraction of the core
samples set, drilled in different areas of the Jabali minesite, have been
published in full by Mondillo et al. (2011). The mean values, calculated
from the same cores are reported in Table 3.
The interval 57.3–62.7 m from core J109, and the interval 68–75 m
from core J138 correspond to the upper zone of the orebody. Dolomite
is the predominant lithology of both cores, while smithsonite has been
detected in an interval 1–2 m-thick (Fig. 5) with an average grade of
6–7 wt.%.
Three mineralized intervals have been sampled from core J125
(Fig. 5): the uppermost interval, immediately below the shales of the
Madbi Formation (Unit 8), is comprised between 50.78 and 80 m, the
middle interval can be traced between 74.5 and 80.5 m and the deepest
255
Table 3
Mineral abundances in the core intervals from X-ray quantitative phase analysis (Mondillo et al., 2011 modified).
Drillcore
Interval
Mineralized zone
Dol
Cal
Sm
Cer
Gp
Ang
Sp
Gn
Cha
Hem
Gth
Kln
Sau
Ilt
Qz
7.4
0.0
40.0
2.4
6.0
7.6
7.3
27.2
10.3
24.0
6.4
18.7
0.1
0.5
–
0.2
0.6
0.3
–
8.4
0.3
–
–
3.7
–
1.0
–
–
–
0.4
0.1
1.2
–
–
0.1
0.5
0.1
0.1
–
–
0.8
0.2
0.1
0.1
–
–
–
0.1
–
0.3
–
0.2
0.2
0.2
1.8
1.6
4.2
3.6
3.0
2.5
0.9
0.5
7.7
0.5
0.3
1.5
0.1
–
0.1
–
–
–
–
–
0.1
–
–
–
–
–
0.1
–
–
–
(wt.%)a
(m)
J109
J125
57.30–62.70
Upper zone
50.78–70.00
Upper zone
74.50–80.50
Middle zone
84.50–91.50
Lower zone
J138
68.00–75.00
Upper zone
Mean value on the entire sample batch
82.1
52.8
37.4
69.2
82.5
61.5
Notes: “–” not detected. Mineral abbreviations mostly after Whitney and Evans (2010). Dol; dolomite; Cal, calcite; Sm, smithsonite; Cer, cerussite; Gp, gypsum; Ang, anglesite;
Sp, sphalerite; Gn, galena; Cha, chalcophanite; Hem, hematite; Gth, goethite; Kln, kaolinite; Sau, sauconite; Ilt, illite; Qz, quartz.
Values derive from a weighted mean calculation referred to the length of core samples and the total length of the mineralized intervals (details in text).
Statistical indicators ranges: Rp 5.11–6.40%, wRp 6.65–8.85%.
one reaches from 84.5 down to 91.5 m. Also in core J125 dolomite is
generally the most abundant mineral phase, though being totally absent
in the best mineralized samples. Smithsonite is abundant in the upper
and lower sections of the core, with average grades between 24 and
27 wt.%. In the middle section of the core, smithsonite has a mean
value of 10.3 wt.%, and is associated with calcite.
With an average amount of 18.7 wt.% calculated over the whole sample set, smithsonite represents the main ore mineral at Jabali. Sphalerite
and galena have very low amounts (below 2 wt.% on average) in all
mineralized intervals. Cerussite and anglesite were found only locally.
All core samples contain goethite, with mean values ranging
between 1.6 and 4.2 wt.%. Only the samples from the upper zone of
the J-125 core, directly below the boundary with Unit 8, contain gypsum
(~ 8.4 wt.%). Sauconite (Zn-smectite) and other clay minerals have a
very limited distribution, as well as Zn–Mn-hydroxides (chalcophanite).
Hydrozincite and hemimorphite are completely absent in drillcores.
X-ray diffraction quantitative mineralogical analyses (Table 4) were
conducted on 17 outcrop samples collected in the area of the Jabali mine
site, where both high-grade ore and gossanous rocks occur. The most
abundant ore mineral in the outcrop samples is again smithsonite,
with a variable percentage. In some samples, smithsonite amounts to almost 100 wt.% of the rock total (JS MON 9, JS MON 10). Hydrozincite, detected only in outcrop samples, can locally reach concentrations around
50–70 wt.%. Hemimorphite is slightly less abundant. Cerussite and galena have been detected only in two specimens, sampled along a fault
zone. Dolomite can be very abundant, and calcite may occur with
maximum amounts of 15 wt.%. Quartz is abundant only locally (JS
MON 19). Goethite is ubiquitous: together with hematite it is the
main “oxide mineral” in the gossanous samples (for example in JS
MON 8). Several sulfate minerals have been identified in the gossan:
plumbojarosite, jarosite, dietrichite, and rozenite.
All minerals detected at Jabali are listed in Table 5.
5.2. Petrography of the nonsulfide orebody
In the previous published papers on the Jabali deposit (Al Ganad
et al., 1994; Mondillo et al., 2011) the petrography of the nonsulfide
mineral association was dealt in a limited way. Here we report the results of a comprehensive study on the secondary mineral assemblage
that documents several characteristics of the deposit never highlighted
before. The classic thin section microscopic observation has been integrated with cathodoluminescence and scanning electron microscopy
(Figs. 6–9). Cathodoluminescence emission spectra of selected minerals
are shown in Fig. 10.
The Jabali host rock consists of two main generations of dolomite: an
early diagenetic dolomite, which corresponds to the D1 and D2 phases
of Al Ganad et al. (1994), and a saddle (hydrothermal) dolomite (D3
and D4 phases of Al Ganad et al., 1994). Saddle dolomite replaces the
previous dolomite phase and occurs as well in veins (Fig. 6A). This
hydrothermal dolomite contains variable amounts of Mn (up to 2 wt.%
MnO) and iron (up to 6 wt.% FeO).
Table 4
Mineral abundances (wt.%) from X-ray diffraction quantitative analysis of outcrop samples.
JS MON 3
JS MON 6
JS MON 8
JS MON 9
JS MON 10
JS MON 13
JS MON 14
JS MON 15
JS MON 18
JS MON 19
JS MON 21
JS MON 22A
JS MON 22B
JS MON 28
JS MON 29
JS MON 32_2
JS MON 33
Sm
Hz
Hm
Cer
Gn
Dol
Cal
Qz
Gth
Hem
Prls
Sau
Dtrch
Pb-jrs
Jrs
Roz
Gp
64.0
96.5
–
97.3
97.5
54.9
36.4
25.3
20.5
–
33.6
69.3
0.9
39.4
19.5
8.4
–
14.6
–
–
–
–
–
21.2
9.3
–
–
33.9
28.8
73.9
50.2
75.5
17.8
–
–
–
–
–
–
–
2.2
3.6
–
–
–
–
7.6
0.4
2.8
4.4
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
0.1
0.5
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
1.0
20.2
–
–
–
–
43.2
38.7
44.7
60.6
–
–
–
–
–
–
67.6
97.5
0.2
–
–
–
–
–
0.5
1.3
15.0
0.8
–
–
1.1
–
1.1
0.7
0.8
–
–
–
–
–
–
–
–
–
55.9
–
–
–
–
–
–
–
0.9
3.2
73.2
1.6
1.6
1.7
0.9
15.4
0.4
30.0
28.3
1.9
16.5
9.8
1.0
0.6
0.1
–
–
26.8
1.1
0.9
–
–
–
2.7
13.2
3.8
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
0.1
–
–
–
–
–
–
0.1
0.3
–
–
–
0.2
0.1
0.4
0.8
–
0.3
–
–
0.2
0.1
0.1
0.1
–
–
–
–
–
–
–
–
–
0.1
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
tr
–
–
–
tr
tr
–
–
–
–
–
–
–
–
–
–
–
tr
–
–
–
–
–
–
–
–
–
–
–
–
–
–
tr
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
0.3
–
Notes: “–” not found, “tr” traces. Mineral abbreviations mostly after Whitney and Evans (2010): Sm, smithsonite; Hz, hydrozincite; Hm, hemimorphite; Cer, cerussite; Gn, galena;
Dol, dolomite; Cal, calcite; Qz, quartz; Gth, goethite; Hem, hematite; Prls, pyrolusite; Sau, sauconite; Dtrch, dietrichite; Pb-jrs, plumbojarosite; Jrs, jarosite; Roz, rozenite; Gp, gypsum.
Statistical indicators ranges: Rp 5.67–8.40%, wRp 8.25–9.40%.
256
Table 5
List of minerals detected at Jabali.
Name
Chemical formula (International Mineralogical Association)
Acanthite
Anglesite
Calcite
Cerussite
Chalcophanite
Covellite
Dietrichite
Dolomite
Galena
Goethite
Greenockite
Gypsum
Hematite
Hemimorphite
Hydrozincite
Illite
Jarosite
Kaolinite
Plumbojarosite
Pyrite/marcasite
Pyrolusite
Pyromorphite
Quartz
Rozenite
Sauconite
Smithsonite
Sphalerite
Ag2S
PbSO4
CaCO3
PbCO3
ZnMn4+ 3O7·3H2O
CuS
(Zn,Fe2+,Mn2+)Al2(SO4)4·22H2O
CaMg(CO3)2
PbS
FeO(OH)
CdS
CaSO4·2H2O
Fe2O3
Zn4(Si2O7)(OH)2·H2O
Zn5(CO3)2(OH)6
(K,Ca,Na)(Mg,Al,Fe3+)2(Si,Al)4O10(OH)2·nH2O
KFe3+3(SO4)2(OH)6
Al2Si2O5(OH)4
Pb0.5Fe3+3(SO4)2(OH)6
FeS2
MnO2
Pb5(PO4)3Cl
SiO2
Fe2+SO4·4H2O
Na0.3Zn3(Si,Al)4O10(OH)2·4H2O
ZnCO3
ZnS
Early diagenetic dolomite is non-luminescent under CL. Saddle
dolomite has generally a bright red luminescence, but it can be locally
characterized by non-luminescent crystal cores and red luminescent
crystal rims (Fig. 6B). Non-luminescent crystal cores are generally associated with opaque Fe-hydroxides. Cathodoluminescence emission
spectra of the saddle dolomite at Jabali have a pronounced emission
maximum at 650–660 nm (Fig. 10A). As reported in the literature
(Götze, 2012), this indicates that luminescence is activated by Mn2+,
and that Mn2+ is hosted in the Mg-positions of the crystal structure.
In fact, when Mn2 + occurs in the Ca-positions, the emission spectra
show a peak (here absent) at 575 nm and yellow-orange CL colors.
The non-luminescence in the crystal cores of saddle dolomite is due to
their high Fe contents (measured by scanning electron microprobe
and testified by the occurrence of Fe-hydroxides), which quench the
Mn-related brightness. The emission spectra of the dark crystal cores
of the dolomite crystals have also a maximum at 650–660 nm, but it is
less pronounced and intense than in crystal rims (Fig. 10A).
All dolomite phases at Jabali show two different types of alteration
or replacement. The first is a dedolomitization (Bischoff et al., 1994;
Coniglio, 2003), with dolomite being replaced by calcite. Manganese
and Fe, previously contained in the dolomite structure, are precipitated
as oxides and hydroxides in the interstices of the crystals and in small
vugs and fissures (Fig. 6C, D, E). The newly formed calcite may contain
Pb (up to ~ 5 wt.% PbO) and Cd (up to ~ 2 wt.%). The second type of
replacement is most evident in wide gradational bands, which mark
the boundary between the host dolomite and replacive smithsonite.
Under the microscope, the altered dolomite crystals appear clear and
transparent along their margins, but are generally opaque and brown
colored towards the center (Fig. 6F, G). These brown colored phases
are mainly Zn-bearing (Figs. 7A, B, C), characterized by ZnO contents
reaching 17–22 wt.% (Mondillo et al., 2011). The ZnO content is inversely
correlated with that of MgO in the Zn-bearing dolomite (Zn2+ replaces
Mg2 + by up to 70%), while the calcium content corresponds to the
values generally measured in stoichiometric dolomites. Zinc-bearing
dolomite preserves limited amounts of Fe and Mn (below 1–2 wt.%)
from the saddle dolomite phase. Cadmium can be present locally (CdO
around 1.5 wt.%).
The Zn-bearing dolomite has a yellowish-orange CL color (Fig. 6H).
The luminescence emission spectra of this dolomite are characterized
by two overlapping peaks: the most prominent occurs between 585
and 600 nm, the second at 645–655 nm (Fig. 10B). No specific literature
exists on the cathodoluminescence behavior of Zn-bearing dolomite.
However, adapting to this case the concepts expressed by Götze
(2012), we assume that the yellowish-orange color is produced by the
preferred occurrence of Mn2 + in the Ca-position instead than in the
Mg-position of the crystal structure.
All earlier formed dolomites and Zn-bearing dolomite are replaced
by smithsonite (smithsonite 1) (Fig. 8A). The boundary between the
original dolomite and smithsonite can be very gradational; it generally
consists of thin concentric belts with a complex assemblage of smithsonite, Zn-bearing dolomite and remnants of original dolomite (preferentially hydrothermal dolomite). Smithsonite 1, together with newly
deposited gypsum, also replaces sphalerite (Fig. 8B). Replacive smithsonite has a very fine texture. Specifically, it consists of microcrystals,
which form agglomerates mimicking/following the original macrocrystalline rhombohedral/saddle habitus of the replaced dolomite
(Fig. 7D, E), and locally appear to fill dissolution cavities inside the dolomite macrocrystals (Fig. 7F). Smithsonite microcrystals assume a
rhombohedral shape (Fig. 9A) and occasionally can occur as tiny
euhedral crystals, showing combination of forms belonging to the
scalenohedral class, i.e. hexagonal dipyramids, rhombohedra,
scalenohedra (Fig. 9B). Smithsonite 1 microcrystals are often chemically
zoned, with Mg contents locally reaching 8–10 wt.% MgO (Fig. 7D), and
variable contents of Mn (up to ~3 wt.% MnO), Fe (up to 5–6 wt.% FeO),
and Ca (up to 6 wt.% CaO). Smithsonite 1 has a red CL color (Fig. 8C, D).
Cathodoluminescence emission spectra have a pronounced maximum
at 650–660 nm, and a broader, though less intense peak around
440 nm (Fig. 10C). In analogy with dolomite, the pronounced
peak indicates that cathodoluminescence is produced by variable
amounts of Mn2 + hosted in the smithsonite structure. The peak at
440 nm is instead related to the presence of crystal lattice defects in
the mineral.
A second type of smithsonite (smithsonite 2) consists of different
cement generations, occurring in vugs, cavities and veins as newly
formed crystals and concretions (Fig. 8E). Crystals are characterized by
a roughly rounded morphology and a size that generally increases
from the earliest generation to the latest. Concretions are prevailingly
botryoidal (Fig. 7G). At microscale, crystals and concretions show
variably developed rhombohedral faces. Smithsonite 2 is commonly
zoned. Generally, it is possible to recognize a first generation of crystals
and concretions with colors variable from brown, to orange and yellow
that precipitated directly over smithsonite 1. Other generations may
follow, generally white colored at first sight, and clear and transparent
under the microscope (Fig. 8E). A variation in color of smithsonite 2
commonly corresponds to a chemical zonation (Fig. 7H). Composition
of smithsonite 2 can vary from pure end-member Zn-carbonate to
terms characterized by over 20 wt.% MgO (substitution of Mg2 + for
Zn2 + of about 70% in the smithsonite structure), ~ 2 wt.% MnO,
~ 2 wt.% CdO, ~ 1.5 wt.% CaO, and ~ 1.5 wt.% PbO. The concretionary
phases of smithsonite 2 have different CL colors: the first phases are
pinkish-red, whereas the latest are blue (Fig. 8F). Luminescence
emission spectra of pinkish-red phases are similar to those of
smithsonite 1. Spectra of blue smithsonite 2 are slightly different and
characterized by two peaks, approximately having the same intensity,
in correspondence to the same wavelengths detected before (~440 nm
and ~ 650 nm) (Fig. 10C). This indicates that the Mn-related redluminescence is strongly quenched (low intensity of peak at 650 nm),
and the blue color is mostly activated by crystal lattice defects.
Smithsonite 2 is precipitated in veins either shortly before, or
together with variable amounts of gypsum, which cut replacive
smithsonite 1 (Fig. 8G). Specifically, gypsum forms fibrous, partially
deformed crystals perpendicular to the veins, suggesting crystal growth
contemporary to vein opening. In outcrop samples, smithsonite 2 is
257
Fig. 6. Thin section petrography: A) J125-2. Sphalerite associated with saddle dolomite; B) J125-2. Zoned saddle dolomite crystals under CL; C) J138-10. Dedolomitization with
dolomite remnants (NII); D) the same as in panel C, under CL; E) J125-19. Dedolomitization with dolomite remnants (NII); F) J125-15. Zn-dolomite replacing dolomite in the crystal
cores; G) J125-15. Zn-dolomite replacing hydrothermal dolomite in the crystal cores; smithsonite concretions at the rims (NII); H) Zn-dolomite (crystal cores) with yellow-orange CL
colors; hydrothermal dolomite in red; smithsonite crusts dark blue. Cal = calcite; Dol = dolomite; Sm 2 = smithsonite 2; Sp = sphalerite; Zn-Dol = Zn-dolomite.
intergrown with clays, Fe-hydroxides and oxides, forming reddish
concretions.
Hydrozincite and hemimorphite have been detected solely in
outcrop samples. Hydrozincite generally shows a fine-grained texture
and acicular crystal habitus; it may occur as smithsonite replacement,
as vein and porosity filling, and in crusts. Petrographic investigation of
samples collected from the mineralized outcrops shows that hydrozincite
occurs as surface coatings, in vugs, as well as filaments replacing smithsonite. Hydrozincite has a blue CL color, and the emission spectra are
dominated by the effects related to the defects of crystal lattice.
Hemimorphite has been detected as cm-sized fan agglomerates of
tabular crystals throughout the deposit and as small crystals in vugs.
Several small veins with crystalline hemimorphite, locally associated
with Fe-hydroxides, are quite widespread.
Lead secondary minerals (cerussite and anglesite) strictly replace
primary galena, and not the host rock (Fig. 11A). Cerussite is quite
pure chemically and shows a green CL color (Fig. 8H).
In the Jabali samples we were able to recognize not only the remnants of the primary sulfide association (sphalerite, galena and pyrite/
marcasite), but also a few secondary sulfides related to the nonsulfide
mineral assemblage. A secondary galena phase precipitated between
the gypsum veins and smithsonite 2 concretions. This secondary galena
may be accompanied by greenockite, and by small crystals of pyromorphite. At the boundary between primary sphalerite and replacing
258
Fig. 7. Backscattered (BSE) images: A) J125-31. Zn-dolomite, patchily replacing dolomite crystals; B) J109-5. Zn-dolomite pervasively replacing the rim of a dolomite crystal; C) J125-15.
Dolomite crystal, partly replaced by Zn-dolomite, at the boundary Fe-hydroxides and smithsonite 2; D) JS MON 10. Smithsonite 1 mimicking the rhombohedral habitus of a dolomite
crystal; E) J125-7. Smithsonite 1, locally cut and replaced by gypsum veins; F) JS MON 2. Zoned smithsonite 1 and Zn-dolomite replacing dolomite; G) J125-15. Concretions of smithsonite
2 at the border of smithsonite 1 replacing Zn-dolomite; H) JS MON 2. Zoned concretions of Mg-bearing smithsonite 2. Dol = dolomite; Fe–Mn-Dol = Fe–Mn-dolomite; Gp = gypsum; Sm
1 = smithsonite 1; Sm 2 = smithsonite 2; Mg–Sm 1 = Mg-bearing smithsonite 1; Mg–Sm 2 = Mg-bearing smithsonite 2; Zn-Dol = Zn-dolomite.
smithsonite, secondary greenockite, galena and covellite were identified (Fig. 11B).
The Ag-sulfide acanthite is always associated at Jabali with the
secondary ores. It occurs together with concretionary smithsonite
(Fig. 11C), and also in association with gypsum (Fig. 11D) and greenockite. Hemimorphite containing small particles of Ag-sulfide has been also
identified in a gypsum vein.
Zn-clays are not very common but sauconite has been detected in
the porosity of the host rock, locally associated with kaolinite and illite.
The Fe-hydroxides do not have a typical goethite composition:
they contain Zn (up to 12 wt.% ZnO), Pb (up to 7 wt.% PbO) and
SiO 2 (up to 6 wt.%). Also most Mn-hydroxides consist not only of
chalcophanite (which contains only Mn and Zn), but of possibly
amorphous phases containing Mn, Pb and Fe in variable proportions (i.e. PbO ~ 20–30 wt.%, FeO ~ 10 wt.%).
5.3. Major and minor element geochemistry
Whole rock chemical analyses by ICP-MS of the samples from the
Jabali drillcores (ZincOx plc.) have been published in Mondillo et al.
(2011); a short resume, from which valuable information on the average geochemistry of the Jabali ores can be inferred, is shown in Table 6.
When correlating the element concentrations of the three drillcores
considered, the upper section of the J125 core is particularly enriched in
Zn compared with the same section of both J109 and J138 cores. The Zn
grade of the upper section of core J125 (15.2 wt.%) is two times higher
259
Fig. 8. Thin section petrography: A) J125-2. Smithsonite 1 replacing a saddle dolomite crystal (NII); B) J125-5. Smithsonite 1 replacing dolomite and sphalerite (NII); C) J125-7. Replacive
smithsonite 1 mimicking dolomite (NII); D) the same as in panel C under CL: smithsonite shows strong red colors; E) J125-15. Zoned concretions of smithsonite 2 at the border of
Zn-dolomite (NII); F) the same as in panel E under CL; G) J125-3. Smithsonite 1 cut by veins bordered by concretionary smithsonite 2, and then filled by gypsum (NII); H) J125-7.
Cerussite directly replacing galena and host dolomite replaced by smithsonite 1 under CL: cerussite is green. Cer = cerussite; Gn = galena; Gp = gypsum; Sm 1 = smithsonite 1; Sm
2 = smithsonite 2: Zn-Dol = Zn-dolomite.
Fig. 9. Secondary electron microscopy micrographs: A) JS MON 3. Smithsonite microcrystals with rhombohedral shape; B) J125-30. Smithsonite crystals, showing a combination of forms
belonging to the scalenohedral class, i.e. hexagonal dipyramids, rhombohedra, scalenohedra.
260
A
5000
intensity (counts)
4000
3000
saddle
dolomite crystal rim
saddle
dolomite crystal core
2000
1000
0
350
500
650
800
wavelength (nm)
B
14000
Zn-dolomite
sulfides and gypsum) have been measured in the topmost zone of the
J125 core.
Silver shows mean relevant values in the upper (141 ppm) and
lower (181 ppm) sections of the J125 drillcore. Cadmium concentration
is on average very high (880 ppm).
The chemical composition of 20 outcrop samples is given in Table 7.
The specimens from the gossan are marked by very high Fe (over
30 wt.%), and locally by anomalous vanadium values. High sulfur
contents are associated with the samples containing supergene sulfates and remnant sulfides. The Zn content can be very high (over
50 wt.% locally) in some outcrop samples collected from the
massive ore body. The highest Zn concentrations (e.g. JS MON 22A,
JS MON 29) are associated with smithsonite, or with a smithsonite–
hydrozincite–hemimorphite assemblage. Samples characterized by
discrete Pb contents (~5–6 wt.% Pb) are instead low in Zn. In outcrop samples, Cd generally occurs in higher amounts than Ag, with values locally
higher than 3000 ppm. Copper can reach values higher than 500 ppm in
a few mineralized samples. A strategic element like Ga reaches an anomalous content in the sample JS MON 22B (68 ppm), but it is not significant in
most other samples.
The environmentally dangerous Tl is not abundant at Jabali, with
maximum concentrations of 50 ppm measured only in one sample.
intensity (counts)
12000
10000
5.4. Carbon, oxygen and sulfur isotopes
8000
6000
4000
2000
0
350
500
650
800
wavelength (nm)
C
20000
intensity (counts)
16000
smithsonite 1
smithsonite 2
12000
8000
4000
0
350
500
650
800
wavelength (nm)
Fig. 10. Cathodoluminescence emission spectra of various Jabali carbonates: A) dolomite;
B) Zn-bearing dolomite; C) smithsonite.
than in the same section of the other two drillcores (averaging
both 7.3 wt.%). The middle section of the J125 drillcore is the less Znrich (mean Zn grade around 10 wt.%), whereas the lowermost section is
the richest (mean Zn grade around 18 wt.%). The Fe contents are similar
in the mineralized sections of all cores, whereas significant Pb amounts
are detected only in the upper zones of J125 and J138. Sulfur
concentrations averaging around 2–3 wt.% (associated to both unaltered
In order to characterize the fluid that precipitated the secondary
mineralization, we conducted stable isotope (C, O) analyses of the
nonsulfide Zn-minerals and host rock. A few S-isotope analyses were
also carried out, to identify the origin of gypsum detected throughout
the cores.
Carbon and oxygen isotope analyses were conducted on Jurassic
limestone samples (bioherms from Unit 7), saddle dolomite (associated
with sulfide mineralization), Zn-dolomite (drillcores), smithsonite
(drillcores and outcrop), and hydrozincite (outcrop) (Table 8, Fig. 12).
Two Jurassic limestone samples from Unit 7 gave different results: in
a coral sample a δ13C value of −1.0‰ and a δ18O value of −6.8‰ have
been measured, while the carbonate matrix between the bioherms
shows δ13C and δ18O values of − 2.3 and − 7.8‰, respectively. The
saddle dolomite shows δ13C and δ18O values between −1.1 and 0.2‰
and −10.7 and −9.1‰, respectively. Zinc-bearing dolomite is characterized by δ13C values ranging from − 1.8 to − 1.0‰. δ18O values are
between −11.9 and −10.4‰.
Smithsonite specimens from drillcore and outcrop samples have
been analyzed, excluding Zn-carbonate strongly intergrown with
gypsum. In some cases, it was possible to sample replacive smithsonite
1 and clear concretionary smithsonite 2 from the same sample. However,
the main difference in the isotope ratios is related to their provenance
from the upper, middle, or lower zones. Smithsonite from the upper
zone is characterized by δ13C values between − 4.7 and − 2.9‰, and
by δ18O values between − 11.4 and − 10.5‰. One sample from the
middle zone has a carbon and oxygen isotope composition of − 4.6‰
and − 9.6‰, respectively. The δ13C and δ18O values of smithsonite
from the lower zone are between − 5.7 and − 4.7‰, and between
− 10.0 and − 9.2‰, respectively. A smithsonite from the outcrop
sample JS MON 22A has a δ13C value of − 6.1‰ and a δ18O value of
−8.8‰, being comparable with the values for smithsonite from subsurface samples.
Hydrozincite has carbon and oxygen isotope ratios between −7.4
and −6.6‰ and −4.9 and −0.8‰, respectively.
Sulfur isotope analyses have been carried out on sphalerite, galena
and gypsum (Table 9). Two sphalerite samples, from interval J125-2 to
J125-5 have sulfur isotope ratios of 7.1 and 6.5‰. Galena, sampled
from an outcrop vein shows a composition of − 3.6‰. Gypsum from
sample J125-3 has a composition of 7.3‰, which is very similar to the
δ34S value of sphalerite.
261
Fig. 11. Backscattered (BSE) images: A) J125-7. Anglesite replacing primary galena; B) J125-3. Secondary sulfides at the border of altered sphalerite: secondary galena, greenockite,
covellite (locally containing Ag); C) J125-32. Ag-sulfide and smithsonite concretions; D) J125-3. Ag-sulfides between smithsonite and gypsum. Ang = anglesite; Cv = covellite; Gn =
galena; Grn = greenockite; Gp = gypsum; Sm = smithsonite; Sp = sphalerite.
6. Discussion
6.1. Relationship between host rock and mineralization
The Jabali secondary deposit was formed through the replacement
process of the dolomite host rock by Zn-carbonates (Zn-bearing dolomite and smithsonite). The original lithotype was a carbonate platform
limestone, which underwent strong dolomitization. Differently from Al
Ganad et al. (1994), we could distinguish only two (not four) dolomite
types: an early diagenetic dolomite and a hydrothermal dolomite. From
our observation it becomes clear that original sulfides (sphalerite, galena,
pyrite/marcasite) were strictly related to hydrothermal dolomite. This
primary mineral association (hydrothermal dolomite and sulfides) was
then widely altered and replaced by secondary phases.
During alteration, dolomite was firstly replaced by calcite, following
a typical de-dolomitization process:
Ca
2þ
ðaqÞ
2þ
þ CaMgðCO3 Þ2ðsÞ → 2CaCO3ðsÞ þ Mg
dolomite
ðaqÞ
:
calcite
In this reaction, Ca2+ is usually supplied by the dissolution of a mineral source, like gypsum or anhydrite (Bischoff et al., 1994; Coniglio,
2003). In some studies (Raines and Dewers, 1997), a supergene nature
of dedolomitization is invoked, possibly related to post-burial surface
weathering, when meteoric waters infiltrate a fractured dolostone, or
during development of karstic processes. Coniglio (2003 and references
therein) evidenced that Fe-bearing dolomite is more susceptible to
dedolomitization than pure dolomite, because in the surface environment Fe2+ is easily altered to Fe3+, forming hydroxides, disrupting the
dolomite structure, and causing the precipitation of calcite. A possible
Ca2+ supply for the dedolomitization of Jabali dolomite, can derive
from the abundant gypsum intercalations occurring in the Madbi and
Sab'atayn Formations, which directly overlay the dedolomitized rock.
The prevailing hydrothermal nature of Jabali dolomite, and its ferromanganoan character facilitated the dedolomitization process.
A second and most important process that affected the dolomite
host rock consists of its partial replacement by Zn-bearing dolomite
phases. Zinc-bearing dolomite is always associated with remnants of
the original dolomite. It is quite common to observe the hydrothermal
dolomite crystals corroded and altered along their rims, along fractures
and vugs by Zn-bearing dolomite patches. Zn-bearing dolomite never
occurs as isolated, newly formed crystals in connection with dissolution.
Chemical analyses of Zn-bearing dolomite indicate that in this phase
the Zn-enrichment is proportional with Mg-decrease. Cathodoluminescence microscopy evidences that the luminescence of Znbearing dolomite is activated by Mn2+ substituting Ca2+ in the mineral
structure, whereas the CL colors of the hydrothermal dolomite associated
Table 6
Major and minor element concentrations of whole rock ICP chemical analyses (Mondillo et al., 2011 modified).
Drillcore
Interval
Mineralized zone
(m)
J109
J125
57.30–62.70
Upper zone
50.78–70.00
Upper zone
74.50–80.50
Middle zone
84.50–91.50
Lower zone
J138
68.00–75.00
Upper zone
Mean value on the entire sample batch
Zn
Fe
Mg
Pb
Ca
Mn
S
(wt.%)a
7.3
15.2
10.3
17.6
7.3
12.7
Ag
Cd
Cu
Ni
P
800
988
673
1164
537
880
34
79
25
45
39
54
10
9
23
16
18
13
96
193
400
343
257
243
(ppm)b
2.6
2.6
3.6
3.9
3.4
3.1
9.1
6.3
4.3
6.7
8.1
6.7
1.3
3.4
0.5
1.4
3.0
2.4
19.5
12.3
21.6
12.9
18.9
15.5
0.6
0.5
0.7
0.5
0.7
0.6
0.2
2.5
0.1
0.1
0.2
1.2
15
141
5
181
76
103
Notes: Values derive from a weighted mean calculation referred to the length of core samples and the total length of the mineralized intervals (details in text).
a
Detection limits (wt.%): Zn 0.0001, Fe 0.01, Mg 0.01, Pb 0.01, Ca 0.01, Mn 0.0001, S 0.05.
b
Detection limits (ppm): Ag 0.5, Cd 0.5, Cu 0.5, Ni 0.5, P 5.
262
Table 7
Major (wt.%) and minor (ppm) element concentrations of outcrop samples from ICP analysis.
Sample JS MON
n.
4
JS MON
6
JS MON
9
JS MON
12
JS MON
13
JS MON
14
JS MON
15
JS MON
17
JS MON
18
JS MON
19
JS MON
20
JS MON
21
JS MON
22A
JS MON
28
JS MON
29
JS MON
32_1
JS MON
32_2
JS MON
33
JS MON
36
41.96
12.79
0.12
1.24
1.16
0.18
0.01
0.16
0.01
b0.01
0.03
0.23
63.6
952.3
2.7
604.5
35.9
5.9
6
3.1
b0.5
29
0.9
11
1.6
5.7
14
1.06
2.8
b0.5
68
23
48.30
5.83
0.2
0.61
0.16
0.39
0
0.07
b0.01
b0.01
0.01
b0.05
13.8
2225
1.9
122.9
63.9
12
5
1.6
b0.5
7
b0.5
11
1.2
3
b5
1.41
1.6
1.1
b5
21
51.04
2.38
0.22
0.02
0.27
0.86
0.01
0.42
0.04
0.04
0.02
b0.05
21.9
603.2
3
4
53.2
12.1
b5
1.1
0.6
9
b0.5
11
9.8
3.6
158
0.22
2
3.5
b5
22
41.58
2.46
0.61
1.23
4.71
1.07
0.01
0.56
b0.01
b0.01
0.02
0.08
15.3
1224
6.2
100.7
11.3
8.2
21
4.6
0.9
26
1.2
14
4.8
7.7
9
0.25
2.8
b0.5
13
16
23.52
3.2
5.1
0.99
12.49
0.42
0
0.16
b0.01
b0.01
0.01
0.17
81.6
1639.5
2.6
86.3
6.4
2.9
6
3.5
b0.5
37
b0.5
b10
9.9
2.9
8
0.17
2.7
b0.5
26
10
1.39
2.43
8.65
7.95
19.87
0.61
b0.001
0.06
b0.01
b0.01
0.01
0.82
350.7
73.8
1.8
7.6
4.5
1.6
5
1.6
b0.5
126
b0.5
13
20
4.1
30
0.37
0.6
0.5
b5
b2
0.13
1.79
2.52
0.00
33.84
0.64
0
0.03
b0.01
b0.01
0.01
b0.05
0.9
b0.5
b0.5
1.3
b0.5
b0.5
b5
0.7
b0.5
72
b0.5
b10
3.3
1.1
9
b0.05
b0.5
b0.5
b5
b2
Zn
3.42
38.29
42.05
2.96
29.26
36.37
19.70
1.15
2.47
1.30
23.37
34.57
47.20
Fe
2.32
13.72
8.54
18.69
12.59
2.73
14.51
34.56
32.3
24.6
35.49
19.81
5.3
Mg
6.63
0.36
0.71
0.16
2.5
2.04
1.91
0.1
0.12
0.18
0.08
0.13
0.22
Pb
0.41
0.71
0.56
7.42
1.06
0.61
1.12
1.17
5.70
0.11
0.55
1.14
0.64
Ca
24.79
0.10
0.09
0.24
3.81
6.41
11.43
11.6
6.72
8.55
0.11
0.15
0.24
Mn
0.66
0.51
0.46
0.05
0.78
0.70
0.49
0.01
0.24
0.28
0.08
0.01
0.08
Ti
0.002
0.01
0
0.04
0.01
0
0.01
0.05
0.01
0.03
0.01
0.01
0
Al
0.05
0.30
0.07
1.2
0.13
0.4
0.32
0.33
2
0.7
0.13
0.34
0.05
Na
0.01
b0.01
b0.01
0.06
0.02
b0.01
b0.01
0.03
0.08
0.18
0.07
0.01
b0.01
K
b0.01
b0.01
b0.01
0.67
0.02
b0.01
b0.01
0.16
0.08
0.02
0.04
b0.01
b0.01
P
0.03
0.02
0.02
0.03
0.02
0.02
0.02
0.06
0.08
0.11
0.03
0.03
0.01
S
b0.05
0.05
b0.05
5.79
b0.05
0.06
0.09
0.47
1.9
0.12
0.2
0.1
b0.05
Ag
9.4
3.4
5.1
37.9
9.6
44.4
58.4
3.9
119.6
3.4
3.6
10.9
108
Cd
319
1439.7
2307.5
22.1
3722.7
1358.2
715.3
546.9
38.5
69.3
185.9
1499.8
2952.6
Mo
1.4
3.6
2.9
3.3
4.4
5.3
6.4
29.8
5.8
8.3
7.5
2.2
0.9
Cu
13.8
49.1
30.4
76.8
25.9
14
9.2
30.7
60
378.9
7.3
79.2
308.1
Ni
9.5
128.3
152.6
23.4
240.6
38.7
43.5
12.5
52
74.4
34.8
19.5
37
Co
1.8
8.4
6.2
3.9
17.8
9.1
4.9
1.3
13.7
166.4
2.2
1.7
4.9
As
6
b5
b5
7
b5
b5
18
8
33
b5
23
34
b5
U
2.4
1.5
1.1
0.9
1.7
4.2
4
1.4
6.1
1.4
2.5
1.9
1.7
Th
b0.5
b0.5
b0.5
2.7
b0.5
b0.5
0.5
0.9
1.4
1.8
b0.5
b0.5
b0.5
Sr
55
b5
b5
117
5
26
134
41
160
96
9
8
6
b0.5
b0.5
b0.5
b0.5
b0.5
b0.5
b0.5
Sb
b0.5
b0.5
b0.5
b0.5
b0.5
b0.5
V
16
11
11
19
10
11
16
180
21
54
15
15
b10
La
5.7
4.9
2.5
10.3
5
3.2
3.7
1.9
14.6
15.7
2.4
1.1
0.6
Cr
9.6
5
3.8
22
8.4
3.6
4.7
48.7
20.9
47.2
6.3
6.4
2.1
Ba
41
b5
b5
400
11
22
21
56
80
75
b5
b5
b5
Hg
0.35
0.19
0.25
1.32
1.04
0.18
0.28
0.2
0.2
b0.05
0.36
0.4
0.44
Sc
1.2
1
1.2
1.9
0.9
2
2.9
2.5
7.1
3.3
4
1.3
1.1
Tl
1.1
0.5
0.7
52.8
5.2
13.2
4.1
b0.5
2.1
2.9
0.9
b0.5
b0.5
Ga
b5
b5
b5
b5
b5
b5
6
16
27
28
b5
10
22
Se
2
17
20
9
9
17
9
b2
b2
2
8
14
22
Detection limits (wt.%): Zn 0.01, Fe 0.01, Mg 0.01, Ca 0.01, Mn 0.0005, Ti 0.001, Al 0.01, Na 0.01, K 0.01, P 0.001, S 0.05.
Detection limits (ppm): Ag 0.5, Cd 0.5, Mo 0.5, Cu 0.5, Ni 0.5, Co 0.5, As 5, Th 0.5, Sr 5, Sb 0.5, V 10, La 0.5, Cr 0.5, Ba 5, Hg 0.05, Sc 0.5, Tl 0.5, Ga 5, Se 2.
JS MON
22B
263
Table 8
Carbon and oxygen isotopes of Jabali carbonates (outcrop and drillcore samples).
Sample n.
Mineral
Description
δ13C (‰ VPDB)
δ18O (‰ VPDB)
−6.83
−7.75
−9.03
−8.80
−0.79
−4.90
−4.82
δ18O (‰ VSMOW)
JS MON 38
JS MON 38
JS MON 3
JS MON 22A
JS MAR 4
JS MON 22A
JS MON 22A
Outcrop
Cal
Cal
Sm
Sm
Hz
Hz
Hz
Coral fossils from limestone
Calcite matrix between bioherms from limestone
Concretionary smithsonite
Hydrozincite crust
Hydrozincite crust
Hydrozincite crust
−0.97
−2.30
−2.61
−6.06
−6.75
−7.37
−6.62
Core sample n.
Depth (m)
Mineral
Description
δ13C (‰ VPDB)
δ18O (‰ VPDB)
δ18O (‰ VSMOW)
J125-6
J125-7
J125-15
J109-5
J125-11
J125-11
J125-15
J125-9
J125-10
J125-10
J125-14
J125-15
J125-21
J125-30
J125-30
J125-30A
J125-30B
J125-32
J125-32
J125-32A
J125-32B
59.2
60.8
69.2
62.6
65.2
65.8
69.8
63.7
64.1
64.2
68.7
69.1
77
85.6
85.65
85.7
85.7
87.6
87.7
87.75
87.75
Dol
Dol
Dol
Zn-Dol
Zn-Dol
Zn-Dol
Zn-Dol
Sm
Sm
Sm
Sm
Sm
Sm
Sm
Sm
Sm
Sm
Sm
Sm
Sm
Sm
Hydrothermal dolomite crystal
Dolomite crystal with Zn-dolomite core
Vacuolar smithsonite
Massive smithsonite
Smithsonite concretion
Massive smithsonite
Mixed massive-concretionary smithsonite
Smithsonite micro-concretions
Massive smithsonite
−1.10
−1.14
0.23
−1.85
−0.97
−1.17
−1.19
−3.75
−3.40
−4.16
−2.88
−4.68
−4.56
−5.48
−4.85
−5.22
−5.47
−4.83
−4.92
−5.71
−4.69
−9.36
−9.12
−10.70
−11.29
−11.95
−11.16
−10.42
−11.19
−10.83
−10.50
−11.41
−11.41
−9.57
−9.85
−9.99
−9.87
−9.93
−9.20
−9.49
−9.78
−9.85
21.21
21.46
19.83
19.22
18.54
19.36
20.11
19.33
19.70
20.03
19.10
19.09
20.99
20.70
20.56
20.69
20.62
21.38
21.08
20.78
20.71
23.82
22.87
21.55
21.79
30.04
25.81
25.89
Notes: Mineral abbreviations mostly after Whitney and Evans (2010). Cal = calcite; Dol = dolomite; Hz = hydrozincite; Sm = smithsonite; Zn-Dol = Zn-dolomite.
with primary sulfides were activated by Mn2 + occupying the Mgposition. This phenomenon suggests that during the secondary process
Zn2 + substitutes for Mg2 + in the dolomite structure, and that this
replacement causes a decrease of Mn in the same structural positions.
The occurrence of Fe- and Mn-hydroxides, associated with Zn-bearing
dolomite, indicates that the Zn2 + substitution for Mg2 + occurred at
the same time as the oxidation of Mn2+ and Fe2+ previously contained
in the dolomite structure.
In our opinion the Zn-rich waters, which caused the replacement of
host dolomite by Zn-bearing dolomite, were not related to the primary
hydrothermal fluids depositing sulfides, but were rather original
groundwaters that altered sphalerite and brought Zn2 + in solution
(Boni et al., 2011). Strictly hydrothermal Zn-bearing dolomites related
to the emplacement of sulfides were described in the Navan Zn–Pb
deposit (Ireland) by Kucha and Wieczorek (1984), at Broken Hill
(New South Wales) by Riaz Khan and Barber (1990), and in the
Silesia–Cracow district (Poland) by Coppola et al. (2009 and references
therein). In all these cases, Zn-bearing dolomite precipitated together
or before sulfides, as newly formed zoned crystals.
In the mining district of southwestern Sardinia (Italy), supergene
alteration and oxidation of sulfides and formation of Zn-oxidized
deposits (Boni et al., 2003) occurred after a regional scale hydrothermal
dolomitization event. In this district, a Zn-dolomite of strict supergene
origin, replacing the previous hydrothermal dolomite, was detected
with characteristics similar to those we have described from Jabali
Table 9
Sulfur isotopes of Jabali sulfides and sulfates.
Sample n.
Mineral
Description
δ34S (‰ VCDT)
JS MAR 2
J125-2
J125-5
J125-3
Gn
Sp
Sp
Gp
Macrocrystalline galena agglomerate
Sphalerite vein
Sphalerite vein
Gypsum vein within smithsonite
−3.58
7.13
6.47
7.34
Abbreviations: Gn = galena; Gp = gypsum; Sp = sphalerite.
(Boni et al., 2013). Another case of Zn-dolomite of supergene origin
was also identified by Boni et al. (2011) and Mondillo et al. (2014) in
the Yanque Zn–Pb deposit (Peru).
The economically most relevant replacement of the original
dolomite host rock resulted in the precipitation of the Zn-carbonate
smithsonite (smithsonite 1), consisting of rhombohedral microcrystals
in agglomerates mimicking the saddle shapes of hydrothermal dolomite
macrocrystals, and also filling dissolution cavities in the dolomite.
The red CL color indicates that smithsonite 1 inherited variable Mn
amounts from original dolomite. Smithsonite 2 forms crystals and
concretions occurring in veins and solution vugs of the host rock. The
blue CL color of smithsonite 2, activated by lattice defects, indicates a
poor crystalline state of this late phase. Gypsum, together with smithsonite 2 occurs in veins cutting smithsonite 1. These textural observations point to: 1) an increase in dolomite instability during the
alteration process, 2) a partial dissolution of the dolomite, followed by
smithsonite 1 precipitation, and 3) a late precipitation of smithsonite
2 and gypsum.
Dolomite is generally unstable in aqueous solutions characterized
by low pH, and goes into solution releasing Ca2 + and Mg2 + to
carbonate and bicarbonate anionic groups (Busenberg and Plummer,
1982; Morse and Arvidson, 2002; Pokrovsky et al., 1999, 2005, 2009).
After Sangameshwar and Barnes (1983), during the oxidation of a
sulfide body by groundwaters, the aqueous solutions can carry metals
(e.g. Zn2+) only if they have an acid pH; in this setting, the acidity of
waters is mainly produced by the oxidation of pyrite, which releases
sulfate anionic groups into solution. In the Jabali case, the acid Znbearing solutions altered and partially dissolved the dolomite of the
host rock, releasing Ca2 + and Mg2 +. The presence of the carbonate
anionic groups in solution together with Zn2+ cations favored firstly
the precipitation of smithsonite 1, followed by the precipitation of
smithsonite 2. Mg2 + was partially incorporated in the smithsonite
structure, whereas Ca2+ precipitated mainly in gypsum from the reaction with sulfate anionic groups in solution.
264
2.0
6.2. Fluids involved in the oxidation of sulfides and formation of secondary
mineralization
13C
(‰ VPDB)
0.0
-2.0
-4.0
-6.0
-8.0
-10.0
16.0
18.0
20.0
22.0
18O
24.0
26.0
28.0
30.0
32.0
(‰ VSMOW)
calcite corals
hydrothermal dolomite
Zn-bearing dolomite J109-5
Zn-bearing dolomite J125-11 to 15
hydrozincite outcrop
smithsonite outcrop
smithsonite J125-9 to 15
smithsonite J125-21
smithsonite J125-30 to 32
Fig. 12. δ18O vs. δ13C compositions of Jabali carbonates (values in Table 8).
This interpretation is supported by the δ34S composition of the late
gypsum veins at Jabali, being very different from the published δ34S
values of Mesozoic marine sulfates (Kampschulte and Strauss, 2004;
Strauss, 1999), but similar to the δ34S ratio measured in the primary
sphalerite. A possible scheme of the alteration processes which affected
the studied deposit can be exemplified by the following chemical
reactions:
oxidation of sphalerite
2−
2þ
1) ZnS þ 2O2 →Zn þ SO4
sphalerite
2−
2þ
þ
−
2) ZnS þ 4H2 O→Zn þ SO4 þ 8H þ 8e
sphalerite
formation of Zn-bearing carbonates
3) CaMgðCO3 Þ2 þ Zn2þ þ SO2−
4 →
dolomite
→Ca Mg1−x; Znx ðCO3 Þ2 þ ð1−xÞZn2þ þ xMg2þ þ SO2−
4
Zn‐dolomite
Ca Mg1−x; Znx ðCO3 Þ2 þ ð1−xÞZn2þ þ xMg2þ þ SO4 2 þ 2H2 O→
4)
Zn‐dolomite
→Zn2þ þ Mg2þ þ Ca2þ þ 2CO3 2 þ SO4 2 þ 2H2 O→
→2ðZn; MgÞCO3 þ
Mg‐smithsonite
CaSO4 2H2 O
gypsum:
The association of secondary Ag- and Cd-sulfides and nonsulfide Zn
phases (smithsonite 2 and hemimorphite) plus gypsum is very characteristic at Jabali. Following Sangameshwar and Barnes (1983), who
traced the chemical reactions resulting in nonsulfide mineralizations
at the Burgin mine (USA) and at Tynagh (Ireland), the co-precipitation
of smithsonite and Ag-sulfides occurs at temperatures between 25°
and 60 °C, only at neutral pH, if the Eh varies between 0 and − 2 V,
because the stability fields of the two phases are very close, but not
superimposed. The mentioned values give a constraint on the pH and
Eh ranges at Jabali during precipitation of smithsonite 2, which followed
the replacement of dolomite by smithsonite 1.
The δ13C values measured in the reef corals from the Jabali
undolomitized limestone are in the range of carbon isotope ratios of
Jurassic (Kimmeridgian–Tithonian) marine carbonates; their δ18O
compositions, instead, are lower than those commonly reported in literature (Jenkyns et al., 2002). The modification in oxygen isotope data can
be related to a diagenetic alteration during the stabilization of the
primarily aragonitic skeletons. The δ13C and δ18O ratios of the saddle
dolomite are typical of other hydrothermal dolomites in the world
(Boni et al., 2013; Diehl et al., 2010; Radke and Mathis, 1980). The
δ13C composition of Zn-bearing dolomite at Jabali is in the same range
of hydrothermal dolomite in the area, from which it is inherited most
likely. The δ18O ratio probably results from a mixing between the isotopic composition of hydrothermal dolomite and that of the Zn-carrying
fluid that deposited smithsonite.
The δ13C values of smithsonite are always negative: they are heavier
in the specimens from the upper zone of the deposit, than in those from
the middle and lower zones. This range in the carbon isotope values is
typical of most supergene nonsulfide Zn deposits in the world and
interpreted as a result of mixing between carbonate carbon from the
host rock and soil/atmospheric CO2 (Gilg et al., 2008). The negative
carbon isotope ratios of the Jabali samples may be explained by a contribution of isotopically light organic carbon from weathered soils or from
the black shales of Unit 8.
The Jabali smithsonite shows δ18O values substantially lower and
with a larger variability in comparison with other smithsonites considered as supergene in the literature (Gilg et al., 2008). As suggested by
Gilg et al. (2008), this variability may indicate effects of temperaturerelated fractionation. If the δ18O value of the solution from which smithsonite was formed can be approximated, the precipitation temperature
can be calculated using the following equation (Gilg et al., 2008):1000
ln αsmithsonite-water = 3.10 (106/T2) − 3.50.
Modern meteoric waters in Yemen have a δ18O composition ranging
from −8.0 to +10.5‰ VSMOW (Al-Ameri, 2011). In the Jabali region,
several springs about 20 km south of the deposit area have an oxygen
isotope composition between − 4.2 and − 3.6‰ VSMOW (Minissale
et al., 2007). The δ18O values of Pleistocene to Holocene speleothems
measured in various caves of Saudi Arabia and Yemen point to an
oxygen isotope composition of groundwater ranging between ~ − 12
and ~0‰ VSMOW (Fleitmann et al., 2004, 2011). Thus both rainwater
and groundwater in the region are characterized by a wide range
in δ18O. However, δ18O compositions of groundwaters are always
negative.
It follows that 1) if the smithsonites from the ore deposit precipitated from a fluid with a δ18O value comparable to modern groundwater,
the precipitation temperature would have been between ~ 55 and
~ 65 °C, corresponding to a low-temperature hydrothermal environment; 2) if the smithsonites precipitated from a fluid with a δ18O
value comparable to Pleistocene to Holocene cave water, the precipitation temperature would have been substantially lower (30–40 °C),
corresponding to a normal weathering environment at this latitude.
Variability in δ18O values throughout the different mineralized zones
could indicate different temperatures or processes active during smithsonite precipitation, and/or different stages of smithsonite formation.
Unusually low and variable δ18O values were measured in the Sierra
Mojada smithsonites by Hye In Ahn (2010) (Fig. 13). The δ18O composition of modern groundwaters in the Sierra Mojada is particularly low
(−8‰ VSMOW) translating into estimated precipitation temperatures
of about 33 °C for Zn-carbonate. In the Angouran Zn deposit (Iran),
unusual δ18O values in smithsonite were related to a hypogene genesis
of the nonsulfide ores (Fig. 13), associated with local travertine deposition (Boni et al., 2007). The mineralizing fluids were low-temperature
hydrothermal, probably circulating during the waning stages of
Tertiary–Quaternary volcanic activity in the area (Boni et al., 2007).
265
8.0
6.0
1
Supergene
smithsonite Gilg et al. 2008
2
Angouran
smithsonite 1 Boni et al. 2007
3
Sierra Mojada
smithsonite Hye In Ahn 2010
2
4.0
13
C (‰ VPDB)
2.0
0.0
3
Jabali
smithsonite this study
-2.0
-4.0
1
-6.0
-8.0
-10.0
-12.0
16.0
18.0
20.0
22.0
18O
24.0
26.0
28.0
30.0
32.0
(‰ VSMOW)
Fig. 13. δ18O vs. δ13C of Jabali smithsonite, compared to smithsonites from other nonsulfide deposits/districts: 1 = supergene smithsonite (Gilg et al., 2008), 2 = Angouran smithsonite 1
(Boni et al., 2007), 3 = Sierra Mojada smithsonite (Hye In Ahn, 2010).
Several travertine successions associated with thermal springs have
been recorded in the Marib volcanic district (about 40 km from Jabali),
as well as south of Jabal as Saad (about 15 km from Jabali). Their occurrence may suggest that a deep hydrothermal circulation involving
oxygenated waters could have been active also in the Jabali area. This
setting, paired with the previously quoted examples, could support a
model in which Jabali smithsonite precipitated from fluids characterized by different temperatures, consisting of local groundwaters
variably mixed with low-temperature hydrothermal waters. The same
mechanism may be assumed for precipitation of Zn-bearing dolomite.
The carbon and oxygen isotope ratios of hydrozincite are both in the
range of supergene Zn-carbonates (Gilg et al., 2008). This is a proof of
Early diagenetic dolomite
the most recent formation of hydrozincite, precipitated from meteoric
waters at the surface (Takahashi, 1960).
The results of our study do not confirm a possible beginning of
the oxidation process during a Cretaceous weathering stage, as stated
by Al Ganad et al. (1994). Instead, our observations are compatible
with two different oxidation processes, probably partially combined
(Fig. 14): 1) supergene weathering, mainly developed from the early
Miocene (~17 Ma) when major uplift and exhumation in Yemen commenced as a result of the main phase of Red Sea extension (Menzies
et al., 1992), and continued until present; 2) oxidation related to lowtemperature hydrothermal circulation in combination with magmaticinduced geothermal activity in the area (Miocene–Holocene).
Late Jurassic
Cretaceous (?)
Final stages of
host rock
deposition
Final Mesozoic
stages of Sab'atayn
basin evolution
Sulfide
formation
Miocene
Beginning of major
uplift phases and
geothermal activity
Present
Arid
climate
Weathering and oxidation
Hydrothermal dolomite
Sphalerite
Galena
Pyrite
Calcite (dedolomitization)
Zn-bearing dolomite
Smithsonite 1
Anglesite - Cerussite
Smithsonite 2
Acanthite - Greenockite
Gypsum
Hemimorphite
Hydrozincite
Fig. 14. Paragenesis of the main mineralogical phases observed at Jabali, framed in the geological evolution of the region.
266
7. Conclusions
This study sheds new light on many characteristics of the Jabali
secondary nonsulfide mineralization and on the related genetic
processes.
Smithsonite is the most abundant economic mineral in the secondary
deposit, where it is associated with minor hemimorphite, hydrozincite
and Ag-sulfides. The secondary mineralization evolved through different
stages: 1) alteration of original sulfides (sphalerite, pyrite and galena),
and release of metals in acid solutions; 2) alteration of dolomite host
rock and formation of Zn-bearing dolomite; 3) partial dissolution of
dolomite by metal-carrying acid fluids and replacement of dolomite
and Zn-bearing dolomite by a first smithsonite phase (smithsonite 1).
To this stage also belong the direct replacement of sphalerite and galena
by secondary minerals (smithsonite and cerussite); 4) precipitation of a
second smithsonite phase (smithsonite 2) in veins and cavities, together
with gypsum and Ag-sulfides.
The Jabali smithsonite has variable δ18O values in different parts of
the orebody, possibly associated with different environments or temperatures during smithsonite precipitation, and/or with different stages
of smithsonite formation. The δ18O values are also generally lower than
those of other smithsonites derived from pure weathering processes,
whereas the carbon isotope composition is in the same range of values
for supergene Zn-carbonates from other nonsulfide ores (Gilg et al.,
2008). With this scenario, and considering the negative δ18O values
of groundwaters and paleo-groundwaters in this area of Yemen, we
argue that the Jabali smithsonite may have precipitated from a combination of fluids, possibly consisting of local groundwaters variably
mixed with low-temperature hydrothermal waters.
The most favorable setting for the development of supergene
deposit with these geochemical characteristics could have been initiated
in the early Miocene (~ 17 Ma) and continued until Recent. Lowtemperature, hydrothermal circulation at Jabali could have been
possibly active through the magmatically-induced geothermal activity
(Miocene–Holocene) in the area.
Acknowledgments
This study is part of the PhD Thesis of N. Mondillo at the University of
Napoli “Federico II”. The authors would like to thank A. Woollett
(ZincOx), and especially B. Grist (former ZincOx) for his help during
the fieldwork and drillcore sampling. We are indebted to R. de' Gennaro
of the CISAG Napoli for his support during SEM analyses. We also thank
J. Götze for his help and discussion on the cathodoluminescence analyses, that were carried out at the TU Bergakademie Freiberg (Germany)
in the frame of a joint Deutsche Forschungsgemeinschaft project between Freiberg and Napoli Universities. The careful comments of D.
Large and of two other anonymous reviewers of Ore Geology Reviews
have greatly improved the quality of the manuscript. A special thank
is due to F. Pirajno for final editing.
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