October 2011 LIP of the Month

The Gawler Range Volcanics-Hiltaba Suite silicic large igneous province and the Cu-U-Au-Ag Olympic Dam deposit – New research developments

Andrea Agangi*, Jocelyn McPhie, Vadim S. Kamenetsky, Sharon Allen, Kathy Ehrig, Adam Bath, Isabelle Chambefort

School of Earth Sciences and CODES, University of Tasmania, Private Bag79, Hobart,
Tasmania 7001, Australia
* Present address: Paleoproterozoic Mineralisation Group, Geology Department, University of Johannesburg, PO BOX 524, South Africa
E-mail address: aagangi@uj.ac.za

The Gawler Range Volcanics (GRV) and Hiltaba Suite (HS) form a Mesoproterozoic silicic-dominated large igneous province (the Gawler SLIP) emplaced in an intraplate setting in the Gawler Craton, South Australia.

The Gawler SLIP has been the subject of several research projects at ARC-CODES in the last several years. In a previous contribution to this website (Allen et al., February 2009 www.largeigneousprovinces.org/09feb), we have outlined the main characteristics of the igneous province and the associated Cu-U-Au-Ag Olympic Dam deposit. Therefore, we will only summarise some key aspects here, and refer to that contribution for more detail. Here, we report the outcomes of our new research, and namely: 1) some of the previously less studied aspects of magma storage and magma chamber dynamics; and 2) the role of magmatic volatile elements in the (re)mobilisation of rare earth and high field strength elements, which can help explaining the formation of the mineral deposits associated with the Gawler SLIP.

The Gawler Range Volcanics and Hiltaba Suite

The Gawler SLIP (Fig. 1) was emplaced in an intracontinental setting, during the Laurentian supercontinent assembly (Allen and McPhie, 2002; Allen et al., 2008; Betts and Giles, 2006; Blissett et al., 1993; Creaser, 1995; Giles, 1988) and is coeval with the 1.3 -1.6 Ga anorogenic magmatic event throughout Laurentia and Baltica (Anderson and Morrison, 2005; Rämö and Haapala, 1995). U-Pb zircon dating of the volcanic units has yielded a narrow age range of 1591-1592 Ma (Creaser, 1995; Creaser and Cooper, 1993; Fanning et al., 1988), whereas ages of the HS granites range from 1583±7 to 1598±2 Ma (Flint, 1993). The Gawler SLIP is associated with a major metallogenic event that affected most of the Gawler Craton (Budd and Fraser, 2004; Fraser et al., 2007; Skirrow et al., 2007; 2002) (Fig. 1).

Figure 1: Geology of the Gawler Craton (after Daly et al., 1998; Betts and Giles, 2006, Hand et al., 2007). Inset shows the location of the Gawler Craton and distribution of Proterozoic.

The GRV include several medium- to large-volume (tens to several hundreds of km3) felsic lavas and ignimbrites (Allen et al., 2008; Blissett et al., 1993) and minor mafic and intermediate units. The GRV have been subdivided into lower and upper sequences (Blissett et al., 1993). The lower GRV consist of thick (up to 3 km) successions, erupted from several discrete volcanic centres. Evenly porphyritic felsic lavas are interbedded with ignimbrites and very minor volcanogenic sedimentary facies. The Chitanilga Volcanic Complex at Kokatha (Blissett, 1975, 1977a, 1977b; Branch, 1978; Stewart, 1994) and Glyde Hill Volcanic Complex at Lake Everard (Blissett, 1975, 1977a, 1977b; Giles, 1977) are the two best exposed parts of the lower GRV. The upper GRV are composed of three large-volume (>500 km3) evenly porphyritic felsic lavas (Allen and McPhie, 2002; Allen et al., 2008; McPhie et al., 2008). The GRV are essentially undeformed and unmetamorphosed and primary textures are well preserved, in spite of the moderate alteration of feldspar. The GRV sequence is cross-cut by numerous porphyritic, rhyolite and andesite dykes. These dykes are up to 100 m wide and 10-20 km long, and mostly trend northwest to north-northeast (Giles, 1977; Blissett et al., 1993). The HS includes large batholiths and smaller intrusions of granite and minor quartz monzodiorite and quartz monzonite (Flint, 1993). Typical of much of the HS is medium-grained, locally porphyritic pink granite composed of quartz, alkali-feldspar, minor plagioclase, biotite, apatite and fluorite.

Magma chamber dynamics

A point of longstanding discussion in the study of silicic-dominated large igneous provinces is the nature of crustal magma storage, including magma chamber geometry and dynamics, and residence time of crystals before eruption. Recent studies have proposed complex models involving zoned magma chambers with variable melt to solid ratio and non-continuous (“waxing and waning”) production of melt (Hildreth, 1981; 2004; Lipman et al., 1997; Charlier et al., 2005). Addition of heat and new magma from the mantle can result in “rejuvenation” of the magma chamber (e.g. Hildreth and Wilson, 2007), causing temperature increase, magma mixing, and remelting of crystal mush (largely solid marginal portions of plutons). These variations in magma composition and temperature are potentially recorded by zoned crystals (e.g. Streck, 2008; Vazquez and Reid, 2002).

In order to address some of these points, we have carried out a study (Agangi et al., 2011) focussed on the characterisation of quartz populations in the Gawler Range Volcanics on the basis of texture, cathodoluminescence, and trace element content. Because of the moderate rock alteration, quartz is the only well preserved phenocryst phase. This study involves a wide array of quartz occurrences in different, but genetically associated, volcanic and intrusive rocks (lavas, ignimbrites, shallow and deeper intrusions) to assess the implications of the characteristics of quartz for the magma dynamics in this large igneous province.

Intragranular textures and chemical zoning of quartz

CL images can highlight cryptic intra-granular textures, undetectable in both optical and back-scattered electron (BSE) microscopy. These textures include: 1) growth-related textures (growth zones), twinning, grain shapes and growth modes (e.g. D'Lemos et al., 1997); 2) resorption-related textures, indicated by intersection relationships between growth surfaces (“unconformity”); 3) healed brittle deformation structures. Other than being an intrinsic characteristic of each mineral, CL is strongly dependent on defects in the crystal lattice, particularly point defects induced by trace element substitutions, or “activators”. Therefore, CL can be used as a proxy for trace element distribution (e.g. Müller et al., 2000; Perny et al., 1992).

Oscillatory and step zones of quartz are primary (magmatic) characteristics defined by similarity with compositional zones in plagioclase (Sibley et al., 1976). Oscillatory zones are periodic, small-scale (µm-scale) and small-amplitude variations in CL and are considered to be due to slow, diffusion-controlled crystallisation under conditions of low oversaturation (Shore and Fowler, 1996; Sibley et al., 1976). Step zones are defined as wide, non-periodic and larger-scale (≥tens of µm) variations in CL intensity. Step zones are interpreted to reflect variations in intensive parameters (P, T) and magma composition caused by processes such as crystal settling, magma convection, mixing, and reservoir replenishment (Shore and Fowler, 1996).

More than 200 crystals have been imaged by scanning electron microscopy cathodoluminescence (SEM-CL). In volcanic units, quartz grains contain three main CL step zones (zones 1-3, Fig. 2). Observed crystals consist of one (2, 3) or two zones (1, 3; 2, 3). Different combinations of these zones cause crystals to have contrasting and non-correlatable textures, even between crystals found side by side. In the dykes, the main step zones are similar and can be correlated between crystals in each dyke, although significant differences can be seen among different dykes. Intragranular growth textures include planar or wavy to deeply embayed (“vermicular”) surfaces, and truncation of growth textures.

Figure 2: Cathodoluminescence textures in the volcanic units and dykes of the lower GRV. a-c Quartz in the Lake Gairdner Rhyolite (sample GH51). A sulfide grain constituted a mechanical growth impediment (e). For a-c growth textures are highlighted. Zones (1), (2) and (3) were not found together in the same grain. Growth textures (oscillatory zones) are parallel to subhedral grain margins in zone (3), except where fractured (top of c), but are truncated by round zone boundaries or grain margins in zones (1) and (2). d-e Round mantle-rim boundary truncates the internal growth textures (arrowed), in e the core is surrounded by a discontinuous mantle.

Trace element concentrations in different CL zones were determined in core-to-rim microprobe profiles (Fig. 3). The different CL zones are characterised by different trace element contents. The total range in Ti concentration is approximately 20 to 130 ppm and Ti abundance shows a positive correlation with CL intensity. The correlation between Ti concentration and the blue ~420-nm CL emission has been found in other studies (e.g. Müller et al., 2002) This is a prominent emission that dominates panchromatic images, and justifies the use of CL brightness as a proxy for Ti distribution (e.g. Müller et al., 2005; Wark and Watson, 2006). Iron content is in the range 10-330 ppm; Al is in the range 100-680 ppm and in places it is above 3000 ppm. Aluminium and Fe abundances are not correlated with CL, and no clear correlation was found between trace elements.

Figure 3: a Quartz crystallisation temperature (TitaniQ geothermometer; Wark and Watson, 2006) compared with rutile solubility model (Hayden and Watson, 2007). Temperature of melt at the moment of trapping modelled based on zircon saturation (Watson and Harrison, 1983). Quartz crystallisation temperature modelled for aTi = 0.6 in the melt. b Reverse zoning suggests a rimwards increase of temperature.

Implications of quartz textures on magma dynamics

Succession of quartz step zones with different compositions and textures (“crystal stratigraphy”) records information on the crystallisation history. Primary (syn-crystallisation) CL textures in quartz are better preserved in rapidly cooled volcanic units and dykes of the lower GRV than in slowly cooled granite samples. Preservation of sharp Ti profiles suggests short residence time of quartz crystals at high temperature: eruption (or shallow emplacement of dykes) occurred shortly (102-103 years) after quartz crystallisation (Cherniak et al., 2007).

Different degrees of complexity can be observed in primary CL textures of quartz phenocryst. The simplest case occurs in the dykes, where zones can be correlated among quartz phenocrysts. The homogeneity of quartz populations in single dykes is interpreted as evidence that quartz crystals shared the same crystallisation history and probably crystallised largely after isolation of these small magma batches in intrusions. Embayments are common in quartz in the dykes and are mirrored by CL textures, suggesting that, in many cases, embayments had a primary (growth-related, rather than resorption-related) origin.

In the volcanic units, multiple quartz populations coexist in the same sample. Each of these populations records a complex history of crystallisation and resorption events. We propose that the volcanic units tapped a larger part of the magma characterised by a dynamic regime, which resulted in juxtaposition of different quartz populations, each with different crystallisation histories. Geothermometric estimates based on Ti content of quartz zones (Wark and Watson, 2006) suggest significant differences of quartz crystallisation temperatures (ΔT up to 70°C in volcanic units) between adjacent zones.

Several pieces of evidence are consistent with non-monotonous thermal evolution of the GRV-HS magma and suggest the occurrence of different thermal “pulses”. These include: alternating events of crystallisation and resorption (truncation of growth textures), reverse zoning (rimwards increase in Ti content) of quartz, and melting of already crystallised portions of the magma chamber (lava-hosted felsic enclaves).

The described textural and microchemical features are best explained by re-heating and convective stirring of the magma chamber (self-mixing; Couch et al., 2001). Heat input represented both the “engine” for convection and the cause of re-melting of previously crystallised magma, and was possibly supplied by underplating of mafic magma (Fig. 4).

Figure 4: Conceptual model for the crystallisation of quartz in the lower GRV magma chamber.

Modern models of felsic igneous systems agree on the fact that magma chambers are mostly composed of largely solid crystal mush with interstitial melt (e.g. Bachmann and Bergantz, 2008), mostly incapable of bulk flow (Vigneresse et al., 1996). Large crustal intrusions are assembled incrementally, via successive injections of magma and do not exist as large volumes of molten rock at one time (Glazner et al., 2004; Lipman, 2007). The mechanism proposed for the lower GRV is only apparently in conflict with existing models. In fact, mixing of crystal populations does not need the entire magma chamber to be largely molten at one time, and may occur locally in hotter volumes of magma located at the top or core of the chamber or in hot, rising plumes.

The high F composition of the Gawler SLIP magmas, and the GRV-Olympic Dam link

As part of our study of the GRV, we have reported evidence for the mobilisation of rare earth elements (REE) and high field strength elements (HFSE) in some un-mineralised rhyolite samples (Agangi et al., 2010). Accessory mineral assemblages contained in micromiaroles, vesicles and lithophysal vugs or occupying interstitial positions in the GRV (Fig. 5) include significant amounts of REE (lanthanides, U, Th), Y, HFSE (Ti, Zr, Nb) and transition metals (Cu, Zn, Mo, W), as well as F.

Figure 5: Vesicle- and micromiarole-filling assemblages in the lower GRV. a, b. Concentric zones with quartz inside the rim. Euhedral quartz and titanite contrast with the anhedral habit of epidote and chlorite (Rhyolite-dacite (Mi2), Kokatha). c. Aggregate of Y-bearing fluorite, euhedral Nb-bearing anatase and biotite infilling a micromiarolitic cavity (the walls are indicated). Assuming a growth from the walls inwards, substantiated by the orientation of crystals of anatase, anatase overgrew fluorite and biotite. d. Zircon is subhedral and completely to partially includes apatite. Needle-like crystals of REE-F-carbonate are distributed around and within fractures in apatite and fluorite. The carbonate (synchysite?) shows a predominance of LREE over HREE and locally contains Th (Moonamby Dyke Suite at Lake Everard).

Laser ablation traverses carried out across the vesicles and into the surrounding groundmass show that vesicles are enriched in REE (especially LREE and MREE), HFSE and transition elements (Cu, Zn, Mo, W), as well as Ca and Fe, compared to both the surrounding groundmass and whole rock (Fig. 6).

Figure 6: a Primitive mantle-normalised trace element composition for the vesicles, groundmass (LA-ICP-MS) and whole rock (non-vesicular sample GH34, ICP-MS). b Whole-rock-normalised base metal content of vesicles and groundmass.

These elements are mainly concentrated in REE-fluoro-carbonate (synchysite), zircon, niobian Ti oxide, fluorite, titanite and, to a lesser extent, apatite and allanite, thus indicating a correlation with F-bearing minerals. Textures and mineral associations indicate that these accessory minerals formed at a late-magmatic stage, rather than in super-solidus conditions and that they crystallised from a volatile (F, H2O, CO2, ±P, ±S)-rich low-viscosity (fluid) phase.

Melt inclusion analyses carried out in several units of the GRV indicate that the melt was enriched in F (≤1.3 wt%) in comparison with the average crust, thus suggesting a likely magmatic origin for this element.

The role of halogens (F and Cl) as complexing agents for REE, Y and HFSE has been pointed out by several studies (e.g. Keppler and Wyllie, 1990; 1991; Webster et al., 1989; Charoy and Raimbault, 1994; Audétat et al., 2000; Schönenberger et al., 2008; Pan and Fleet, 1996). A mobility of these elements is shown by both experimental studies and evidence from natural systems. Hydrothermal systems enriched in these elements occur in different geological environments and the enrichment can reach ore grades (e.g. Metz et al., 1985; Gieré and Williams, 1992; Oreskes and Einaudi, 1990; Monecke et al., 2002). Furthermore, solubility of REE in aqueous fluid has been found to increase in presence of CO2 (Wendlant and Harrison, 1979), and mobility of trace elements (including REE, Ta, Nb and Y) during brittle deformation of carbonates has been shown (Pili et al. (2002). Thus, we infer that F and CO2 have played a role in the mobilisation of these trace elements by forming complexes.

In a recent paper (McPhie et al., 2011), we show that high F contents at Olympic Dam can be traced back to high F in the Gawler SLIP magmas. We argue that this F-rich setting was crucial because any fluid involved in the Olympic Dam ore-forming hydrothermal system would have been F-rich and capable of transporting metals and REEs on the scale required to create this supergiant.

The Olympic Dam ore deposit is famous for its supergiant size and spectacular array of elements present at concentrations (0.87 wt% Cu, 0.27 kg/t U3O8, 0.32 g/t Au, 1.50 g/t Ag) (BHP Billiton, 2010) well above the crustal average (Cu 158 times, U 92 times, Au 80 times, Ag 300 times) (cf. Wedepohl, 1995). It contains in excess of 70 × 106 t of Cu alone. In addition, the combined La and Ce concentration in the ore is ~5000 ppm. The immediate host to the ore is hydrothermal breccia within granitic and volcanic rocks of the Gawler SLIP (Fig. 1) (Fanning et al., 1988; Reeve et al., 1990; Johnson and Cross, 1995; Allen et al., 2008). The connection between this SLIP and Olympic Dam has not been adequately considered in studies of ore genesis (e.g., Reeve et al., 1990; Oreskes and Einaudi, 1990; Johnson and Cross, 1995; Haynes et al., 1995).

Olympic Dam is enriched in F (Roberts and Hudson, 1983; Reeve et al., 1990; Oreskes and Einaudi, 1990; Haynes et al., 1995), and there is an intimate association between Cu sulfides and fluorite in the hematite-rich breccias. Fluorite is common, being present as disseminated crystals, crystal fragments, and in veins. Some 2.5 wt% of the ore mined consists of fluorite which means that the F content due to fluorite alone is ~108 t (22 times the crustal average of 550 ppm; Wedepohl, 1995). In addition, the principal alteration minerals have high F content (chlorite ~0.4 wt%, sericite ~0.8 wt%, apatite ~5.3 wt%). On a deposit scale, the hematite-rich breccias that contain high amounts of fluorite also have the highest abundance of Cu sulfides (Reeve et al., 1990), and Cu sulfide and fluorite typically occur together (Oreskes and Einaudi, 1990). The main REE-bearing minerals in the ore are bastnäsite, synchysite, and florencite, all of which contain F, and fluid inclusion analyses suggest the presence of F in the hydrothermal fluid (Oreskes and Einaudi, 1990, 1992).

The huge amount of F contained by this SLIP meant the Olympic Dam hydrothermal system, and others in the region, had to be F-rich. Fluorine-rich hydrothermal fluids were substantially responsible for the high metal content (including REEs) of the Olympic Dam ore deposit and its polymetallic character, and contributed to the formation of the hydrothermal breccia that hosts the ore through dissolution of volcanic and plutonic wall-rocks. The Olympic Dam case demonstrates that while LIP magmas in general are implicated as fluid and metal sources for supergiant ore deposits, simply because of the huge volumes of magma from which fluids and metals can be derived, F-rich SLIPs have superior ore potential.

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