Origin of Mineralizing Fluids of the Sediment-Hosted Navachab Gold ...

©2010 Society of Economic Geologists, Inc. Economic Geology, v. 105, pp. 285–302

Origin of Mineralizing Fluids of the Sediment-Hosted Navachab Gold Mine, Namibia: Constraints from Stable (O, H, C, S) Isotopes KATHARINA WULFF,1 ANNIKA DZIGGEL,1,† JOCHEN KOLB,2 TORSTEN VENNEMANN,3 MICHAEL E. BÖTTCHER,4 AND F. MICHAEL MEYER1 1Institute

of Mineralogy and Economic Geology, RWTH Aachen University, Wüllnerstrasse 2, 52062 Aachen, Germany

2Geological 3Institute

Survey of Denmark and Greenland (GEUS), Øster Voldgade 10, DK-1350 Copenhagen K, Denmark

of Mineralogy and Geochemistry, University of Lausanne, Anthropole, CH-1015, Lausanne, Switzerland

4Leibniz-Institute

for Baltic Sea Research, Geochemistry and Isotope Geochemistry Group and Marine Geology Section, Seestraße 15, D-18119 Rostock, Germany

Abstract The Navachab gold mine in the Damara belt of central Namibia is characterized by a polymetallic Au-Bi-AsCu-Ag ore assemblage, including pyrrhotite, chalcopyrite, sphalerite, arsenopyrite, bismuth, gold, bismuthinite, and bismuth tellurides. Gold is hosted by quartz sulfide veins and semimassive sulfide lenses that are developed in a near-vertical sequence of shelf-type metasedimentary rocks, including marble, calcsilicate rock, and biotite schist. The sequence has been intruded by abundant syntectonic lamprophyre, aplite, and pegmatite dikes, documenting widespread igneous activity coeval with mineralization. The majority of quartz from the veins has δ18O values of 14 to 15 per mil (V-SMOW). The total variations in δ18O values of the biotite schist and calcsilicate rock are relatively small (12−14‰), whereas the marble records steep gradients in δ18O values (17−21‰), the lowest values being recorded at the vein margins. Despite this, there is no correlation between δ18O and δ13C values and the carbonate content of the rocks, indicating that fluid-rock interaction alone cannot explain the isotopic gradients. In addition, the marble records increased δ13C values at the contact to the veins, possibly related to a change in the physicochemical conditions during fluid-rock interaction. Gold is interpreted to have precipitated in equilibrium with metamorphic fluid (δ18O = 12−14‰; δD = −40 to −60‰) at peak metamorphic conditions of ca. 550°C and 2 kbars, consistent with isotopic fractionations between coexisting calcite, garnet, and clinopyroxene in the alteration halos. The most likely source of the mineralizing fluid was a midcrustal fluid in equilibrium with the Damaran metapelites that underwent prograde metamorphism at amphibolite- to granulite-facies grades. Although there is no isotopic evidence for the contribution of magmatic fluids, they may have been important in contributing to the overall hydraulic regime and high apparent geothermal gradients (ca. 80°C/km−1) in the mine area.

Introduction THE NAVACHAB gold deposit is situated in the southern Central zone of the Pan-African Damara orogen, Namibia, ca. 120 km northwest of Windhoek (Fig. 1). The mine belongs to a small group of known economic gold deposits in Namibia and has an estimated total gold production of about 1.2 million ounces (Moz) Au (Steven and Badenhorst, 2002; Steven, pers. commun., 2008). The mine is run as an open-pit operation, with an average gold grade of 1.6 to 2.0 g/t (Dziggel et al., 2009a). Gold is hosted by quartz sulfide veins and semimassive sulfide lenses that are developed in a sequence of amphibolites-facies shelf-type metasedimentary rocks, including biotite schist, calcsilicate rock, and marble (e.g., Nörtemann et al., 2000; Kisters, 2005). The metasedimentary rocks are intruded by abundant syntectonic lamprophyre, aplite, and pegmatite dikes (Kisters, 2005; Dziggel et al., 2009a). The dikes commonly crosscut mineralization but locally are themselves crosscut by auriferous veins, indicating igneous activity coeval with vein formation. On a regional scale, the southern Central zone represents a typical mediumto high-temperature, low-pressure metamorphic terrane that is characterized by intrusion of voluminous syn- to late tectonic granitoids (Puhan, 1983; Jung and Mezger, 2003). † Corresponding

author: e-mail, [email protected]

0361-0128/10/3873/285-18

Although the deposit shares many characteristics with structurally controlled orogenic gold deposits (sensu Groves et al., 1998), the close spatial and temporal association of the deposit to igneous intrusions, the presence of skarn-type alteration assemblages, as well as an unusual metal association of Au-Bi-As-Cu-Ag (Dziggel et al., 2009a) has been used to suggest a possible genetic relationship to intrusion-related gold deposits (Thompson et al., 1999; Thompson and Newberry, 2000). As a result, most previously proposed genetic models have emphasized the role of igneous activity, either (1) as an agent for concentrating and redistributing gold and other metals into the ore bodies (Piranjo and Jacob, 1991), (2) by acting as a redox trap for the gold-bearing fluid (Nörtemann et al., 2000), or (3) by contributing to elevated geothermal gradients and overall hydraulic regime in the mine area (Dziggel et al., 2009a). As such, the Navachab gold deposit provides an ideal site to study the importance of magmatic fluids in a high-temperature hydrothermal gold deposit of this type. In this study, we use stable isotope analysis to constrain the source of the mineralizing fluid and to assess the processes of fluid-rock interaction and ore deposition. This study builds on previous work on the structural controls of quartz vein formation and fluid flow in the mine’s deposits (Kisters, 2005; Kolb, 2008) and the dynamics and petrologic

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Submitted: June 19, 2009 Accepted: January 20, 2010

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Kamanjab Inlier

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Omaruru Lineament

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Fau rg rbe e e t n Wa Zo al ntr e nC

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Ce n er h rt No USAKOS Navachab e on

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Swakopmund

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Walvis Bay

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WINDHOEK

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So uth

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Ma rgi na

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Northern Central Zone Southern Central Zone 0

Pre-Damara Basement

50

100 km

FIG. 1. Tectonostratigraphic zones of the Damara orogen (modified from Miller, 1983).

and geochemical consequences of fluid rock interaction at a deposit scale (Nörtemann et al., 2000; Dziggel et al., 2009a, b). Our new data support earlier studies, which favor a metamorphic origin of the mineralizing fluid (Dziggel et al., 2009a, b) and suggest that the most likely source of the mineralizing fluid was deep-crustal equivalents of the Damaran metasedimentary rocks that underwent prograde metamorphism at amphibolite- to granulite-facies grades. Regional Geology The Damara belt in central Namibia is part of the PanAfrican Damara orogen, which formed due to the collision of the Kongo and Kalahari cratons during the Neoproterozoic (e.g., Miller, 1983). Based on differences in lithology, structure, and metamorphism, the belt has been subdivided into several distinct tectonostratigraphic zones, including the Northern, Central, Southern, and Southern marginal zones (Fig. 1). The Navachab gold mine is situated in the southern part of the Central zone, called the southern Central zone (Kisters, 2005). The southern Central zone is a typical lowpressure, medium- to high-temperature metamorphic terrane made up of supracrustal rocks of the Damara sequence, various generations of I- and S-type granitoids (Jung and Mezger, 2003; Jung et al., 2003), and Mesoproterozoic gneisses, which are generally referred to as the Abbabis Metamorphic Complex (Fig. 2; Miller, 1983). The structural pattern of the southern Central zone is dominated by kilometer-scale, elongate, 0361-0128/98/000/000-00 $6.00

and northwest-vergent noncylindrical anticlines (Fig. 2; Kisters et al., 2004). The cores of these dome structures consist commonly of older gneisses and/or Pan-African granitoids, which are surrounded by, and infolded with, high-grade metamorphic rocks of the Damara sequence (Poli and Oliver, 2001). The metamorphic grade increases from amphibolite facies in the east (ca. 550°C and 2 kbars in the mine area; Puhan, 1983; Dziggel et al., 2009b) to granulite facies in the west near the Atlantic coast (ca. 750°C and 5 kbars; Masberg et al., 1992; Ward et al., 2008; Kisters et al., 2009). The peak of metamorphism in the granulite-facies rocks has been dated at ca. 525 and 504 Ma and was followed by several high-temperature metamorphic events that lasted until ca. 470 Ma (Nex et al., 2001; Jung and Mezger, 2003; Jung et al., 2003). The latter have been interpreted to be related to emplacement of late- to posttectonic granitoids (e.g., Nex et al., 2001; Jung and Mezger, 2003; Jung et al., 2003). Geology of the Navachab Gold Deposit The Navachab gold deposit is hosted by a near-vertical sequence of carbonate-bearing metasedimentary rocks, including biotite schist, marble, and calcsilicate rock of the Spes Bona, Okawayo, and Oberwasser Formations (Figs. 2, 3; Nörtemann et al., 2000; Kisters et al., 2004). The peak metamorphic mineral assemblages of the biotite schist and marble mainly comprise biotite + K-feldspar + plagioclase + quartz ± actinolite ± muscovite and calcite ± dolomite ± graphite

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ORIGIN OF THE NAVACHAB Au DEPOSIT, NAMIBIA: CONSTRAINTS FROM STABLE (O, H, C, S) ISOTOPES

287

15°45´

N Regional stri ke/dip of S0 Mon Repos Thrust Zone MRTZ Sand/ calcrete cover Rote Kuppe granite Mon Repos granodiorite

Navachab open pit

Oberwasser Formation (Daheim member)

22°00´

Okawayo Formation

Kari bib Dome

Spes Bona Formation Chuos Formation

Damara Sequence

Kari bib Formation

Usakos Dome

Etusis Formation

MRTZ

Mon Repos granodiorite 546 ± 6 Ma

Abbabis Metamorphic Complex

Rote Kuppe granite 539 ± 6 Ma

Km 0

1

2

3

4

5

FIG. 2. Simplified geologic map of the Navachab area, showing the Karibib and Usakos domes (after Kisters et al., 2004). The Navachab gold deposit is located at the northwestern limb of the Karibib dome.

(Table 1). The calcsilicate rock consists of alternating layers of clinopyroxene- and calcite-rich lithologic units, made up of variable proportions of calcite, clinopyroxene, anorthite, Kfeldspar, and quartz. The metasedimentary sequence is situated on the northwestern limb of a northwest-vergent, noncylindrical anticline known as the Karibib dome. The rocks were intruded by abundant lamprophyre, aplite, and pegmatite dikes, which show concordant as well as discordant contact relationships (Fig. 3a, b; Kisters et al., 2004; Kisters, 2005; Dziggel et al., 2009a). The aplite dikes are characterized by high K2O and SiO2 concentrations, pointing to a sedimentary origin (Wulff, 2009). Gold is hosted by quartz sulfide veins crosscutting the metasedimentary sequence, as well as cigar-shaped, bedding-parallel, and shallowly plunging, semimassive sulfide lenses that are hosted by the layered calcsilicate rocks of what is commonly known as a marble calcsilicate unit (Fig. 3; Kisters, 2005; Kolb, 2008; Dziggel et al., 2009a, b). The semimassive sulfide lenses can be traced downplunge for at least 1,800 m (Kisters, 2005). The quartz sulfide veins can be subdivided into a conjugate set of sheeted quartz veins, which dip at shallow angles to the northwest and northeast, and at least three minor vein sets of various orientations (Kisters, 2005). The intersection lineation of the sheeted quartz veins, as well as the semimassive sulfide lenses plunge at ca. 20° NNE, parallel to the fold axis of the Karibib dome in the mine (Fig. 3a; Kisters, 2005; Dziggel et al., 0361-0128/98/000/000-00 $6.00

2009a; Wulff, 2009). The geometry of the system indicates that mineralization is related closely to folding of the Karibib dome. The age of mineralization, however, has not been resolved unequivocally. Based on structural data and previously published U-Pb zircon dates on the Rote Kuppe granite and Mon Repos granodiorite (Fig. 2; Jacob et al., 2000), the age of gold mineralization has been estimated at ca. 550 to 540 Ma (Kisters, 2005). This age is also consistent with an upper intercept zircon age of 544 ± 39 Ma for an aplite dike that crosscuts the mineralization. However, U-Pb titanite dating of a quartz sulfide vein and metalamprophyre dike gives considerably younger ages of 494 ± 8 and 494 ± 12 Ma, respectively (Jacob et al., 2000). It has been suggested that this young age may reflect metamorphic crystallization during a later thermal event. Details about the petrology and geochemistry of the deposit are presented elsewhere (Dziggel et al., 2009a, b). In general, gold occurs as free gold, and the mineralized zones are characterized by a polymetallic ore assemblage of pyrrhotite-chalcopyrite-sphalerite-arsenopyrite-native bismuthgold-bismuthinite-bismuth telluride (Nörtemann et al., 2000; Dziggel et al., 2009a; Wulff, 2009). Mass-balance calculations point to the addition of several orders of magnitude of Au, Bi, As, Ag, Cu, Fe, and Mn (Dziggel et al., 2009a). Reconnaissance fluid inclusion analysis reveals that the ore fluid was a H2O-CO2-NaCl-CaCl2 mixture of low to moderate salinity

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Okawayo Formation

Karibib Formation

A)

N

Spes Bona Formation

Oberwasser Formation

Ore shoot

Semi-massive sulfide lens

Metalamprophyre

Breccia Marble and Banded Marble

Qtz-sulfide veins MC unit

MC un

ble ma rb le

ed

am ci

Aplite dyke

Ba nd

ec Br

Spes Bona Fm.

it

Okawayo Fm.

Oberwasser Fm.

ar

B)

Lamprophyre dyke

FIG. 3. A. Schematic block diagram of the Navachab open pit. The exposed lithologic units include the Spes Bona, Okawayo, Oberwasser, and Karibib Formations. The mineralization is hosted by quartz sulfide veins crosscutting the lithologic units and bedding-parallel semimassive sulfide lenses that plunge at 15°−20° NNE. B. Photographs of the Navachab open pit (view to the NE). The sequence is intruded by abundant lamprophyre, pegmatite, and aplite dikes, which show concordant as well as discordant contact relationships.

(Wulff, 2009). Three types of fluid inclusions have been identified, including (1) aqueous two-phase (L + V) inclusions, (2) three-phase H2O (L) and CO2 (L + V) inclusions, and (3) carbonic two-phase (L + V) inclusions. However, the inclusions are characterized by transposition fabrics, and it is therefore 0361-0128/98/000/000-00 $6.00

not possible to draw any conclusions on the conditions at the time of ore mineralization and hydrothermal alteration. The field relationships and alteration systematic of the quartz sulfide veins and associated alteration zones are illustrated in Figure 4. The quartz sulfide veins vary in thickness

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ORIGIN OF THE NAVACHAB Au DEPOSIT, NAMIBIA: CONSTRAINTS FROM STABLE (O, H, C, S) ISOTOPES TABLE 1. Summary of the Wall-Rock and Alteration Assemblages at the Navachab Gold Deposit (modified from Dziggel et al., 2009a)1 Wall rock

Style of mineralization

Peak metamorphic mineral assemblage

Alteration assemblage (low F/R ratio)

Alteration assemblage (high F/R ratio)

Biotite schist

Qtz sulfide vein

Bt–Kfs–Pl-Qtz ± Act ± Ms

Act-Qtz ± Cal ± Tit

Grt-Bt

Calcsilicate rock

Qtz sulfide vein

Cal-Cpx-Qtz-Kfs-An ± Tit (Cal-rich layer) Cpx-Qtz-Kfs-An-Cal ± Tit (Cpx-rich layer)

Grt-Cpx-Qtz-Kfs

Grt-Bt

Grt-Cpx-Qtz-Kfs

Grt-Bt

Grt-Cpx ± Qtz + Kfs

(Grt-Bt)

Grt-Cpx ± Qtz + Kfs

(Grt-Bt)

Semimassive sulfide lens Marble

Qtz sulfide vein

1 Mineral

Cal ± Gr (banded marble) Cal-Dol (breccia marble)

abbreviations are from Kretz (1983)

and mineralogy depending on the composition and rheological behavior of the wall rocks (Dziggel et al., 2009a). In the marble, the veins are only a few millimeters thick, strongly folded, and dominated by sulfides. In the biotite schist and calcsilicate rock, the same veins have a thickness of up to several decimeters and are mainly made up of quartz. The associated alteration halos comprise an actinolite-quartz alteration in the biotite schist, a garnet-clinopyroxene-K-feldspar-quartz alteration in the marble and calcsilicate rock, and a garnet-biotite alteration that is developed mainly in biotite schist and calcsilicate rock (Fig. 4, Table 1; Wulff et al., 2004; Dziggel et al., 2009a). However, recent field work shows that in rare cases, the garnet-biotite alteration may also be developed in the marble (Table 1). The garnet-clinopyroxene-K-feldspar-quartz and garnet-biotite alterations also characterize the semimassive sulfide lenses in the banded calcsilicate rock of the marble calcsilicate unit (Figs. 3, 4; Table 1). A characteristic feature of the deposit is the asymmetric development of alteration halos around subhorizontal veins. In addition to thin alteration halos on both sides of the veins, the alteration halos are much more pronounced facing upward, where they follow lithologic conMarble

MC unit

Grt-Cpx-Kfs-Qtz alteration

Bt-schist

z-s Qt

in ve ide f l u Act-Qtz alteration

Grt-Bt alteration

Semi-massive sulfide lens

FIG. 4. Schematic cross section, illustrating the alteration systematic around quartz sulfide veins and semimassive sulfide lenses at the Navachab gold deposit. See text for discussion. 0361-0128/98/000/000-00 $6.00

tacts and foliation planes away from the mineralization (Fig. 4; Dziggel et al., 2009a). This points to buoyancy-driven fluid flow during the mineralization event and that the banding of the rocks was near vertical when veining occurred. Petrologic and geochemical investigations show that the calcsilicate-dominated types of alteration formed in equilibrium with a fluid (super-) saturated in silica (Table 1), whereas the garnet-biotite alteration formed by interaction with a fluid undersaturated in silica (Dziggel et al., 2009a). The alteration systematics are consistent with fluid pressure fluctuations during fault valve action. Pressure-Temperature Conditions of Ore Formation Conventional thermobarometry on ore and alteration assemblages together with phase diagram modeling of the unaltered wall rocks indicate that ore mineralization occurred at or close to the peak of regional metamorphism, which has been estimated at ca. 500° to 650°C and 2 to 3 kbars (Puhan, 1983; Dziggel et al., 2009b). For example, sphalerite geothermobarometry yields pressure-temperature conditions of 590°C and 2 to 2.5 kbars, consistent with temperatures of 575° ± 15°C deduced from arsenopyrite thermometry (Nörtemann et al., 2000). In order to provide further constraints on the temperature conditions during ore formation, Dziggel et al. (2009b) carried out garnet-clinopyroxene thermometry on alteration assemblages from the semimassive sulfide lenses and quartz sulfide veins. The temperature estimates (at 2 kbars) yield ca. 480° to 670°C, with most values ranging between 480° and 600°C. The relatively large range in estimated temperatures is mainly due to compositional zoning in garnet (see Dziggel et al., 2009b for a detailed discussion). A careful reevaluation of the compositions used in these calculations shows that temperature estimates >600°C correlate with the highest Ca concentrations in garnet, interpreted to reflect disequilibrium conditions (Dziggel et al., 2009b). Excluding these values, the best estimate for the average temperature of ore formation is 550° ± 50°C. Samples and Methods The O, H, and C isotope compositions were measured on 37 samples of unaltered wall rocks, quartz sulfide veins, semimassive sulfide lenses, and associated alteration halos. Thirtythree samples were analyzed as whole-rock samples, and 24

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mineral separates of quartz, garnet, clinopyroxene, and biotite were prepared. In addition, six whole-rock samples of relevant igneous rocks, including the Rote Kuppe Granite, a pegmatite dike crosscutting the Rote Kuppe Granite, the Mon Repos Granodiorite, as well as the metalamprophyre, aplite, and pegmatite dikes from the Navachab open pit were analyzed for stable isotopes. The O, H, and C isotope composition of

all samples was measured at the University of Lausanne, Switzerland. The δ18O and D values are given in per mil (‰) V-SMOW; δ13C values are in per mil V-PDB. A compilation of all measured isotopic compositions is given in Table 2. The C and O isotope compositions of the carbonates were measured with a GasBench II connected to a Finnigan MAT DeltaPlus XL mass spectrometer, using a He carrier gas

TABLE 2. Oxygen, Hydrogen, and Carbon Isotope Data from the Navachab Gold Deposit Whole rock Sample no. Rock type/host rock

Alteration assemblage

Unaltered host rocks N123c Bt schist N135b Bt schist N4b Bt schist N7d Bt schist N86a Calcs. rock N86b Calcs. rock N9d Banded marble N20c Banded marble N22 Banded marble N44c Banded marble N140 Banded marble N149 Breccia marble

Formation

δ18O

δD

Spes Bona Spes Bona Oberwasser Oberwasser Okawayo Okawayo Okawayo Okawayo Okawayo Okawayo Okawayo Okawayo

11.6 (Silic.) 12.5 (Silic.) 13.6 (Silic.) 13.8 (Silic.) 16.1 (Carb.) 13.7 (Silic.) 19.7 (Carb.) 17.1 (Carb.) 21.4 (Carb.) 17.5 (Carb.) 21.4 (Carb.) 17.9 (Carb.)

−86 −82 −80 −8

Intrusive rocks N23 Metalamprophyre N95 Aplite N104 Pegmatite N208 Rote Kuppe Granite N207 Pegmatite N209 Mon Repos Granodionite

10.2 13.7 12.2 10.9 9.6 7.4

Bt schist Bt schist Bt schist Calcs. rock Calcs. rock

Grt-Bt Act-Qtz Act-Qtz Grt-Bt Grt-Bt

Oberwasser Oberwasser Oberwasser Okawayo Okawayo

13.8 (Silic.) 13.0 (Silic.) 14.2 (Silic.)

–73 −88 −89 −80

N84c N84d

Calcs. rock Calcs. rock

Grt-Cpx-Kfs-Qtz Grt-Cpx-Kfs-Qtz

N9b N9c N20b N101b

Banded marble Banded marble Banded marble Banded marble

Grt-Cpx (±Kfs±Qtz) Recrystallized calcite Grt-Cpx (±Kfs±Qtz) Grt-Cpx (±Kfs±Qtz)

Okawayo Okawayo Okawayo Okawayo

18.4 (Carb.) 18.3 (Carb.) 18.1 (Carb.) 17.1 (Carb.)

Mineralized zones N7a Qtz vein, Bt schist N8 Qtz vein, Bt schist N16a Qtz vein, Bt schist N146a Qtz vein, Bt schist N123a Qtz vein, Bt schist N135a Qtz vein, Bt schist N77a Qtz vein, Calcs. rock N9a Qtz vein, marble N101a Qtz vein, marble N84a MSL, Calcs. rock

Grt-Cpx-Kfs-Qtz

Oberwasser Oberwasser Oberwasser Oberwasser Spes Bona Spes Bona Okawayo Okawayo Okawayo Okawayo

10.0 (Bt)

−81

−87

N92 N16b N7b N77b1 N84b

Okawayo Okawayo

−84

−85

−86

Grt-Cpx-Kfs-Qtz Grt-Cpx-Kfs-Qtz

8.8 (Bt)

–77

9.5 (Silic.)

Calcs. rock Calcs. rock

δD

11.8 (Cpx)

Spes Bona

N85a N85b

δ18O

5.0 0.5 −1.3 −0.6 −2.2 −1.5

Grt-Bt

Okawayo Okawayo

δ13C

−9.0

Alteration zones N123b Bt schist

12.1 (Silic.) 17.1 (Carb.) 16.3 (Carb.) 13.3 (Silic.) 16.9 (Carb.) 16.3 (Carb.) 12.9 (Silic.)

Mineral separate

8.9 (Grt) 7.5 (Bt)

−80

11.8 (Grt) 11.0 (Grt) −6.2 −8.3 −6.7 −8.9 4.6 3.5 −0.4 −4.9

11.3 (Grt) 12.2 (Cpx) 11.6 (Grt) 12.1 (Cpx) 12.6 (Cpx) 12.4 (Grt) 14.5 (Qtz) 15.3 (Qtz) 14.9 (Qtz) 16.7 (Qtz) 14.6 (Qtz) 15.1 (Qtz) 16.5 (Qtz) 14.7 (Qtz) 17.9 (Qtz) 12.2 (Grt) 12.7 (Cpx)

Notes: 18O and D are given in ‰ relative to V-SMOW; abbreviations: Bt = biotite, Calcs. = calcsilicate, Carb. = carbonate portion, Cpx = clinopyroxene, Grt = garnet, Silic. = silicate portion 0361-0128/98/000/000-00 $6.00

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system according to methods adapted after Spoetl and Vennemann (2003). Samples are normalized using an inhouse standard calibrated against δ13C and δ18O values of NBS-19 (+1.95 and –2.20‰, relative to V-PDB). External reproducibility for the analyses estimated from replicate analyses of the inhouse standard is ±0.07 per mil for δ13C and ±0.08 per mil for δ18O. Measurements of the hydrogen isotope compositions of minerals were made using high-temperature (1,450°C) reduction methods with He carrier gas and a TC-EA linked to a Delta Plus XL mass spectrometer from Thermo-Finnigan on 2- to 4-mg-sized samples according to a method adapted after Sharp et al. (2001). The precision of the inhouse kaolinite standard and NBS-30 Bt for hydrogen isotope analyses was better than ±2 per mil for the method used; all values were normalized using a value of −125 per mil for the kaolinite standard and −65 per mil for NBS-30 analyzed during the same period as the samples. The oxygen isotope composition of the silicate separates and whole-rock samples was measured using a method similar to that described by Sharp (1990) and Rumble and Hoering (1994). A detailed description is given by Kasemann et al. (2001). Between 0.5 to 2.0 mg of sample was loaded onto a small Pt sample holder and pumped out to a vacuum of about 10–6 mbars. After preflourination of the sample chamber overnight, the samples were heated with a CO2 laser in 50 mbars of pure F2. Excess F2 is separated from the O2 produced by conversion to Cl2 using KCl held at 150°C. The extracted O2 is collected on a molecular sieve (5A) and subsequently expanded into the inlet of a Finnigan MAT 253 isotope ratio mass spectrometer. Replicate oxygen isotope analyses of the standards used (NBS-28 Quartz and UWG-2 Grt; Valley et al., 1995) generally have an average precision of ± 0.1 per mil for δ18O. The accuracy of δ18O values is commonly better than ±0.2 per mil compared to accepted δ18O values for NBS-28 of 9.64 per mil and UWG-2 of 5.8 per mil. Sulfur isotope compositions were determined for individual sulfides for eight samples of pyrrhotite, chalcopyrite, and pyrite (Table 3). Sulfides were obtained by microdrilling from polished sections after careful microscopic observation of the ore paragenesis. Analysis was performed on sulfide powders by combustion isotope ratio-monitoring mass spectrometry, using a Thermo Finnigan Delta+ isotope mass spectrometer coupled to a Eurovector elemental analyzer via a Finnigan TABLE 3. Sulfur Isotope Composition of Sulfides from the Navachab Gold Deposit Sample no. N26 N29 N8 N9 N12 N33 N7 N15

Mineral Po Po Po Po Po Po Ccp Py

Mineralization type

Formation

δ34S

Semimassive sulfide lens Semimassive sulfide lens Qtz sulfide vein Qtz sulfide vein Qtz sulfide vein Qtz sulfide vein Qtz sulfide vein Qtz sulfide vein

Okawayo Okawayo Oberwasser Okawayo Karibib Spes Bona Oberwasser Oberwasser

2.8 1.3 8.3 4.5 3.7 1.0 4.7 5.9

Conflo III split interface at the Max Planck Institute for Marine Microbiology, Bremen, as described by Butler et al. (2004). The isotopic data are reported relative to the V-CDT standard (Ding et al., 2001). IAEA-S-1, IAEA-S-2, and NBS127 were used for calibration. The precision of sulfur isotope analyses was ±0.4 per mil. Stable Isotopes O, H, and C isotope composition of the unaltered host rocks The range of whole-rock δ18O values of the biotite schist (VSMOW) is 11.6 to 13.8 per mil (n = 4; Table 2), within the typical range of clastic metasedimentary rocks (Hoefs, 1987). The biotite schist of the Spes Bona Formation is characterized by δ18O values of 11.6 to 12.5 and D values of −84 to −86 per mil. Those of the Oberwasser Formation have slightly higher δ18O values (13.6−13.8‰) but similar D values (Table 2). Two biotite separates have δ18O values of 8.8 and 10 per mil. The silicate portion of the clinopyroxene-rich calcsilicate rock (marble calcsilicate unit) has a δ18O value of 13.7 per mil (n = 1). A somewhat lower value (11.8‰) was obtained for a clinopyroxene separate. The values are consistent with the REE signatures, which indicate a clastic sedimentary origin (Dziggel et al., 2009a). The carbonate portion in the unaltered calcite-rich calcsilicate rock has a δ18O value of 16.1per mil V-SMOW and a δ13C value of −9 per mil V-PDB (Table 2). These values differ markedly from compositions of typical marine carbonates but are characteristic for the presence of 13C-depleted organic matter (Fig. 5; Veizer and Hoefs, 1976; Hudson, 1977; and below). The isotopically least altered samples from the banded marble are characterized by a δ18O value of 21.4 per mil VSMOW and δ13C values of −1.3 and −2.2 per mil V-PDB (Table 2; Fig. 5). Similar to the rocks of the marble calcsilicate unit, the δ18O values of calcite from most of the unaltered samples vary considerably. Samples taken proximal to quartz sulfide veins (i.e., within about 1 dm from the vein but outside the macroscopically visible alteration) are characterized by lower δ18O and δ13C values, ranging from 17.1 to 17.9 per mil V-SMOW and −1.5 and +0.5 per mil V-PDB, respectively (Fig. 5; Table 2). This indicates that the isotopic composition of carbonate has been reset on a larger scale than the macroscopically visible alteration (Fig. 4). Similar low δ18O and δ13C values were also obtained from the unaltered breccia marble (Table 2; Fig. 5). O isotope composition of the intrusive rocks The δ18O values of the Mon Repos and Rote Kuppe granitoids (Fig. 2) and the metalamprophyre dike have a range between 7.4 and 10.9 per mil V-SMOW. The δ18O values of the aplite and pegmatite dikes (13.7 and 12.2‰ V-SMOW, Table 2) are significantly higher. Similar high δ18O values of up to 15.1 per mil have been reported from other S-type granites from the Damara belt (Haack et al., 1983).

Notes: δ34S is given in ‰ relative to V-CDT; minerals analyzed: Ccp = chalcopyrite, Po = Pyrrhotite, Py = pyrite 0361-0128/98/000/000-00 $6.00

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O and H isotope composition of the hydrothermal alteration in the biotite schist The range of whole-rock δ18O values of the altered biotite schist is between 9.5 and 14.2 per mil V-SMOW (n = 4; Table

291

292

WULFF ET AL. 6 Marine carbonate

MC unit (unaltered)

4

MC unit (altered)

13

‰ δ C PDB

2

Banded marble (unaltered) Banded marble (altered)

0

Breccia marble (unaltered)

-2 -4 least altered samples

-6 -8

unaltered MC unit

-10 0

5

10

15

20

25

‰ δ O SMOW 18

FIG. 5. δ18O vs. δ13C plot for carbonates analyzed from the unaltered wall rocks and alteration zones of the Navachab gold deposit. Field representing δ18O and δ13C values of marine carbonates has been adopted from Veizer and Hoefs (1976).

2). The δ18O values of the actinolite-quartz alteration have a range between 13.0 and 14.2 per mil V-SMOW, overlapping with the values of the unaltered wall rock (Table 2). In sample N7, where it was possible to directly compare the wall rock (N7d) and the adjacent alteration zone (N7b), an increase in δ18O values from 13.8 to 14.2 per mil was observed. The δ18O values of the garnet-biotite alteration in the Oberwasser Formation are similar to the unaltered biotite schists, whereas the δD value (−73) differs significantly (Table 2). The garnet-biotite alteration in the Spes Bona Formation (N123b) has a lower δ18O value than the unaltered biotite schist (9.5‰), while the δD value is similar. The garnet and biotite separates of this sample have δ18O values of 8.9 and 7.5 per mil V-SMOW (Table 2).

O, C, and H isotope composition of the hydrothermal alteration in the calcsilicate rock The isotopic composition of the altered calcsilicate rocks was investigated on a profile through the semimassive sulfide lens and associated hydrothermal alteration halo (Fig. 6). The whole-rock silicate δ18O values range between 12.1 and 13.3 per mil (n = 3; Table 2). Carbonate whole-rock δ18O values have a restricted range in composition of 16.3 to 16.9 per mil (n = 4). The δ13C values are between −8.9 and −8.3 per mil V-PDB, indistinguishable from the unaltered calcsilicate rock (Fig. 5). The clinopyroxene separates have δ18O values of 12.1 and 12.2 per mil V-SMOW, similar to the unaltered calcsilicate rock (Fig. 6). The garnet separates have lower δ18O

18 17 16 15 14 13 12 11 10

MSL

Grt-Bt

N84a

N84b

Grt-Cpx-Kfs-Qtz alteration

unalt. wall rock

9

Garnet

N84c

Clinopyroxene

N84d

N85a

Whole rock silicate

N85b

N86a

N86b

Whole rock carbonate

FIG. 6. Variations in δ18O values from a profile through a semimassive sulfide lens and associated alteration zone (see Dziggel et al. (2009b) for a detailed petrographic description). 0361-0128/98/000/000-00 $6.00

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ORIGIN OF THE NAVACHAB Au DEPOSIT, NAMIBIA: CONSTRAINTS FROM STABLE (O, H, C, S) ISOTOPES

values (11.6−11.3‰) toward the semimassive sulfide lens (Fig. 6). It should be noted, however, that garnet in this profile records complex major and trace element zoning patterns (Dziggel et al., 2009b), and it is therefore likely that the δ18O values also vary within the grains. An isotopic disequilibrium between garnet and calcite is also reflected by the low ∆18O garnet-calcite values (Fig. 7). In contrast to the ∆18O clinopyroxene-calcite values, which are consistent with the estimated temperatures of metamorphism and mineralization (ca. 550°C), the garnet values plot somewhat below the 500°C isograd. The garnet-biotite alteration in contact with the semimassive sulfide lens has slightly lower whole-rock silicate δD and δ18O values (Table 2; Fig. 6). Analysis of the carbonate portion gives a δ18O value of 17.1 per mil V-SMOW and a δ13C value of 6.2 per mil V-PDB, similar to the garnet-clinopyroxene-K-feldspar-quartz alteration (Fig. 6). The garnet separate has a slightly lower δ18O value than any other garnet in the profile. However, garnet separated from a garnet-biotite alteration around a quartz sulfide vein has a δ18O value of 11.8 per mil V-SMOW (sample N77b2; Table 2). O and C isotope composition of the hydrothermal alteration in the marble The carbonate portion from the garnet-clinopyroxene-Kfeldspar-quartz alteration in the banded marble has δ18O values between 17.1 and 18.4 per mil V-SMOW (n = 4; Table 2). The samples have a large variation in their δ13C values, ranging between −4.9 and +4.6 per mil V-PDB (Fig. 5). The garnet and clinopyroxene separates have δ18O values of 12.4 and 12.6 per mil V-SMOW. The garnet-calcite ∆18O value in sample N101b indicates isotopic equilibrium (Fig. 7), while the clinopyroxene-calcite value plots below the range in estimated temperatures. O and S isotope composition of quartz sulfide veins and semimassive sulfide lenses The total range in δ18O values for the quartz sulfide veins is between 14.5 and 17.9 per mil V-SMOW (n = 9). The majority

15 14 13

N101b unaltered calcs. rock

18.4

12 11 10

0°

60

9

C

°C

N84d N84b

0 50

8 Clinopyroxene

7

Garnet

6 5 9

10 11 12 13 14 15

16

17 18

FIG. 7. Plot of δ18O (Grt) and (Cpx) vs. δ18O (Cal) for the Grt-Cpx-(KfsQtz) alteration in the calcsilicate rock and marble (modified from Cartwright et al., 1997). 0361-0128/98/000/000-00 $6.00

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of quartz sulfide veins crosscutting the biotite schists have δ18O values between 14.5 and 14.9 per mil (Table 2). One sample from a quartz sulfide vein in the Oberwasser Formation taken from close to the contact with the overlying marble (N146a) has a considerably higher δ18O value of 16.7 per mil. A similar value has also been obtained from a quartz sulfide vein crosscutting the marble calcsilicate unit (Table 2). In the banded marble, a ca. 15-cm-thick quartz sulfide vein (N9) has a δ18O value of 14.7 per mil, comparable to the values obtained from veins in the biotite schists. In contrast, quartz from a few millimeter-thick vein in the marble (N101a) has a much higher δ18O value of 17.9 per mil. The S isotope composition of sulfides from the quartz sulfide veins range between 1.0 and 8.3 per mil V-CDT (n = 6; Table 3). Due to the high sulfide content of the semimassive sulfide lenses only mineral separates could be analyzed. A garnet separate has a δ18O value of 12.2 per mil V-SMOW, while the associated clinopyroxene has a value of 12.7 per mil. These values are somewhat higher than those of the calcsilicate rocks (Fig. 6). The S isotope composition of sulfides from the semimassive sulfide lens has a range between 1.3 and 2.8 per mil V-CDT (n = 2; Table 3). Discussion Isotopic variations in the carbonates The O and C isotope variations observed in the marble and calcsilicate rock suggest that these rocks were exposed to significant postdepositional isotopic changes (Fig. 5). In order to investigate if the low δ18O and δ13C values of the carbonates were caused by decarbonation reactions during metamorphism and/or fluid infiltration, a Rayleigh-type decarbonation model was calculated (Fig. 8). In contrast to batch volatilization, where all CO2 is produced before leaving the system, each molecule of CO2 produced during Rayleigh volatilization is immediately removed (Valley, 1986). At Navachab, the phase relationships suggest a reaction-induced fluid immiscibility and thus, a continuous removal of CO2 during fluid-rock interaction (Dziggel et al., 2009a, b). For the model, an initial carbonate composition of 21.4 per mil V-SMOW and 1.3 per mil V-PDB was chosen, as these values represent the isotopic composition of the least altered marble (Table 2). The calculations were conducted over a temperature range of 300° to 550°C, the estimated range in temperature from the onset to the peak of metamorphism (Dziggel et al., 2009b). The calculated decarbonation curves predict decreasing δ18O and δ13C values with decreasing carbonate content (Fig. 8a, b). However, there is no correlation between wt percent carbonate and δ18O values in the samples investigated (Fig. 8a), and the model also fails to explain the variations in δ13C values (Fig. 8b). The lack of correlation between the δ18O values and the carbonate content may be interpreted in two ways: the O isotope composition of the carbonates resulted from isotopic exchange between carbonates and silicates prior to, or during, metamorphism (Fourcade et al., 1996), or it reflects an exchange with a fluid of a lower 18O/16O compared to the rocks (see below). It is likely that both processes were active. The carbon isotope compositions are more complex to interpret as the δ13C values differ between the different units

293

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WULFF ET AL. 22

a) 20

18

16

14

550°C

12

300°C

500°C

10 0

20

40

60

80

100

80

100

Carbonate content wt. % 6

b) 4

MC unit (unaltered) MC unit (altered) Banded marble (unaltered)

2

Banded marble (altered) Breccia marble (unaltered)

0

550°C

-2 -4

-6 -8 -10 0

20

40

60

Carbonate content wt. % δ18O

FIG. 8. A. vs. wt percent carbonate, and B. δ13C vs. wt percent carbonate for marbles and calcsilicate rocks of the Navachab gold deposit. The curves correspond to a Rayleigh distillation model, calculated using the calibration of Bottinga (1968).

(Fig. 8b). This points to a partial lithologic control on the C isotope compositions. In addition, the δ13C values of the marble calcsilicate unit (−9.0 to −6.2‰ V-PDB; Fig. 8b) are unusually low, the lowest values being recorded from the unaltered calcite-rich calcsilicate rock (Table 2). Two main hypotheses are considered to explain this phenomenon: the low δ13C values in the marble calcsilicate unit may be due to the original presence of 13C-depleted, synsedimentary organic matter (e.g., Oberthür et al., 1996; Salier et al., 2005) that has exchanged with or provided a source for the carbonate carbon, or they are caused by isotopic exchange with magmatic CO2 (e.g., Taylor and O’Neil, 1977). However, an igneous origin for the low δ13C values appears to be in conflict with the large differences in δ13C values between the different units. 0361-0128/98/000/000-00 $6.00

Therefore, a synsedimentary or early diagenetic and/or metamorphic origin of the low δ13C values is preferred. This is also supported by the presence of graphite in the layered calcsilicate rocks (e.g., Nörtemann et al., 2000), as well as stable isotope analyses of regional equivalents of the marble calcsilicate unit, which also have low δ13C values (K.-H. Hoffmann, pers. commun., 2008). It should be noted, however, that similar large negative excursions in carbonate δ13C values (≤−5‰ VPDB) have also been recorded in marine carbonates of the Damara sequence that are interpreted to have been affected by glacial events. In the area around Navachab, two Neoproterozoic “snowball earth” ice ages, namely the Sturtian and Marinoan glaciations, have been documented in the Chuos and upper Oberwasser Formations to the foot- and hanging

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ORIGIN OF THE NAVACHAB Au DEPOSIT, NAMIBIA: CONSTRAINTS FROM STABLE (O, H, C, S) ISOTOPES

wall of the marble calcsilicate unit (cf. Fig. 2; e.g., Hoffman et al., 1998). To date, however, there is no geologic evidence to support a glacial origin for the low δ13C values. The δ13C values in the marble contrast strongly with those of the marble calcsilicate unit (Fig. 8b). Samples that are macroscopically unaltered but that have been taken in close (cm) proximity to the veins have higher δ13C values than their unaltered counterparts (Fig. 8b). This suggests that the carbon isotope variations in the marble are linked directly to the interaction with the mineralizing fluid. Overall, the large spread in values points to very local and probably complex processes that are related to carbonate dissolution and/or precipitation. However, the fluid/rock ratios in the marble must have been low as the values differ significantly from the hostrock composition, and there are large spatial variations. A possible explanation for such higher values may be the isotopic exchange with CO2 released during decarbonation reactions in the alteration zones. For CO2 in equilibrium with the unaltered banded marble, δ13C values of up to 2.5 per mil VPDB were calculated using the calibration of Scheele and Hoefs (1992) at a temperature of 550°C. If exchange with CO2 was the underlying process, this value should not be exceeded, indicating that CO2 release alone cannot explain the observed high δ13C values. An alternative explanation for the combined carbon and oxygen isotope compositions in these samples is a change in the physicochemical conditions such as pH, fO2, or temperature (e.g., Ohmoto, 1972). At Navachab, there is no measurable change in temperature (Dziggel et al., 2009b), and the presence of graphite (e.g., Nörtemann et al., 2000), absence of oxides and sulfates and/or Fe3+-bearing minerals, as well as the ore assemblage of pyrrhotite-arsenopyrite (Dziggel et al., 2009a) all point to a generally low oxygen fugacity of ca. 10−20 atmospheres (i.e., below the QFM and possibly along the CO2-CH4 buffers; Mikucki and Ridley, 1993). This suggests that the increase in δ13C values may be controlled by a change in the redox conditions during fluid-rock interaction. The predominance of sulfides in veins crosscutting the marble also indicates that sulfide deposition in the veins was caused by fluid-rock interaction (Fig. 4; e.g., McCuaig and Kerrich, 1998). This requires an acidic or mildly acidic silica- and sulfur-bearing fluid that reacted with the marble to form calcsilicate minerals, eventually resulting in an increase in the pH and the precipitation of ore minerals due to the consumption of protons by carbonate (e.g., Reed, 1997). At 250°C and a neutral or mildly alkaline pH, calcite in equilibrium with graphite may record increasing δ13C values at an oxygen fugacity less than 10−38 atmospheres (e.g., Rye and Ohmoto, 1974). Extrapolating this value to ca. 550°C by keeping the oxidation state constant gives a value of ca. 10−20 atmospheres. This value is consistent with the above estimates and also agrees well with estimated oxygen fugacity from reduced amphibolites-facies gold deposits elsewhere (Mikucki and Ridley, 1993; McCuaig and Kerrich, 1998; Kolb et al., 2000). Another process that may have influenced the increasing δ13C values is the presence of methane. At Navachab, the phase relationships and alteration systematics indicate that the XCO2 of the mineralizing fluid was ≤0.2 (Dziggel et al., 2009b). When in contact with the marble, this value must have increased considerably due to decarbonation reactions 0361-0128/98/000/000-00 $6.00

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and/or mixing with metamorphic CO2, thereby creating a local sink for hydrogen (e.g., Mumin et al., 1996). This would lead to the production of methane via reactions such as the following: CO2 + 4H2 = CH4 + 2H2O.

(1)

Because methane is strongly enriched in 12C, changes in the CH4 abundance have a dramatic effect on the δ13C values of coexisting carbonate species in the fluid, such that calcite may be strongly enriched in 13C (Bottinga, 1969; Ohmoto, 1972). However, the presence of methane is not supported by our preliminary fluid inclusion data, which indicate that CH4 was not a major component in the ore fluid. Composition of the ore fluid In order to provide constraints on the origin of the mineralizing fluid, the O and H isotope composition of an aqueous fluid in equilibrium with the mineral separates and wholerock samples was calculated. The equations used for calculating the fractionation factor for the O isotope composition for the silicate-water and calcite-water systems are from Zheng (1993) and Zheng (1999), respectively, whereas the H isotope values were calculated using the calibrations of Suzoki and Epstein (1976). For whole-rock samples, the fractionation factors for the rock-forming minerals were weighted according to the modal proportions of the minerals. The calculations assumed a temperature of equilibration of 550°C, the estimated average temperature of mineralization (Dziggel et al., 2009b; Wulff, 2009), except for the igneous rocks, where a minimum solidification temperature of 600°C was assumed. The δ18O values of a fluid coexisting with vein quartz and alteration minerals in the biotite schist are within a narrow range of 11.7 to 14.2 per mil V-SMOW (Fig. 9; Table 4). Similar δ18O values (11.7−14.3) have also been obtained from the unaltered samples of the Spes Bona and Oberwasser Formations, indicating that the mineralizing fluid was in isotopic equilibrium with the biotite schists. The D values of a fluid in equilibrium with the unaltered and altered biotite schist have a range from −39 to −60 per mil V-SMOW, with most values between −50 and −60 per mil (Table 4). The δ18O values of a fluid in equilibrium with whole-rock samples and mineral separates from the calcsilicate rock of the marble calcsilicate unit are frequently higher, with most values ranging between 14.0 and 15.1 per mil V-SMOW (Fig. 9; Table 4). A considerably lower δ18O value (13.0‰) has been obtained from the garnet-biotite alteration in contact with the semimassive sulfide lens (Fig. 6). Together with the associated D value of −42 per mil V-SMOW (Table 4), the data are in good agreement with the values obtained from the biotite schist. A fluid in equilibrium with the least altered banded marble has a significantly higher δ18O value than that of any other sample (19.4‰ V-SMOW; Fig. 9), whereas the remaining data record a large spread in δ18O values. The fluid in equilibrium with a ca. 15-cm-thick quartz vein crosscutting the banded marble (sample N9) has a δ18O value of 12.2 per mil, similar to fluid coexisting with vein quartz in the biotite schist. In contrast, a fluid coexisting with quartz from the relatively thin vein in the marble (sample N101a) has a higher δ18O value (Fig. 9, Table 4). We interpret this higher δ18O value to

295

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WULFF ET AL. 21

19

Vein quartz Garnet Biotite

Whole rock silicate Whole rock carbonate Clinopyroxene

17

15

Fluid composition

13

11

Rote Kuppe Granite 9

Mon Repos Granodiorite 7

Spes Bona

MC unit

Banded marble

Breccia marble

Oberwasser

Magmatic rocks

FIG. 9. Oxygen isotope composition of an aqueous fluid in equilibrium with the mineralized, altered, and unaltered samples from the Navachab open pit. The calculations at 550°C were carried out based on the calibrations of Zheng (1993, 1999) and Suzoki and Epstein (1976).

be related to a more pronounced buffering of the oxygen isotope composition of the mineralizing fluid with the wall-rock carbonate. A similar scenario may also account for the higher δ18O values in the marble calcsilicate unit (Fig. 9). For an aqueous fluid in equilibrium with the lamprophyre, aplite, and pegmatite dikes in the Navachab open pit, δ18O values of 12.2 to 12.8 per mil V-SMOW were calculated (Table 4, Fig. 9). These values are indistinguishable from the calculated composition of the vein quartz fluids. In contrast, fluids coexisting with the Rote Kuppe Granite and the Mon Repos Granodiorite have considerably lower δ18O values (Fig. 9). Even though the δ18O value of a fluid in equilibrium

with the granitoid dikes is indistinguishable from that of vein quartz, the δD values of the granitoid samples are considerably lower (Fig. 10). In contrast to the igneous rocks, most of the other samples plot in a narrow range, consistent with an origin as metamorphic fluid (Taylor, 1974). However, the δ18O values overlap with the field for magmatic fluids if the field for peraluminous Cornubian granite-derived fluids is considered (Sheppard, 1986). The high δ18O values of the Cornubian fluids have been interpreted to be related to melting, assimilation and/or exchange with argillaceous metasedimentary rocks, hence they may be typical for S-type granites. One might, therefore, argue that the δD values of the altered

20

Grt-Bt alteration (Bt schist) SMOW

0

metamorphic water

lin e

‰ δD SMOW

-20

unaltered Bt schist, Grt-Bt alteration (MSL)

ter

-40

wa

Cornubian magmatic H2O

Me

teo

ric

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Act-Qtz alteration primary magmatic water Navachab granitoids

-100 -120 -20

-15

-10

-5

0

5

10

15

20

25

30

‰ δ O SMOW 18

FIG. 10. δD vs. δ18O diagram of fluid compositions in equilibrium with altered and unaltered metasedimentary rocks from the Navachab gold deposit compared to magmatic rock types. Only samples containing hydrous minerals are shown (symbols as in Fig. 9). The field for primary magmatic water is from Taylor (1974), the Cornubian magmatic H2O is from Sheppard (1986), and data for the field of metamorphic water are from Taylor (1974) and Sheppard (1981). Meteoric water line is from Epstein (1970). 0361-0128/98/000/000-00 $6.00

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TABLE 4. Oxygen and Hydrogen Isotope Composition of an Aqueous Fluid in Equilibrium with the Samples Investigated1 Whole rock Sample no. Rock type/host rock

Alteration assemblage

Unaltered host rocks N123c Bt schist N135b Bt schist N4b Bt schist N7d Bt schist N86a Calcs. rock N86b Calcs. rock N9d Banded marble N20c Banded marble N22 Banded marble N44c Banded marble N140 Banded marble N149 Breccia marble

Mineral separate

Formation

δ18O

δD

δ18O

δD

Spes Bona Spes Bona Oberwasser Oberwasser Okawayo Okawayo Okawayo Okawayo Okawayo Okawayo Okawayo Okawayo

11.7 (Silic.) 12.5 (Silic.) 14.3 (Silic.) 13.7 (Silic.) 14.1 (Carb.)

–52 −50 −45 −54

11.3 (Bt)

−48

12.5 (Bt)

−47

Intrusive rocks N23 Metalamprophyre N95 Aplite N104 Pegmatite N208 Rote Kuppe Granite N209 Mon Repos Granod.

13.9 (Cpx) 17.7 (Carb.) 15.1 (Carb.) 19.4 (Carb.) 17.5 (Carb.) 19.4 (Carb.) 16.0 (Carb.) 12.4 12.8 12.2 10.1 7.7

−65 −66

−67

Alteration zones N123b Bt schist

Grt-Bt

Spes Bona

11.7 (Silic.)

−50

N92 N16b N7b N77b1 N84b

Bt schist Bt schist Bt schist Calcs. rock Calcs. rock

Grt-Bt Act-Qtz Act-Qtz Grt-Bt Grt-Bt

Oberwasser Oberwasser Oberwasser Okawayo Okawayo

−39 −60 −59

N84c N84d

Calcs. rock Calcs. rock

Grt-Cpx-Kfs-Qtz Grt-Cpx-Kfs-Qtz

Okawayo Okawayo

N85a N85b

Calcs. rock Calcs. rock

Grt-Cpx-Kfs-Qtz Grt-Cpx-Kfs-Qtz

Okawayo Okawayo

13.7 (Silic.) 13.4 (Silic.) 13.7 (Silic.) 14.8 (Grt) 13.0 (Silic.) 15.1 (Carb.) 14.3 (Carb.) 14.4 (Silic.) 14.5 (Carb.) 14.4 (Carb.)

N9b N9c N20b N101b

Banded marble Banded marble Banded marble Banded marble

Grt-Cpx (±Kfs±Qtz) Recrystallized calcite Grt-Cpx (±Kfs±Qtz) Grt-Cpx (±Kfs±Qtz)

Okawayo Okawayo Okawayo Okawayo

Grt-Cpx-Kfs-Qtz

Oberwasser Oberwasser Oberwasser Oberwasser Spes Bona Spes Bona Okawayo Okawayo Okawayo Okawayo

Mineralized zones N7a Qtz vein, Bt schist N8 Qtz vein, Bt schist N16a Qtz vein, Bt schist N146a Qtz vein, Bt schist N123a Qtz vein, Bt schist N135a Qtz vein, Bt schist N77a Qtz vein, Calcs. rock N9a Qtz vein, marble N101a Qtz vein, marble N84a MSL, Calcs. rock

16.4 (Carb.) 16.3 (Carb.) 16.1 (Carb.) 15.1 (Carb.)

−46

11.9 (Grt) 10.0 (Bt)

−46

14.0 (Grt) 14.2 (Grt) (Cpx) 14.5 (Grt) 14.1 (Cpx) 14.6 (Cpx) 15.4 (Grt) 12.0 (Qtz) 12.8 (Qtz) 12.2 (Qtz) 14.2 (Qtz) 12.0 (Qtz) 12.6 (Qtz) 13.9 (Qtz) 12.2 (Qtz) 15.4 (Qtz) 15.2 (Grt) 12.7 (Cpx)

Notes: 18O and D are given in ‰ relative to V-SMOW; abbreviations: Bt = biotite, Calcs. = calcsilicate, Carb. = carbonate portion, Cpx = clinopyroxene, Grt = garnet, Silic. = silicate portion; calculations were carried out after the calibrations of Zheng (1993, 1999) and Suzoki and Epstein (1976) at 550°C

samples are not a primary metamorphic feature, but rather the result of an isotopic exchange of preexisting hydrous minerals with an infiltrating magmatic fluid of S-type granitic origin. This is particularly relevant as the unusual metal association of Au-Bi-As-Cu-Ag is typically associated with intrusionrelated gold systems (e.g., Thompson et al., 1999; Müller et al., 2004), although it has also been recorded from a number 0361-0128/98/000/000-00 $6.00

of amphibolites-facies orogenic gold deposits (e.g., Kolb et al., 2000; Ridley et al., 2000). A modification of the H isotope composition during fluid-rock interaction, however, does not seem to be recorded by the δD values of the altered samples, which are either indistinguishable from those of the unaltered biotite schist or have somewhat higher δD values (Fig. 10). Only samples of the actinolite-quartz alteration have δD

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values in between those of the igneous samples and the biotite schists, but actinolite has a fractionation factor relative to water that is different and normally smaller than that of biotite (Suzoki and Epstein, 1976). Taken in conjunction, the simplest interpretation for the O and H isotope data is that the mineralizing fluid (δ18O = 12−14‰; δD = −40 to −60‰) originated from a metamorphic source compositionally similar to the biotite schist, or alternatively, were at least partially buffered by the rocks of the surrounding metasedimentary sequence, most notably the carbonates that may have reacted with the mineralizing fluids. However, the stable isotope compositions of a magmatic fluid, if derived from S-type, high δ18O granites, may well be indistinguishable from such a metamorphic fluid (e.g., Ridley and Diamond, 2000) and/or may also have been at least partially buffered by exchange with the wall rocks. Therefore, the possibility that the mineralizing fluid was originally an external (i.e., magmatic) fluid that resided long enough in the biotite schist to attain isotopic equilibrium cannot be excluded. The possible role of this igneous activity in the genesis of the Navachab gold deposit will be discussed further below. S isotopes Figure 11 shows the S isotope composition of the analyzed sulfides in comparison to a typical range of δ34S values previously found for orogenic gold deposits (McCuaig and Kerrich, 1998; Kolb et al., 2004), intrusion-related skarn deposits (Ohmoto and Goldhaber, 1997), and clastic metasedimentary rocks (Chambers, 1982). The δ34S values in orogenic gold deposits are generally interpreted to be indicative of sulfur sources that have been derived from magmatic fluids or from the mobilization of sulfur in igneous, metamorphic, or sedimentary rocks (McCuaig and Kerrich, 1998). Although the δ34S values of the present study cannot be unequivocally allocated to a specific sulfur source, the data (δ34S between

1−8‰ V-CDT) are within the typical range reported from orogenic gold deposits. They point to an isotopically uniform S source and are consistent with reducing conditions of the aqueous fluids during the time of sulfide mineral formation (Ohmoto and Goldhaber, 1997; McCuaig and Kerrich, 1998; Ridley and Diamond, 2000; Kolb et al., 2004). Origin of the ore fluid and a model for the fluid evolution The stable isotope data presented in this study suggest that the mineralizing fluid at the Navachab gold deposit is compatible with a metamorphic fluid origin. The geological, structural, and geochemical characteristics of the deposit complement such interpretation, as they are consistent with typical orogenic gold deposits elsewhere (Kisters, 2005; Kolb, 2008; Dziggel et al., 2009a). However, the crosscutting relationships between the quartz sulfide veins and intrusive rock types indicate that magmatic activity overlapped with the mineralization, and regional plutonism was coeval with deformation (Kisters et al., 2004). Plutonism thus must have played a fundamental role for the hydraulic regime in the southern Central zone of the Damara belt, even though the geochemical, fluid inclusion, and stable isotope data do not provide direct evidence for this igneous activity (Dziggel et al., 2009a). The most likely explanation is that the ore fluids originated from deep-crustal equivalents of the biotite schists and were released by dehydration reactions during prograde metamorphism. At metamorphic degrees and typical bulk compositions prevailing in the Damara belt, biotite is the dominant hydrous mineral in the clastic metasedimentary rocks. Fluid released at higher metamorphic grades should, therefore, mainly be generated by the breakdown of biotite. Migmatization would generally set an upper temperature limit on this type of fluid release, because fluids present during partial melting tend to be partitioned into the melt (e.g., McMillan and Holloway, 1987; Holtz et al., 1995).

Host rock:

Mineral analyzed:

Spes Bona Fm.

Pyrrhotite

Okawayo Fm.

Chalcopyrite

Oberwasser Fm.

Pyrite

Karibib Fm.

Clastic sediments S-type source skarns I-type source skarns Orogenic gold deposits Qtz-sulfide veins Massive sulfide lenses

-4

-2

0

2

4

6

8

10

δ S sulfide 34

FIG. 11. Sulfur isotope composition of the samples analyzed. The range of δ34S values for orogenic gold deposits are from McCuiag and Kerrich (1998), those from skarn deposits associated with I- and S-type granites from Ohmoto and Goldhaber (1997). 0361-0128/98/000/000-00 $6.00

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In a recent study, Ward et al. (2008) reported fluid-present melting at pressure-temperature conditions of ca. 750°C and 5 kbars in the deeper crustal metasedimentary rocks of the southern Central zone, some 100 km to the southwest of the Navachab gold deposit. Melting experiments at water-saturated conditions suggest that the rocks underwent incongruent melting via the reaction biotite + quartz + plagioclase + H2O = melt + garnet + cordierite (Ward et al., 2008). These data indicate that significant amounts of fluid were present during partial melting of the deeper crustal rocks. This is also supported by δ18O values of 14 to 16 per mil V-SMOW of quartz from these higher grade metamorphic rocks, which are consistent with an internally derived fluid (Hoernes and Hoffer, 1985). However, Sm-Nd garnet and U-Pb monazite ages indicate that high-grade metamorphism and partial melting occurred between ca. 525 ± 2 and 504 ± 3 Ma and, therefore, postdate the mineralization and peak of metamorphism at Navachab (Jung and Mezger, 2003). The ages also show that the S-type granites in the study area were not derived from these migmatites but from a deeper and unexposed part of the orogen, while the clastic metasedimentary rocks were still below their solidus and exposed to prograde dehydration. Nevertheless, assuming an average crustal density of 2.7 g/cm−3, the data indicate that the fluids generated during prograde metamorphism may have formed over a crustal profile of at least ca. 10 km (i.e., from ca. 2 kbars and 550°C at Navachab, to 750°C and 5 kbars—the minimum pressuretemperature conditions of partial melting). It is interesting to note that these pressure-temperature estimates point to distinctly different apparent geothermal gradients, ranging from ca. 40°C/km−1 at the sites of partial melting, to almost 80°C/ km−1 in the mine area. Thus, granitoid plutonism at mid- to

upper crustal levels appears to have influenced the overall geothermal regime considerably. Fluids released from normal crustal rocks at relatively high temperature and/or low pH have the potential to form ore fluids (Yardley, 2005). For the Damara belt, Haack et al. (1984) showed that prograde metamorphism between the biotite isograd and the onset of partial melting was associated with a substantial loss of metals, including 61 percent of Cu and 86 percent of Bi. These data are consistent with our mass-balance calculations, which indicate the addition of several orders of magnitute of Au, Bi, As, Ag, and Cu (Dziggel et al., 2009a). The study of Haack at et al. (1984) also shows that the mobilization of metals in the Damara belt was related to progressive dehydration of the clastic metasedimentary rocks, signifying their potential to form hydrothermal deposits. Similar results were also obtained by Pitcairn et al. (2006), who related the metal inventory of orogenic gold deposits in the Otago and Alpine schists to metamorphic dehydration reactions and breakdown of sulfides at higher metamorphic grades. Because of the low melting temperature of native bismuth (melting point 271°C) and bismuth-rich polymetallic assemblages (e.g., Au-Bi eutectic at 241°C), much of the metal assemblage produced during prograde metamorphism should be incorporated into a bismuth or sulfide melt (e.g., Tomkins and Mavrogenes, 2003; Tooth et al., 2008). Ciobanu et al. (2006) and Tooth et al. (2008) showed that bismuth melts coexisting with a hydrothermal fluid are able to scavenge gold from hydrothermal fluids such that even a fluid undersaturated with respect to native gold may form economic gold deposits. Our proposed genetic model for the gold mineralization at Navachab is shown in Figure 12. Fluids generated at deeper crustals levels are derived from metamorphic dehydration Mon Repos Granodiorite

Navachab Usakos dome Karibib dome

-1

80°C/km

2 kbar

δ18O (fluid) = 12-14 0 V -SMOW

metamorphic dehydration -Bi, Au, Cu onset of partial melting

5 kbar

40°C/km-1 FIG. 12. Sketch illustrating the proposed origin of the mineralizing fluid at the Navachab gold deposit (not to scale). Fluid is produced by metamorphic dehydration reactions at mid-crustal levels and then focused into suitable fluid pathways, i.e., high-permeability zones such as fractures and shear zones. At even deeper crustal levels, the rocks of the Damara sequence underwent partial melting (Ward et al., 2008). Far-field transport of melt likely followed the same structures (Kisters et al., 2009) and provided the heat for the elevated geothermal gradients in the mine area. See text for discussion. 0361-0128/98/000/000-00 $6.00

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reactions of the biotite schists at subsolidus conditions. In the mid- and lower crust (>10 km), the permeability of metamorphic rocks is very low (e.g., Brace, 1980), and the fluid pressure is typically at or close to lithostatic values (Etheridge et al., 1984). At these deeper crustal levels, the initial fluid release from the rocks was likely controlled by gradients in fluid pressure created in dilational fractures during ongoing deformation (cf. Kisters et al., 2009). Far-field transport of the fluid is interpreted to have occurred along high-permeability structures such as fractures and shear zones or regional-scale lithologic contacts within the Damara sequence (Fig. 12). The melts generated at even deeper crustal levels likely ascended along the same ductile structures (e.g., Brown, 2007), and the elevated geothermal gradients in the mine area are related to the advection of heat by these melts. It has been argued that if magmatic fluids contributed to the fluid source of orogenic gold deposits, they may not be recorded due to fluid-rock interactions along the fluid pathways and/or mixing of compositions of the source and the wall rock (Ridley and Diamond, 2000). In the case of Navachab, however, the ultimate source in both metamorphic and intrusion-related models may have been identical, as both metamorphic fluids and S-type granites (e.g., the aplite dikes at Navachab) originated from clastic metasedimentary rocks. Therefore, although the production and ascent of magmas into upper-crustal levels may have contributed in concentrating and transporting gold and other metals from the sites of fluid production to that of ore deposition, the currently available petrological and geochemical data are most consistent with metamorphic dehydration as the dominant process in extracting fluids and metals from midcrustal levels. At the deposit site, the geometry and progressive deformation of the quartz sulfide veins indicate that bedding-parallel fluid infiltration and ore deposition were coeval with northwest-directed thrusting and folding of the Karibib dome (Kisters, 2005; Kolb, 2008). At the elevated temperature conditions during ore formation, bismuth would still be present as melt and only solidify after ore formation (e.g., Tooth et al., 2008). Fluid-rock interaction was probably an important mechanism for ore deposition in the carbonate-bearing lithologic units (McCuaig and Kerrich, 1998), leading to the formation of the semimassive sulfide lenses and the sulfidedominated veins in the marble. Acknowledgments This study was supported by the German Science Foundation (grant ME 1425/13-1). We are grateful to Anglogold Ashanti Namibia and the staff of the Navachab gold mine for logistical and financial support during mine visits. Sulfur isotope measurements were carried out by M.E.B. when he was a member of the Max Planck Institute for Marine Microbiology, Bremen. Nick Steven and Frik Badenhorst are thanked for many fruitful discussions during the fieldwork. Reviews by K.L. Shelton and S. Hagemann substantially improved the manuscript and are gratefully acknowledged. L. Meinert is thanked for many useful comments and editorial handling. REFERENCES Bottinga, Y., 1968, Calculation of fractionation factors for carbon and oxygen isotopic exchange in the system calcite-carbon dioxide-water: Journal of Physical Chemistry, v. 72, 3, p. 800−808. 0361-0128/98/000/000-00 $6.00

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