Fluctuations in late Neoproterozoic atmospheric oxidation — Cr isotope chemostratigraphy and iron speciation of the late Ediacaran lower Arroyo del Soldado Group (Uruguay)

Gondwana Research 23 (2013) 797–811

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Fluctuations in late Neoproterozoic atmospheric oxidation — Cr isotope chemostratigraphy and iron speciation of the late Ediacaran lower Arroyo del Soldado Group (Uruguay) Robert Frei a, d,⁎, Claudio Gaucher b, d, Daniel Stolper c, d, Don E. Canfield d a

Institute of Geography and Geology and Nordic Center for Earth Evolution (NordCEE), University of Copenhagen, Øster Voldgade 10, 1350 Copenhagen, Denmark Departamento de Geologia, Facultad de Ciencias, Igua 4225, 11400 Montevideo, Uruguay c Division of Geological and Planetary Science, California Institute of Technology, Pasadena 91125, CA, USA d Nordic Center for Earth Evolution (NordCEE) and Institute of Biology, University of Southern Denmark, Campusvej 55, 5230 Odense, Denmark b

a r t i c l e

i n f o

Article history: Received 11 April 2012 Received in revised form 31 May 2012 Accepted 18 June 2012 Available online 23 June 2012 Handling Editor: M. Santosh Keywords: Chromium isotopes Iron speciation Ediacaran Arroyo del Soldado Group Uruguay Atmospheric oxidation

a b s t r a c t Oxygenation of the Earth's atmosphere occurred in two major steps, near the beginning and near the end of the Proterozoic Eon (2500 to 542 Ma ago), but the details of this history are unclear. Chromium isotopes in iron-rich chemical sediments offer a potential to highlight fine-scale fluctuations in the oxygenation of the oceans and atmosphere and to add a further dimension in the use of redox-sensitive tracers to solve the question regarding fluctuations of atmospheric oxygen levels and their consequences for Earth's climate. We observe strong positive fractionations in Cr isotopes (δ53Cr up to +5.0‰) in iron-rich cherts and banded iron formation horizons within the Arroyo del Soldado Group (Ediacaran; Uruguay) that can be explained by rapid, effective oxidation of Fe(II)-rich surface waters. These fluctuations are correlated with variations in ratios of highly reactive iron (FeHR) to total iron (Fetot) which indicate a predominance of anoxic water columns (FeHR/Fetot >0.38) during the onset of oxidation pulses. We favor the scenario by which isotopically heavy Cr(VI) entered the basin after pulses of oxidative weathering on land and in which Fe(II) accumulated in the water column. Neodymium isotopes reveal that these oxygenation pulses were followed by increased influxes to the basin of continental crust-derived detrital components of Paleoproterozoic (Nd TDM model ages=2.1–2.2 Ga) provenance typical of the Rio de la Plata Craton. The association of positive δ53Cr–ferruginous (FeHR/Fetot >0.38) stratigraphic intervals with low-diversity acritarch assemblages dominated by Bavlinella faveolata strongly support models postulating a stratified, eutrophic Neoproterozoic ocean. Thus, even within a few million years of the Precambrian–Cambrian boundary, paleoceanographic conditions resembled more those of Paleoproterozoic oceans than Phanerozoic and present oceans. This highlights the sheer magnitude of ecological changes at the Precambrian–Cambrian transition, changes which ultimately led to the demise of the Precambrian world and the birth of the metazoan-dominated Phanerozoic. © 2012 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

1. Introduction Because of the strong relationship between atmospheric oxygen and the evolution of life there is an interest in correlating major biodiversification events with indicators of chemical changes on land and in the hydrosphere. The late Neoproterozoic era is one of the key periods when drastic changes in ocean chemistry, caused by geodynamic changes related to the break-up and accretion of supercontinents and by concomitant climatic changes, predicted the evolution of large motile animals during the Precambrian–Cambrian transition (Knoll et al., 2006; Canfield et al., 2007). Indeed, numerous lines of ⁎ Corresponding author at: Institute of Geography and Geology and Nordic Center for Earth Evolution (NordCEE), University of Copenhagen, Øster Voldgade 10, 1350 Copenhagen, Denmark. Tel.: +45 35322450; fax: +45 3532250. E-mail address: [email protected] (R. Frei).

evidence point to an increase in levels of atmospheric oxygen during the later Neoproterozoic (Canfield and Teske, 1996; Fike et al., 2006; Canfield et al., 2008; Scott et al., 2008; Frei et al., 2009; Dahl et al., 2011). Although it becomes more clear that this second large rise of oxygen in the atmosphere is somehow connected with several glacial events of regional to global extents (Hoffman and Schrag, 2002), the causes and details of these oxygenation events remain ambiguous and relatively poorly understood. A number of trace elements (e.g., Sr, C), in conjunction with their stable and radiogenic isotope properties, have been used in the past to unravel in detail the dynamics and nature of sedimentary successions deposited during Neoproterozoic glacial intervals (Kaufman et al., 1991; Kaufman and Knoll, 1995; Misi et al., 2007). Many trace-element proxies, for example concentrations of molybdenum, rhenium, and uranium concentrations in black shales deposited under anoxic water conditions, rely on the strong redox character of these elements and thereby on the

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R. Frei et al. / Gondwana Research 23 (2013) 797–811

could be used to indicate glaciation-related climatic changes on land during this time. We combine our Cr isotope record with iron speciation data (Poulton and Canfield, 2005) to gain an insight on the anoxic– oxic state of the surface waters, and with neodymium isotope data to obtain information on the average crustal residence time of the source rock in the hinterland and on whether or not the type of eroded crust in the hinterland has changed during the sedimentation of the basin succession studied. 2. Geological setting The Arroyo del Soldado Group (ASG) crops out over an area in excess of 30,000 km 2 in the Nico Pérez Terrane, part of the Río de la Plata Craton (Fig. 1). The ASG represents a marine platform succession of more than 5000m thick (Gaucher, 2000), which unconformably overlies Archaean and Proterozoic rocks of the Nico Pérez Terrane. The unit is composed, from base to top of the Yerbal, Polanco, Barriga Negra, Cerro Espuelitas, Cerros San Francisco and Cerro Victoria Formations (Gaucher, 2000; Fig. 2). 2.1. Lithostratigraphy of the Arroyo del Soldado Group The main characteristics of the ASG have been described in detail (Gaucher, 2000; Gaucher et al., 2004a, 2004b, 2007, 2008a) and are summarized below: (a) The Yerbal Formation is a mainly siliciclastic fining-upward sequence representing the transgression of the Ediacaran sea onto the Nico Pérez Terrane. Sandstones occur at the base and

N

Key

NICO

+

SYSZ

+ + + + + + + +

oxygen sensitivity of patterns of release of these elements during weathering of exposed continental crust and their removal into marine sediments. For example, increases in the Mo concentrations in black shales have been shown to follow the Paleoproterozoic (Great Oxidation Event; GOE) and the Neoproterozoic acceleration in atmospheric oxidation. These patterns are interpreted to reflect the balance between the increased release of Mo as the oxidized molybdate form from the weathering continents as atmospheric oxygen concentrations increase and the enhanced removal of Mo during times of expanded sulfidic conditions (Scott et al., 2008). Using iron speciation analyses (Poulton and Canfield, 2005) of sediments deposited during the late Neoproterozoic, Canfield et al. (2007) were able to trace fluctuations in the chemistry of the ocean deep-water, and they demonstrated the transition from ferruginous to oxic deep waters just following the Gaskiers glaciation. Recently, chromium stable isotopes in banded iron formations and iron-rich cherts have been shown to reflect the redox sensitivity of the Cr(III)– Cr(VI) pair in geological and oceanographic processes (Frei et al., 2009). The transfer of Cr, bound in oxides and silicates as Cr(III) to the aqueous-soluble form Cr(VI) as CrO4−, or HCrO42−, is strongly dependent on the presence of MnO2 which only forms in an oxidized atmosphere. Although quantification of the isotopic fractionation during this oxidation is somewhat difficult, the transfer of this signal into the oceans links an isotopic signature in a Fe-rich chemical sediment directly to the oxidation processes on land. The aim of this study is to test the stable Cr isotope system in a late Neoproterozoic (Ediacaran) basin succession deposited at low latitude after the Gaskiers glacial event (Gaucher et al., 2008a; Rapalini et al., 2008), in an attempt to trace fine-scale fluctuations of Cr species in the contemporaneous water column that potentially

+

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PEREZ

+ + + + + + +

PLATA

PIRIAPOLIS

+ + + +

+ +

Archean rocks 2.6-3.4Ga

SAL CH CC

+ ++ + + ++ ++

BRAZIL

ARGENTINA

+ +

+

+ + +

++

SBS Z

SAL +

++

CUCHILLA

TERRANE

MINAS +

++

SYSZ

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TANDILIA

í

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+

TERRANE

+ + 34° + CC + + ++ + + + ++ + CH + + + ++

CSZ

é

+

+

PIEDRA ALTA TERRANE

URUGUAY

35°S

Fig. 1. Geological sketch map of the southern Nico Pérez Terrane and surrounding areas showing sampled sections. SYSZ: Sarandí del Yí Shear Zone (shear sense indicated for the main Mesoproterozoic event), SBSZ: Sierra Ballena Shear Zone; and CSZ: Colonia Shear Zone. Modified from Gaucher et al. (2008b) and references therein.

Yerbal

13 C ‰VPDB -4 -2 0 2 4 6

N4 N3

4

P4 ?

Sr / 86 Sr

>532 11

Key Conglomerate/breccia

87

Group Soldado d el

km

Sandstone

?

Pelite P3

3

N2 ?? ?

2

Limestone 0 0 07 08 0.7 0.7

P2 N1

799

Dolostone

+

583 7

Radiometric Thalassinoides Cloudina riemkeae Bavlinella faveolata

P1

BN

Arroyo

Polanco Cerro Espuelitas CSF CV Fm.

R. Frei et al. / Gondwana Research 23 (2013) 797–811

1

0

+++++

<566 8 583 7

-4 -2 0 2 4 53

Cr ‰

Fig. 2. Synthetic stratigraphic column of the Arroyo del Soldado Group showing sampled stratigraphic intervals, modified from Gaucher and Poiré (2009) and references therein. Radiometric ages according to Kawashita et al. (1999), Gaucher et al. (2008b) and Blanco et al. (2009). Sr isotope ratios from Gaucher and Poiré (2009) and Frei et al. (2011). Cr isotope ratios from Frei et al. (2011) and Frei et al. (2009) and this work. BN: Barriga Negra Formation, CSF Fm: Cerros San Francisco Formation, CV: Cerro Victoria Formation. “P” and “N” denote positive and negative excursions in the δ13C record, respectively, as introduced by Gaucher et al. (2009a). Unit B in the middle Polanco Formation is characterized by a negative δ53Cr excursion accompanied by the N1 negative δ13C excursion (Frei et al., 2011).

(b)

(c)

(d)

(e)

banded siltstones dominate up section. At the top, subordinate banded iron formation (BIF), chert and dolostone interbeds occur (Fig. 2). The Polanco Formation conformably overlies the Yerbal Formation representing a large carbonate ramp. The Polanco Formation is composed of up to 1 km-thick limestones and limestone–dolostone rhythmites. The terrigenous components of the Polanco Formation are mainly composed of well rounded, sand-sized quartz grains. The Barriga Negra Formation has been traditionally placed stratigraphically above the Polanco Formation on top of an erosional unconformity. The unit consists of conglomerates and breccias with carbonate clasts of underlying carbonates, passing into conglomerates and sandstones that include basement clasts up section. The Barriga Negra Formation represents stream-dominated alluvial fans passing into fan-deltas and finally into glauconite-bearing siltstones and marls. Paleocurrents are predominantly to the east. Recent geological mapping and Sr compositions of the limestones below the Barriga Negra Formation (Gaucher et al., 2011) suggest that the limestones are considerably older than the Polanco Formation (Gaucher, 2000), meaning that the Barriga Negra Formation rests on pre-Neoproterozoic basement. Thus, the latter unit is correlated with the lower Yerbal Formation (Fig. 2), sharing the same deepening-upward trend, the same microbiotas and clay-mineral assemblages (Gaucher et al., 2011). It is beyond the scope of this paper to discuss the new stratigraphic scheme in more detail. The Cerro Espuelitas Formation conformably overlies the Polanco Formation and is made up of an alternation of black shales and siltstones, BIF, cherts and thin carbonate beds at the base (Fig. 2). According to Gaucher (2000), the Cerro Espuelitas Formation represents a marine sedimentary environment, corresponding to a moderately deep shelf. The Cerros San Francisco Formation overlies the Cerro Espuelitas Formation and the pre-Ediacaran basement with an erosional unconformity (Fig. 2). It consists of quartz arenites and subordinate subarkoses (Gaucher, 2000; Gaucher et al., 2008b; Blanco et al., 2009), showing well-preserved sedimentary structures, such as ripples, trough and hummocky cross-stratification indicative of deposition above storm-wave base level.

(f) The Cerro Victoria Formation conformably overlies the Cerros San Francisco Formation and is characterized by intercalated micritic and oolitic dolostones, passing into stromatolitic dolostones up-section (Gaucher et al., 2007, Fig. 2). Trace fossils (Thalassinoides) occur in the dolostones, and are indicative of a lower Cambrian age (Sprechmann et al., 2004). The sedimentary environment is interpreted as a shallow, well-oxygenated carbonate shelf. We studied three sections located in the southern Nico Pérez Terrane, namely: La Salvaje Farm, Cerro de la Higuerita and Cerro Carreras (Fig. 1). Whereas at La Salvaje Farm the upper Yerbal and lower Polanco Formations are exposed (Gaucher et al., 2004a), at Cerro de la Higuerita only the upper Yerbal Formation occurs. At Cerro Carreras, BIF and chert of the Cerro Espuelitas Formation are found (Gaucher, 2000). It is worth noting that the BIFs explored in the present study are probably the youngest so far described, just within 20Ma of the Precambrian–Cambrian boundary (see below). 2.2. Age and evolution of the Arroyo del Soldado Group The ASG is folded and intruded by several plutons of lower Cambrian age (Fig. 1). Gaucher (2000) used mineral paragenesis, palynomorph-maturity, illite-crystallinity (Kübler Index) and calcite-twin morphology to show that the regional metamorphic grade of the ASG never exceeded very low-grade conditions with maximum palaeotemperatures of 200 °C (Gaucher, 2000; Blanco et al., 2010). The ASG overlies the Puntas del Santa Lucía Batholith (Fig. 1), which yielded a U–Pb SHRIMP age of 633±11Ma (Hartmann et al., 2002). Gaucher et al. (2008b) reported a U–Pb SIMS age of 583±7 Ma for the Mangacha Granite, which is overlain by the Cerros San Francisco Formation in the Cerro de la Sepultura area. The youngest detrital zircons of the Yerbal, Cerros San Franscisco and Barriga Negra Formations yielded U–Pb ages of 664±14 Ma, 605±8 Ma and 566±8 Ma (Fig. 2), respectively (Blanco et al., 2009). The new stratigraphic assignment of the Barriga Negra Formation (see above) in fact implies that the 566±8Ma age is the maximum age of the whole Arroyo del Soldado Group (Fig. 2). The Guazunambí Granite intrudes the ASG and yielded an Rb–Sr whole rock isochron age of 532± 11Ma (Sri =0.70624; Kawashita et al., 1999). Cingolani et al. (1990)

800

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dated illite recrystallization in pelites of the ASG to between 532±16 and 492±14Ma. The occurrence of the skeletal fossil Cloudina riemkeae in the Yerbal Formation along with well-preserved palynomorph assemblages indicates a late Ediacaran age for the lower and middle ASG (Gaucher, 2000; Gaucher and Poiré, 2009). C-, O-, and Sr-isotopic data also support a late Ediacaran age for the Yerbal, Polanco, Barriga Negra and Cerro Espuelitas Formations and a lower Cambrian age for the Cerro Victoria Formation (Gaucher et al., 2004b, 2007; Gaucher, 2009). Therefore, deposition of the ASG took place between ca. 565 and 535 Ma ago. On the basis of litho-, bio- and chemostratigraphic data, the ASG can be correlated with other Neoproterozoic–early Paleozoic successions of South America (Corumbá Group in Mato Grosso, Brazil; Sierras Bayas Group, Argentina) and Southern Africa (Gaucher et al., 2003, 2005; Gaucher and Germs, 2006), indicating a large shelf margin along the eastern side of the Río de la Plata Craton (Gaucher et al., 2008b). 3. Analytical procedures Powders of individual mesobands of BIF and Fe-rich chert samples, as well as entire rock samples, were prepared from one-centimeter-thick slices of hand specimens and subsequently milled in an agate mortar. For trace elemental analyses, the rock powders were attacked in HBr, dissolved with HF and HNO3 during addition of HBO3 (Connelly et al., 2006), and then dried and re-dissolved in HNO3. Trace-element concentrations were determined by solution ICP-MS (inductively coupled plasma-mass spectrometry) with a Perkin Elmer ELAN 6100 DRC spectrometer at the Geological Survey of Denmark and Greenland (GEUS), using international standards for calibration. For a comparison of GEUS analytical results on some standards with published values, refer to table 1 in Kalsbeek and Frei (2006). Rock powder aliquots (amounts adjusted to yield 2–5 μg Cr in the final separate) were spiked with an adequate amount of a 50Cr– 54Cr double spike and digested in HF:HNO3 mixtures in closed PFA vials on a hot plate at 150 °C. After heating to dryness, the residues were taken up in aqua regia and reheated to 170 °C for 2 h to destroy any fluoride complexes that may have formed during the digestion. After renewed drying, the sample was then taken up in 6 M hydrochloric acid for the Cr extraction. We employed an anion exchange chromatography technique adapted from previously published methods (Ball and Bassett, 2000; Frei and Rosing, 2005) with a few modifications to separate chromium of natural samples from the other matrix elements. Firstly, because of the high iron contents of some of the samples, we passed the solutions through a cation exchange column charged with 6 mL Dowex AG 1×12 in 6 M HCl to remove Fe. Sometimes we passed the samples twice over this column to ensure that Fe was removed quantitatively. In a second chromatographic separation over 1 mL stem columns charged with Dowex AG 1×8 anion resin, we cleaned the Cr fractions from rock matrices collected from the Fe-cleanup columns in dilute 0.2 M HCl. This separation method is based on the exchange of chloride ions on the Dowex AG 1×8 resin by the Cr(VI)-oxyanions (Schoenberg et al., 2008). Since Cr is present in its trivalent (CrIII) form, after sample digestion and the first cationic exchange, oxidation of Cr(III) to Cr(VI) was achieved using (NH4)S2O8 as oxidizing agent (Ball and Bassett, 2000) on a hot plate at 130 °C. Release of Cr from the anion resin was achieved by reduction to Cr(III) with the help of 2 M HNO3 and H2O2. The procedure yields for Cr in this separation method varied from 80 to 90%, and Cr procedure blanks were in the order of 5–10ng, which is negligible compared to the amount of chromium separated from the samples studied herein. The addition of a 50Cr–54Cr double spike of a known isotope composition to a sample before chemical purification allowed accurate correction of both the chemical and the instrumental shifts in Cr isotope abundances (Ellis et al., 2002; Schoenberg et al., 2008). With this

method, we achieve a 2σ external reproducibility of the δ53Cr value with 1.5 μg Cr loads of the NIST SRM 3112a standard on our IsotopX/ GV IsoProbe T thermal ionization mass spectrometer (TIMS) of +/− 0.05‰ with 52Cr signal intensities of 1 V and of +/−0.08‰ for 52Cr beam intensities of 500 mV. Since the double spike correction returns Cr isotope compositions of samples as the per mil difference to the isotope composition of the NIST SRM 3112a Cr standard (which was used for the spike calibration Schoenberg et al., 2008), to maintain inter-laboratory comparability of Cr isotope data, we recalculated our data of natural samples relative to the certified Cr isotope standard NIST SRM 979 as follows (Eq. (1)): 53

d CrsampleðSRM979Þ ¼

h


53

 52 53 52 Cr= Crsample = Cr= CrSRM979 −1  1000: ð1Þ

All Cr isotope measurements were performed on an IsotopX/GV IsoProbe T TIMS equipped with eight Faraday collectors that allow simultaneous collection of all four chromium beams ( 50Cr +, 52Cr +, 53Cr +, 54 Cr +) together with 49Ti +, 51V+, and 56Fe+ as monitors for small interferences of these masses on 50Cr and 54Cr. Cr separates were measured from Re filaments at 1000–1100 °C and loaded with ultraclean water into a mixture of 3 μl silica gel, 0.5μl 0.5M H3BO3 and 0.5 μl 0.5 M H3PO4. Every separate was analyzed at least one to six times with minimum 52Cr beam intensities of 400mV, allowing within-run precisions of the δ53Cr value of +/−0.09‰ or better. To achieve this, we ran the sample over 120cycles (grouped into 24 blocks of 5cycles each) in static mode, and integrated over 10s with 20s background (baseline) collection at 0.5AMU on either side of the peaks. This led to an average analysis time of ca. 1 1/2h. The final δ53Cr value of a sample was then calculated as the average of the repeated analyses. We spiked our samples with an aliquot of the double spike used by (Schoenberg et al., 2008) in their study of silicates and oxides of magmatic and metamorphic rocks, and employed the double spike correction developed by their group. The average 53Cr/52Cr and 50Cr/52Cr ratios for the NIST SRM 3112a standard are 0.113452+/−50ppm (n=200, 2σ) and 0.0282095+/ 151ppm (n=200; 2σ), respectively, using the 50Cr/54Cr ratio of (Shields et al., 1966) for mass bias correction, and measured during the years 2010/2011. These values are indistinguishable from those of the SRM 979 isotopic standard, which we reproduce at a 53Cr/52Cr ratio of 0.1134502+/−78ppm (n=100; 2σ) and a 50Cr/52Cr ratio of 0.0282089+/−161ppm (n=100; 2σ). The average δ53CrNIST SRM 3112a is −0.019+/−0.050‰ (n=32, 2σ). The small deviation from the nominal value of 0‰ is most likely due to a slight inaccuracy in the calibration of the Cr isotope composition of the double spike (Schoenberg et al., 2008). Sequential iron extraction procedures followed the protocols developed and described by (Poulton and Canfield, 2005). Highly reactive iron (FeHR) includes iron oxide, carbonate, and sulfide minerals. This pool of iron represents the iron that is geochemically and biologically active during early sediment diagenesis (Canfield, 1989). Sediments deposited from an oxygen containing water column are characterized by a ratio of highly reactive iron to total iron (FeHR/FeT) below 0.38. In contrast, sediments deposited from an anoxic water column may obtain additional reactive iron from iron mineral formation (such as iron oxyhydroxides or Fe sulfides) in the water column and therefore the FeHR/FeT may exceed 0.38. 4. Cr isotope systematics The mobile Cr(VI) anion (HCrO4−) is the most thermodynamically stable Cr form in equilibrium with present-day air. Oxidation of Cr(III) to Cr(VI) in soils depends on the co-occurrence of Cr(III) (bound most commonly as FeCr2O4) and manganese oxides (catalyzing Cr(III) oxidation). Once mobilized during oxidative weathering, Cr(VI) is mobile as either chromate (CrO42−; alkalic pH) or bichromate (HCrO4−; acidic pH) ions which enter the oceans via riverine transport (Oze et al., 2007).

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There is a considerably smaller input of Cr from atmospheric and hydrothermal vent sources. In today's oceans, total dissolved Cr concentrations are in the range of 2 to 10nM with a relatively short residence time of ca. 2.5 to 4×104 years (Campbell and Yeats, 1984). Cr(VI) can be reduced to Cr(III) by microbes (Sikora et al., 2008) and by aqueous Fe(II) or Fe(II)-bearing minerals (Ellis et al., 2002). Indeed, the oxidation of Fe(II) (aq) by Cr(VI) is more rapid than by oxygen, even under well-aerated, high pH conditions (Eary and Rai, 1989). This means that in the presence of Fe(II), Cr(IV) is efficiently reduced to Cr(III). The Cr(III) is subsequently and effectively scavenged into Fe(III)–Cr(III) oxyhydroxides (Fendorf, 1995) due to the extremely low solubility of Fe,Cr(OH)3 solids (Sass and Rai, 1987). Some Cr(III) can be regenerated and lost from sediments as a result of Fe oxide reduction but, as on land, the Cr(III) is reoxidized rapidly (Eary and Rai, 1988) to Cr(VI) in a catalytic reaction (Eq. (2)) with MnO2 (Oze et al., 2007). CrðIIIÞðaqÞ þ 3FeðIIIÞðaqÞ→CrðVIÞðaqÞ þ 3FeðIIÞðaqÞ:

ð2Þ

At equilibrium, the Cr(VI)O42− anion is enriched by up to 7‰ at room temperature in 53Cr compared to coexisting compounds containing Cr(III) (we use the delta notation relative to the certified National Bureau of Standards Cr reference standard SRM 979, defined as δ53Cr=1000× [(53Cr/52Cr)sample /(53Cr/52Cr)SRM 979]−1) (Schauble et al., 2004). Therefore, subsurface aqueous environments will have positive δ53Cr values (Izbicki et al., 2008). Although the isotopic composition of Cr in seawater has not yet been measured, the positive groundwater Cr(VI) signal should be transferred to the sea, as subsequent adsorption of Cr onto particles (as might occur in soils and rivers) produces no isotope effect (Ellis et al., 2004). The microbial reduction of Cr(VI) generates isotopic shifts of up to −4.1‰, comparable to those produced during abiotic reduction (Ellis et al., 2002; Sikora et al., 2008). This will potentially enrich the heavier isotope in the remaining, unreacted dissolved Cr(VI). However, because of the efficient sequestration of Cr(VI) during Cr reduction and subsequent precipitation of Cr(III) with Fe-oxyhydroxides, the stable Cr isotope signatures of chemically precipitated Fe(III)-rich sediments should mirror the seawater from which the Fe oxides precipitated. The prerequisite for Cr isotopes to record the presence of Cr(VI) in seawater is a predominance of dissolved Fe(II) which acts as a reductant. Therefore, the isotopic composition of Cr in ancient iron-rich sediments should provide a first-order proxy for the presence of Cr(VI) in ancient surface waters, and thus the evolution of oxidative weathering of Cr on land. This approach should be relatively insensitive to the type of iron-rich chemical sediments and the paleoenvironment in which these were deposited. Since oxidation and solubilization of Cr from soils are strongly dependent on the presence of MnO2, which is stable under elevated oxygen fugacities, its pathway to the oceans in the early Precambrian would have been limited by the absence of Mn(IV) under low atmospheric oxygen pressures. The geochemical behavior of Cr in seawater is therefore highly sensitive to levels of atmospheric oxygen. The elemental concentration of total dissolved chromium in present-day seawater falls in the range of 2 to 10nM (~0.2ppb) (Nriagu and Nieboer, 1988). At pH 8.1 and pE (redox potential) of 12.5, the expected ratio of Cr(VI) to Cr(III) should be 1020; Cr(VI) should predominate overwhelmingly (Elderfield, 1970; Pettine and Millero, 1990). At pE 6.5, consistent with control of seawater pE by the O2–H2O2 couple and [H2O2]=0.1μM (Moffett and Zika, 1983), the ratio falls to 10 2 [99% Cr(VI) and 1% Cr(III)]. Literature values of Cr(VI)/Cr(III) in seawater range from b1 to 70 for oxic conditions (Pettine and Millero, 1990, and references therein). Because it is known that the kinetics of oxidation of Cr(III) to Cr(VI) with O2 are slow (Schroeder and Lee, 1975; Cranston and Murray, 1978; Emerson et al., 1979), levels of Cr(VI) in surface seawater are thought to be directly linked to riverine input and so to reflect the oxidative removal by weathering of the continental

801

surface (e.g. Fantoni et al., 2002; Izbicki et al., 2008; Oze et al., 2007). For example, Cranston and Murray (1980) reported that Columbia River water contained an average of 3.2 nM dissolved Cr, composed of 98% CrO42−. A study of six off-shore California coastal seawater samples by Jan and Young (1976) revealed median concentrations for dissolved trivalent and hexavalent chromiums of 0.045 and 0.14 pg/L, respectively (Cr(VI)/Cr(III)~3.1). Similar values have been reported for the N.W. Mediterranean (Cr(VI)/Cr(III)~3.3–6; Boussemart et al.(1992)), for the North Atlantic Ocean and Meditteranean (Cr(VI)(Cr(III)~2.8–3.6; Achterberg and Van den Berg, 1994) and for the Pacific Ocean (Cr(VI)/ Cr(III)~1–4; Nakayama et al., 1981). The only study which reports Cr isotope data for natural groundwater in alluvial aquifers is by Izbicki et al. (2008). These authors reported δ53Cr values ranging from +0.7 to +5.1‰ in native groundwaters, consistent with the addition of Cr(VI) that was fractionated on mineral surfaces prior to entering the solution. Lithogenic sources of Cr on the modern oxygenated Earth, including their weathering products, have isotopic compositions similar to those of the major igneous silicate reservoirs with a common Cr isotope ratio of −0.12+/−0.10‰ (2σ; Schoenberg et al., 2008). Theoretical estimates of equilibrium chromium isotope fractionation predict that the highly oxidized [Cr 6+O4] 2− anion will tend to have higher 53Cr/ 52 Cr than coexisting compounds containing Cr3+ and these are predicted to reach ca. 6 to 7‰ at 298K (Schauble et al., 2004). This theoretical estimate agrees qualitatively with the fractionation of 3.3–3.5‰ observed between Cr6+ in [CrO4]2− and Cr3+ in either [Cr(H2O)6]3+ or Cr2O3 during experimental reduction of [CrO4]2− in solution (Ellis et al., 2002). It is also consistent with our maximum δ53Cr value of +4.9‰ measured on Fe-rich cherts from the Arroyo del Soldado (Uruguay) profile (table S1 in Frei et al., 2009). Finally, while Cr(VI) in nature is not adsorbed onto negatively charged surfaces of soils, clays or sediments (Fendorf, 1995; Kent et al., 1995; Buerge and Hug, 1999), its effective adsorption on the positively charged surfaces of e.g., alumina (y-Al2O3), goethite or other solids with similar properties may complicate efforts to quantify redox levels in natural waters. However, sorption processes are not accompanied by significant Cr stable isotope fractionation (Ellis et al., 2004). 5. Results Cr stable isotope ratios, FeHR/FeT ratios as well as Cr and Sc concentrations of samples from the Arroyo del Soldado Group are listed in Table 1; Sm–Nd isotope data are contained in Table 2 and plotted in Fig. 3; and results (Nd TDM model ages, εNd, δ53Cr and FeHR/FeT values) are plotted as sample averages in Figs. 4–5 together with the corresponding stratigraphic columns. 5.1. Sm–Nd isotopes TDM model ages (Goldstein et al., 1984) scatter from sample to sample but lie consistently between 1.9 and 2.5 Ga in samples from the La Salvaje section (Yerbal and Polanco Formations; Figs. 3–4), defining an average of 2.22+/−0.18 Ga (2σ). In the Cerro de la Higuerita log (Yerbal Formation, Fig. 5), the TDM model ages range from 1.7 to 2.9 Ga (average=2.05+/−0.37Ga; 2σ), which makes the Sm–Nd model age constraints of these two Yerbal Formation sections statistically indistinguishable. The iron-rich sediments from the Cerro Carreras section (Cerro Espuelitas Formation) exhibit similar TDM ages in the range from 1.97 to 3.01Ga (average=2.28+/−0.49Ga; 2σ), implying that the type of eroded crust in the hinterland did not change during the transition from the Polanco Formation into the Cerro Espuelitas Formation. On closer inspection (Fig. 4), fluctuations in TDM and εNd (570Ma) values in the La Salvaje profile seem to correlate with lithology (e.g., chemical sediments such as BIF and chert vs. detrital sediments, such as siltstones). This can be demonstrated by plotting scandium (Sc) concentrations vs. the εNd (570 Ma) values (Fig. 6A) of the samples. We use Sc to indicate terrigenous inputs to the sediments

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R. Frei et al. / Gondwana Research 23 (2013) 797–811

Table 1 Chromium and scandium concentrations and stable chromium isotope compositions of sediments from the Arroyo del Soldado profile. Age Samplea (Ga)

Location

Unit

Cerro Carreras

Cerro Espuelitas Fm. 0.55 0.55 0.55 0.55 0.55 0.55 0.55

Cerro de la Higuerita Yerbal Fm.

La Salvaje Farm Polanco Fm.

Yerbal Fm.

Crb δ53Cr (ppm) (‰)

±2σmc nd δ53Cr ± 2σmc FeHR/Fetot FeHR/Fetot Sc Reference (averages; ‰) (averages) (ppm) (for Cr isotopes)

CC-1A CC-1A r CC-1B 1 CC-1B 2 CC-1B A CC-1B B CC-1B C

3.6 3.3 9.9 9.4 1.4 9.4 1.3

1.34 1.32 0.52 0.44 0.64 0.62 0.24

0.14 0.04 0.13 0.11 0.09 0.11 0.09

1 3 1 1 3 2 3

0.55 0.55 0.55 0.55 0.55 0.55 0.55 0.55 0.55 0.55 0.55

CH-1 A CH-1 B CH-1 C CH-2A A CH-2A B CH-2B A CH-2B B CH-2C A CH-2D A CH-3 CH-5A

12.4 11.1 6.6 16.9 37.8 3.6 3.4 10.0 3.8 68.9 74.9

1.17 3.64 0.99 0.27 0.11 1.13 1.24 1.74 0.63 −0.04 −0.15

0.12 0.13 0.11 0.09 0.07 0.13 0.12 0.14 0.09 0.09 0.09

4 3 3 3 4 3 3 2 3 5 5

0.55 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57

CH-5C SAL-11 A SAL-11 B SAL-11 C SAL-1A A SAL-1A A d SAL-1A B SAL-1A B d SAL-1A D SAL-1A D d SAL-1A E SAL-1A E d SAL-1A E d SAL-1B A SAL-1B B SAL-2 SAL-3 SAL-3 d SAL-4A SAL-4A d SAL-4A d SAL-4B SAL-4B r SAL-4C SAL-4D SAL-4E SAL-4E d SAL-4E d SAL-4E d SAL-4E d SAL-4E d SAL-4F SAL-5A SAL-5A d SAL-5A r SAL-5A r SAL-5A r SAL-5B SAL-5B d SAL-5B r SAL-5C SAL-5C d SAL-5D SAL-5D d SAL-5D r SAL-6 A SAL-6 B SAL-6C SAL-6 D SAL-7A A SAL-7A A d SAL-7A B SAL-7BB A SAL-7BB B

70.4 2.0 2.4 2.5 80.9

−0.21 −0.11 −0.21 −0.17 −0.24 −0.22 −0.16 −0.20 −0.11 −0.10 −0.24 −0.14 −0.23 −0.25 −0.24 −0.24 −0.09 −0.06 3.22 3.39 3.25 4.92 5.00 4.51 3.12 2.09 2.16 1.90 2.16 1.89 2.03 0.31 0.98 1.10 1.10 0.96 0.85 0.29 0.32 0.27 0.65 0.62 1.06 1.19 1.21 0.52 −0.06 0.09 −0.12 −0.22 −0.06 −0.19 −0.23 −0.17

0.08 0.02 0.01 0.03 0.07 0.02

5 2 2 3 4 4 6 6 6 6 3 6 4 3 6 2 2 1 5 3 4 5 5 5 5 5 5 5 5 3 5 5 7 5 5 3 2 6 5 2 5 5 7 5 3 4 4 4 4 5 5 2 1 3

4.3 29.5 83.0

37.7 40.3 3.9 1.3 31.5

18.0 18.1 23.9 21.1 56.4

53.4 5.5 7.2 7.2 7.6 7.4 7.3 5.1 6.1 5.4 6.5 6.5 6.7 1.6 12.2 11.5 8.1 67.9 65.4 53.0 116.3

2.73

Frei et al. (2009)

0.23

1.53 1.59 1.38

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1.93

1.48

3.35 2.24 3.20

0.19

0.11

1.19

0.08

0.63 −0.10

0.09 0.09

−0.18

0.04

1.33

0.02

0.48

0.06

0.50

−0.16

−0.18

0.05

0.06 0.02

−0.08

0.02

0.82

0.21 0.19 0.21

1.32

28.84

0.36

24.64

0.32

25.04 0.33

0.36 0.43 0.85 0.35

0.29

−0.16

0.08

0.39 0.85 0.35

24.51 25.04 3.78 0.52

0.51

1.03

0.54

0.93

0.55 0.71 0.65

0.87 0.67 1.73

0.56 0.66

0.59

2.85 0.91

0.68

0.94

0.54

0.94

0.57

0.85

0.35

0.11

0.20

29.47 0.27 0.45 0.58 22.78

0.31

0.06

−0.25 −0.24

2.85

5.12 2.34 1.74 1.58 26.32 28.27

0.61 0.72 0.33 0.44 0.60 0.24 0.40 0.09 0.09

0.52

0.32

0.77 4.56 3.97 3.44 11.20 15.69 17.21 19.08

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R. Frei et al. / Gondwana Research 23 (2013) 797–811

803

Table 1 (continued) Location

Unit

Age Samplea (Ga) 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57 0.57

a b c d

Crb δ53Cr (ppm) (‰)

SAL-7BB C 79.8 SAL-7BB D 107.9 SAL-7BB E 72.3 SAL-7BB E d 72.1 SAL-8 A 99.8 SAL-8 B 52.9 SAL-8 A A 73.5 SAL-8 A B 79.6 SAL-8C A 73.7 SAL-8C B 72.0 SAL-8C C 96.8 SAL-8C D 69.6 SAL-9A A 2.0 SAL-9A B 1.9 SAL-9A C 0.9 SAL-9A D 0.7 SAL-9A E 0.5 SAL-9B A 15.5 SAL-9B A d SAL-9B B 31.3

±2σmc nd δ53Cr ± 2σmc FeHR/Fetot FeHR/Fetot Sc Reference (averages; ‰) (averages) (ppm) (for Cr isotopes)

−0.27 −0.13 0.10 0.12 −0.13 0.09 −0.06 −0.13 −0.07 −0.09 −0.14 −0.10 0.60 0.67 1.05 1.08 1.18 −0.10 −0.11 −0.17

1 4 5 5 3 2 2 3 3 4 2 5 2 2 2 1 1 6 3 2

0.09 0.14 0.16 −0.10

0.17

−0.02

0.16

−0.09

0.04

−0.10

0.03

0.92

0.26

−0.13

0.04

20.25 22.82 16.74 0.11

0.21 0.46 0.18 0.17 0.23 0.19 0.27 0.31 0.40 0.40 0.44 0.43 0.43 0.37 0.40

0.33 0.18

0.25

0.42

0.39

24.56 15.74 12.35 15.16 7.89 6.66 15.20 7.70 3.17 6.78 0.39 0.18 0.11 7.06 7.26

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d = duplicated analysis; r = repeated (chemically processed) analyses. Concentrations calculated by double spike isotope dilution. Error is the two standard error of the mean of number (n) of mass spectrometrical runs of an individual sample. Number of mass spectrometrical runs.

studied here, similar to the approach applied to BIFs from the Pongola Supergroup (South Africa) by Alexander et al. (2008) and the classification used by Frei et al. (2009) for the same lithologies. Sc concentrations b5 ppm thereby indicate small to insignificant detrital (lithogenic) components, and such values are exhibited by carbonates and (iron-rich) cherts at La Salvaje section (Table 1, Fig. 4). Regression through the data points in Fig. 6A results in a significantly correlated trend line (r2 =0.7) which we interpret as a mixing line. The two end-members of this mixing line are constrained by (1) a low Sc concentration (b2 ppm) and an εNd (570Ma) value of ca. −11.5, and by (2) an elevated Sc concentration (30 ppm) and an εNd (570Ma) value of ~−17.5. We associate the first end-member with seawater/basin water, and the second with the average Nd isotopic characteristics of the detrital source (Archaean–Paleoproterozoic hinterland). Iron-rich chert and BIF samples from the Cerro Carreras and Cerro de la Higuerita sections (Cerro Espuelitas and Yerbal Formations, respectively) plot along the mixing array defined by the Salvaje samples, supporting the idea that the detrital loads contained in the sediments of these three outcrops were well mixed and originated from a hinterland whose eroded crust did not change significantly. 5.2. Iron speciation data Ratios of highly reactive iron to total iron (FeHR/Fetot) are listed in Table 1 together with the stable Cr isotope data. Average values of individual samples are plotted along the stratigraphic profile in Fig. 4. FeHR/Fetot values fluctuate around 0.38; when greater than this a value deposition of sediments in anoxic water columns (FeHR / Fetot >0.38) is indicated (Poulton and Canfield, 2005). The lower, carbonate-dominated portion of the profile exhibits FeHR/Fetot values close to and slightly below 0.38. In these sediments a clear anoxic signal is not obvious. Three excursions from this signature can be distinguished in the overlying sediment packages. The first excursion, comprising the lower horizon of Fe-bearing cherts and BIF (samples SAL-4, 5 and 6) is defined by ratios exceeding 0.38. The maximum values measured in this excursion is FeHR /Fetot =0.71 (sample 4D; Table 1), and thus, not surprisingly, the iron-rich cherts and BIFs deposited from an anoxic water column. An excursion of FeHR/Fetot to values significantly below 0.38 follows, and this is recorded in a package of gray siltsones (samples SAL-7 and 8; Fig. 4). This negative excursion (lowest FeHR /Fetot =0.09) is indicative of deposition in oxic

conditions, and the values measured are compatible with the average FeHR /Fetot =0.15+/−0.06 for the Phanerozoic excluding modern sediments (Poulton and Raiswell, 2002). The final change in FeHR/ Fetot values occurs in the transition to the second Fe-rich chert horizon in the La Salvaje section (Fig. 4). As with samples SAL-4 and 5, samples SAL-9A and 9B exhibit FeHR/Fetot values slightly higher than 0.38, again pointing to deposition from a weakly anoxic water reservoir. This positive excursion is followed by a possible oxygenated depositional environment, as indicated by a single FeHR/Fetot value of 0.20 for the carbonate sample SAL-11 at the base of the Polanco Formation (Fig. 4). 5.3. Cr isotope data Cr stable isotope data are listed in Table 1 and plotted as sample averages in the La Salvaje stratigraphic column in Fig. 4. Most of the samples along the profile are characterized by δ53Cr values of ca. −0.15‰ which is compatible with chromium from magmatic and volcanic rocks and signals an origin for chromium from high-temperature sources (Schoenberg et al., 2008; Frei et al., 2009). Two excursions from this “magmatic” value exist. The first one, characterized by highly fractionated δ53Cr values in samples SAL-4, 5 and 6 (maximum δ53Cr=+5.00‰; sample SAL-4 B) coincides with iron-bearing cherts and BIF in the middle of the section (Fig. 4). The second positive excursion, less pronounced, is defined by δ53Cr values up to +1.08‰ (sample SAL-9A D). This excursion also coincides with the presence of Fe-rich cherts in the uppermost Yerbal Formation. The two positive δ53Cr excursions correspond with the two FeHR/Fetot excursions toward elevated values (>0.38), and to some degree with εNd (570Ma) minimum values in the Nd isotope record (Table 2, Fig. 4). Cr isotope data of Fe-cherts within the Cerro de la Higuerita profile, regarded as equivalent in time to the La Salvaje Fe-cherts (Gaucher et al., 2004a), are also characterized by strongly positive δ53Cr values (maximum value of +3.58‰; sample CH-1B; Table 1). These values correspond in their magnitude of fractionation to those of the Fe-cherts and BIF horizons of the Yerbal Formation in the middle part of the La Salvaje profile. These highly fractionated values from Fe-rich chemical sediments stand in marked contrast to the values measured from associated organic-rich shales (samples CH-3, 5A,B), which revealed δ53Cr values around −0.09 to −0.15‰. These values probably reflect the composition of detrital minerals with a magmatic, high-temperature Cr

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R. Frei et al. / Gondwana Research 23 (2013) 797–811

Table 2 Sm and Nd concentrations and Sm–Nd isotopic data of sediments from the Arroyo del Soldado Group. TDMa

εNd (0)

εNd (570)

143

4 4 5 4

3.13 n.m. 2.93 n.m.

−30.1 −13.2 −14.6 −16.4

−24.0 −20.1 −11.1 −21.0

0.510674 0.510872 0.511338 0.510828

0.511681 0.511675 0.511679 0.511879 0.511803 0.512066 0.512080 0.511919 0.511668 0.511927 0.511324 0.511282 0.511255

6 6 8 9 12 6 6 6 6 11 8 12 7

1.88 1.83 1.93 2.01 1.85 1.75 1.75 1.76 1.91 2.02 2.51 2.88 2.62

−18.7 −18.8 −18.7 −14.8 −16.3 −11.2 −10.9 −14.0 −18.9 −13.9 −25.6 −26.4 −27.0

−11.2 −11.1 −11.5 −9.0 −9.4 −5.5 −5.3 −7.5 −11.5 −8.4 −18.8 −20.5 −20.2

0.511328 0.511337 0.511315 0.511441 0.511422 0.511622 0.511632 0.511520 0.511313 0.511475 0.510942 0.510855 0.510871

0.11788 0.11844 0.11806 0.11864

0.511709 0.511675 0.511716 0.511694

6 8 8 9

2.30 2.37 2.29 2.34

−18.1 −18.8 −18.0 −18.4

−12.4 −13.1 −12.3 −12.7

0.511269 0.511233 0.511275 0.511251

0.09544 0.10005 0.10271 0.10532 0.09475 0.10754 0.10132 0.10001 0.10116 0.10139 0.10963 0.10724 0.11543 0.12133 0.12247 0.12218 0.12116 0.12702 0.12440 0.11026 0.11380 0.10865 0.11811 0.12684 0.10967 0.12760 0.12217 0.10634 0.10626 0.11013 0.10140 0.09950 0.09434 0.09339 0.11215 0.09701 0.10023 0.10097 0.10484 0.10022 0.10556 0.10292 0.10258 0.10532 0.10622 0.09967

0.511317 0.511367 0.511333 0.511302 0.511302 0.511299 0.511694 0.511695 0.511321 0.511338 0.511638 0.511672 0.511682 0.511779 0.511743 0.511713 0.511705 0.511747 0.511770 0.511717 0.511738 0.511700 0.511708 0.511799 0.511676 0.511749 0.511727 0.511647 0.511631 0.511671 0.511576 0.511575 0.511527 0.511480 0.511603 0.511590 0.511595 0.511576 0.511614 0.511548 0.511606 0.511611 0.511588 0.511587 0.511594 0.511575

17 13 5 5 6 6 9 7 11 9 7 7 6 8 6 6 8 8 6 5 7 12 8 8 6 5 7 6 6 5 4 5 7 8 8 13 19 5 6 7 4 6 7 5 7 5

2.37 2.40 2.51 2.61 2.38 2.67 1.98 1.96 2.49 2.47 2.22 2.12 2.29 2.27 2.36 2.40 2.39 2.48 2.36 2.12 2.16 2.11 2.31 2.38 2.17 2.49 2.38 2.14 2.16 2.19 2.15 2.11 2.08 2.12 2.33 2.05 2.10 2.14 2.16 2.16 2.18 2.13 2.15 2.21 2.21 2.11

−25.8 −24.8 −25.5 −26.1 −26.1 −26.1 −18.4 −18.4 −25.7 −25.4 −19.5 −18.8 −18.7 −16.8 −17.5 −18.0 −18.2 −17.4 −16.9 −18.0 −17.5 −18.3 −18.1 −16.4 −18.8 −17.3 −17.8 −19.3 −19.6 −18.9 −20.7 −20.7 −21.7 −22.6 −20.2 −20.4 −20.3 −20.7 −20.0 −21.3 −20.1 −20.0 −20.5 −20.5 −20.4 −20.7

−18.4 −17.8 −18.6 −19.4 −18.7 −19.7 −11.5 −11.4 −18.8 −18.4 −13.2 −12.3 −12.7 −11.3 −12.1 −12.6 −12.7 −12.3 −11.7 −11.7 −11.5 −11.9 −12.4 −11.3 −12.5 −12.3 −12.4 −12.8 −13.1 −12.6 −13.8 −13.7 −14.2 −15.1 −14.1 −13.2 −13.3 −13.8 −13.3 −14.3 −13.5 −13.2 −13.6 −13.9 −13.8 −13.7

0.510960 0.510993 0.510949 0.510909 0.510948 0.510897 0.511316 0.511321 0.510943 0.510959 0.511229 0.511272 0.511251 0.511326 0.511285 0.511256 0.511252 0.511273 0.511306 0.511306 0.511313 0.511294 0.511267 0.511325 0.511266 0.511272 0.511270 0.511250 0.511234 0.511260 0.511197 0.511203 0.511175 0.511131 0.511184 0.511228 0.511221 0.511199 0.511223 0.511173 0.511212 0.511227 0.511205 0.511193 0.511198 0.511203

Formation

Sm (ppm)

Nd (ppm)

147

Cerro Espuelitas Fm.

3.38 0.31 1.17 0.32

18.07 0.64 4.77 0.76

0.11319 0.29213 0.14795 0.25885

0.511097 0.511963 0.511890 0.511795

Cerro de la Higuerita CH-1 A Yerbal Fm. CH-1 B CH-1 C CH-2B A CH-2B B CH-2C A CH-2C B CH-2C C CH-2D A CH-2D B CH-3 CH-5 A CH-5 C

3.48 3.31 2.89 0.48 0.29 1.57 4.56 2.22 1.72 0.57 3.77 5.22 6.06

22.27 22.16 17.95 2.49 1.73 8.01 22.97 12.61 10.96 2.84 22.29 27.61 35.60

0.09443 0.09034 0.09729 0.11710 0.10205 0.11876 0.12016 0.10680 0.09516 0.12107 0.10230 0.11442 0.10296

La Salvaje Farm SAL-11 A Polanco Fm. SAL-11 B SAL-11 B SAL-11 C

0.27 0.27 0.26 0.26

1.39 1.36 1.32 1.30

La Salvaje Farm SAL-1A A Yerbal Fm. SAL-1A A SAL-1A B SAL-1A C SAL-1A D SAL-1A E SAL-1A/C SAL-1A/C SAL-1B A SAL-1B B SAL-2 SAL-2 A SAL-3 SAL-4 A SAL-4 B SAL-4 C SAL-4 D SAL-4 E SAL-4 F SAL-5 A SAL-5 B SAL-5 C SAL-5 D SAL-6 A SAL-6 B SAL-6 C SAL-6 D SAL-7A A SAL-7A B SAL-7A C SAL-7BA A SAL-7BA B SAL-7BA C SAL-7BA D SAL-7BB A SAL-7BB B SAL-7BB C SAL-7BB D SAL-7BB E SAL-8 A SAL-8 B SAL-8A A SAL-8A B SAL-8C A SAL-8C B SAL-8C C

6.03 6.49 5.98 6.45 5.22 5.42 0.21 0.22 5.42 4.62 1.41 0.33 0.48 0.42 0.43 0.42 0.25 0.87 1.52 0.46 0.54 0.53 0.56 0.25 4.15 1.77 1.32 5.05 6.31 9.88 5.07 4.15 8.96 5.50 3.50 4.64 4.45 5.36 5.46 9.90 5.28 5.29 6.98 6.80 6.76 7.90

38.20 39.26 35.24 37.04 33.35 30.52 1.26 1.32 32.41 27.59 7.80 1.88 2.51 2.12 2.13 2.06 1.22 4.16 7.39 2.51 2.88 2.97 2.84 1.19 22.90 8.39 6.56 28.75 35.91 54.28 30.27 25.24 57.48 35.62 18.87 28.97 26.85 32.10 31.48 59.76 30.26 31.12 41.16 39.08 38.51 47.96

Sample Cerro Carreras CC-1A CC-1B A CC-1B B CC-1B C

Sm/144Nd

143

Nd/144Nd

±2σ mean last digits

Nd/144Nd ini

R. Frei et al. / Gondwana Research 23 (2013) 797–811

805

Table 2 (continued) Sample

Formation

SAL-8C D SAL-9A A SAL-9A A SAL-9A B SAL-9A B SAL-9A C SAL-9A D SAL-9A E SAL-9B A SAL-9B B SAL-9B C

Sm (ppm)

Nd (ppm)

147

Sm/144Nd

8.45 0.11 0.11 0.09 0.09 0.07 0.05 0.04 1.83 3.27 1.60

56.88 0.64 0.64 0.55 0.56 0.42 0.30 0.24 10.23 19.56 8.03

0.08990 0.10023 0.10044 0.10288 0.10165 0.10186 0.10288 0.10955 0.10796 0.10125 0.12037

143

Nd/144Nd

0.511553 0.511709 0.511736 0.511767 0.511750 0.511801 0.511639 0.511676 0.511736 0.511602 0.511759

±2σ mean last digits

TDMa

εNd (0)

εNd (570)

143

Nd/144Nd ini

9 12 8 8 12 15 8 6 5 6 5

1.97 1.94 1.91 1.91 1.91 1.85 2.09 2.17 2.05 2.11 2.28

−21.2 −18.1 −17.6 −17.0 −17.3 −16.3 −19.5 −18.8 −17.6 −20.2 −17.1

−13.4 −11.1 −10.6 −10.2 −10.4 −9.4 −12.7 −12.4 −11.2 −13.3 −11.6

0.511217 0.511335 0.511361 0.511383 0.511370 0.511420 0.511255 0.511267 0.511332 0.511224 0.511310

n.m. = no meaningful age (post-depositional disturbance?). a TDM after Goldstein et al. (1984).

isotope signature. This is demonstrated by the relationship between Sc concentration and δ53Cr (Fig. 6B), showing that only samples with b6ppm Sc exhibit positive δ53Cr values. As we will discuss below, the most likely process capable of oxygenating surface waters and the atmosphere and at the same time driving the redox state of deep waters to anoxic (i.e. ferruginous) conditions is primary productivity. Photosynthesis, and particularly the burial of organic matter, releases O2 to atmosphere, while the decomposition of the organic matter as it sinks to the bottom consumes dissolved O2 which leads to anoxic bottom waters. Higher up in the Arroyo del Soldado Group, δ 53Cr recently reported for the Polanco Formation (Fig. 2) are characterized by positive values of up to +0.3‰ at the base, followed by a negative excursion down to −0.2‰ in the middle and again positive δ53Cr of +0.2‰ at the top (Frei et al., 2011). The latter positive excursion likely continues in Fe-rich cherts of the overlying Cerro Espuelitas Formation (Cerro Carreras section), which exhibit positively fractioned δ53Cr signatures, with values up to +1.34‰ (sample CC-1A; Table 1). These trends therefore complement the general picture of positively fractioned δ53Cr values in the investigated iron-rich chemical sediments deposited in the Arroyo del Soldado basin.

area of these samples has been interpreted as including early Neoproterozoic light rare earth depleted rocks of the São Gabriel Arc in southernmost Brazil (Blanco et al., 2009). Thus we confirm that the Yerbal basin to the south (as exemplified by the La Salvaje and Cerro de la Higuerita sections) did not experience input from these younger basement sources. The scatter of εNd values of the Yerbal, Polanco and Cerro Espuelitas Formations corresponds to the scatter of U–Pb detrital zircon data recorded in these rocks (Gaucher et al., 2008b; Blanco et al., 2009). These authors reported a predominance of Paleoproterozoic (1.9–2.2 Ga and 2.5Ga) ages, followed by a small Mesoproterozoic population (1.0–1.1 Ga) and finally by a few Neoproterozoic as well as Archaean (ca. 3.0Ga) zircons. The dominant, major igneous events are recorded by the TDM ages and/or by the detrital zircon distribution around 2.1Ga, which is typical for the core of the Rio de la Plata Craton (Pankhurst et al., 2003; Rapela, 2007). The provenance of Mesoproterozoic zircons remains speculative, but Gaucher et al. (2008b) postulated a proto-Andean margin as the main source of Mesoproterozoic detritus, in contrast to Rapela (2007) who favored a probable source in the Kalahari Craton in their reconstruction of SW Gondwana. Our data also support the assignment of the Nico Pérez Terrane to the Río de la Plata Craton (Bossi and Cingolani, 2009; Gaucher et al., 2011), despite recent proposals to the contrary (Oyhantcabal et al., 2010).

6. Discussion 6.2. Oxygenation of hydrosphere/atmosphere and Cr isotopes 6.1. Provenance of the ASG A detailed petrographic, (isotope-) geochemical and U–Pb detrital zircon dating study by Blanco et al. (2009) aimed to define the provenance for sediment sources to the Arroyo del Soldado Group. From petrographic observations, these authors concluded that the clastic rocks of the ASG were mainly derived from granitic and felsic metamorphic rocks with only minor contributions from sedimentary sources. Based on Th/Sc and Zr/Sc relationships, these authors also concluded that the provenance of detrital material of the ASG sediments (in this case from the Yerbal and Cerro Espuelitas Formations) is typically a mixture of recycled metasediments (Zr/Sc>10, Th/Sc>1) with typical unrecycled upper continental crustal material. The excessive scatter of data in tectonic discrimination diagrams (e.g., in a La/Sc vs. Ti/Zr plot as in fig. 9 of Blanco et al., 2009) can be interpreted as to show a poor mixing of components which prevents a classification of their tectonic setting. The Sm–Nd isotope systematics of sample studies here from the ASG are comparable to those published by Blanco et al. (2009) for the Yerbal, Barriga Negra, Cerro Espuelitas and Cerros San Francisco Formations (Fig. 3). εNd (570) values between ca. −24 and −7 are compatible with reworked Archaean–Paleoproterozoic crust of the Rio de la Plata Craton in Uruguay. We did not, however, obtain εNd (570) values for the Yerbal Formation that are less negative than −7, in contrast to Blanco et al. (2009) who report two values of −0.9 and −0.4 for siltstone samples of the Yerbal Formation in the northernmost outcrops. They clearly discriminate these values from their other data. The source

It is widely accepted that the oxidation and solubilization of Cr from modern soils is strongly dependent on the presence of MnO2 (Eq. (3)) The reaction necessitates that Cr(III) comes into direct contact with Mn oxides (Fendorf, 1995): 2þ

CrðOHÞ

þ 1:5MnO2 →CrO4

2−

2þ

þ 1:5Mn

þ

þH :

ð3Þ

At the pH and Eh values characteristic of the modern Earth's surface the activation energy of Mn(II) to Mn(IV) is much larger than that of Fe(II) to Fe(III) and, even under fully anoxic conditions at neutral pH, the inorganic oxidation reaction proceeds prohibitively slowly. A number of bacteria and fungi are capable of catalyzing the oxidation of Mn(II) to MnO2. Today this is seen as the primary mechanisms by which Mn(IV) is formed in natural waters and sediments (Nealson et al., 1988; Murray and Tebo, 2007). Biological Mn(II) oxidation rates have been shown experimentally to be directly proportional to dissolved O2 concentrations, with oxidation rates decreasing fivefold (from 0.005 to 0.001 μM mg−1 min−1) when O2 concentrations are reduced from 250 μM to 15μM (Zhang et al., 2002). These basic constraints should be set in relation to the observation of fractionated δ53Cr values in many late Neoproterozoic, iron-rich chemical sediments (Frei et al., 2009), including those reported here for the ASG. Indeed, the observation of positively fractionated δ 53Cr values implies and necessitates the existence of an oxidation process (i.e. Cr(III) to Cr(VI)) during the transfer of chromium from the

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10.0

DMM CHUR

eNd(570)

0.0

-10.0

-20.0 Salvaje Cerro de la Higuerita Cerro Carreras

-30.0 0

500

1000

1500

2000

2500

3000

3500

T (Ma) Fig. 3. Nd isotopic evolutionary diagrams for samples of the three studied sections. The dot-outlined rectangle encompasses data reported by Blanco et al. (2009) of the Cerros San Francisco Fm., the Cerro Espuelitas Fm., the Barriga Negra Fm. and the Yerbal Fm. The gray field shows the evolution of Archaean–Paleoproterozoic crustal rocks of the Río de la Plata Craton basement as shown by Gastal et al. (2005). DMM denotes the depleted mantle evolution curve of Goldstein et al. (1984). CHUR: chondrite uniform reservoir (DePaolo and Wasserburg, 1976).

continent to sediment deposition. Simple leaching under more acidic conditions would be predicted to liberate the lighter isotopes (due to weaker bonds). Thus, the prediction would be that acid leaching would transfer chromium with negatively fractionated δ 53Cr values into the aqueous phase. None of the Precambrian BIFs studied so far (Frei et al., 2009) have returned negatively fractionated chromium (i.e., δ 53Cr values are more negative than the −0.12±0.1‰ magmatic average value), which supports our contention that an oxidative process was responsible for chromium mobilization. Mobilization of Cr(III) is furthermore readily and effectively prevented in weathered soils on land due to the immediate adsorptive capture of Cr(III) by energetically preferentially formed FeOOH under oxidative weathering conditions. We also note that in today's rivers, under circum-neutral (pH ca. 6.5–7.5) to the slightly basic pH (ca. 8), condition corresponding to that recorded in modern seawater, Cr(VI) is at least as abundant or, in the basic pH range, even more abundant than Cr(III) (Yusof et al., 2007). In another study aimed at defining Cr species in natural, alkaline, and oxic groundwater in alluvial aquifers in the western Mojave Desert, southern California, high Cr(VI)>50 mg/L was measured in aquifers from mafic rock, but Cr(VI) as high as 27 mg/L were measured in aquifers from granitic rock (Izbicki et al., 2008). The δ53Cr values in these natural groundwaters ranged from 0.7 to 5.1%, supporting a scenario whereby Cr(VI) was fractionated on mineral surfaces prior to entering solution. This study also revealed that only

Key

Fig. 4. Stratigraphic column of the La Salvaje Farm section. The shaded horizons are characterized by positive δ53Cr values and anoxic (ferruginous) conditions, as indicated by iron speciation (FeHR/Fetot) data. Modified from Gaucher et al. (2004a).

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807

Fig. 5. Stratigraphic column of the Cerro de la Higuerita section. Note positive δ53Cr values in BIFs (shaded interval).

small concentrations of Cr(VI)b5 mg/L were present in waters with pHb7.5, supporting the findings of Yusof et al. (2007) for Cr in river waters. Together, at least today, Cr(VI) is available for exchange with the aqueous phase and is the predominant form of Cr in water under oxidizing conditions at elevated pH. Although in principle oxidation is inhibited by the formation of mixed Fe–Cr oxyhydroxides, field studies suggest that oxidation of Cr(III) to Cr(VI) nevertheless occurs at the longer time scales available in natural settings (Bartlett and James, 1979). The fluctuations of Cr isotope ratios recorded in the ASG, in combination with iron speciation data, reveal the following: (1) The correlation of positively fractionated δ 53Cr values in Fe-rich cherts and BIF of the Yerbal Formation with FeHR/Fetot values>0.38 (Fig. 4) indicate that while deep basin waters remained anoxic, surface waters were oxic enough to allow for the existence of Cr(VI) species (which upon reduction by upwelling of Fe 2+ waters were transferred as mixed Fe–Cr oxyhydroxides into the chemical precipitates). (2) The existence of Cr(VI) in the shallow basin waters (indicated by positively fractionated δ53Cr values) signals periods of increased oxidative atmosphere levels during intervals of BIF deposition, which allowed mobilization of Cr bound as Cr(III) in minerals exposed to weathering on land (Oze et al., 2007; Frei et al., 2009, 2011; Fig. 8). (3) The FeHR/Fetot valuesb0.38, in sediments deposited after the iron-rich sediments, mark the onset of oxic conditions of deep and shallower basin waters (Figs. 4, 8).

The redox decoupling of deep and surface oceans highlights the stratified nature of Neoproterozoic oceans (Fig. 8), as also suggested by other lines of evidence (Gaucher, 2000; Shen et al., 2005; Canfield et al., 2007). It is important that the positive δ53Cr signal cannot be recovered from the siliciclastic intervals; these are dominated by detrital Cr and therefore nothing can be said about the redox state of the surface ocean and atmosphere just by looking at their δ 53Cr values. Iron speciation data suggest that the water column became progressively more oxygenated in the upper Yerbal Formation and into the lower Polanco Formation. Frei et al. (2011) report δ53Cr values up to +0.3‰ for the lower Polanco Formation (Unit A) and trace-element evidence of an oxygenated water column, in accordance with the data presented here. Other redox-sensitive trace elements and biomarkers confirm an oxygenated environment for the lower Polanco Formation (Velasquez, 2010). An impressive δ 53Cr negative excursion down to magmatic values of −0.17‰ occurs in the overlying Unit B of the Polanco Formation (Frei et al., 2011). No iron speciation data are available for this stratigraphic interval, but redox-sensitive trace elements strongly suggest a return to anoxic conditions in the water column (Velasquez, 2010). A positive δ53Cr excursion follows (Frei et al., 2011), and continues into the overlying Cerro Espuelitas Formation, with values reaching +1.34‰ in BIF of that unit. The general picture is of positive δ53Cr values in the ASG, punctuated by a distinct negative excursion in the middle Polanco Formation (Fig. 2). Iron speciation and trace element data of the studied sections strongly suggest that oxygen-deficient, ferruginous conditions dominated in the deeper Arroyo del Soldado shelf, except for the upper Yerbal and

808

A

R. Frei et al. / Gondwana Research 23 (2013) 797–811

-8 -10 -12

Nd(570Ma)

-14 -16 -18 -20 -22

Salvaje Cerro de la Higuerita Cerro Carreras

-24

R2 =0.73

-26 0

5

10

15

20

25

30

Sc(ppm)

B

6 5

53

Cr‰

4 3 2 53 1 0 0

5

10

15

20

25

30

Fig. 7. Bavlinella faveolata from thin sections of cherts of the Yerbal Formation. (A) Two large colonies and also loose, individual microcells (arrowed). (B) Three well-preserved, discrete colonies. Scale bars represent 10μm.

-1

Sc(ppm) Fig. 6. (A) Sc concentration vs. εNd diagram. (B) Sc concentration vs. δ53Cr diagram. Note how the δ53Cr signal in sediments is effectively suppressed beyond 6 ppm Sc.

lower Polanco Formations. Further geological indications of ferruginous conditions are the several BIF horizons occurring in both the Yerbal and Cerro Espuelitas Formations. Strong positive fractionations of Cr isotopes in late Neoproterozoic, iron-rich chemical sediments have been reported by Frei et al. (2009), for instance, for the 755–730Ma Rapitan BIF deposited in a glaciomarine setting during the early Cryogenian (“Sturtian”) glaciation. First results were obtained for the ~740 Ma Chuos BIF (Namibia), the ~580Ma Jakkalsberg BIF (Namibia), and the latest Cryogenian or middle Ediacaran Urucum/Mutun BIFs (Brazil and Bolivia), all associated with diamictites and dropstones and showing positive δ53Cr values. This may point to a causal relationship between Neoproterozoic BIF deposition and oxygenation of the surface environments.

6.3. Cr and C isotopes, acritarchs and Neoproterozoic paleoceanography The number of glacial events that occurred during the Neoproterozic remains contentious. BIFs, however, were deposited worldwide during the last 200 Ma of the Proterozoic, together with manganese deposits and evaporites, and the BIFs are associated with glacial deposits. It has been suggested that the return of phosphorite deposition in the late Neoproterozoic may have been due to the increased size of acritarchs and other protistan morphotypes, e.g. Butterfield and Rainbird (1998), whose remains sank into the deep ocean. The increased downward transport of organic matter was apparently sufficient to return to anoxic conditions in the deep ocean for at least parts of the late Neoproterozoic. This has been comprehensively demonstrated by Canfield et al. (2008)

who came to the conclusion that deeper water masses, despite limited areas with sulfidic conditions, were characteristically Fe2+-enriched (ferruginous), similar to conditions prevailing in the Paleoproterozoic. The close association of BIFs and many of the manganese deposits with glacial deposits is striking, as the recurrence of iron formations occurred during a long-term increase in seawater oxygenation. Examples are known from all three major ice ages, i.e. middle Cryogenian, end-Cryogenian and Gaskiers events. These iron formations are thought to represent the accumulation of Fe2+ in an ice-capped anoxic ocean (Hoffman et al., 1998). Ferruginous deep-ocean waters were also dominant during the latest of the Neoproterozoic glaciations, namely the Gaskiers ice age (580Ma; (Canfield et al., 2007)). Oxygenation of the deep ocean, timed with emergence of large animals by the end of the Neoproterozoic Era, has been reported from the Avalon Peninsula, Newfoundland (Canfield et al., 2007), but details on whether such an oxygenation was a local phenomena or whether it was a globally occurring feature are unknown. Low FeHR/Fetot values were also identified in the deep-water sediments from the Ediacaran upper Kaza Group, Caribou Mountains, western Canada (Canfield et al., 2008); they add to the growing body of evidence for widespread oxygenation of the global deep ocean around 580 to 560Ma (Fike et al., 2006; Scott et al., 2008). The Arroyo del Soldado Group overlies the Las Ventanas Formation, which contains a distinct glaciogenic unit at the base composed of diamictites and dropstone-bearing shales (Gaucher et al., 2008a). Whereas a U–Pb zircon age of gabbros below the diamictites yielded 590±2 Ma (Mallmann et al., 2007), acid pyroclastics above the glacial horizon gave a U–Pb zircon age of 573±11 Ma (Oyantcabal et al., 2009). These ages confirm the assignment of the Las Ventanas Formation to the Gaskiers Glaciation, first made on the basis of acritarch biostratigraphy (Gaucher et al., 2008a). No glacial rocks sensu stricto occur in the overlying ASG. However, a series of positive and negative δ 13 C excursions in carbonates of the ASG that correlate with

R. Frei et al. / Gondwana Research 23 (2013) 797–811

Oxidative weathering

BIF deposition

Iron fertilization

high phytoplankton productivity

CO2

Cr (VI)

O2

oxygenated layer

upw

elli

ng

anoxic layer organic C burial (13 C-depleted)

Fig. 8. Conceptual model for the deposition of BIFs of the Arroyo del Soldado Group. Whereas BIFs deposited in this scenario are characterized by positive δ53Cr values and FeHR/Fetot >0.38, carbonates deposited concomitantly elsewhere in the basin yield positive δ13C ratios. Note the critical role of the planktic microbiota in producing both oxygen and sinking organic particles. The acritarch blooms may have been enhanced by iron “fertilization” (see text for details). Modified after Gaucher et al. (2004b).

palaeobathymetric fluctuations and microfossil diversity changes attest to palaeoclimatic oscillations that affected the basin (Gaucher et al., 2004b). Above the Yerbal Formation, carbonates of the lower Polanco Formation (Unit A) record positive δ 13C and δ 53Cr values, the latter up to +0.3‰ (Frei et al., 2011). Up section, Unit B of the Polanco Formation records a negative δ 13C excursion down to −4.5‰; it is concomitant to sea-level fall and has been interpreted as a glacial event coeval with the Shuram–Wonoka anomaly (Gaucher et al., 2009b), which is comparable in age to the recently described Fauquier glacials (Hebert et al., 2010). A markedly negative δ 53Cr excursion to −0.2‰ characterizes Unit B (Frei et al., 2011), paralleling δ 13C values. This is the only stratigraphic interval (ca. 200 m thick) in the ASG that was probably characterized by anoxia in both deep and surface waters. The absence of glaciogenic rocks associated with Unit B in the Polanco Formation may be explained by the tropical setting of the Río de la Plata Craton during this time (Rapalini et al., 2008). Positive δ 13C and δ 53Cr values return in the upper Polanco and Cerro Espuelitas Formations, of latest Ediacaran age. The iron-rich chemical deposits of the upper Yerbal Formation have been interpreted by Gaucher et al. (2004b) as reflecting upwelling of nutrient-rich (P,N,Fe) waters during a high bioproductivity period characterized by positive δ13C values (Gaucher et al., 2004b, 2009b). This is indicated by the occurrence of a low-diversity, high-abundance microflora dominated by B. faveolata recording eutrophic conditions (Gaucher, 2000; Gaucher et al., 2004b; Gaucher, 2009) (Fig. 7). B. faveolata (=Sphaerocongregus variabilis) is probably a planktic cyanobacterium comparable to extant Microcystis (Mansuy and Vidal, 1983). From an ecological point of view, Microcystis is characterized by blooms under eutrophic conditions with the exclusion of eukaryotic plankton by the production of toxins (Nasri et al., 2007). These ecological characteristics are also typical of Bavlinella, which occurs in almost monospecific, depauperate assemblages and in large concentrations (Fig. 7) leading to high TOC values (Mansuy and Vidal, 1983; Gaucher et al., 2003). Seawater-like REE+Y patterns of the iron-rich chemical sediments of the Yerbal Formation with negative Ce anomalies, positive Eu anomalies and positive Y anomalies (Frei et al., 2009) imply that these elements and probably also iron were dominantly derived from subaqueous, distal hydrothermal fluids, similar to many Archaean and Proterozoic BIFs worldwide, and that they were (in contrast to Archaean and early Paleoproterozoic BIFs) deposited from an oxygenated shallow water mass (allowing Ce(IV) to be partially stable) as also indicated by the strongly fractionated Cr isotope signatures (see above). These

809

iron-rich fluids may have “fertilized” surface waters and initiated the acritarch blooms, as suggested by Gaucher (2000) and Gaucher et al. (2004b) for the Arroyo del Soldado basin. We envisage that primary productivity (i.e. the “biological pump”) was responsible for the simultaneous establishment of oxygenated conditions in the surface environments (positive δ 53Cr values) and anoxia in deeper waters (high FeHR/Fetot values; Fig. 8). Release of oxygen during photosynthesis to the surface ocean (photic zone) was balanced by sinking organic particles, and these maintained low oxygen levels in the deeper waters (Fig. 8). It is also possible that Fe and other reductants sourced in the mid-ocean ridges contributed significantly to the establishment and continuity of deep anoxia. The system shifted to fully anoxic conditions (deep and surface water) in the middle Polanco Formation (Unit B), which also records a negative δ13C excursion (see above). This can be interpreted as a biotic crisis leading to lowered photosynthetic oxygen production and widespread anoxia. In the upper Polanco and Cerro Espuelitas Formations, the basin returns to a redox-stratified system similar to the Yerbal Formation. As for the cause of all these oscillations, palaeoenvironmental changes as they occurred throughout the Neoproetrozoic, are suspected as the ultimate controlling factor (e.g. Gaucher et al., 2009a).

7. Conclusions Strong positive fractionations of Cr isotopes (δ 53Cr up to +5.0‰) are reported here for iron-rich cherts and BIF horizons within the Arroyo del Soldado Group of Uruguay. These BIFs are probably the youngest so far reported, and may be within 20 to 40 Ma of the Precambrian–Cambrian boundary. The δ53Cr fluctuations are correlated with variations in iron speciation; these variations indicate a predominance of anoxic water columns (FeHR /Fetot >0.38) during deposition of BIF and led to elevated δ53Cr values. These apparently contradictory observations can be reconciled if we accept a stratified water column, with a large anoxic, ferruginous water mass overlain by a much thinner layer of oxygenated water. As recorded in the Yerbal and Polanco Formations, the depth of the redoxcline varied quite significantly, perhaps due to variations in photosynthetic oxygen production or fluctuations in the rate of reductant input, or both. The nature of co-occurring acritarch assemblages, characterized by low-diversity, high-abundance B. faveolata-dominated palynofloras, strongly suggests eutrophic (i.e. high-nutrient) conditions. Redox stratification is a hallmark of eutrophic basins, thus independent lines of evidence point to the same conclusions. In this sense, it must be noted that even within a few million years of the Precambrian–Cambrian boundary, paleoceanographic conditions resembled more those of Paleoproterozoic oceans than Phanerozoic and present oceans. The paleooceanographic conditions prevailing during the Precambrian–Cambrian transition possibly were vital for the explosion of metazoan life forms later dominating the Phanerozoic eon.

Acknowledgments We would like to thank Toni Larsen for help in the ion chromatographic separation of chromium from the samples, and Toby Leeper for always maintaining the mass spectrometer in optimal running condition. John Bailey is thanked for his comments and an initial internal review. Financial support was through the Danish Agency for Science, Technology and Innovation grant nr. 272-07-0244 to RF as well as by the Nordic Center for Earth Evolution (NordCEE), a center of excellence established by the Danish National Research Foundation. We are thankful for the critical and constructive comments of Hartwig Frimmel and an anonymous reviewer, which helped to improve the final version of this manuscript.

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