Earth and Planetary Science Letters 428 (2015) 139–150

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Earth and Planetary Science Letters

www.elsevier.com/locate/epsl

Gold in the oceans through time

Ross R. Large a,∗, Daniel D. Gregory a, Jeffrey A. Steadman a, Andrew G. Tomkins b, Elena Lounejeva a, Leonid V. Danyushevsky a, Jacqueline A. Halpin a, Valeriy Maslennikov c, Patrick J. Sack d, Indrani Mukherjee a, Ron Berry a, Arthur Hickman e

a ARC Centre of Excellence in Ore Deposits (CODES), School of Physical Sciences, University of Tasmania, Private Bag 79, Hobart, Tasmania 7001, Australiab School of Geosciences, Monash University, Melbourne, Victoria 3800, Australiac Russian Academy of Science, Urals Branch, Miass, Russiad Yukon Geological Survey, Whitehorse, Yukon, Canadae Geological Survey of Western Australia, Perth, Western Australia, Australia

a r t i c l e i n f o a b s t r a c t

Article history:Received 7 July 2014Received in revised form 15 June 2015Accepted 9 July 2015Available online xxxxEditor: G.M. Henderson

Keywords:Au in seawatersedimentary pyriteorogenic Au depositssediment-hosted Auboring billionbanded iron formation

During sedimentation and diagenesis of carbonaceous shales in marine continental margin settings, Au is adsorbed from seawater and organic matter and becomes incorporated into sedimentary pyrite. LA-ICPMS analysis of over 4000 sedimentary pyrite grains in 308 samples from 33 locations around the world, grouped over 123 determined ages, has enabled us to track, in a first order sense, the Au content of the ocean over the last 3.5 billion years. Gold was enriched in the Meso- and Neoarchean oceans, several times above present values, then dropped by an order of magnitude from the first Great Oxidation Event (GOE1) through the Paleoproterozoic to reach a minimum value around 1600 Ma. Gold content of the oceans then rose, with perturbations, through the Meso- and Neoproterozoic, showing a steady rise at the end of the Proterozoic (800 to 520 Ma), which most likely represents the effects of the second Great Oxidation Event (GOE2). Gold in the oceans was at a maximum at 520 Ma, when oxygen in the oceans rose to match current maximum values. In the Archean and Proterozoic, the Au content of seawater correlates with the time distribution of high-Mg greenstone belts, black shales and banded iron formations, suggesting that increases in atmospheric oxygen and marine bio-productivity, combined with the higher background of Au in komatiitic and Mg-rich basalts were the first order causes of the pattern of Au enrichment in seawater. We suggest the lack of major Au deposits from 1800 to 800 Ma, is explained by the low levels of Au in the oceans during this period.

Crown Copyright © 2015 Published by Elsevier B.V. All rights reserved.

1. Introduction

Over the last 50 years, research on the origin of orogenic and sediment-hosted Au deposits has mainly focused on three ma-jor Au sources: 1) magmas (e.g., Muntean et al., 2011), 2) de-volatilization of the deep crust (15–20 km) during metamorphism of hydrated mafic rocks at the amphibolite–greenschist bound-ary (e.g., Phillips and Powell, 2010), and 3) the mantle (e.g., Hronsky et al., 2012). Gold from sedimentary rocks, particularly carbonaceous shales, has recently been considered a viable alter-native source for some Au deposit styles (Pitcairn et al., 2006;Large et al., 2007; Tomkins, 2013). In these models the Au is ultimately sourced from seawater, either as dissolved complexes, or adsorbed on particulate organic matter, Fe and Mn (hydr)ox-

* Corresponding author.E-mail address: ross.large@utas.edu.au (R.R. Large).

http://dx.doi.org/10.1016/j.epsl.2015.07.0260012-821X/Crown Copyright © 2015 Published by Elsevier B.V. All rights reserved.

ides and clays, and becomes concentrated in the shales through growth of syngenetic and diagenetic pyrite (Large et al., 2011;Tomkins, 2013). In this paper we use a novel approach to provide an indication of the changing Au content of seawater through time and suggest how this may relate to the cyclic periods of Au ore formation in greenstone belts and sedimentary basins. Our find-ings support recent work by Tomkins (2013) who identified the importance of ocean oxidation on the cycles of Au ore distribution.

2. Methodological approach

2.1. Au in sedimentary pyrite as a proxy for Au in seawater

Experimental and field studies have shown that trace elements (TE) are absorbed from seawater and local pore waters during pyrite growth, either in the water column, and/or during diage-nesis on the seafloor (Lyons, 1997), and that the level of enrich-ment is controlled by the amount of pyrite produced and the

140 R.R. Large et al. / Earth and Planetary Science Letters 428 (2015) 139–150

amount of trace metals available (Huerta-Diaz and Morse, 1992;Gregory et al., 2014). These results indicate that for a constant amount of pyrite in the shales (commonly between 1 and 4 wt%; Rickard, 2012), the TE content in sedimentary pyrite is propor-tional, in a first order sense, to their concentration in seawater (Large et al., 2014; Swanner et al., 2014). This concept has been tested for a suite of TE in sedimentary pyrite from the present day Cariaco Basin on the Venezuela shelf by Large et al. (2014). They show that the TE concentrations in pyrite correlate positively with the composition of mean global ocean water, and that TE concen-trations in sedimentary pyrite are 5–8 orders of magnitude higher than in coeval seawater. The mean Au concentration in sedimen-tary pyrite from the Cariaco Basin (313 ppb, Table S3) relative to a mean Au concentration in the present day deep oceans (0.02 ppt; http://www.mbari.org/chemsensor/summary.html) yields a concen-tration factor of about 107 for Au in sedimentary pyrite compared to coeval seawater. Thus by analyzing the Au content of sedimen-tary pyrite at any given time through Earth history, it is possible to indicate first order trends in seawater concentration through time. However, it should be noted that different depositional areas will have different chemical conditions and the 107 value should not be taken as an absolute concentration factor of Au in sedimentary pyrite globally and at all points in time throughout Earth’s history.

2.2. Effects of diagenesis and metamorphism on sedimentary pyrite textures and Au concentration

In the proof of concept paper, Large et al. (2014) demon-strated that processes during early diagenesis do not substantially change the mean first order TE content of sedimentary pyrite. However, late diagenetic and medium to high grade metamorphic processes that cause recrystallization of pyrite and conversion to pyrrhotite, may be accompanied by major changes in many TE concentrations, including Au, Ag, Te and Hg (Pitcairn et al., 2006;Large et al., 2007). In this study we have analyzed early-formed syngenetic and early diagenetic pyrite in order to evaluate the variation in seawater TE concentrations. Pyrites with a framboidal texture or clusters of micro-crystals (Fig. 1A, B) were the preferred textural type as these are considered to form either in the water column in euxinic environments, or in the top few centimeters of muds in anoxic environments (Wilkin et al., 1996). If these two styles were unavailable for analysis, then patches or nodules of porous or inclusion-rich pyrite (Fig. 1C–E), with little evidence of recrystallization, were chosen for analysis. Clearly recrystallized eu-hedral crystals or euhedral rims overgrowing porous pyrite cores (Fig. 1D–F) were avoided, as previous studies have demonstrated that recrystallized sedimentary pyrite (late diagenetic or metamor-phic) is typically inclusion poor and has a lower Au and other TE content than framboidal and porous pyrite types in the same sam-ple.

2.3. Variability of Au in the modern ocean

An assumption we make here is that the first order Au content of the past oceans has been globally homogeneous at any given time interval. Very little data are available to validate this assump-tion. Recent research suggests an average open ocean mean for the north Pacific of 0.03 ppt with a range from 0.01 to 0.06 ppt (Koide et al., 1988). These data show little difference between the deep and shallow ocean, although coastal seawater shows a marginally elevated mean of 0.04 ppt. The difference between filtered and unfiltered seawater concentrations was found to be within analyt-ical error, indicating that most of the Au is in solution or as nano particles. Falkner and Edmond (1990) compared Au ocean profiles from the North Atlantic and North Pacific and found little differ-ence with a range of 0.01 to 0.03 ppt for both profiles. However a

third profile in the Mediterranean showed Au enriched by a factor of two in the deep waters, compared with the open ocean, reach-ing a maximum of 0.04 ppt Au.

The residence time of Au in the oceans is estimated at around 1000 yrs (Falkner and Edmond, 1990), which is less than most other redox sensitive TE (e.g. Mo 760,000 yrs; As 39,000 yrs; Se 26,000 yrs; Cu 5000 yrs), but more than others; Co (340 yrs), Hg (350 yrs) and Te (100 yrs). Since the Au residence time is the same order as the mixing time of the ocean, then a reasonably homogeneous distribution of Au throughout the open ocean may be expected. However higher Au concentrations of up to an order of magnitude have been measured in seawater close to continental margins and source river systems, compared to open ocean values (Nekrasov, 1996).

2.4. Forms of Au in oceans and rivers

Although Au is assumed to be present predominantly as dis-solved complex ions in hydrothermal ore forming fluids (Seward, 1989), it may be present in seawater as a mixture of dissolved complexes, colloids, nano-particles, aqueous clusters, absorbed onto detrital clays and as Au–organic complexes. Past research has focused on the chloride complexes of Au, AuCl−4 and AuCl−2 , as the dominant species in the open ocean (e.g., Krauskopf, 1951), but recent thermodynamic considerations and measurements indicate that AuOH(H2O) is the stable species in the modern oxygenated ocean (Vlassopoulos and Wood, 1990). In river systems Au is more likely carried as suspended particles and in the colloidal state (Nekrasov, 1996). Luther III and Rickard (2005) highlight the signif-icance of aqueous clusters of metal sulfides and argue they form a major fraction of the metal load in the modern oceans and rivers. For example, Canadian rivers have reported Au concentrations from 2 to 4700 ppt, compared with open ocean concentrations at around 0.02 ppt (Falkner and Edmond, 1990). Rivers draining Au mining provinces have the highest Au concentrations. Chibisov(1964) reports average values of 10 ppt Au in the Kolyma River, draining the Shrednekansky Au district, Russian Far East, and es-timated a discharge rate of 4 tons of Au per year into the ocean. In this case the Au was transported as suspended particles ≤20 to 30 μm in size, much of which is deposited on the river delta and adjacent shelf (Chibisov, 1964). The observation that a signif-icant Au flux from rivers draining mineral districts is carried as particulate rather than dissolved Au, means that a significant por-tion of Au input to the ocean is likely deposited on continental shelves and marginal basins, with only the colloidal and dissolved Au extending into the open ocean. Therefore, since the black shales sampled in this study are principally from continental margin basins, the Au concentration in sedimentary pyrites is likely to be higher than in pyrite formed in deep ocean sediments. This sug-gests that our estimates of Au content of the paleo-oceans may be up to an order of magnitude higher than the open ocean values. However, this effect may be partially offset by factors that reduce our measured Au concentrations in sedimentary pyrite. These in-clude the effect of diagenesis and metamorphism, which both are shown to reduce the measured Au content of pyrite by up to an order of magnitude or more (Large et al., 2007).

2.5. Au contribution from seafloor hydrothermal vents

Previous studies have concluded that the TE flux from rivers is the dominant source of elements in the ocean; however a small number of elements, in particular Mn, Fe, Li and Rb may have a significant or even dominant contribution from hydrothermal vents (Elderfield and Schultz, 1996). The metal flux from seafloor hy-drothermal vents is poorly understood, and whether hydrothermal

R.R. Large et al. / Earth and Planetary Science Letters 428 (2015) 139–150 141

Fig. 1. Various textural forms of sedimentary pyrite. A) Three large framboidal pyrites (70 μm) and fine pyrite microcrystals in an organic-rich matrix, typical of syngenetic and early diagenetic pyrite. Aravalli Group, Rajasthan, India (∼1700 Ma), B) Aggregates of pyrite microcrystals and small euhedral pyrites (<20 μm) of early diagenetic origin, from Skillogalee Dolomite, South Australia (790 Ma), C) Early diagenetic nodule of pyrite composed of pyrite microcrystals intergrown with matrix clays. Blake River Group, Abitibi, Canada (∼2690 Ma), D) compacted early inclusion-rich diagenetic pyrite nodule with late diagenetic or early metamorphic overgrowth of inclusion-poor euhedral pyrites, from Roberts Mountain Formation, Nevada (420 Ma), E) Aggregates of fine microcrystals, same as (B), but partially overgrown by large euhedral metamorphic pyrites, from Skillogalee Dolomite, South Australia (790 Ma), F) Euhedral late diagenetic or metamorphic pyrites with a core of inclusion-rich early diagenetic pyrite from Khomolkho Formation, Siberia (600 Ma). The black circles are laser pits from the LA-ICPMS analysis.

activity is a net source or net sink for particular elements is un-resolved (Von Damm, 2010). This is particularly the case for Au where very little data are available. Measurements of vent fluids at the 21◦N site on the East Pacific Rise (Falkner and Edmond, 1990) returned values of 7 ppt and 50 ppt Au for two different samples. These authors estimated that seafloor hydrothermal vent fluids maybe enriched up to 1000 times with respect to normal seawater. More recent workers report values of 50 to 1400 ppt in vent fluids from various black smoker systems (e.g., Hannington et al., 2005). In modern and ancient seafloor massive sulfide deposits formed at vent sites, Au is commonly concentrated in the sulfides (particularly pyrite), and varies from 0.02 to 28 ppm Au with a mean massive sulfide grade of around 1 ppm Au (Hannington et

al., 2005). The hydrothermal plumes emanating from the black and grey smoker vents carry sulfide particles (1 to 100 μm), and the smallest of these (<10 μm) are quickly oxidized under current ocean conditions, whereas the larger particles fall out into met-alliferous ferruginous sediments at distances of 40 to 80 m from the vent (e.g., Bogdanov et al., 2006). Yucel et al. (2011) estimate that nanoparticles of pyrite make up to 10% of the filterable iron discharged by vent fluids. Plume particles of colloform pyrite, chal-copyrite and marcasite may carry minor Au and contribute to the ferruginous sediments which have been measured at sites along the Mid-Atlantic Ridge to contain 8 to 20 ppt Au (Cherkashev, 1992). The strong affinity of Au for sulfides (e.g., Reich et al., 2005), and the relatively high Au content of vent field sulfide deposits,

142 R.R. Large et al. / Earth and Planetary Science Letters 428 (2015) 139–150

suggests that most Au is probably precipitated in and around the vent sites and little is dispersed in the ocean. However a small component of soluble or colloidal Au may become adsorbed onto organic matter or taken up by bacteria (e.g., Ivanov, 1997).

2.6. Sampling and analytical methods

Laser ablation-inductively coupled plasma mass spectrometry (LA-ICPMS) was used to analyze over 4000 individual pyrite grains for TE concentrations in 308 samples of marine sediments col-lected from 33 locations, grouped over 123 determined ages, dis-tributed from the Archean (∼3515 Ma) to recent (Cariaco Basin). Table S1 provides details of each sample.

The instrumentation involved a New Wave UP-213 Nd:YAG laser microprobe coupled to an Agilent 7500a ICP-MS and a New Wave UP-193ss Nd:YAG laser microprobe coupled to an Agilent 7700s ICP-MS both housed at the University of Tasmania. Laser beam size varied between 10 and 105 μm depending on the size of the pyrite grains. Laser fluence was ∼2.5 J/cm2 for the UP193ss and ∼3.5 J/cm2 for the UP213. Laser repetition rate was 5 Hz. Each analysis involved collection of 30 s of instrument background (laser off) to properly assess detection limits (Table S4) followed by 40–60 s signal acquisition time in time-resolved mode. An in-house reference material (STDGL2b2; Danyushevsky et al., 2011) was used as the primary calibration standard for quantification of siderophile and chalcophile elements; a USGS reference mate-rial (GSD-1G; Jochum et al., 2005) was used as the primary cal-ibration standard for quantification of lithophile elements, and a natural pyrite (PPP-1; Gilbert et al., 2014) was used for quantifi-cation of sulfur. Reference materials were analyzed, each twice, consequently, every 1–1.5 h during analytical sessions to correct for instrumental drift and perform quantification using standard methods (Longerich et al., 1996). Factors contributing to preci-sion of LA-ICPMS analyses are described in detail in Gilbert et al.(2013). For this study, the main component is the heterogeneity of Au distribution within the analyzed volume. In general, the total uncertainty of individual analyses varied between 20 to 50%, which is insignificant given the overall range of concentrations between repeat analyses within the same sample. The analytical method we developed involves subtraction of the silicate matrix from the raw data to determine the composition of pyrite (Large et al., 2014). The Au concentration was measured on early-formed pyrites, such as framboids, disseminated single grains, and nodule cores. Five to twelve LA-ICPMS spot analyses were performed on each 2 cm shale fragment, for every individual sample.

3. Results

The 4006 individual pyrite analyses for Au, As, Ag, Se, Mo and Te are given in Table S2, mean data for fixed ages are in Ta-ble S3 and detection limits in Table S4. Gold concentrations were above the detection limits (0.001–0.2 ppm, Table S4) in 2620 pyrite grains varying from 0.001 to 9.90 ppm Au with an arithmetic mean of 0.26 ppm and geometric mean of 0.09 ppm (Fig. 2, Table 1). This database is an extension of that presented in Large et al. (2014)and Gregory et al. (2015b). Using the 107 concentration factor esti-mated above, the first order measured Au variation in sedimentary pyrite equates to variations in seawater Au concentration through time of <10−3 ppt to around 0.4 ppt with a geometric mean of 0.009 ppt Au, which is about one third of the modern ocean value.

The pyrite TE data obtained in this study is not normally dis-tributed in a statistical sense, rather it has an approximate log–normal distribution (see Supplementary Information, Fig. S1). As a result the arithmetic mean and standard deviation have little meaning for our data set. Instead we prefer to present the geo-

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R.R. Large et al. / Earth and Planetary Science Letters 428 (2015) 139–150 143

Fig. 2. A) Histogram of LA-ICPMS Au content in sedimentary pyrite from black shales ranging in age from 3515 to present. The Au concentration shows a log–normal relationship. B) Au–Te relationship in sedimentary pyrite.

metric mean and multiplicative standard deviation, which better describes data with log–normal distributions (Limpert et al., 2001).

3.1. Variability of Au in sedimentary pyrite from individual samples and sedimentary formations

The Au concentration in sedimentary pyrite from four drill holes through black shale formations of differing ages from Archean to Triassic have been measured in detail to determine Au variability both within individual samples and across multiple samples through different black shale formations. Two of the drill hole datasets are discussed below (Fig. 3) and the other two in the Supplementary Information (Fig. S2). It is clear from the figures that some individual samples show considerable Au variability; however, taken as a whole, the variability is consistent within each sedimentary formation. The high standard deviation of LA-ICPMS Au analyses in sedimentary pyrite from a single sample is likely due to a number of factors: 1) irregular distribution of Au in the structure of pyrite (e.g., Deditius et al., 2011), 2) presence of dis-crete Au particles in pyrite, commonly when the mean content of Au is above 10 ppm (e.g., Reich et al., 2005), 3) the analytical er-rors associated with the LA-ICPMS technique (Danyushevsky et al., 2011), and 4) variable recrystallization of sedimentary pyrite re-

Fig. 3. Downhole variation in Au content of sedimentary pyrite in two separate black shale sequences. A) DDH RI08-24 Selwyn Basin, Yukon, B) DDH ABDP9 Hamersley Basin, WA (data from Gregory et al., 2015a). Each data point represents a separate LA-ICP-MS Au analysis of pyrite. The red lines are the range of one standard de-viation from the mean. Black lines join the geometric means of each sample. Two more drill hole Au profiles are given and discussed in the Supplementary Informa-tion (Fig. S2). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

144 R.R. Large et al. / Earth and Planetary Science Letters 428 (2015) 139–150

lated to diagenesis and low grade metamorphism, which causes release of Au to the fluid phase (Large et al., 2007).

Drill hole RI08, in the Richardson Mountains, Selwyn Basin, Yukon, intersects Middle Devonian black shales of the Imperial and underlying Canol Formations (Fig. 3A; Supplementary Infor-mation Table S1). The complete pyrite dataset in this drill hole has a geometric mean pyrite Au value of 69 ppb with a multiplicative standard deviation of 2.2. The Canol Formation in the lower sec-tion of the hole (below 400 m) has pyrite with a geometric mean Au content of 62 ppb and a multiplicative standard deviation of 2.8. The overlying Imperial Formation contains pyrite with a ge-ometric mean of 71 ppb and a multiplicative standard deviation of 2.0. Although the sedimentary formation Au means are similar, the stratigraphic trends in individual sample means are reasonably distinct (Fig. 3A). For example the up-hole decreasing gold trend towards the contact between that Canol and Imperial Formations is clear, as is the general up-hole increase in gold, through the Im-perial Formation.

The second drill hole section (Fig. 3B) is from Neoarchean black shales in the Hamersley Basin of the Pilbara region in Western Australia. Drill hole ABDP9 intersects a sequence of Neoarchean (2580 to 2500 Ma) carbonaceous shales and sandstones of the Hamersley Group, below the Brockman Iron Formation (Anbar et al., 2007; Gregory et al., 2015a). The sedimentary pyrite analyzedin this sequence has a geometric mean of 176 ppb Au with a multiplicative standard deviation of 2.9. There is a clear trend of upwardly increasing Au in pyrite in the lower section of the stratigraphy (geometric mean of 50 to 550 ppb) from the top of the Paraburdoo Member to upper Bee Gorge Member, followed by a declining trend through to the Mt McRae Shale (geometric mean of 560 to 30 ppb).

3.2. Global concentration–time curves for Au

The full set of LA-ICPMS data for Au in sedimentary pyrite from global black shales over the past 3.5 billion years is shown in Fig. 4A (analytical data are presented in Table S2). The number of samples, pyrite analyses, geometric means and standard deviations, for each plotted time interval are given in Table S3. In Fig. 4A each data point is a separate pyrite analysis and the geometric mean of pyrites of a particular age are superimposed and joined by a con-tinuous line. This method of presentation was chosen as it shows all data points, the variability of the data for each plotted inter-val and the geometric mean value for each interval. A comparison of plotted Au trends using arithmetic means, geometric means and medians for each time interval demonstrates very little difference (Fig. 4E).

In developing our pyrite database and the ocean-Au time curve we have attempted to acquire an even distribution of samples through the Precambrian, with a target of at least one black shale formation sampled every 200 million years. However, this has not yet been achieved, with sample gaps from 3400 to 3000, from 2500 to 2200 and from 1250 to 1000 Ma. Notwithstanding these gaps, there are significant first order trends in the time series curve for Au. Peak Au concentrations are recorded in the Archean at around 3000, 2700 and 2550 Ma where geometric mean values exceed 400 ppb Au in pyrite. The elevated Au in the Archean is followed by a generally decreasing first order trend through the Paleoproterozoic to reach a minimum of less than 10 ppb around 1640 Ma (Fig. 4A). Between 1640 and 980 Ma the Au concen-tration in sedimentary pyrite gradually rises to a peak around 600 ppb, then falls abruptly until ca. 800 Ma. Through the Cryoge-nian and Ediacaran (850–540 Ma; Fig. 4A), Au concentration rises again to reach a maximum of over 1000 ppb in the early Phanero-zoic (520 Ma). Gold in sedimentary pyrite through the Phanerozoic seems to be cyclical and different to the patterns in the Archean

and Proterozoic, though this maybe partly due to the higher num-ber of analyses. In general terms (Fig. 4A), Au geometric mean val-ues decline from Cambrian toward present-day values, but in detail the pattern is more complex (Supplementary Information, Fig. S3). Geometric mean Au concentrations and standard deviations for the broad time intervals outlined above are presented in Table 1, to show that the trends in the Fig. 4A gold time series have a statis-tical basis. Application of the t-test to our database demonstrates that gold concentrations in pyrite in the time intervals 66–250, 540–900, 900–1400, 1400–2000, 2800–3200 and 3200–3600 Ma are statistically different at the 95% confidence level from the pre-ceding intervals. This confirms that the trends of high levels of gold in the Archean, generally low levels in the Proterozoic, returning to high levels in the Early Phanerozoic are statistically meaningful.

The geometric means and temporal trends of Te, Sb and Ag in sedimentary pyrite are compared with Au in Table S3 and Fig. 4. These three elements are chosen as they have the highest Spear-man correlation coefficients with Au (Table 2) – Au:Te (0.67), Au:Sb (0.62) and Au:Ag (0.53) – out of the 14 TE measured by LA-ICP-MS (Large et al., 2014). Selenium, Cu and Bi have positive correlation coefficients between 0.4 and 0.5, whereas all other TE have corre-lation coefficients less than 0.4 (Table 2). The temporal trends in Au, Te, Sb and Ag for the Precambrian (Fig. 4) are somewhat simi-lar. Tellurium and Sb, like Au, are generally elevated in the Meso-and Neoarchean oceans, and drop through the Paleoproterozoic to reach a minimum in the early Mesoproterozoic, rising again, but reaching a second minimum in the early Neoproterozoic. Signifi-cant temporal trends in Ag (Fig. 4D) are not so clear, and there appears to be little difference between the Archean and Protero-zoic values of Ag in sedimentary pyrite. In very general terms, Sb and Ag like Au drop by about an order of magnitude through the Phanerozoic (Fig. 4 and Supplementary Information Fig. S3).

4. Discussion

4.1. Effect of pyrite content of shales on Au concentration in sedimentary pyrite

Recent research by Gregory et al. (2015a) has demonstrated that high levels of sedimentary pyrite in black shales leads to a decrease in the Au content of individual pyrite grains compared to adjacent samples containing less pyrite. This effect is demonstrated in Fig. S4 (Supplementary Information) where samples from two drill holes RB2 and ABDP9 (Fig. 3) are compared. Although there are insufficient data to be confident of the trends, there is a gen-eral decrease in the mean Au concentration in pyrite with increas-ing pyrite content in samples from both drill hole datasets. For the Archean shales in drill hole ABDP9 there is a marked drop in Au concentration of pyrite, when pyrite content exceeds 4 wt% (2 wt% S). For the Permian/Triassic black shales the drop in Au content occurs at about 20 wt% pyrite (10 wt% S). The reason for the decrease is likely to be related to a limiting amount of Au in seawater, such that pyrites in pyrite-rich sediments (greater than 4–20 wt% pyrite), are limited by the available Au that can be taken up during their growth. We have avoided this issue as much as possible by selecting shales for analysis with less than 5 wt% pyrite; however, this is not always achievable due to the very fine grained nature of pyrite in some samples, and thus approxi-mately 10% of our sample set has exceeded the 5 wt% limit, based on microscopic studies and available bulk S analyses. A plot of the temporal Au curve (Supplementary information, Fig. S5) when the top 10% of pyrite-rich sediments (above estimated 5 wt% py) have been removed from our dataset, indicates very similar trends compared to the curve using the full dataset. In particular the Au enrichment in the Archean, as well as the minima in the Late Pale-oproterozoic and Neoproterozoic, are still clearly evident (Fig. S5).

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Fig. 4. Temporal trends in A) Au, B) Te, C) Sb and D) Ag in sedimentary pyrite. Each small dot is an LA-ICP-MS pyrite analysis. Red dots are the geometric means for each time interval. E) Comparison of arithmetic mean, geometric mean and median for Au concentration at each time interval.

Table 2Spearman correlation coefficients between Au and other trace elements in sedimentary pyrite, n = 2620.

Au Te Sb Ag Se Cu Bi As Ni Cd Co Pb Tl Zn

1 0.67 0.62 0.53 0.45 0.44 0.43 0.39 0.36 0.26 0.19 0.19 0.18 0.16

146 R.R. Large et al. / Earth and Planetary Science Letters 428 (2015) 139–150

On this basis we are confident that the first order trends in Au through time shown in Fig. 4A are not significantly affected by the variation in pyrite content of individual samples

4.2. Local versus global patterns

An important issue to test for our pyrite dataset is whether the measured Au content of sedimentary pyrite of a particular age, and by inference the Au concentration in the ocean at that time, is simply a measure of a local concentration related to lo-cal sources, or a measure of a homogeneous global ocean value at a particular time. Two factors suggest we may be measuring local concentrations. Firstly Au is considered to have a relatively short ocean residence time compared to most other redox sen-sitive TE, estimated at around 1000 years (Falkner and Edmond, 1990), and thus mixing and homogenization of Au throughout the global ocean may not be readily achieved. Secondly, our samples are from continental margin basins where seawater Au concen-trations are likely to be variable and affected by local sources. However, a close look at our dataset may suggest otherwise. The best test of a global ocean measure, is if pyrite in samples from widely geographically dispersed basins, which formed at around the same time, return Au concentrations that are similar. About 50% of our samples are from Australia, and although supercon-tinent reconstructions need to be considered, we have sufficient samples from other continental plates to be able to test the local versus global issue. In the Archean, our samples are from West-ern Australia and South Africa. Two samples of geologically very similar ages from the Hamersley Basin, Mt McRae Shale (2504 Ma) Pilbara Craton and the Transvaal Basin, Upper Nauga Formation, N2 Member (2521 Ma), Kaapvaal Craton, give geometric mean Au in pyrite values of 50 ppb and 86 ppb, respectively. However, it has been argued that the super-craton Vaalbara, may have linked the Kaapvaal and Pilbara cratons in the Neoarchean (e.g., Smirnov et al., 2013). In the mid Proterozoic where our Au time curve shows an all time minimum, samples from central Australia in the McArthur Basin, Barney Creek Formation (1640 Ma), gave mean values of 8 ppb, compared with 13 ppb from pyrite in the Satka Formation (1550 Ma), Southern Urals, Russia. In the Mesoprotero-zoic, where our Au curve (Fig. 4A) shows a rise out of the trough at 1600 Ma, two samples from different global locations show similar Au values in pyrite; Western USA, Belt Basin, Newland Formation (1470 Ma) and Northern Australia, McArthur Basin, Velkerri Forma-tion (1360 Ma), returned means of 25 and 72 ppm Au. Close to the Meso–Neoproterozoic boundary Au reaches a high with samples from Central Australia, Lillian Formation (1040 Ma) and South-ern China, Meidang Formation (980 Ma) having means of 160 ppb and 580 ppb Au respectively. The final comparison is a series of samples near the Neoproterozoic–Cambrian boundary where Au and other TE in pyrite rise again. Sample locations are Southern China, Doushantuo Formation (550 Ma) mean 146 ppb Au; West-ern Tasmania, Togari Group, Salmon River Siltstone (540 Ma) mean 87 ppb Au; Central Asia, East Tuva, Tumattaigiskaya black shale (540 Ma) mean 134 ppb Au. However two further examples show significantly different Au concentrations in similarly aged samples; 1) north Western Australia, Jeerinah Shale (2660 Ma) mean 35 ppb Au compared with central Western Australia, Early Black Flag Beds (2680 Ma) mean 288 ppb Au, and 2) Tasmania, Benjamin Lime-stone (450 Ma) mean 38 ppb compared with Scotland, Moffat Shale (444 Ma) mean 220 ppb Au in pyrite. Not withstanding these two exceptions, the marked similarity in mean Au values for most examples of globally diverse samples, at distinct time intervals, and at critical parts of the time series Au curve, is good evidence that sedimentary pyrite is preserving a first order global ocean Au sig-nal.

4.3. Source controls on Au concentration in seawater

The first order controls on the mean global Au content of sea-water, over the 10–100 million-year time spans considered here, are likely to be: 1) the Au content of continental source rocks being eroded and supplying Au to the oceans; 2) the ratio of dis-solved Au to particulate and absorbed Au in seawater; 3) the Au concentration of seafloor hydrothermal vent fluids and their activ-ity through time; 4) the oxygen concentration of the atmosphere, which effects rates of oxidative erosion releasing Au from conti-nental source sulfides; 5) chemical conditions in the ocean that control the solubility of Au complexes (e.g. pH, Eh, temperature, salinity, oxygenation, temperature, etc.); and 6) bio-productivity in the oceans that controls pyrite-forming sulfate reduction, draw-down of Au onto pyrite and organic matter and sedimentation in carbonaceous muds on the seafloor. Some of these source con-trols have been already discussed and others are considered be-low.

The primary source of most TE in seawater is continental ero-sion, whereby elements are transported in dissolved form, col-loidal form and as suspended particles via river systems, or as wind blown dust, to the ocean (e.g., Falkner and Edmond, 1990;Elderfield and Schultz, 1996). Thus the variation in the composi-tion and Au content of exposed rocks during Earth’s history should have a primary control on Au content of the oceans. Most crustal rock types average less than 2 ppm Au (Pitcairn, 2011), with the exception of carbonaceous black shales, komatiites, back arc basin basalts, some S-undersaturated continental flood basalts, and lay-ered intrusions, all of which contain up to 50 ppb Au (Keays and Scott, 1976; Pitcairn, 2011). In addition to crustal rocks, some widely dispersed ore deposit types may also provide a local source of Au from erosion. Under the reduced ferruginous conditions of the Archean ocean, hydrothermal vents may have dispersed more Au into the ocean, compared with the modern situation. Active venting between periods of submarine volcanism on the Archean mafic volcanic dominated seafloor environment would release Au as the soluble Au(HS)−2 complex into an anoxic ocean and thus likely maintain a relatively high level of dissolved Au in seawater around the vents.

The time period of maximum Au in sedimentary pyrite(>500 ppb) between 3000 and 2500 Ma, coincides with the age of abundant komatiitic volcanics (Fig. 5A, B), development of mid-to upper crustal orogenic Au deposits, and seafloor black smoker massive sulfide deposits. The general decrease in Au content of seawater through the Neoarchean and Paleoproterozoic, followed by a rise in the Mesoproterozoic, matches the peaks and troughs in the distribution of greenstone belts through this period (Fig. 5B). Tholeiites (or high Mg-basalts), which are common in greenstones in the Proterozoic, may also contain a high background of Au and thus are a potential source rock, in addition to komatiites. From 1800 to 1200 Ma, komatiites are virtually non existent (Fig. 5B), and the Au content of the ocean appears to have dropped substan-tially. This period of low Au in sedimentary pyrite (<50 ppb) in the late Paleoproterozoic and Mesoproterozoic overlaps the period termed the “Boring Billion” when very few Au deposits formed (Fig. 5D). This also corresponds with a time when the Sr and C isotopes of seawater were fairly constant indicating a balance be-tween continental and mantle sources (Condie, 2005) and a period of low tectonic activity and continental erosion. This suggests a decreased supply of Au to the oceans, with drawdown of Au into seafloor muds exceeding supply to the oceans.

4.4. Relationship of Au in seawater to Au ore deposits

The similarity in the peaks and troughs of our interpreted time series curve for Au concentration in seawater (Fig. 5A) and the

R.R. Large et al. / Earth and Planetary Science Letters 428 (2015) 139–150 147

Fig. 5. Temporal trends in key parameters related to Au content of seawater. (A) Variation of Au content in sedimentary pyrite. Each small dot represents a pyrite analysis. (B) Distribution of greenstone belts and komatiites (after Condie, 1994; Isley and Abbott, 1999; modified from Bradley, 2011). (C) Distribution of banded iron formations (Bekker et al., 2010). (D) Distribution of host rocks for sediment- and greenstone-hosted Au deposits (confined to deposits with >200 tonnes Au).

age of host rocks to sediment-hosted and orogenic Au deposits (Fig. 5D) suggests a common factor. The high levels of Au in sea-water in the Meso- and Neoarchean corresponds with the greatest periods of Au ore formation in sedimentary basins and greenstone belts. The low period of Au in seawater in the mid Proterozoic cor-responds with a period of virtually no Au deposits, and the return to high Au in seawater in the Paleozoic matches the second great period of sediment-hosted Au ore deposits (Fig. 5A, D). There are two possible interpretations for this coincidence: 1) the high Au in the oceans is due to the erosion of the newly formed major Au districts, or 2) the high Au in the oceans is the ultimate source of Au in the time related deposits. It is tempting to conclude that the first explanation is the most likely as it is simple and elegant. However there are several other factors that suggest both interpre-tations need consideration.

Fig. 5 compares the age of the Au deposit host rocks rather than the age of Au mineralization; in many cases these ages are close together, but in other cases they are not. For example, black shale samples in this study from the period 2900–3000 Ma come from the lower stratigraphy of the Witwatersrand Basin, South Africa. This includes samples of marine pyrites from the Promise, Corona-tion and Reitkuil Formations in the West Rand Group (Guy et al., 2010; Table S1). These samples have some of the highest gold in marine pyrite measured in this study (mean of 790 ppb Au), how-

ever they occur in shales stratigraphically well below the major gold reefs in the basin. Consequently erosion of the gold deposits could not have led to the high gold in these sedimentary pyrite samples. A second example are marine pyrites in black shales from the Late Ordovician Castlemaine Group in Central Victoria. The anomalous gold measured in these pyrites has been demon-strated by Thomas et al. (2011) to be of early diagenetic age (∼490 Ma; Table S1), well before the timing of gold mineraliza-tion in the Victorian Goldfields, dated in two events at 445 Ma and 380–370 Ma (Phillips et al., 2012). A third example is from the Carlin District in the Great Basin in Nevada. Our sedimentary pyrites in this locality are from the Popovich and Roberts Moun-tain Formations ranging in age from 420 to 390 Ma (Table S1). However the main gold mineralization in this district is of Tertiary age, dated at 42 to 36 Ma (Cline et al., 2005). In all three cases it is impossible for the gold in marine pyrite to be formed by ero-sion of the local gold deposits, as in each case, the gold deposits are significantly younger than the gold enrichment in the marine pyrites.

In summary, although the authors favor the second hypothesis, that the high Au in the oceans is the ultimate source of Au in the time related deposits, it is most likely that both processes have acted to enrich gold in the oceans in a cyclical manner.

148 R.R. Large et al. / Earth and Planetary Science Letters 428 (2015) 139–150

4.5. Relationship of Au in seawater to atmosphere and ocean oxygenation

Thermodynamic modeling (Vlassopoulos and Wood, 1990;Rickard and Luther III, 2006) suggests that dissolved Au in nat-ural Archean waters was predominantly as the Au(HS)−2 species, whereas in the more oxygenated Phanerozoic ocean the most sta-ble species was Au(OH)(H2O). In near shore environments and proximal to volcanic activity, particulate gold may have been a significant component. The deep oceans in the Archean were likely dominated by ferruginous waters, characterized by exten-sive amounts of dissolved Fe2+ (Holland, 1984; Canfield et al., 2008). Aqueous sulfide species are not stable in ferruginous wa-ters because pyrite forms upon interaction between these compo-nents (Planavsky et al., 2011). Therefore, since aqueous Au-chloride species are only stable at moderately oxidized and extremely acidic conditions, and Au-hydroxide is insoluble at the reduced condi-tions required for ferruginous waters (Vlassopoulos and Wood, 1990), the deep oceans away from hydrothermal vent sites are unlikely to have contained significant dissolved Au during the Archean. However, several lines of evidence suggest that episodic oxygenation of the Archean atmosphere may have commenced during the period 3.2–2.5 Ga prior to the GOE1 (e.g. Anbar et al., 2007; Large et al., 2014; Gregory et al., 2015a). These oxygenation pulses may have been short lived, but appear to have been of suf-ficient duration to effect chemical processes in the oceans. Where parts of the shallow to middle levels of the Archean oceans under-went periods of oxidation, the reduced aqueous iron species were oxidized to generate banded iron formations (BIF), which stripped the ocean of dissolved Fe2+, creating the conditions necessary for persistence of aqueous sulfide species. Thus, there were likely some shallow to middle level ocean domains, particularly in conti-nental margin basins, where Au was concentrated as the Au(HS)−2species during the Archean. This is the likely scenario in the Wit-watersrand Basin, South Africa, which hosts the greatest accumula-tion of known Au at a time (2900–3000 Ma) when a significant rise in atmosphere oxygen has been speculated based on Cr isotopes and pyrite chemistry in equivalent aged strata (Crowe et al., 2013;Large et al., 2014). Partial oxygenation of relatively shallow con-tinental marginal basins would have made ideal conditions for a change from ferruginous to sulfidic bottom waters and trapping of Au in organic- and pyrite-rich muds, before reworking along ancient shorelines to develop the highly Au enriched conglomer-ate reefs of the Witwatersrand (Large et al., 2013). The Hamersley Basin, Western Australia, is another example, at 2700–2500 Ma, where Au shows a gradual increase in pyrite in the Bee Gorge member of Wittenoom Formation shales (Fig. 3B) associated with pulses or whiffs of atmosphere oxidation (Anbar et al., 2007;Gregory et al., 2015a) and deposition of associated BIF.

The pulses of atmosphere oxygenation may also have led to increased oxidative erosion of Au enriched komatiites, and gold-bearing sulfide mineralization, resulting in increased levels of dis-solved Au contributed to the oceans, thus maintaining Au sup-ply through the late Archean. GOE1 established minor sulfate in the near-surface regions of the oceans (H2S was the dominant S species in the deep oceans; Planavsky et al., 2011). If these upper reaches of the oceans were only mildly oxidized, as might be ex-pected given that the atmosphere is considered to have contained <1% O2 after GOE1 (Canfield, 2005), the dominant oxidation state may well have been within the window of minimum Au solubil-ity (the red outline at 4 in Fig. 6). This suggestion is consistent with our data showing a significant drop in the Au concentrations in black shale pyrite at this time. This time also corresponds to a drop in the abundance of komatiitic volcanism (Fig. 5B).

Bekker and Holland (2012) suggest that following the GOE1, atmosphere oxygen reached a peak around 2300–2100 Ma be-

Fig. 6. Stability of Au and aqueous Au complexes in seawater at 25 ◦C as a function of redox potential (pE) and pH (modified from Vlassopoulos and Wood, 1990). The numbered points refer to specific sections in the text. Archean ferruginous waters were likely approximately at the position of point 1, though with salinities up to twice as high as the present ocean value (Knauth, 2005). The modern ocean varies from surface waters, which have high reduction potential (i.e., high potential to be reduced), to localized domains at point 2, where redox is controlled by sulfate re-duction. The red outline highlights an island of solid Au stability within the range of ocean chemistry.

fore dropping to a constant background Proterozoic level. Based on Cr isotopes in BIF, Frei et al. (2009) argue for a rise in atmo-sphere O2 at 1840 Ma. which coincides with another episode of BIF and minor Au deposits between 1900 and 1800 Ma (Fig. 5C). Then followed roughly a billion years with no significant BIF or Au deposits, when oxygen levels in the ocean remained below that required for the saturation of Fe(OH)3, and ferruginous to sulfidic conditions prevailed in the deep oceans for most of the period (e.g., Canfield et al., 2008; Planavsky et al., 2011). Our data in-dicate that Au build-up in the oceans did not occur again until around 1000 Ma, which was followed by renewed BIF sedimenta-tion and Au ore formation from 750 to 520 Ma, when oxygen in the atmosphere–ocean system had begun to rise again (Fig. 5).

During the second Great Oxidation Event (GOE2) at the end of the Neoproterozoic, sulfate became stable in the deep ocean (Canfield et al., 2008). Under these significantly more oxidized conditions, Au was likely present in the ocean as the more solu-ble Au(OH)(H2O) species (Vlassopoulos and Wood, 1990). Our data (Fig. 4A) show a gradual rise in Au in sedimentary pyrite by about two orders of magnitude over the period 800 to 520 Ma. The peak in gold concentration of over 1000 ppb, at 520 Ma, corresponds with the first time oxygen levels in the oceans are considered to have reached the equivalent of modern levels (Chen et al., 2015).

5. Conclusions

Here we suggest, based on the composition of sedimentary pyrite in marine black shales, that the Au content of syngenetic to diagenetic pyrite is a good proxy for the Au content of sea-water through time. Analysis of over 4000 sedimentary pyrites indicates that the Archean ocean was relatively enriched in Au compared to present day seawater, but decreased progressively through the Paleoproterozoic, with a rise and then fall through

R.R. Large et al. / Earth and Planetary Science Letters 428 (2015) 139–150 149

the Mesoproterozoic to the mid-Neoproterozoic. In the late Neo-proterozoic, Au content of the oceans rose in steps parallel with the rise in atmosphere–ocean oxygenation to be at a maximum in the early Cambrian around 520 Ma. The global coincidence in temporal trends in the Precambrian between Au in seawater and BIF deposition suggests that variations in oxygenation of the atmosphere–ocean system was the driver for the variation in Au content of seawater. Our favored model is that Au in the oceans was principally sourced from erosion of continental rocks and their associated gold mineral provinces. The oceanic gold was deposited with organic matter and trapped in syngenetic and diagenetic pyrite in seafloor continental-margin organic-bearing muds. These muds, when lithified, likely became the Au-enriched source rocks for sediment-hosted Au deposits, which formed during basin in-version some 10’s to 100’s of million years after sedimentation. Finally, the lack of Au deposits during the boring billion (1800 to 800 Ma) may be explained by low levels of Au in the oceans, lead-ing to the deposition of Au-poor black shale source rocks.

Acknowledgements

We wish to acknowledge colleagues who have supplied sam-ples for this study: S Johnson, C. Makoundi, T. Lyons, P. McAurick, S. Bull, P. Haines, C. Calver, B. Guy, R. Scott, R. Coveney Jr., J. Abbott, D. Huston, A. Lambeck, R. Batchelor, S. Smith, M. Krupenin, J. Shar-rock, P. Sorjonen-Ward, L. Leonova, V. Murzin, K. Ivanov, R. Stein, J. Slack, K. Kelley V. Lisenko and S. Karpov. Thanks to the core li-brary staff of the Geological Survey of Western Australia (GSWA) who kindly assisted in our sampling from drill core. Thanks also to David Cooke, David Rickard and an anonymous EPSL reviewer for their comments that helped to improve the manuscript. A. Hick-man publishes with permission of the Executive Director, GSWA. Funding was provided by an ARC Centre of Excellence grant to RRL and ARC Discovery Grant DP150102578 to RRL, JAH and LVD.

Appendix A. Supplementary material

Supplementary material related to this article can be found on-line at http://dx.doi.org/10.1016/j.epsl.2015.07.026.

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