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Earth and Planetary Science Letters 255 (2007) 243–256 www.elsevier.com/locate/epsl

Suboxic deep in the late : Evidence from hematitic and formation related to seafloor-hydrothermal deposits, central , USA ⁎ J.F. Slack a, , T. Grenne b, A. Bekker c, O.J. Rouxel d, P.A. Lindberg e

a U.S. Geological Survey, National Center, MS 954, Reston, VA 20192, USA b Geological Survey of Norway, Leiv Eirikssons vei 39, 7491 Trondheim, Norway c Geophysical Laboratory, Carnegie Institution of Washington, 5251 Broad Branch Road N.W., Washington, D.C. 20015, USA d Department of Marine Chemistry and Geochemistry, Woods Hole Oceanographic Institution, MS 25, Woods Hole, MA 02543, USA e 205 Paramount Drive, Sedona, AZ 86336, USA Received 27 September 2006; received in revised form 11 December 2006; accepted 13 December 2006 Available online 22 December 2006 Editor: H. Elderfield

Abstract

A current model for the evolution of deep seawater composition involves a change from anoxic sulfide-free to sulfidic conditions 1.8 Ga. In an earlier model the deep ocean became oxic at that time. Both models are based on the secular distribution of (BIF) in shallow marine sequences. We here present a new model based on rare earth elements, especially redox-sensitive Ce, in hydrothermal silica–iron sediments from deeper-water, open-marine settings related to volcanogenic massive sulfide (VMS) deposits. In contrast to , Paleozoic, and modern hydrothermal sediments, 1.74 to 1.71 Ga hematitic chert () and iron formation in central Arizona, USA, show moderate positive to small negative Ce anomalies, suggesting that the redox state of the deep ocean then was at a transitional, suboxic state with low concentrations of dissolved O2 but no H2S. The presence of jasper and/or iron formation related to VMS deposits in other volcanosedimentary sequences ca. 1.79–1.69 Ga, 1.40 Ga, and 1.24 Ga also reflects oxygenated and not sulfidic deep ocean waters during these time periods. Suboxic conditions in the deep ocean are consistent with the lack of shallow-marine BIF ∼1.8 to 0.8 Ga, and likely limited nutrient concentrations in seawater and, consequently, may have constrained biological evolution. © 2006 Elsevier B.V. All rights reserved.

Keywords: Proterozoic; deep ocean redox state; hematitic chert; iron formation; rare earth elements; Ce anomalies

1. Introduction terrestrial sulfate flux to the ocean due to oxidative continental weathering. The redox state of the deep Atmospheric levels rose twice in the Prote- ocean during this interval is still debated. Holland [3] rozoic Eon, between 2.47 and 2.32 Ga [1] and after used the disappearance of banded iron formation (BIF) 0.8 Ga [2]. The intervening time period is marked by in shallow-marine settings ca. 1.8 Ga to argue that the moderate levels of atmospheric oxygen and a substantial deep ocean was fully oxygenated after the rise of at- mospheric oxygen, whereas Canfield [4] suggested it ⁎ Corresponding author. was anoxic and sulfidic (euxinic) based on three-box E-mail address: [email protected] (J.F. Slack). modeling of the oxygen content of deep ocean water and

0012-821X/$ - see front matter © 2006 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2006.12.018 244 J.F. Slack et al. / Earth and Planetary Science Letters 255 (2007) 243–256 on isotope data for sedimentary . The two iron formation that can provide a window into the redox models are based on the assumption that insoluble iron state of the deep ocean during this time period. Recent oxyhydroxides or iron sulfides, respectively, formed in studies [16,17] show that such laterally persistent jasper the deep ocean and prevented the transport of dissolved formed from silica–iron gels that precipitated largely iron of seafloor-hydrothermal origin to shallow-marine from the non-buoyant parts of submarine-hydrothermal settings. plumes, due to the oxidation of dissolved hydrothermal The sulfidic deep ocean model has found increasing ferrous iron (Fe2+) to insoluble ferric iron (Fe3+) oxyhy- support from recent studies of S and Mo isotopes, Fe droxide that promoted flocculation of seawater-derived speciation, and biomarkers in late Paleoproterozoic amorphous silica. The formation of such -rich and sedimentary successions of North seafloor deposits implies that the redox state of ambient America and northern [5–11]. Biomarker seawater was at or above the oxidation state of ferrous studies [11] suggest that sulfidic conditions extended iron. VMS-related hematitic iron formations also likely even into the . This model has critical impli- formed by the precipitation of iron oxyhydroxide par- cations for biological evolution because under sulfidic ticles from submarine-hydrothermal plumes [18]. The conditions, some trace metals (e.g., Fe, Mo, and Cu) are setting of these volcanic-hosted, deep-water “Algoma- efficiently scavenged with precipitating sulfides to sedi- type” iron formations [19] is distinctly different from ments, leaving seawater depleted in elements that are that of the sediment-hosted, “Superior-type” BIFs, which essential for enzymes involved in the biological nitrogen consist mainly of interlayered chert and iron in cycle [12]. Because eukaryotes cope poorly with N- shallow-water sedimentary sequences [19]. limitation and would be particularly affected by these In this paper we investigate the geochemistry of late conditions, the ca. 1 b.y. biological stasis in their evol- Paleoproterozoic jasper and iron formation in the Je- ution ca. 1.8 to 0.8 Ga has generally been interpreted rome mining district of Arizona, with a focus on redox- [12] as reflecting deep-ocean sulfidic conditions; alter- sensitive Ce. These data are used for testing current natively, this stasis might reflect intrinsic evolutionary models for the existence of sulfidic deep seawater conservatism in the absence of natural selection im- during the Proterozoic, and for introducing a new model posed by later animals [13]. that argues for suboxic conditions in the deep part of the Most geochemical arguments for marine sulfidic late Paleoproterozoic ocean. conditions during the late Paleoproterozoic and Meso- proterozoic are based on studies of sedimentary se- 2. The Paleoproterozoic Jerome mining district of quences that were deposited in intracratonic settings, Arizona including S isotope data for sedimentary sulfate and sulfide that are consistent with low seawater Jasper and iron formation form beds that are spatially sulfate contents [4]. Significantly, direct indicators of and temporally related to seafloor VMS deposits within such conditions in deep-water, clearly open-marine set- a late Paleoproterozoic submarine volcanosedimentary tings during this time period are lacking. sequence of oceanic arc affinity in the Jerome mining isotope data [10] potentially constrain Mo burial in oxic district of central Arizona (Figs. 1 and 2) [22,23,27]. versus suboxic/sulfidic settings and, consequently, This sequence comprises a 1.0 to 2.5-km thick suc- ocean redox conditions. However, Mo isotope data do cession of submarine volcanic and volcaniclastic rocks. not distinguish between sulfidic and suboxic deep-water Basal strata consist of the Shea Basalt together with conditions [10,14]; suboxic waters have low concen- andesite, dacite, and rhyolite, which are overlain by trations of dissolved O2 (b5 μmol/l [15] in contrast to rhyolite flows and domes of the Deception Rhyolite, typical values of 150–200 μmol/l in modern oxygenated rhyolite flows, domes, , and tuffs of the Cleo- deep oceans), but no H2S. Rare earth elements (REE), patra Rhyolite, and an upper unit of volcaniclastic tur- and specifically Ce anomalies, in chemical sediments bidites and minor rhyolites of the Grapevine Gulch like BIF and can provide important con- Formation. A U–Pb age of 1738.5±0.5Ma has been straints on the redox state of seawater, but such sedi- determined recently by S.A. Bowring on igneous zir- ments are rare in open-marine sequences between 1.8 cons from the upper part of the Cleopatra Rhyolite (see and 0.8 Ga. Supplementary Table 1, Supplementary Fig. 1). Jasper Some deep-ocean volcanogenic massive sulfide and iron formation occur along the contact of the Shea (VMS) deposits of Paleoproterozoic and Mesoproter- Basalt and the Deception Rhyolite, and along the base of ozoic age have related chemical sediments including and interbedded with the Cleopatra Rhyolite, where extensive beds of hematitic chert (jasper) and hematitic jasper locally forms rip-up clasts within volcaniclastic J.F. Slack et al. / Earth and Planetary Science Letters 255 (2007) 243–256 245

Fig. 1. Sketch map showing the distribution of volcanogenic massive sulfide (VMS) deposits and jasper and iron formation beds in late Paleoproterozoic volcanosedimentary sequences of central Arizona, USA. Proterozoic intrusions and rocks are not shown. Geology and VMS deposits from [20] and [21], respectively; distribution of jasper and iron formation from [22–25] and [26], respectively. . Extensive beds of interlayered jasper and hematitic iron-oxidizing bacteria preserved in Ordovician and iron formation up to 3 m thick also occur in the lower younger hydrothermal jasper and Si–Fe oxyhydroxide part of the Grapevine Gulch Formation. Regional meta- deposits [28,29]. Based on their stratiform geometry and morphic grade is lower greenschist facies [20]. The close association with submarine volcanic rocks, the large ca. 110 million metric tonne (Mt) United Verde massive sulfide and related deposits of jasper and iron Cu–Zn–Au–Ag deposit, consisting of with vari- formation in the Jerome district formed contemporane- able amounts of and above a ously with their host volcanic and sedimentary rocks, chalcopyrite-rich footwall feeder zone, has locally asso- like syngenetic hydrothermal deposits in other ancient, ciated jasper in beds b1 m thick above and marginal to and modern, volcanosedimentary sequences [30]. the pyritic top of the orebody [22]. The much smaller Water depth of the Jerome and iron forma- (b1Mt) Chief Cu–Au–Ag deposit, at an equi- tions can be inferred from the Cu-rich composition of valent stratigraphic level 5 km to the south, contains the associated sulfide and from the nature of sur- jasper beds in the immediate footwall and marginal rounding volcanic and sedimentary rocks. The solubility to the sulfides, and interlayered with iron formation in of Cu in seafloor-hydrothermal fluids is strongly b1-m thick beds that extend nearly 1 km laterally along temperature-dependent, and precipitation of Cu-rich strike [23]. Jasper beds at the contact of the Deception sulfide deposits like those of the Jerome district [22] on Rhyolite and overlying Cleopatra Rhyolite (Fig. 2) are at or near the seafloor is practically limited to hydrother- the same stratigraphic level as the nearby Verde Central mal solutions with temperatures of N300 °C [31,32]. Cu–Au–Ag massive sulfide deposit [22]. Assuming a minimum temperature of 300 °C and a Jasper and iron formation comprise essentially seawater salinity of 3.2 M NaCl for the hydrothermal and hematite. They are little recrystallized, and primary solutions, which is similar to the salinities of most depositional and early diagenetic textures are generally modern seafloor-hydrothermal systems [32], a hydro- well preserved, including hematitic filaments 1 to 3 μm static pressure equivalent to ∼850 m water depth is in diameter and 30–50 μm long (Fig. 3). These fila- required to prevent fluid boiling and near-total loss of ments are morphologically very similar to remains of Cu by precipitation of chalcopyrite at depth in the 246 J.F. Slack et al. / Earth and Planetary Science Letters 255 (2007) 243–256

Fig. 2. Schematic geological section of Paleoproterozoic rocks in the Jerome district and vicinity, Arizona, showing stratigraphic relationship of jasper and iron formation to volcanic and sedimentary rocks and coeval . Left and right panels show features for southern and northern parts of the district, respectively, which are 5 km apart. Note that jasper and/or iron formation in the district occur at four stratigraphic levels. Total thickness of the exposed Paleoproterozoic sequence ranges from ∼1 to 2.5 km; thickness from upper part of the Cleopatra Rhyolite to lower part of the Grapevine Gulch Formation is b100 m. Size and thickness of massive sulfide, jasper, and iron formation are exaggerated for clarity. Stratigraphic relations and nomenclature shown here are based on published and unpublished geological mapping by P.A.L., and supersede those of previous studies [22,23]. hydrothermal upflow zone [31,32]. In modern seafloor 3. Geochemistry of Jerome district jasper and iron hydrothermal systems on mid-ocean ridges and intra- formation oceanic arcs, high-temperature (N300 °C) venting and significant Cu-rich mineralization are unknown at All analyzed samples come from surface outcrops depths of less than 1000 m [30,31]. Thus 850 m is a and are distal (N200 m) from inferred hydrothermal conservative estimate of the minimum water depth for vents on the paleoseafloor. Jasper consists mainly of deposition of massive sulfide, jasper, and iron formation silica (SiO2 =87.3–93.4wt.%) and iron (∑Fe2O3 =3.4– in the Jerome district. Although 850 m is shallower than 10.8wt.%), with low concentrations of lithophile ele- the abyssal depths of N2000 m that dominate the mod- ments such as Al, Ti, Zr, Th, and Sc that characterize ern oceans, it is deeper than photic zones (b200 m) and detrital sediments (Table 1)1. Hematitic iron formation continental shelves (b600 m). A deep-water origin for is less siliceous (SiO2 =11.2–51.7wt.%) and has higher the Jerome sequence is also consistent with the abun- concentrations of iron (∑Fe2O3 =45.3–86.5wt.%) and dance of distal arc-supplied turbidites and the absence of lithophile elements. If the concentrations of these continentally-derived sediments and shelf deposits in lithophile elements represent detrital material only and the Grapevine Gulch Formation. A possible modern the detrital component is similar in composition to as- analogue of the Jerome sequence is the Brothers sub- sociated tuffaceous argillite and rhyolite tuff (Table 1), marine volcano, which is part of an oceanic arc in the the lithophile element contents would reflect b3wt.% southwestern Pacific Ocean where Zn–Cu–Pb massive sulfides occur at a water depth of 1850 m together with 1 Supplementary Table 1 contains geochemical data for all samples active hydrothermal vents and related particle-rich including detailed information on location, lithology, and stratigraphic plumes in the water column [33]. setting. J.F. Slack et al. / Earth and Planetary Science Letters 255 (2007) 243–256 247

tive to moderate positive Ce anomalies (Fig. 4; Table 1). Eu anomalies, expressed as Eu/Eu⁎, are calculated here ⁎ using Eu/Eu =2EuSN /(SmSN +GdSN), where SN repre- sents normalization to Post-Archean Australian (PAAS) [38]. Ce anomalies (Ce/Ce⁎) are calculated as ⁎ Ce/Ce =2CeSN /(LaSN +PrSN) with normalization also to PAAS. Consideration must be given to whether the REE patterns reflect seawater composition or post-deposi- tional processes and redox states. For ferric-oxide sedi- ments, significant mobilization and fractionation of REE may occur in response to reductive dissolution of the carrier oxyhydroxides during anoxic or suboxic diagen- esis, leading to a preferential release of the LREE, and in Fig. 3. Photomicrograph of hematitic filaments in jasper from the particular Ce, to pore waters [41] and a corresponding Jerome district; sample is from the contact of the Deception Rhyolite and the lower unit of the Cleopatra Rhyolite (Fig. 2). These filaments depletion in the sediment. A positive correlation be- are similar to those in Phanerozoic and modern iron oxide sediments of tween REE and total Fe contents in our samples, com- submarine-hydrothermal origin [28,29]. bined with similar LREE/HREE and REE/Fe ratios for jasper and hematitic iron formation (Table 1), is con- sistent with a model in which REE abundances are contribution of detrital components to the iron forma- controlled primarily by iron oxyhydroxide scavenging tions and b1.5wt.% to the jaspers. These are maximum [42] with most samples showing insignificant diagenetic estimates because in hydrothermally-derived, iron- modification of REE patterns. This interpretation is oxyhydroxide sediments lithophile elements such as in accord with the predominance of ferric iron oxide Al and Th also have a significant hydrogenous com- in the majority of samples, which implies that reduc- ponent that reflects adsorption of oxyanions and hy- tive dissolution of iron was minimal. Samples that show drolyzed ionic species from seawater [17,34]. textural or geochemical evidence of Fe and REE mob- MnO concentrations are uniformly low (b0.1wt.%) ilization have been rejected from the data set (see in both jasper and iron formation. Mn/Fe ratios range Supplementary Methods). from 6×10− 4 to 2×10− 2. These ratios are lower than Alteration of the REE during post-diagenetic pro- the Mn/Fe ratios of most hydrothermal plume particles cesses is also considered unlikely, given the generally and plume-related, iron oxyhydroxide sediments in accepted low mobility of REE under low water/rock modern oxygenated oceans including those of Mid- conditions during low-grade regional metamorphism Atlantic Ridge vent sites [35,36] that have the lowest [43], and the preservation of abundant primary textures Mn/Fe ratios for such precipitates. Concentrations of in the jaspers that exclude the possibility of late fluid- chalcophile elements are very low (e.g., Cub1–30 ppm; rich alteration. Moreover, the REE pattern of associated avg=8 ppm) in jasper and iron formation of the Grape- rhyolite at Jerome (Supplementary Table 2) is similar to vine Gulch Formation where no associated massive those of modern submarine rhyolites in having negative sulfide deposits are known, whereas jasper and iron Eu anomalies and no Ce anomalies, thus arguing against formation up to 1 km laterally from the Copper Chief extensive alteration of REE concentrations during re- massive sulfide deposit have somewhat higher Cu gional metamorphism. In addition, relatively high Y/Ho values (13–176ppm; avg=61 ppm). This pattern is (wt) ratios of ca. 33–42 are closer to those of seawater consistent with a plume model whereby the proportion (Y/Ho ∼40–80) than to those of shale and continental of large and fast-settling sulfides relative to fine-grained, crust (Y/Ho ∼27) [44], thus providing further support iron oxyhydroxide particles decreases away from the for a seawater-dominated source of the REE. REE hydrothermal vents [37]. patterns in the Jerome district samples are therefore REE patterns for jasper and iron formation from the considered primary, recording interaction of hydrother- Jerome district normalized to Post-Archean Australian mal plume particles with seawater. Indeed, except for Shale [38] show REE abundances of 0.01 to 0.7× shale the Ce anomalies, the patterns closely resemble those of values, moderate depletion in the light REE (LREE), modern plume-related iron oxyhydroxide particles and enrichment of the middle to heavy REE (HREE), insig- sediment as well as Phanerozoic hydrothermal jasper nificant to large positive Eu anomalies, and small nega- beds (Fig. 4). 248 J.F. Slack et al. / Earth and Planetary Science Letters 255 (2007) 243–256

Table 1 Compositions of representative samples of late Paleoproterozoic jasper and iron formation from central Arizona⁎ Sample no. JS-05-37 JS-05-39 JS-05-32 JS-05-42 JS-05-48A-1 JS-05-48C JS-05-55 JS-05-58 JS-05-60 BE-05-7 Rock type Jasper Jasper Iron fm Jasper Jasper Iron fm Jasper Jasper Jasper Iron fm Location C. Chief C. Chief C. Chief C. Chief M. Mtn M. Mtn Mf. Cyn. H. Mtns H. Mtns H. Mtns

SiO2 (wt.%) 87.30 91.75 11.24 91.27 91.98 23.25 86.48 88.72 93.46 19.98 TiO2 0.014 0.007 0.014 0.011 0.007 0.047 0.003 0.003 0.003 0.051 Al2O3 0.16 0.21 0.31 0.31 0.14 1.42 0.08 0.22 0.06 1.38 T Fe2O3 8.80 3.41 86.55 6.90 7.62 71.50 8.87 8.65 5.60 75.28 MnO 0.058 0.066 0.050 0.035 0.021 0.044 0.176 0.565 0.205 0.303 MgO 0.07 0.42 0.04 0.11 b0.01 0.21 0.25 0.02 b0.01 b0.01 CaO 1.24 1.57 1.18 0.18 0.11 1.48 1.25 0.41 0.03 0.10

Na2O 0.08 0.07 0.08 0.10 0.07 0.07 0.07 0.08 0.06 0.12 K2O 0.06 0.05 0.06 0.08 0.05 0.16 0.02 0.10 0.08 0.45 P2O5 0.04 b0.01 0.77 0.03 0.01 0.81 0.02 0.05 0.02 0.04 LOI 1.4 2.42 −0.66 0.80 −0.26 0.99 2.86 0.58 0.26 1.41 Total 99.22 99.98 99.63 99.83 99.75 99.98 100.08 99.39 99.78 99.11 Cu (ppm) 22 13 52 176 30 b1905b1 Th b0.05 b0.05 0.13 b0.05 b0.05 0.22 b0.05 b0.05 b0.05 0.69 Sc 0.2 0.2 0.4 0.5 0.3 2.0 0.1 0 0.1 1.5 Y 2.4 1.2 12.3 2.6 1.5 22.7 2.4 2.3 1.9 15.8 Zr 5.7 4.5 16.2 3.8 3.4 20.9 3.3 1.5 1.5 17.0 La 1.02 0.35 6.88 0.91 0.91 11.42 1.30 1.09 0.72 5.88 Ce 2.52 0.76 17.11 1.60 1.82 33.86 3.96 3.05 1.61 13.78 Pr 0.24 0.09 1.48 0.17 0.19 2.89 0.35 0.25 0.17 1.17 Nd 0.98 0.40 5.74 0.71 0.74 11.58 1.42 0.99 0.70 4.42 Sm 0.22 0.11 1.21 0.16 0.16 2.60 0.32 0.22 0.16 0.92 Eu 0.090 0.084 0.356 0.062 0.046 0.767 0.093 0.066 0.041 0.271 Gd 0.27 0.13 1.47 0.23 0.18 3.24 0.30 0.23 0.16 1.27 Tb 0.045 0.023 0.253 0.043 0.033 0.539 0.051 0.043 0.026 0.277 Dy 0.28 0.16 1.58 0.30 0.21 3.28 0.30 0.26 0.17 1.93 Ho 0.060 0.036 0.332 0.070 0.044 0.656 0.061 0.054 0.048 0.43 Er 0.185 0.12 0.97 0.24 0.14 1.80 0.18 0.15 0.18 1.28 Tm 0.028 0.021 0.129 0.042 0.021 0.238 0.026 0.021 0.033 0.187 Yb 0.18 0.15 0.68 0.30 0.14 1.30 0.16 0.12 0.22 1.06 Lu 0.024 0.027 0.087 0.046 0.021 0.160 0.019 0.017 0.033 0.143 Mn/Fe (×103) 7.3 21 0.6 5.6 3.1 0.7 22 73 41 4.5 Sm/Fe (×106) 3.6 4.4 2.0 3.4 2.9 5.2 5.1 3.6 4.2 1.8 Y/Ho 40 34 37 37 34 35 40 42 39 37

(La/Yb)SN 0.43 0.17 0.75 0.23 0.48 0.65 0.62 0.66 0.24 0.41 Ce anomaly 1.16 0.95 1.24 0.94 1.00 1.36 1.34 1.33 1.05 1.21 Eu anomaly 1.73 3.33 1.24 1.47 1.29 1.22 1.41 1.37 1.20 1.15 ⁎Abbreviations: C. Chief, Copper Chief mine area; M. Mtn, area (Grapevine Gulch Formation), Mf. Cyn., MacFarland Canyon T (), H. Mtns, Hieroglyphic Mountains; Fe2O3, total iron as Fe2O3. Mn/Fe, Sm/Fe, and Y/Ho are weight ratios.

Jasper and iron formation both have mostly positive dilution of the hydrothermal fluids by ambient seawater Eu anomalies. Jasper in the Copper Chief mine area away from the vents with dilution factors on the order of shows the largest anomalies including one sample with a 104 in the non-buoyant plume [34,45]. Eu/Eu⁎ value of 3.33; the majority have anomalies of b1.4 (Supplementary Table 2). By analogy with iron 4. Ce anomalies and implications for the deep ocean oxyhydroxide precipitates on the modern seafloor redox state (Fig. 4), such positive Eu anomalies imply contributions of Eu2+ from reduced, high-temperature hydrothermal All samples of iron formation from the Jerome fluids [35,39]. The range of anomalies is comparable to district, and one of the jasper samples, show positive Ce that of oxide particles in the buoyant (near-vent) and anomalies of 1.16–1.45 (Table 1; Supplementary non-buoyant (more distal) parts of modern plumes Table 2). Notably, most jasper lacks analytically sig- [36,39,45], and is consistent with a model of increasing nificant Ce anomalies; two samples have small negative J.F. Slack et al. / Earth and Planetary Science Letters 255 (2007) 243–256 249

Fig. 4. Plots of rare earth element data normalized to post-Archean Australian shale (PAAS). Jasper is shown by open symbols and iron formation by closed symbols. A. Jasper and hematitic iron formation from the Jerome district, Arizona. Samples are from the contact between the lower and upper units of the Cleopatra Rhyolite (Copper Chief mine area), and from the overlying Grapevine Gulch Formation (Mingus Mountain area). B. Jasper and hematitic iron formation from the Mazatzal Mountains (MacFarland Canyon area) and the Hieroglyphic Mountains, Arizona. C. Hydrothermal plume-derived iron oxide sediment from the TAG hydrothermal , Mid-Atlantic Ridge [35] and the Rainbow vent field of the Mid-Atlantic Ridge [39], Ordovician jasper from the Løkken district, Norway [16,17], and Jurassic jasper from ODP Site 801C [40] in the western Pacific Ocean (unpublished ICP-MS data of O.J.R.). D. Modern deep ocean water [41], and hydrothermal plume particles from the Rainbow vent field of the Mid- Atlantic Ridge [36] and the East Pacific Rise at 9°45′N [42]. PAAS data from McLennan [38].

Ce anomalies (0.80 and 0.83) that are both valid on the that characterize seafloor-hydrothermal jaspers deposit- basis of the discrimination criteria of Bau and Dulski ed in oxic Phanerozoic oceans [16,17] and modern [47]. Samples of iron formation uniformly show higher plume-related iron oxyhydroxide particles and sediment Ce anomaly values than associated jasper from the same (Fig. 4). The Ce anomalies in the distal jaspers from locality. The uniformly positive Ce anomalies in iron Jerome reflect a dominant seawater signature with a formation from the Jerome district contrast with those negligible (b1%) contribution of Ce derived from the of hematitic iron formation from other fluids, based on a comparison of the sequences [47–50], which show no, or rarely, small uniformly small Eu anomalies that occur in these sam- negative Ce anomalies. The few exceptions that have ples and the REE patterns for varying proportions of positive Ce anomalies are single samples of iron for- modern submarine vent fluids and seawater [17,45]. mation from the ca. 2.74 Ga Michipicoten greenstone Positive Ce anomalies are known in modern alkaline belt in , [48], the ca. 2.68 Ga Bjørnevann lakes [47] and could have existed in a hypothetical, pre- Group in northern Norway [48], and the ca. 1.88 Ga 1Ga“soda ocean” as suggested by Kempe and Degens Biwabik Iron Formation in , USA [49]. Sig- [51]. However, a “soda ocean” model is considered nificantly, jaspers of the Jerome district also lack the unlikely for this time period because deposits moderate to large negative Ce anomalies (∼0.3–0.6) of late Paleoproterozoic age contain Ca sulfate minerals 250 J.F. Slack et al. / Earth and Planetary Science Letters 255 (2007) 243–256

[52,53] instead of the expected trona. The high bicar- ∼0.6), attributed to redox cycling of both Mn and Ce at bonate content of an alkaline soda ocean would have led the oxic/suboxic boundary [15].Incontrasttothese to Ca depletion in seawater due to precipitation of elements, concentrations of dissolved Fe are retained at minerals, with the evaporation of seawater low levels throughout the suboxic zone; a large increase forming trona rather than gypsum or anhydrite [51,54]. in concentration is observed only below 110 m (Fig. 5) Albite after trona have been reported due to reductive dissolution of Fe within the sulfidic deep from playa-lake in Namibia waters. This is accompanied by a further increase in the [55], but to our knowledge no evidence exists for the Ce anomaly with depth (Ce/Ce⁎ from ∼0.8 to 1.0). former presence of trona deposits in coeval or older The Ce anomalies of sulfidic waters in the Black Sea marine strata. are comparable to the small negative anomalies of the The formation of Ce anomalies in oxide particles Jerome jaspers. However, the hematite-rich nature of the reflects the redox-sensitive nature of Ce in contrast to Jerome jaspers and iron formations argues against the other REE that, except for Eu2+, are strictly trivalent anoxic, and particularly sulfidic, ambient seawater [41]. The decoupling of Ce4+ from REE3+ in seawater during deposition of their precursors. In addition, anoxic requires a minimum redox potential between that of conditions do not explain the systematically larger Ce Fe3+ and Mn4+ [41]. In modern, fully oxygenated deep anomalies of iron formation relative to jasper. This oceans, dissolved REE concentrations have large discrepancy is interpreted to reflect fractionation of Ce negative Ce anomalies (∼0.06–0.16 [41]; Fig. 4)due from the trivalent REE during scavenging from seawater to oxidation of Ce and removal in Mn nodules and Mn– by iron oxyhydroxide precursors of the hematite in jasper Fe oxide particles. In suboxic and anoxic waters, Ce and iron formation (Fig. 6). For oxides it has anomalies are smaller or absent due to reductive long been established that such fractionation is respon- dissolution of settling Mn- and Fe-rich particles [15]. sible for the positive Ce anomalies of hydrogenous deep- For example, the Black Sea shows large negative Ce sea nodules and crusts. Oxidation of dissolved Ce3+ to anomalies in the oxic upper part of the stratified water insoluble Ce4+ oxide is kinetically slow, unless the cat- column (Fig. 5) due to Ce uptake by Mn oxide particles alytic effects of Mn-oxide surfaces or bacterial mediation [15]. Below an oxic/suboxic interface at 76 m there is a [41] are involved. However, oxidation on Mn-oxide sharp increase in dissolved Mn concentrations accompa- surfaces is precluded by the very low MnO contents in nied by a shift from large to moderate negative Ce jasper and iron formation from the Jerome district. anomalies (with Ce/Ce⁎ values increasing from 0.05 to Bacterially-mediated Ce oxidation could have been im- portant in the Paleoproterozoic ocean but it does not explain the consistent variation in Ce anomalies between jasper and iron formation. Experimental work [57] indicates that Ce oxidation may also be mediated by newly-formed iron oxyhydr- oxide surfaces. In this process, a relative enrichment of Ce develops in Fe-rich chemical sediments by the adsorption of seawater-dissolved REE, including Ce3+, onto the iron oxyhydroxide particles followed by Ce oxidation coupled with partial desorption of the trivalent REE. In modern, fully oxygenated deep oceans this oxidative Ce scavenging to a significant enrich- ment of Ce relative to the other REE in plume-derived iron oxyhydroxide particles, which typically have less- negative Ce anomalies of 0.3 to 0.8 [36,42] in contrast to values of 0.06 to 0.16 for ambient seawater [41].In suboxic deep waters where the precipitation of Ce-rich Fig. 5. Profiles of dissolved oxygen and trace elements in the upper Mn oxides is precluded, only moderate negative Ce part of the Black Sea, showing variations with respect to upper oxic, anomalies would develop in seawater together with intermediate suboxic, and deep sulfidic zones at site BS3-6. Data for small positive Ce anomalies on the iron oxyhydroxide dissolved Fe and Mn, and the Ce anomaly (Ce/Ce⁎), are from [15]; particles. dissolved O2 concentrations are from [46]; data for NO3 are from [56], adjusted from the profile for site BS3-2. Values for Ce/Ce⁎ are The level of Ce enrichment with respect to seawater normalized to average shale [15]. produced by this process depends on how long the J.F. Slack et al. / Earth and Planetary Science Letters 255 (2007) 243–256 251

Fig. 6. Model for the origin of jasper and iron formation, and related REE scavenging in the late Paleoproterozoic submarine-hydrothermal plumes of the Jerome district. In buoyant part of the plume where ambient, silica-rich seawater is entrained at high rates, dissolved ferrous iron in high- temperature hydrothermal fluid is oxidized to form ferric iron oxyhydroxide particles that promote flocculation with amorphous silica, thus forming the gel protolith of jasper [17]. Rapid silica–iron–oxyhydroxide particle aggregation minimizes further interaction with seawater, resulting in jasper with REE patterns that closely resemble the composition of the seawater-dominated plume waters (small to no Ce anomaly). In more distal, non- buoyant parts of the plume where entrainment of silica-rich seawater is limited, continued iron oxidation and Si–Fe particle formation intermittently leads to silica undersaturation and the precipitation of iron oxyhydroxide particles with only minor silica, thus forming protolith of iron formation. This process allows continued adsorption/desorption of REE, including oxidative Ce scavenging mediated by the Fe oxyhydroxide surfaces [57] and the formation of positive Ce anomalies in iron formation. VMS=volcanogenic massive sulfide. Note that this model is valid only for ancient seafloor- hydrothermal systems; modern systems are unlikely to form seafloor jasper beds due to the silica-depleted nature of post-Cretaceous seawater [16]. particles are exposed to seawater, with the result that the that at Rainbow are more than four orders of magnitude process has the potential to develop variably positive Ce longer than the ca. 6 min measured in the experiments. anomalies in jasper and iron formation that formed from Because the iron oxyhydroxide particles in the hydrothermal plumes if exposure times were different. Jerome area hydrothermal plumes formed mainly in Experimental results by Bau [57] suggest that exchange non-buoyant parts distal from vents and were likely equilibrium for the trivalent REE and Y may be reached comparable in size to those of modern plumes (b3 μm), within 6.5 min during reaction with freshly precipitated the particles that settled to form iron formation probably iron oxyhydroxide (at pH 3.6–6.2), whereas Ce4+/Ce3+ had similarly long residence times in seawater and redox equilibrium requires 120 min. In modern hence sufficient time to acquire positive Ce anomalies. examples, suspension times are 50–150 days for oxide Accordingly, uniformly positive Ce anomalies in the particles in the non-buoyant part of plumes [58,59], iron formation samples reflect greater exchange of iron which is far longer than needed for Ce redox oxyhydroxide particles with seawater while settling equilibration. Based on Bau's experiments [57],it through the plume and ambient water column [36] and might be argued that “superchondritic” Y/Ho ratios of at the sediment–water interface [34]. By contrast, iron N28 (by wt), like those in jasper and iron formation from oxyhydroxides in jasper precursor particles had short the Jerome district (Table 1), reflect exposure times exposure times due to rapid flocculation with amor- of less than ca. 6 min. However, such an approach is phous silica and the formation of larger, fast-settling problematic since data for modern plume particles from aggregate particles that limited further exchange with the Rainbow vent field on the Mid-Atlantic Ridge [36] ambient waters [16,17]. The lack of positive Ce ano- also show superchondritic Y/Ho wt ratios (avg=30), malies in most jasper from the district and specifically even though suspension times for modern plumes like the presence of a small negative Ce anomaly in only one 252 J.F. Slack et al. / Earth and Planetary Science Letters 255 (2007) 243–256 sample therefore are considered the best proxies for the interpreted [10] as reflecting widespread sulfidic condi- composition and redox state of deep seawater 1.74 Ga. tions in the deep ocean at 1.73 and 1.64 Ga, respectively, The presence of abundant hematite in jasper and but these results are equally compatible with suboxic iron formation in the Jerome district requires oxy- conditions [10,14], like those proposed here for the genated seawater during the precipitation of primary Jerome district. Significantly for our model, iron is iron oxyhydroxides. However, fully oxic conditions are oxidized under suboxic conditions thus preventing its inconsistent with the uniformly low Mn concentrations transport to shallow-marine settings and the deposition in both sample types, and by the absence of large of Superior-type BIFs. Suboxic conditions therefore negative Ce anomalies that characterize seafloor- also explain the absence of these BIFs ca. 1.8 to 0.8 Ga. hydrothermal deposits in modern and Phanerozoic Although suboxic environments of the modern ocean oxygenated seawater (Fig. 4; see also [17]). In deep- are generally restricted to narrow transitional layers ocean settings, suboxic conditions best explain the small between oxic and anoxic waters, they were probably positive to small negative Ce anomalies in plume- much more extensive when atmospheric and ocean derived, Si–Fe and iron oxyhydroxide precursors to the oxygen levels were lower. Furthermore, the ocean redox jasper and iron formation (Fig. 6). Recent studies in the state 1.8 to 0.8 Ga was likely affected by higher con- modern ocean indicate that suboxic (microaerobic) centrations of organic carbon in the upper part of the environments are prolific sites for bacteria that mediate water column [61,62]. Both of these factors would have iron oxidation as their primary energy source, because in limited oxygen export to the deep ocean, and led to this setting there is no competition with inorganic stable suboxic conditions over an extended period of oxidation of iron [60]. Hematitic filaments in the Jerome geologic time. district jaspers (Fig. 3) are possible vestiges of such Oceanic residence times for REE are determined by bacteria that lived under suboxic conditions in Paleo- the rate of Fe oxyhydroxide and Mn oxide burial and proterozoic seawater and/or on the seafloor, and that input from the continents, and are estimated at 50 to could have contributed to iron oxide precipitation by 2890yr in the modern oceans [63]. This limits time and catalyzing the oxidation of Fe2+. space extrapolations from the Jerome district, even if REE residence times were somewhat longer in a less- 5. Redox state of the late Paleoproterozoic deep oxygenated Paleoproterozoic ocean. However, jasper ocean and iron formation in other late Paleoproterozoic terranes of central Arizona also show similar REE patterns and Our proposed suboxic redox state for the 1.74 Ga Ce anomalies, ranging from no to small positive ano- deep ocean does not agree with previously suggested malies in jasper to moderate positive anomalies in iron models that infer either oxic [3] or sulfidic [4–12] deep formation (Fig. 4; Table 1). These iron-rich chemical seawater during the late Paleoproterozoic. Iron specia- sediments occur within a 1728Ma volcanosedimentary tion and sulfur isotope data for black of the Rove sequence [64,65] in the Hieroglyphic Mountains and a Formation in Ontario, Canada, were attributed to initi- ∼1710 Ma sequence [20,25] in the MacFarland Canyon ation of sulfidic conditions in the deep ocean ap- area of the Mazatzal Mountains (Fig. 1). Oxic chemical proximately 1.84 Ga [8], 100 m.y. before the suboxic sediments of hydrothermal origin (jasper and hematite or depositional conditions suggested here for strata of the iron formation) are also associated with Jerome district. Similar data and biomarker analyses for several late Paleoproterozoic VMS deposits 1792 to the 1.73 Ga Wollogorang Formation and the 1.64 Ga 1718Ma elsewhere in the western and Reward Formation in the McArthur Basin of northern 1765 to 1690Ma in eastern Bolivia, and with 1395 and Australia [6,11] also imply deposition in sulfidic basins, 1241Ma Mesoproterozoic VMS systems in southern as do results for the 1.49 to 1.43 Ga Roper Formation in Brazil and South Africa, respectively (Table 2). These northern Australia [7] and the 1.47 to 1.44 Ga Newland occurrences, together with those described here from Formation (lower Belt Supergroup) in Montana [5]. Arizona, imply that the deep Proterozoic ocean from 1.8 However, these Paleoproterozoic and Mesoproterozoic to 1.2 Ga was not uniformly sulfidic as argued by sequences were deposited within intracratonic basins other workers [4,6,8], but contained at least some dis- or shallow shelf environments that probably had lim- solved oxygen, either persistently or episodically. Al- ited communication with the global ocean, and as a though positive carbon isotope values in 1250 to 850Ma result may not record the redox state of deep seawater. shallow-marine and increased seawater Molybdenum isotope data for the Wollogorang and sulfate contents at ca. 1.2 Ga [9] are consistent with Reward formations of northern Australia have been the transition to a more oxidized deep ocean, full J.F. Slack et al. / Earth and Planetary Science Letters 255 (2007) 243–256 253

Table 2 Late Paleoproterozoic and Mesoproterozoic VMS deposits with associated jasper and/or iron formation⁎ VMS deposit/district Location⁎⁎ Seafloor hydrothermal deposit# Host rock age& References Encampment/Pearl Southern WY–Northern CO (USA) Ferruginous chert; iron formation 1792±15 Ma [66,67] Gunnison Southern CO (USA) Hematitic chert; magnetite-rich chert 1776±7 Ma [68,69] Jerome Central AZ (USA) Jasper; iron formation 1738.5±0.5 Ma [this study] Bagdad Western AZ (USA) Hematitic chert 1721±6 Ma [25,70] Jones Hill/Pecos Northern NM (USA) Iron formation; jasper 1718±5 Ma [71,72] Miguela Eastern Bolivia Iron formation 1765–1690 Ma [73,74] Itaberaba Southern Brazil Iron formation 1395±10 Ma [75,76] Boksputs Cape Province of South Africa Iron formation 1241±12 Ma [77,78] ⁎With one exception, all of the listed VMS deposits contain abundant Cu in chalcopyrite; the Itaberaba area lacks massive sulfide but shows oxygen isotope evidence for high-temperature, VMS-type alteration prior to regional metamorphism [75].Zn–Pb deposits are not included here because they could have formed in shallow waters and/or intracratonic basins that lacked communication with the global ocean. ⁎⁎AZ=Arizona; CO=Colorado; NM=New ; WY=Wyoming. #Terms used are those reported in the cited papers. &All ages are by U–Pb methods on igneous zircons in volcanic sequences that host the VMS deposits and systems, except for the age range of volcanic host rocks to the Miguela (Bolivia) VMS deposit, which is based on U–Pb zircon ages of older and younger igneous rocks and of detrital zircons in associated sedimentary rocks [74]. oxygenation of the deep ocean was not attained until the in the ocean. Future evaluation of the redox state of other late Neoproterozoic [b580 Ma; 79]. late Paleoproterozoic and Mesoproterozoic open-marine While suboxic conditions apparently were wide- sequences is needed for better understanding the effects spread in the late Paleoproterozoic ocean, the bottom of moderate atmospheric oxygen levels on Proterozoic waters of some basins may have been sulfidic. Deep- ocean chemistry and biological evolution [3,4,8–12,52]. water facies of the Wollogorang Formation in northern A fruitful approach for achieving this goal is through Australia were deposited under sulfidic conditions multidisciplinary studies of jasper and iron formation 1.73 Ga [6], approximately coeval with deep ocean that formed during this time period in association with suboxic deposition of jasper and iron formation in central seafloor-hydrothermal VMS mineralization. Arizona 1.74 to 1.71 Ga. Globally, single basins may have contained multiple water layers with different redox Acknowledgements states that expanded or contracted over time [10]. Sub- oxic conditions in the deep ocean, for example, would We are especially grateful to S.A. Bowring for pro- have promoted expansion of sulfidic conditions within viding the U–Pb age data on the Cleopatra Rhyolite. intracratonic and silled basins like those in northern Geological advice on the Bagdad area was supplied by Australia and , and in upwelling zones on C.M. Conway and P.M. Blacet; on the Mazatzal Moun- open continental shelves. Sulfur isotope data, including tains by C.J. Eastoe; on the Gunnison area by L.P. James those for sulfide minerals in VMS deposits of the Jerome and C.J. Nelsen; and on the Itaberaba area by C. Juliani. district [5–9,80], imply low sulfate levels in the late We thank A.P. Schultz and the USGS Resources Paleoproterozoic ocean. Our data, however, suggest that Program for financial support to J.F.S. and T.G. Funding neither sulfidic nor fully oxic conditions were wide- for A.B. is from the Carnegie Institution of Washington, spread in the deep ocean 1.74 to 1.71 Ga. Suboxic deep the Carnegie Astrobiology Institute, NSF grant EAR- seawater is consistent with low sulfate levels, and im- 05-45484, and NASA Astrobiology Institute award plies that the late Paleoproterozoic ocean was low in NNA04CC09A. Support for O.J.R. was provided by a nitrate due to denitrification at and below the oxic– postdoctoral fellowship from the Deep Ocean Explora- suboxic boundary (Fig. 5) and likely sustained only tion Institute at WHOI and a NASA Astrobiology limited primary organic productivity and eukaryote Institute grant to K.J. Edwards. Appreciation is diversity. Finally, in a low sulfate suboxic ocean, a extended to W.H. Wilkinson and Phelps Dodge Mining larger portion of the hydrothermal Fe flux would extend Corporation for access to mine properties in the Jerome farther from submarine vents, thus contributing to more and Bagdad districts. We also thank H.D. Holland who extensive scavenging of bioessential trace metals from suggested study of seafloor-hydrothermal jasper and seawater by hydrothermal particles. Suboxic deep sea- iron formation for evaluating the redox state of the water therefore would impose harsh conditions for Proterozoic ocean. J.R. Hein and M.B. Goldhaber biological productivity by nitrate and nutrient limitation provided helpful comments on an early version of the 254 J.F. Slack et al. / Earth and Planetary Science Letters 255 (2007) 243–256 manuscript; constructive reviews by two journal refer- [17] T. Grenne, J.F. Slack, Geochemistry of jasper beds from the ees are appreciated. Ordovician Løkken ophiolite, Norway: origin of proximal and distal siliceous exhalites, Econ. Geol. 100 (2005) 1511–1527. [18] A.E. Isley, Hydrothermal plumes and the delivery of iron to Appendix A. Supplementary materials banded iron formations, J. Geol. 103 (1995) 169–185. [19] G.A. Gross, in: O.R. Eckstrand, W.D. Sinclair, R.I. Thorpe Supplementary data associated with this article (Eds.), Geology of Canadian Mineral Deposit Types, Geol. – can be found, in the online version, at doi:10.1016/j. Survey Canada, vol. 8, 1995, pp. 41 80. [20] K.E. Karlstrom, S.A. Bowring, Styles and timing of Early epsl.2006.12.018. Proterozoic deformation in Arizona: constraints on tectonic models, Ariz. Geol. Soc. Dig. 19 (1991) 1–10. References [21] M.E. Donnelly, C.M. Conway, Metallogenic Map of Volcano- genic Massive-Sulfide Occurrences in Arizona, U.S. Geol. [1] A. Bekker, H.D. Holland, P.-L. Wang, D. Rumble III, H.J. Stein, Survey Misc. Field Studies Map MF-1853-B, 1988. J.L. Hannah, L.L. Coetzee, N.J. Beukes, Dating the rise of [22] P.A. Lindberg, M.S. Gustin, Field-trip guide to the geology, atmospheric oxygen, Nature 427 (2004) 117–120. structure, and alteration of the Jerome, Arizona deposits, [2] L.A. Derry, A.J. Kaufman, S.B. Jacobsen, Sedimentary cycling Ariz. Bur. Mines and Mineral Technol. Geol. Survey Branch and environmental change in the Late Proterozoic: evidence from Spec. Paper, vol. 5, 1987, pp. 176–181. stable and radiogenic isotopes, Geochim. Cosmochim. Acta 56 [23] P.A. Lindberg, Geology of the Copper Chief mine, Jerome (1992) 1317–1329. district, Arizona, Ariz. Geol. Soc. Dig. 16 (1986) 343–349. [3] H.D. Holland, The Chemical Evolution of the Atmosphere and [24] T.J. Connelly, C.M. Conway, Massive sulfide deposits at Bagdad, Oceans, Princeton University Press, Princeton, 1984. Arizona: end-products of a recurrent Early Proterozoic volcanic [4] D.E. Canfield, A new model for Proterozoic ocean chemistry, hydrothermal system, Soc. Econ. Geol. Guidebook Ser., vol. 1, Nature 396 (1998) 450–453. 1987, pp. 116–123. [5] T.W. Lyons, J.J. Luepke, M.E. Schreiber, G.A. Zieg, Sulfur [25] R.L. Orr, Geology and mineralization associated with the Early geochemical constraints on Mesoproterozoic restricted marine Proterozoic Alder Group, the Sunflower mining district, deposition: lower Belt Supergroup, northwestern United States, Maricopa and Gila Counties, Arizona, M.S. thesis, University Geochim. Cosmochim. Acta 64 (2000) 427–437. of Arizona, Tucson (1990). [6] Y. Shen, D.E. Canfield, A.H. Knoll, Middle Proterozoic ocean [26] P.A. Lindberg, Precambrian ore deposits of Arizona, Ariz. Geol. chemistry: evidence from the McArthur Basin, northern Soc. Dig. 17 (1989) 187–210. Australia, Am. J. Sci. 302 (2002) 81–109. [27] P. Anderson, Proterozoic plate tectonic evolution of Arizona, [7] Y. Shen, A.H. Knoll, M.R. Walter, Evidence for low sulphate and Ariz. Geol. Soc. Dig. 17 (1989) 17–55. anoxia in a mid-Proterozoic marine basin, Nature 423 (2003) [28] T. Grenne, J.F. Slack, Bedded jaspers of the Ordovician Løkken 632–635. ophiolite, Norway: seafloor deposition and diagenetic maturation [8] S.W. Poulton, P.W. Fralick, D.E. Canfield, The transition to a sul- of hydrothermal plume-derived silica-iron gels, Miner. Depos. 38 phidic ocean ∼1.84 billion ago, Nature 431 (2004) 173–177. (2003) 625–639. [9] L.C. Kah, T.W. Lyons, T.D. Frank, Low marine sulphate and [29] C.T.S. Little, S.E.J. Glynn, R.A. Mills, Four-hundred-and-ninety- protracted oxygenation of the Proterozoic biosphere, Nature 431 million- record of bacteriogenic iron oxide precipitation at sea- (2004) 834–838. floor hydrothermal vents, Geomicrobiol. J. 21 (2004) 415–429. [10] G.L. Arnold, A.D. Anbar, J. Barling, T.W. Lyons, Molybdenum [30] M.D. Hannington, C.E.J. de Ronde, S. Petersen, in: J.W. isotope evidence for widespread anoxia in mid-Proterozoic Hedenquist, J.F.H. Thompson, R.J. Goldfarb, J.P. Richards oceans, Science 304 (2004) 87–90. (Eds.), 100th Anniversary Volume, Society [11] J.J. Brocks, G.D. Love, R.E. Summons, A.H. Knoll, G.A. Logan, of Economic Geologists, Littleton, Colorado, 2005, pp. 111–141, S.A. Bowden, Biomarker evidence for green and purple sulphur includes Appendix in CD. bacteria in a stratified Palaeoproterozoic sea, Nature 437 (2005) [31] M.D. Hannington, I.R. Jonasson, P.M. Herzig, S. Petersen, 866–870. Physical and chemical processes of seafloor mineralization at [12] A.D. Anbar, A.H. Knoll, Proterozoic ocean chemistry mid-ocean ridges, Geophys. Monogr. 91 (1995) 115–157. and evolution: a bioinorganic bridge? Science 297 (2002) [32] K.L. Von Damm, Controls on the chemistry and temporal 1137–1142. variability of seafloor hydrothermal fluids, Geophys. Monogr. 91 [13] N.J. Butterfield, A vaucheriacean alga from the middle (1995) 222–247. Neoproterozoic of Spitsbergen: implications for the evolution [33] C.E.J. de Ronde, M.D. Hannington, P. Stoffers, I.C. Wright, R.G. of Proterozoic eukaryotes and the explosion, Paleo- Ditchburn, A.G. Reyes, E.T. Baker, G.J. Massoth, J.E. Lupton, S.L. biology 30 (2004) 231–252. Walker, R.R. Greene, S.W.R. Soong, J. Ishibashi, G.T. Lebon, [14] A.D. Anbar, G.L. Arnold, T.W. Lyons, J. Barling, Response to Evolution of a submarine magmatic-hydrothermal system: Comment on “Molybdenum isotope evidence for widespread Brothers Volcano, southern Kermadec arc, New Zealand, Econ. anoxia in mid-Proterozoic oceans”, Science 309 (2005) 1017. Geol. 100 (2005) 1097–1133. [15] C.R. German, B.P. Holliday, H. Elderfield, Redox cycling of rare [34] C.R. German, K.L. Von Damm, in: H. Elderfield (Ed.), The earth elements in the suboxic zone of the Black Sea, Geochim. Oceans and Marine Geochemistry, Treatise on Geochemistry, Cosmochim. Acta 55 (1991) 3553–3558. vol. 6, Elsevier, Amsterdam, 2004, pp. 181–222. [16] T. Grenne, J.F. Slack, Paleozoic and Mesozoic silica-rich [35] C.R. German, N.C. Higgs, J. Thomson, R. Mills, H. Elderfield, seawater: evidence from hematitic chert ( jasper) deposits, J. Blusztajn, A.P. Fleer, M.P. Bacon, A geochemical study of Geology 31 (2003) 319–322. metalliferous sediment from the TAG hydrothermal mound, J.F. Slack et al. / Earth and Planetary Science Letters 255 (2007) 243–256 255

26°08′N, Mid-Atlantic Ridge, J. Geophys. Res. 98 (1993) [55] H. Porada, H.J. Behr, Setting and sedimentary facies of Late 9683–9692. Proterozoic alkali lake (playa) deposits in the southern Damara [36] H.N. Edmonds, C.R. German, Particle geochemistry in the belt of Namibia, Sediment. Geol. 58 (1988) 171–194. Rainbow hydrothermal plume, Mid-Atlantic Ridge, Geochim. [56] J.W. Murray, L.A. Codispoti, G.E. Friederich, in: C.P. Huang, Cosmochim. Acta 68 (2004) 759–772. C.R. O'Melia, J.J. Morgan (Eds.), Aquatic Chemistry, Advances in [37] R.A. Feely, J.F. Gendron, E.T. Baker, G.T. Lebon, Hydrothermal Chemistry Series, vol. 244, American Chemical Society, Washing- plumes along the East Pacific Rise, 8°40′ to 11°50′N: particle ton, D.C., 1995, pp. 157–176. distribution and composition, Earth Planet. Sci. Lett. 128 (1994) [57] M. Bau, Scavenging of dissolved yttrium and rare earths by 19–36. precipitating iron oxyhydroxide: experimental evidence for Ce [38] S.M. McLennan, Rare earth elements in sedimentary rocks: oxidation, Y–Ho fractionation, and lanthanide tetrad effect, influence of provenance and sedimentary processes, Rev. Miner. Geochim. Cosmochim. Acta 63 (1999) 67–77. 21 (1989) 169–200. [58] J. Dymond, S. Roth, Plume dispersed hydrothermal particles: a [39] V. Chavagnac, C.R. German, J.A. Milton, M.R. Palmer, Sources time-series record of settling flux from the Endeavor Ridge using of REE in sediment cores from the Rainbow vent site (36°14′N, moored sensors, Geochim. Cosmochim. Acta 52 (1988) MAR), Chem. Geol. 216 (2005) 329–352. 2525–2536. [40] J.C. Alt, D.A.H. Teagle, Hydrothermal alteration of upper [59] M.D. Rudnicki, H. Elderfield, A chemical model of the buoyant oceanic crust formed at a fast-spreading ridge: mineral, chemical, and neutrally buoyant plume above the TAG vent field, 26 and isotopic evidence from ODP Site 801, Chem. Geol. 201 degrees N, Mid-Atlantic Ridge, Geochim. Cosmochim. Acta 57 (2003) 191–211. (1993) 2939–2957. [41] R.H. Byrne, E.R. Sholkovitz, in: K.A. Gschneidner Jr., L. Eyring [60] K.J. Edwards, D.R. Rogers, C.O. Wirsen, T.M. McCollom, (Eds.), Handbook on the Physics and Chemistry of the Rare Isolation and characterization of novel psychrophilic, neutrophilic, Earths, Elsevier, Amsterdam, 1996, pp. 497–593. Fe-oxidizing, chemolithoautotrophic α-andγ-Proteobacteria from [42] R.M. Sherrell, M. Field, G. Ravizza, Uptake and fractionation of the deep sea, Appl. Environ. Microbiol. 69 (2003) 2906–2913. rare earth elements on hydrothermal plume particles at 9°45′N, [61] G.A. Logan, J.M. Hayes, G.B. Hieshima, R.E. Summons, East Pacific Rise, Geochim. Cosmochim. Acta 63 (1999) Terminal Proterozoic reorganization of biogeochemical cycles, 1709–1722. Nature 376 (1995) 53–56. [43] M. Bau, Effects of syn- and post-depositional processes on the [62] D.H. Rothman, J.M. Hayes, R.E. Summons, Dynamics of the rare-earth element distribution in Precambrian iron-formations, Neoproterozoic carbon cycle, Proc. Natl. Acad. Sci. 100 (2003) Eur. J. Mineral. 5 (1993) 257–267. 8124–8129. [44] Y. Nozaki, J. Zhang, H. Amakawa, The fractionation between Y [63] Y. Nozaki, in: J. Steele, S. Thorpe, K.K. Turekian (Eds.), and Ho in the marine environment, Earth Planet. Sci. Lett. 148 Encyclopedia of Ocean Sciences, vol. 4, Academic Press, San (1997) 329–340. Diego, 2001, pp. 2354–2366. [45] C.R. German, G.P. Klinkhammer, J.M. Edmond, A. Mitra, H. [64] J.L. Burr, Proterozoic stratigraphy and structural geology of the Elderfield, Hydrothermal scavenging of rare-earth elements in Hieroglyphic Mountains, central Arizona, Ariz. Geol. Soc. Dig. the ocean, Nature 345 (1990) 516–518. 19 (1991) 117–133. [46] G.W. Luther III, T.M. Church, D. Powell, Sulfur speciation and [65] J.L. Burr, Constraints on timing of Early Proterozoic deformation sulfide oxidation in the water column of the Black Sea, Deep-Sea and metamorphism in the Hieroglyphic Mountains, central Res. 38 (Suppl. 2) (1991) S1121–S1137. Arizona, Abstr. Programs - Geol. Soc. Am. 26 (1994) 405. [47] M. Bau, P. Dulski, Distribution of yttrium and rare-earth elements [66] P.D. Klipfel, Geology and metallogeny of the southern portion of in the Penge and Kuruman iron-formations, Transvaal Super- the Encampment district, Colorado and Wyoming, Ph.D. Thesis, group, South Africa, Precambrian Res. 79 (1996) 37–55. Colorado School of Mines, Golden (1992). [48] L.A. Derry, S.B. Jacobsen, The chemical evolution of Precam- [67] W.R. Premo, W.R. Van Schmus, Zircon geochronology of brian seawater: evidence from REEs in banded iron formations, Precambrian rocks in southeastern Wyoming and northern Geochim. Cosmochim. Acta 54 (1990) 2965–2977. Colorado, Geol. Soc. Am., Spec. Pap. 235 (1989) 13–32. [49] B.J. Fryer, Rare earth evidence in iron-formations for changing [68] P.A. Drobeck, Proterozoic syngenetic massive sulfide deposits in Precambrian oxidation states, Geochim. Cosmochim. Acta 41 the Gunnison gold belt, Colorado, New Mexico Geol. Soc. (1977) 361–367. Guidebook, vol. 32, 1981, pp. 279–286. [50] R. Bolhar, M.J. Van Kranendonk, B.S. Kamber, A trace element [69] M.E. Bickford, R.D. Shuster, S.J. Boardman, U–Pb geochronol- study of –jasper banded iron formation in the 3.45 Ga ogy of the Proterozoic volcano-plutonic terrane in the Gunnison Warrawoona Group, craton—formation from hydrother- and Salida areas, Colorado, Geol. Soc. Am., Spec. Pap. 235 (1989) mal fluids and shallow seawater, Precambrian Res. 137 (2005) 33–48. 93–114. [70] B. Bryant, J.L. Wooden, J.D. Nealey, Geology, geochronology, [51] S. Kempe, E.T. Degens, An early soda ocean? Chem. Geol. 53 geochemistry, and Pb-isotopic compositions of Proterozoic (1985) 95–108. rocks, Poachie region, west-central Arizona—a study of the [52] D.L. Huston, G.A. Logan, Barite, BIFs and bugs: evidence for east boundary of the Proterozoic Mojave crustal province, U. S. the evolution of the Earth's early hydrosphere, Earth Planet. Sci. Geol. Surv. Prof. Pap. 1639 (2001). Lett. 220 (2004) 41–55. [71] W.D. Riesmeyer, Precambrian geology and ore deposits of the [53] A. Bekker, J.A. Karhu, A.J. Kaufman, Carbon isotope record for the Pecos mining district, San Miguel and Santa Fe Counties, New onset of the Lomagundi carbon isotope excursion in the Great Mexico, M.S. Thesis, University of New Mexico, Albuquerque Lakes area, North America, Precambrian Res. 148 (2006) 145–180. (1978). [54] L.A. Hardie, H.P. Eugster, The evolution of closed-basin brines, [72] J.M. Robertson, K.C. Condie, Geology and geochemistry of Min. Soc. Am. Spec. Paper, vol. 3, 1970, pp. 273–290. Early Proterozoic volcanic and subvolcanic rocks of the Pecos 256 J.F. Slack et al. / Earth and Planetary Science Letters 255 (2007) 243–256

, Sangre de Cristo Mountains, New Mexico, Geol. implications for the age of the overlying São Roque Group, Soc. Am. Spec. Pap. 235 (1989) 119–146. Revista Brasil, Geociencias 30 (2000) 82–86. [73] M. Biste, A.W. Gourlay, in: R. Sherlock, M.A.V. Logan (Eds.), [77] G.J. Geringer, J.J. Pretorius, F.H. Cilliers, Strata-bound copper– VMS Deposits of Latin America, Geol. Assoc. Canada, Mineral iron sulfide mineralization in a Proterozoic front arc setting at Deposits Div., Spec. Publ., vol. 2, 2000, pp. 359–374. Boksputs, northwest Cape Province, South Africa—a possible [74] S.D. Boger, M. Raetz, D. Giles, E. Etchart, C.M. Fanning, U–Pb Besshi-type deposit, Miner. Depos. 22 (1987) 81–89. age data from the Sunsas region of eastern Bolivia: evidence for [78] D.H. Cornell, Å. Pettersson, Timing of collisions and thermal the allochthonous origin of the Paragua block, Precambrian Res. events in the ∼1.0 Ga Namaqua-Natal Province of southern 139 (2005) 121–146. Africa, Supercontinents and Earth Evolution Symposium, Free- [75] A. Pérez-Aguilar, C. Juliani, L.V.S. Monteiro, A.E. Fallick, J.S. mantle, , Abs., vol. 18, 2005. Bettencourt, Stable isotope constraints on Kuroko-type paleohy- [79] D.E. Canfield, S.W. Poulton, G.M. Narbonne, Late-Neoproter- drothermal systems in the Mesoproterozoic Serra do Itaberaba ozoic deep-ocean oxygenation and the rise of animal life, Science Group, São Paulo State, Brazil, J. South Am. Earth Sci. 18 (2005) 315 (2007) 92–95. 305–321. [80] C.J. Eastoe, M.S. Gustin, D.F. Hurlbut, R.L. Orr, Sulfur isotopes [76] C. Juliani, P. Hackspacher, E.L. Dantas, A.H. Fetter, The in Early Proterozoic volcanogenic massive sulfide deposits: new Mesoproterozoic volcano-sedimentary Serra do Itaberaba data from Arizona and implications for ocean chemistry, Group of the central Ribeira belt, São Paulo State, Brazil: Precambrian Res. 46 (1990) 353–364.