Gold in the oceans through time

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, Australia School of Geosciences, Monash University, Melbourne, Victoria 3800, Australia c Russian Academy of Science, Urals Branch, Miass, Russia d Yukon Geological Survey, Whitehorse, Yukon, Canada e Geological Survey of Western Australia, Perth, Western Australia, Australia b

a r t i c l e

i n f o

Article history: Received 7 July 2014 Received in revised form 15 June 2015 Accepted 9 July 2015 Available online 31 July 2015 Editor: G.M. Henderson Keywords: Au in seawater sedimentary pyrite orogenic Au deposits sediment-hosted Au boring billion banded iron formation

a b s t r a c t 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. LAICPMS 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 major Au sources: 1) magmas (e.g., Muntean et al., 2011), 2) devolatilization of the deep crust (15–20 km) during metamorphism of hydrated mafic rocks at the amphibolite–greenschist boundary (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 alternative 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: [email protected] (R.R. Large).

http://dx.doi.org/10.1016/j.epsl.2015.07.026 0012-821X/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 findings 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 diagenesis on the seafloor (Lyons, 1997), and that the level of enrichment is controlled by the amount of pyrite produced and the

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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 proportional, 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 concentrations in sedimentary pyrite are 5–8 orders of magnitude higher than in coeval seawater. The mean Au concentration in sedimentary 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 concentration factor of about 107 for Au in sedimentary pyrite compared to coeval seawater. Thus by analyzing the Au content of sedimentary 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) demonstrated 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 euhedral 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 metamorphic) is typically inclusion poor and has a lower Au and other TE content than framboidal and porous pyrite types in the same sample. 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 assumption. 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 analytical 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 difference 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, reaching 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 dissolved 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 AuCl2 , as the dominant species in the open ocean (e.g., Krauskopf, 1951), but recent thermodynamic considerations and measurements indicate that AuOH(H2 O) 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 significance 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 estimated 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 significant Au flux from rivers draining mineral districts is carried as particulate rather than dissolved Au, means that a significant portion 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 suggests 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 include 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 hydrothermal vents is poorly understood, and whether hydrothermal

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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 unresolved (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 metalliferous 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, chalcopyrite 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,

3200

3600

2800

3200

Notes: 1. T -test assumes normally distributed data. Therefore the sample data has been log transformed so that it is approximately normal in its distribution. 2. Original reports on T test are on work Sheet l (Supplementary Information).

Statistically different 15.7 137 0.25

−1.06 0.27 0.03 3.16 0.50

2800 2500

137

−2.01

0.086

169 0.25

−0.15 2.21 0.22 3.19 0.693 0.50

−2.04

2500 2000

169

0.54

2000 1400

340

−1.92

0.146

3.44

0.04

0.50

−0.84

0.29

340

13.7

1.25

1.97

Statistically different

1.98

1.96

May be same

Statistically different 5.41 26 0.29

−0.97 0.37 0.03 3.44 0.54

1400 900

26

−2.07

0.106

106 0.25

−1.58 0.08 0.01 3.17 0.027 0.50

−2.11

900 540

106

0.46

540 250

140

−0.67

0.202

2.92

0.07

0.59

−0.7

0.22

140

14.2

9.27

1.97

1.98

Statistically different

Statistically different 1.97

May be same 0.72 115 0.18

−1.22 0.16 0.02 2.69 0.060 0.43

−1.52

250 66

115

0.60

66 0

1316

−2.36

0.195

3.99

0.05

0.78

−1.17

0.36

1316

2 .1

0.43 266 0.43

−1.09 0.37 0.02 4.58 0.66

−1.69

0.081

5 0.57

−1.22 0.34 0.01 5.71 0.76 5

−1.15

0.060

1.96

1.96

May be same 1.97

Result t crit t stat Count Variance log Au Mean log Au

One std dev above geometric mean One std dev below geometric mean Multiplicative std dev Geometric mean Au (ppm) Std dev log Au log Au n

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 Table 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 estimated 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 distributed 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-

Max age (Ma)

3. Results

Min age (Ma)

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 collected from 33 locations, grouped over 123 determined ages, distributed 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 inhouse reference material (STDGL2b2; Danyushevsky et al., 2011) was used as the primary calibration standard for quantification of siderophile and chalcophile elements; a USGS reference material (GSD-1G; Jochum et al., 2005) was used as the primary calibration standard for quantification of lithophile elements, and a natural pyrite (PPP-1; Gilbert et al., 2014) was used for quantification 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 precision 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.

Table 1 Statistical analysis of Au concentration in sedimentary pyrite at different time periods in earths history, with a T -test measure of statistical significance.

2.6. Sampling and analytical methods

Results of T -test when both samples have the same variance

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).

Statistically different

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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 discrete Au particles in pyrite, commonly when the mean content of Au is above 10 ppm (e.g., Reich et al., 2005), 3) the analytical errors 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 deviation 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 Information (Fig. S2). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

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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 Information 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 section 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 geometric 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 Imperial 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 analyzed in 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).

and Proterozoic, though this maybe partly due to the higher number of analyses. In general terms (Fig. 4A), Au geometric mean values 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 statistical 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 preceding 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 Spearman 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 correlation coefficients less than 0.4 (Table 2). The temporal trends in Au, Te, Sb and Ag for the Precambrian (Fig. 4) are somewhat similar. Tellurium and Sb, like Au, are generally elevated in the Mesoand 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. Significant temporal trends in Ag (Fig. 4D) are not so clear, and there appears to be little difference between the Archean and Proterozoic 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).

3.2. Global concentration–time curves for Au 4. Discussion 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 continuous line. This method of presentation was chosen as it shows all data points, the variability of the data for each plotted interval 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 concentration in sedimentary pyrite gradually rises to a peak around 600 ppb, then falls abruptly until ca. 800 Ma. Through the Cryogenian and Ediacaran (850–540 Ma; Fig. 4A), Au concentration rises again to reach a maximum of over 1000 ppb in the early Phanerozoic (520 Ma). Gold in sedimentary pyrite through the Phanerozoic seems to be cyclical and different to the patterns in the Archean

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 general decrease in the mean Au concentration in pyrite with increasing 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 approximately 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 Paleoproterozoic 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 2 Spearman 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

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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 local 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 sensitive 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 concentrations 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 supercontinent 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 Western 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 Mesoproterozoic, 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 Formation (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 Southern 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; Western 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 Limestone (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 signal.

4.3. Source controls on Au concentration in seawater The first order controls on the mean global Au content of seawater, 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 dissolved Au to particulate and absorbed Au in seawater; 3) the Au concentration of seafloor hydrothermal vent fluids and their activity through time; 4) the oxygen concentration of the atmosphere, which effects rates of oxidative erosion releasing Au from continental 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, drawdown of Au onto pyrite and organic matter and sedimentation in carbonaceous muds on the seafloor. Some of these source controls have been already discussed and others are considered below. The primary source of most TE in seawater is continental erosion, whereby elements are transported in dissolved form, colloidal 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 composition 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 layered 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 midto 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 substantially. 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 between 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

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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 seawater 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 corresponds 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 interpretations 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, Coronation 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 demonstrated by Thomas et al. (2011) to be of early diagenetic age (∼490 Ma; Table S1), well before the timing of gold mineralization 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 Mountain 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 erosion 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.

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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 natural Archean waters was predominantly as the Au(HS)− 2 species, whereas in the more oxygenated Phanerozoic ocean the most stable species was Au(OH)(H2 O). 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 extensive amounts of dissolved Fe2+ (Holland, 1984; Canfield et al., 2008). Aqueous sulfide species are not stable in ferruginous waters because pyrite forms upon interaction between these components (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 conditions 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 sufficient duration to effect chemical processes in the oceans. Where parts of the shallow to middle levels of the Archean oceans underwent 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 continental margin basins, where Au was concentrated as the Au(HS)− 2 species during the Archean. This is the likely scenario in the Witwatersrand Basin, South Africa, which hosts the greatest accumulation 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 continental 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 conglomerate 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 goldbearing sulfide mineralization, resulting in increased levels of dissolved Au contributed to the oceans, thus maintaining Au supply through the late Archean. GOE1 established minor sulfate in the near-surface regions of the oceans (H2 S 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 expected 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 solubility (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 reduction. 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 atmosphere 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 indicate that Au build-up in the oceans did not occur again until around 1000 Ma, which was followed by renewed BIF sedimentation 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 soluble Au(OH)(H2 O) 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 seawater 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

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the Mesoproterozoic to the mid-Neoproterozoic. In the late Neoproterozoic, 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 inversion 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, leading to the deposition of Au-poor black shale source rocks. Acknowledgements We wish to acknowledge colleagues who have supplied samples 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. Sharrock, P. Sorjonen-Ward, L. Leonova, V. Murzin, K. Ivanov, R. Stein, J. Slack, K. Kelley V. Lisenko and S. Karpov. Thanks to the core library 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. Hickman 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 online at http://dx.doi.org/10.1016/j.epsl.2015.07.026. References Anbar, A.D., Duan, Y., Lyons, T.W., Arnold, G.L., Kendall, B., Creaser, R.A., Kaufman, A.J., Gordon, G.W., Scott, C., Garvin, J., Buick, R., 2007. A whiff of oxygen before the great oxidation event? Science 317, 1903–1906. Bekker, A., Holland, H.D., 2012. Oxygen overshoot and recovery during the early Paleoproterozoic. Earth Planet. Sci. Lett. 317–318, 295–304. Bekker, A., Slack, J.F., Planavsky, N., Krapež, B., Hofmann, A., Konhauser, K.O., Rouxel, O.J., 2010. Iron formation: the sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric processes. Econ. Geol. 105, 467–508. Bogdanov, Y.A., Lisitsin, A.P., Sagalevich, A.M., Gurvich, E.A., 2006. Hydrothermal Ore Genesis in Ocean Floor. Nauka, Moscow, 527 pp. (in Russian). Bradley, D.C., 2011. Secular trends in the geologic record and the supercontinent cycle. Earth-Sci. Rev. 108, 16–33. Canfield, D.E., 2005. The early history of atmospheric oxygen: homage to Robert M. Garrels. Annu. Rev. Earth Planet. Sci. 33, 1–36. Canfield, D.E., Poulton, S.W., Knol, A.H., Narbonne, G.M., Ross, G., Goldberg, T., Strauss, H., 2008. Ferruginous conditions dominated later Neoproterozoic deepwater chemistry. Science 321, 949–952. Chen, X., et al., 2015. Rise to modern levels of ocean oxygenation coincide with the Cambrian radiation of animals. Nat. Commun. 6, 7142. http://dx.doi.org/ 10.1038/ncomms8142(2015). Cherkashev, G.A., 1992. Geochemistry of metalliferous sediments from areas of ore formation in the ocean. In: Gramberg, I.S., Ainemer, A.I. (Eds.), Hydrothermal Sulfide Ores and Metalliferous Sediments of the Ocean. Nedra, St. Petersburg, pp. 138–152 (in Russian). Chibisov, N.A., 1964. Gold migration in the waters of the drainage system of Kolyma–Indgirka region. In: Geology and Economic Minerals of Yakutsk Autonomous Soviet Socialist Republic, vol. 6, pp. 112–128 (in Russian).

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