Andy W Heard
email: andyheard@uchicago.edu

Nicolas Dauphas

Romain Guilbaud

Olivier J Rouxel

Ian B Butler

Nicole X Nie

Andrey Bekker

Triple iron isotope constraints on the role of ocean iron sinks in early atmospheric oxygenation

come   L'archive ouverte pluridisciplinaire

Main Text

Irreversible changes of oxic and euxinic sedimentary iron sinks during the Archean and Paleoproterozoic were intimately linked with the oxygenation of Earth's atmosphere during the Great Oxygenation Event (GOE), beginning ca. 2.43 Ga [START_REF] Holland | The Chemical Evolution of the Atmosphere and Oceans[END_REF][START_REF] Gumsley | Timing and tempo of the Great Oxidation Event[END_REF]. Early episodes of oxygenation coincided with enhanced burial of iron sulfide (pyrite) in sediments [START_REF] Scott | Late Archean euxinic conditions before the rise of atmospheric oxygen[END_REF][START_REF] Reinhard | A Late Archean Sulfidic Sea Stimulated by Early Oxidative Weathering of the Continents[END_REF][START_REF] Stüeken | Contributions to late Archaean sulphur cycling by life on land[END_REF]. This limb of the pre-GOE Fe cycle may have directly contributed to early atmospheric oxygenation, as enhanced volcanic SO2 fluxes accompanying Neoarchean continental emergence may have contributed to net burial of reducing power in sedimentary pyrite, following microbial sulfate reduction [START_REF] Kump | Increased subaerial volcanism and the rise of atmospheric oxygen 2.5 billion years ago[END_REF][START_REF] Gaillard | Atmospheric oxygenation caused by a change in volcanic degassing pressure[END_REF][START_REF] Olson | Volcanically modulated pyrite burial and ocean-atmosphere oxidation[END_REF].

Early environmental oxidation should have been buffered by reduced species in the ocean and atmosphere, principally Fe 2+ . Therefore, the marine iron cycle potentially played roles on both sides of Earth's early oxygenation. Major shifts in Fe isotopic records occur across the GOE (9, 10) (Fig. 1A), and reflect evolution of the Fe, S, and O cycles through the Archean and Paleoproterozoic [START_REF] Heard | Constraints on the coevolution of oxic and sulfidic ocean iron sinks from Archean-Paleoproterozoic iron isotope records[END_REF]. Interpretation of the Fe isotopic composition of pre-GOE pyrites, which can be depleted in 56 Fe/ 54 Fe ratio by as much as -3.5‰, a degree unseen in the post-GOE rock record [START_REF] Rouxel | Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF], is however not straightforward. The reason is that these pyrite Fe isotopic compositions could be controlled by (i) the size of oxidizing iron sinks that removed isotopically heavy Fe 3+ oxides, leaving a pool of isotopically light dissolved Fe 2+ from which pyrite could have formed [START_REF] Rouxel | Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF][START_REF] Planavsky | Iron isotope composition of some Archean and Proterozoic iron formations[END_REF]; (ii) microbial dissimilatory Fe 3+ reduction (DIR) that preferentially releases an isotopically light Fe 2+ pool [START_REF] Beard | Iron Isotope Biosignatures[END_REF][START_REF] Archer | Coupled Fe and S isotope evidence for Archean microbial Fe(III) and sulfate reduction[END_REF]; and (iii) kinetic isotope effects (KIEs) accompanying partial pyrite precipitation, which can produce isotopically light pyrite [START_REF] Guilbaud | Abiotic Pyrite Formation Produces a Large Fe Isotope Fractionation[END_REF][START_REF] Rolison | Iron isotope fractionation during pyrite formation in a sulfidic Precambrian ocean analogue[END_REF]. The relative importance of these processes remains heavily debated [START_REF] Rouxel | Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF][START_REF] Planavsky | Iron isotope composition of some Archean and Proterozoic iron formations[END_REF][START_REF] Heard | Constraints on the coevolution of oxic and sulfidic ocean iron sinks from Archean-Paleoproterozoic iron isotope records[END_REF][START_REF] Beard | Iron Isotope Biosignatures[END_REF][START_REF] Archer | Coupled Fe and S isotope evidence for Archean microbial Fe(III) and sulfate reduction[END_REF][START_REF] Guilbaud | Abiotic Pyrite Formation Produces a Large Fe Isotope Fractionation[END_REF][START_REF] Rolison | Iron isotope fractionation during pyrite formation in a sulfidic Precambrian ocean analogue[END_REF][START_REF] Dauphas | Iron Isotope Systematics[END_REF][START_REF] Marin-Carbonne | Coupled Fe and S isotope variations in pyrite nodules from Archean shale[END_REF][START_REF] Marin-Carbonne | Situ Fe and S isotope analyses in pyrite from the 3.2 Ga Mendon Formation (Barberton Greenstone Belt, South Africa): Evidence for early microbial iron reduction[END_REF], and this uncertainty has hindered quantitative interpretation of the ancient iron cycle, exemplified by the fact that δ' 56 Fe records have not yet constrained the degree to which Fe sedimentation on highly productive continental margins was a net sink or source for early O2 [START_REF] Olson | Volcanically modulated pyrite burial and ocean-atmosphere oxidation[END_REF].

Here, we report triple-Fe-isotopic ratio measurements that allow us to remove ambiguities in previous interpretations of the pre-GOE iron cycle. This approach relies on our discovery that the main isotopic fractionation processes implicated in the formation of pre-GOE pyrites follow distinct isotopic mass fractionation laws (MFLs), which describe how different isotopic ratios of the same element covary [START_REF] Young | Kinetic and equilibrium mass-dependent isotope fractionation laws in nature and their geochemical and cosmochemical significance[END_REF][START_REF] Dauphas | Mass Fractionation Laws, Mass-Independent Effects, and Isotopic Anomalies[END_REF]. To resolve MFLs, measurement of Fe isotopic ratios must be of a higher precision than has previously been achieved in analysis of ancient sediments. A similar approach has been used once in igneous geochemistry to demonstrate that Fe isotopic variations in magmatic olivine followed a kinetic MFL for diffusive transport [START_REF] Mccoy-West | The Fe and Zn isotope composition of deep mantle source regions: Insights from Baffin Island picrites[END_REF]. For a given MFL, the ratio of 56 Fe/ 54 Fe to that of 57 Fe/ 54 Fe defines the slope θ 56/57 = δ' 56 Fe/δ' 57 Fe (Eq. 1)

where δ' x Fe = 1,000 ln[( x Fe/ 54 Fe)sample/( x Fe/ 54 Fe)IRMM-014] (IRMM-014 is a standard reference material consistent with bulk planetary composition) [START_REF] Dauphas | Mass Fractionation Laws, Mass-Independent Effects, and Isotopic Anomalies[END_REF][START_REF] Nie | Iron and oxygen isotope fractionation during iron UV photo-oxidation: Implications for early Earth and Mars[END_REF]. Isotopic trends following an array of MFLs are by definition mass-dependent; but they manifest as apparent departures, ϵ' 56 Fe, from an arbitrary reference MFL [START_REF] Dauphas | Mass Fractionation Laws, Mass-Independent Effects, and Isotopic Anomalies[END_REF][START_REF] Dauphas | Iron 60 Evidence for Early Injection and Efficient Mixing of Stellar Debris in the Protosolar Nebula[END_REF][START_REF]Materials and methods are available as supplementary materials on Science Online[END_REF], which we choose here to be the high-temperature equilibrium limit law with θ 56/57 =0.678, where ϵ' 56 Fe = [δ' 56 Fe -0.678 × δ' 57 Fe] × 10 (Eq. 2)

In ϵ' 56 Fe vs. δ' 57 Fe space, MFLs form straight lines whose slopes can be related to θ 56/57 through ϵ' 56 Fe = 10 × [θ 56/57 -0.678] × δ' 57 Fe (Eq. 3)

In order to establish the values of θ 56/57 corresponding to the two endmember hypotheses that have been put forward to explain the ' 56 Fe pyrite record (9, 10, 14), we measured:

(i) A suite of iron formation (IF) samples that show a large range in δ' 56 Fe values, including Mn-rich IFs from the Hotazel Formation in the Griqualand West sub-basin, South Africa, dated at ca. 2.43 Ga (2), that were presumably precipitated from an iron pool that had experienced extensive oxidative iron removal, resulting in IFs characterized by anomalously low δ' 56 Fe values [START_REF] Tsikos | Iron isotopes constrain biogeochemical redox cycling of iron and manganese in a Palaeoproterozoic stratified basin[END_REF].

(ii) Experimental products of pyrite synthesis via the FeS-H2S pathway, which produced pyrite with δ' 56 Fe values as low as -2.4‰ relative to the initial iron pool (Figs. 1B, S2) [START_REF]Materials and methods are available as supplementary materials on Science Online[END_REF]. In these experiments, we precipitated pyrite in anoxic conditions from an FeS precursor [START_REF]Materials and methods are available as supplementary materials on Science Online[END_REF] following a protocol established by Guilbaud et al. [START_REF] Guilbaud | Abiotic Pyrite Formation Produces a Large Fe Isotope Fractionation[END_REF][START_REF]Materials and methods are available as supplementary materials on Science Online[END_REF].

The IFs define a slope of θ 56/57 ox=0.6776±0.0004 for the oxidizing (ox) iron sink (here and elsewhere, the error bars are 95% confidence intervals), which agrees with theoretical expectations for equilibrium isotope exchange that dominates during iron oxidation [START_REF] Dauphas | Mass Fractionation Laws, Mass-Independent Effects, and Isotopic Anomalies[END_REF][START_REF] Nie | Iron and oxygen isotope fractionation during iron UV photo-oxidation: Implications for early Earth and Mars[END_REF], while the pyriteprecipitation experiments gave θ 56/57 KIE=0.6743±0.0006 (Fig. 1B). Triple-Fe-isotopic slopes for the two endmember scenarios are clearly distinct, allowing us to use these slopes to address what caused Fe isotopic variations in natural pyrites before the GOE.

We analyzed a suite of pre-GOE Neoarchean-Paleoproterozoic pyrites with highly depleted δ' 56 Fe values (as low as -3.1 ‰) and a few bulk black shales from the same formations. The pyrite and shales fall in an intermediate space in the triple Fe isotope diagram between the endmember MFLs for Fe 2+ oxidation and pyrite precipitation. A linear regression through these data would have a slope of 0.6761±0.0006, but we do not interpret the apparent linear trend as an MFL, because the pyrite and shale array is composed of samples from several distinct formation and each sample requires contributions from more than one fractionation process (with distinct MFLs) and therefore has no mechanistic significance. More likely, pre-GOE pyrite δ' 56 Fe values record a two-step process; partial marine Fe 2+ oxidation during non-steady-state upwelling of Fe 2+ -rich deep waters [START_REF] Rouxel | Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF], and subsequent KIEs during partial, sulfide-limited pyrite formation from the remaining Fe 2+ reservoir [START_REF] Guilbaud | Abiotic Pyrite Formation Produces a Large Fe Isotope Fractionation[END_REF][START_REF] Rolison | Iron isotope fractionation during pyrite formation in a sulfidic Precambrian ocean analogue[END_REF][START_REF] Rouxel | Response to Comment on "Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF]. In this model, oxic and pyrite sinks sequestered iron upwelling from deep ocean basins lacking a discrete redoxcline, towards black shale depositional settings (Fig. S7) [START_REF] Planavsky | Iron isotope composition of some Archean and Proterozoic iron formations[END_REF][START_REF]Materials and methods are available as supplementary materials on Science Online[END_REF][START_REF] Bekker | Iron Formation: The Sedimentary Product of a Complex Interplay among Mantle, Tectonic, Oceanic, and Biospheric Processes[END_REF] .

For any isotopically light pyrite sample, we can estimate contributions to the δ' 56 Fe value from prior iron-oxidation, and the KIE during pyritization in the porewater iron reservoir. We calculate contributions of Fe-oxidation to δ' 56 Fe values of the water mass (δ' 56 Few) from intercepts with the oxidation MFL, and determined the Fe isotopic fractionation imparted by pyritization by taking the difference in δ' 56 Fe values between those of pyrite and δ' 56 Few (Figs. 2A, S5) [START_REF]Materials and methods are available as supplementary materials on Science Online[END_REF]. An underlying assumption in this calculation is that sedimentary pyrite formed from a pool of dissolved Fe 2+ that sampled the marine Fe 2+ reservoir [START_REF] Rouxel | Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF], although the nodular nature of the pyrites indicates that they formed in sediments through diagenesis. A major source of iron in porewaters would presumably have been downward diffusion of overlying Fe 2+ -rich seawater into the sediment [START_REF] Rouxel | Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF]. We cannot exclude however that some porewater Fe 2+ was produced by DIR [START_REF] Beard | Iron Isotope Biosignatures[END_REF][START_REF] Archer | Coupled Fe and S isotope evidence for Archean microbial Fe(III) and sulfate reduction[END_REF]. This is an uncertainty in our model. However, the role of partial (isotopically fractionating) DIR in the host shales of these pyrites is not supported by Fe isotopic compositions of bulk shale material, which lacks the complimentary heavy isotopic signature that would be expected of a DIR residue [START_REF] Rouxel | Response to Comment on "Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF]. Additionally, coupled Fe and C isotope analyses of Archean sediments suggests that DIR occurred in a closed system with respect to isotopes, with diagenetic carbonate minerals taking on the composition of their precursor oxides [START_REF] Craddock | Iron and carbon isotope evidence for microbial iron respiration throughout the Archean[END_REF][START_REF] Heimann | O isotope compositions of banded iron formation carbonates demonstrate a major role for dissimilatory iron reduction in ~2.5 Ga marine environments[END_REF]. While we cannot rule out the additional influence of DIR in the origin of pre-GOE pyrite Fe isotopic signatures, we anticipate that any uncertainty it introduces would be transferred to interpretations of δ' 56 Few values, and not the departures of triple Fe isotopic values from the Fe 2+ oxidation MFL. The reason for this is that experiments to date suggest that the isotopic fractionation during DIR is an expression of the equilibration of Fe 2+ and Fe 3+ after the reduction step [START_REF] Heidi | The mechanisms of iron isotope fractionation produced during dissimilatory Fe(III) reduction by Shewanella putrefaciens and Geobacter sulfurreducens[END_REF], and therefore we expect it would fall into the same class of Fe redox equilibrium processes that define the Fe 2+ oxidation MFL.

The fraction of Fe 3+ -(oxyhydr)oxide removed to give the δ' 56 Few value on the intercept (Fox = Fe in oxide sink/total Fe), and the fraction of pyrite removed from that remaining Fe 2+ pool (fpy = Fe in pyrite/Fe remaining after Fe removal in the oxide sink), are both calculated assuming Rayleigh fractionation conditions (Figs. 2, S5; Table S4) [START_REF]Materials and methods are available as supplementary materials on Science Online[END_REF] following upwelling of Fe 2+ -rich deep water across the redoxcline (Fig. 2B). The fractional pyrite sink Fpy for iron in the whole depositional system is Fpy = fpy × (1 -Fox). Values of Fox increase heading to more negative δ' 57 Fe and ϵ' 56 Fe values, whereas fpy decrease away from the oxidation MFL, reflecting preservation of a larger KIE at lower degrees of pyritization (Fig. 2B). Fpy decreases strongly with decreasing δ' 57 Fe, due to the combined effects of decreasing fpy and increasing Fox in this direction (Fig. S5). Contours of Fox/Fpy are sub-horizontal, making ϵ' 56 Fe measurements highly diagnostic of the relative size of the oxic and pyritic sinks (Fig. 2C).

To fully propagate the effect of uncertainties in sample measurements; θ 56/57 values for the endmember processes; and model fractionation factors; on uncertainties in Fox, fpy, and Fpy, we also used a Monte-Carlo simulation (Fig. 3). Estimates for Fpy are 10 -80 % of the upwelled iron pool (within 95 % C.I.) among the low δ' 56 Fe pyrites we studied. With initial pre-GOE deep-water [Fe 2+ ] concentrations ~50 µM (1), the pyrite sink could have removed 5-40 µM of dissolved iron.

This requires ~10-80 µM of seawater-dissolved sulfate to be microbially reduced to sulfide, ~350 to 1,400 times lower than the modern seawater sulfate concentration of 28 mM, and within recent estimates for Archean seawater sulfate based on S isotopic modelling [START_REF] Jamieson | Neoarchaean seawater sulphate concentrations from sulphur isotopes in massive sulphide ore[END_REF][START_REF] Crowe | Sulfate was a trace constituent of Archean seawater[END_REF]. For the ~2.65 Ga Jeerinah and Lokammona formations, we infer that as little as 10 % of iron upwelled onto the shelf was deposited as pyrite in euxinic sediments (Figs 2B, 3, S6) [START_REF]Materials and methods are available as supplementary materials on Science Online[END_REF].

When volcanic SO2 is the primary sulfur source, burial of reduced S in pyrite has a net oxidative effect on Earth's surface [START_REF] Kump | Increased subaerial volcanism and the rise of atmospheric oxygen 2.5 billion years ago[END_REF][START_REF] Gaillard | Atmospheric oxygenation caused by a change in volcanic degassing pressure[END_REF][START_REF] Olson | Volcanically modulated pyrite burial and ocean-atmosphere oxidation[END_REF]. For example, the reaction 2SO2 + H2O + Fe 2+ → FeS2 + 2H + + 2.5O2, describing the net effect of SO2 photolysis and hydrolysis, cyanobacterial photosynthesis, microbial sulfate reduction, and pyrite precipitation, indicates that pyrite burial can directly promote the net export of O2 to the atmosphere-ocean system [START_REF] Olson | Volcanically modulated pyrite burial and ocean-atmosphere oxidation[END_REF]. The reaction is a maximum estimate for O2 export during pyrite burial, because a more reduced original sulfur source would weaken the net oxidative effect of pyrite burial; and other types of primary productivity could have contributed organic matter for sulfate reduction, but only cyanobacterial activity could have produced O2. A more realistic estimate of volcanic H2S/SO2 emission ratios ~1 [START_REF] Gaillard | Atmospheric oxygenation caused by a change in volcanic degassing pressure[END_REF] would instead imply a net O2 yield of 1 mole per mole of pyrite buried.

To oxygenate the atmosphere via pyrite burial, the produced O2 would also need to overcome O2 buffers in the ocean, primarily the upwelled Fe 2+ flux. O2-driven Fe 2+ oxidation consumes 0.25 moles of O2 per mole of Fe 3+ buried, so net O2 sources and sinks will be balance when Fox/Fpy = 4 (10) for volcanic H2S/SO2 emission ratios = 1 (0). Lower (higher) Fox/Fpy ratios indicate that Fe sedimentation in these settings was a net source (sink) of O2 to the oceanatmosphere system. Triple-Fe-isotopic systematics are strongly diagnostic for Fox/Fpy ratios, particularly at high values of this ratio where the switch from net O2 source to sink behavior occurs (Fig. 2C). Pre-GOE pyrite data all fall at Fox/Fpy < 4 contour, and the Fox/Fpy < 10 falls outside of the error bar on individual pyrite ϵ' 56 Fe values, so we can robustly rule out net O2 sink-like behavior in the case where SO2 dominated Neoarchean volcanic emissions. Even with conservative H2S/SO2 ratios [START_REF] Gaillard | Atmospheric oxygenation caused by a change in volcanic degassing pressure[END_REF] 3). This raises the possibility that the real (non-idealized) threshold value for net O2 release during Fe upwelling and pyrite burial lies on a Fox/Fpy contour separating 2.65 Ga pyrites (represented by the four most isotopically depleted pyrite samples in this study) from younger, higher δ' 56 Fe samples (Fig. 2C).

The triple-Fe-isotope proxy provides a new and greater insight to the iron cycle in the early

Earth's oceans. Before the GOE, large and probably fluctuating continental and hydrothermal iron fluxes to the oceans [START_REF] Bekker | Iron Formation: The Sedimentary Product of a Complex Interplay among Mantle, Tectonic, Oceanic, and Biospheric Processes[END_REF] were removed to two sedimentary sinks (Fig. S7) [START_REF]Materials and methods are available as supplementary materials on Science Online[END_REF]. The major sink was Fe 3+ -(oxyhydr)oxides that were deposited from upwelling water masses in the oceans that lacked a discrete redoxcline and allowed protracted partial iron oxidation [START_REF] Rouxel | Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF][START_REF] Planavsky | Iron isotope composition of some Archean and Proterozoic iron formations[END_REF]. The secondary iron sink was on highly productive continental margins, where deposition of pyrite-rich sediments was generally sulfate-limited. Small relative changes in iron removal to these oxide and sulfide sinks potentially led to perturbations in the net O2 supply to the atmosphere, and surface ocean oxygenation (6, 7), triggering short-lived, episodic whiffs of O2 [START_REF] Olson | Volcanically modulated pyrite burial and ocean-atmosphere oxidation[END_REF]. The loss of extremely negative δ' 56 Fe values in sedimentary pyrite formed after the GOE shows how this transition irreversibly changed the early Earth's iron cycle by decreasing iron flux to the oceans following the emergence of oxidative continental weathering; developing a discrete water-column redoxcline; and increasing supply of sulfate from the continents, which resulted in a shift in the major iron sink to extensively developed euxinic settings. for H2S/SO2 input ratios of 1 (0). Blue (dashed) lines and shaded areas: mean molar O2 yield for H2S/SO2 input ratios of 1 (0) and 95 % confidence interval for pyrites grouped by age interval.

Supplementary Materials for

Materials and Methods

Methods

Pyrite synthesis experiments

To constrain the triple Fe isotopic expression of pyrite precipitation, pyrite was synthesized at the University of Edinburgh, via the FeS-H2S pathway following the methodology of Guilbaud

Analytical methods

Analytical procedures for iron purification and isotopic measurements followed standard procedures used at the Origins Laboratory at the University of Chicago [START_REF] Nie | Iron and oxygen isotope fractionation during iron UV photo-oxidation: Implications for early Earth and Mars[END_REF][START_REF] Dauphas | Routine isotopic analysis of iron by HR-MC-ICPMS: How precise and how accurate?[END_REF][START_REF] Tang | Abundance, distribution, and origin of 60Fe in the solar protoplanetary disk[END_REF][START_REF] Tang | OF METEORITES AND TERRESTRIAL ROCKS: 60Fe DISTRIBUTION AND MASS FRACTIONATION LAWS[END_REF][START_REF] Craddock | Iron Isotopic Compositions of Geological Reference Materials and Chondrites[END_REF]. Samples were prepared from powders of black shale and IF material, and hand-picked pyrite grains. Sample masses ranged from 12-22 mg, 2-6 mg, and 13-84 mg for black shale, IF, and pyrite grains respectively. Samples were digested in clean Savillex Teflon beakers. First, 1 ml 28 M HF + 0.5 ml 15 M HNO3 + a few drops 11 M HClO4 was added, and closed beakers were heated at 130 ºC.

Samples were evaporated to dryness and re-dissolved in Aqua Regia (0.75 ml 11 M HCl + 0.25 ml 15 M HNO3) and a few drops 11 M HClO4, before heating and evaporation was repeated. The Aqua Regia + HClO4 step was repeated 3 times to release all iron to solution. Samples were evaporated to dryness and 0.5 ml of 6 M HCl or 10 M HCl was added, depending on the purification procedure to be used. Larger volumes of the same acid were used for digestion of pyrite grains, which contained greater masses of Fe. Iron purification made use of both the standard 'short column' procedure, which is now routine in the Origins Laboratory [START_REF] Dauphas | Routine isotopic analysis of iron by HR-MC-ICPMS: How precise and how accurate?[END_REF][START_REF] Craddock | Iron Isotopic Compositions of Geological Reference Materials and Chondrites[END_REF], and a 'long column' procedure designed to more effectively eliminate Cu from the matrix [START_REF] Tang | Abundance, distribution, and origin of 60Fe in the solar protoplanetary disk[END_REF][START_REF] Tang | OF METEORITES AND TERRESTRIAL ROCKS: 60Fe DISTRIBUTION AND MASS FRACTIONATION LAWS[END_REF], which was a concern for sulfide samples.

Short column iron purification: Disposable Bio-Rad Poly-Prep polyethylene columns were filled with 1 ml of AG1-X8 200-400 mesh Cl-form anion exchange resin. The resin was preconditioned with 10 ml of MilliQ H2O, 5 ml of 1 M HNO3, 10 ml of MilliQ H2O, 9 ml of 0.4 M HCl, 5 ml of MilliQ H2O, and 2 ml of 6 N HCl. Samples were loaded onto columns in 0.25 ml of 6 M HCl. Matrix and interfering elements were eliminated by passing 8 ml of 6 M HCl through the column. Iron was eluted with 9 ml of 0.4 M HCl and recovered in clean Teflon beakers.

Samples were evaporated to dryness and redissolved in 0.25 ml of 6 M HCl, before repeating the column procedure a second time with new resin. All experimentally synthesized pyrite and FeSm samples were also purified using this procedure.

Long column iron purification: This alternative iron purification procedure was developed to eliminate Cu as a potentially significant matrix element associated with sulfide phases. Reusable 30 ml Savillex Teflon columns with a 0.64 cm ID capillary cut to 10.5 cm length were loaded with 3 ml of AG1-X8 anion exchange resin. The resin was preconditioned with 10 ml MilliQ H2O, 10 ml 0.4 M HCl, 5 ml MilliQ H2O, 10 ml 0.4 M HCl, and 4 ml of 4 ml of 10 M HCl. Samples were loaded onto columns in 0.25 ml of 10 M HCl. Matrix and interfering elements were eliminated by passing 4.5 ml of 10 M HCl, and 30 ml of 4 M HCl, the latter to eliminate Cu in particular. Iron was eluted with 9 ml of 0.4 M HCl and recovered in clean Teflon beakers. Samples were evaporated to dryness and redissolved in 0.25 ml of 10 M HCl, before repeating the column procedure with new resin.

Iron isotopic compositions were measured on a Neptune MC-ICP-MS at the University of Chicago. Analyses were made of the extent of isotopic fractionation (δ values), and the departure from a reference mass-dependent fractionation law (ϵ). The Fe isotopes at masses 54, 56, 57, and 58 were measured simultaneously along with 53 Cr and 60 Ni for correction of 54 Cr and 58 Ni interferences on 54 Fe and 58 Fe respectively. The 53

⁄ )

Bracketing standards are also internally normalized using the same exponential law.

Subsequent to measurements, data were renormalized to the high-temperature equilibrium limit law for display, to fit the common convention used with other isotope systems. Data normalized to the exponential law are also presented in Tables S1 andS2, and a version of Figure 1B using this normalization is also presented in Figure S1.

Samples and standards were measured in 0.3 M HNO3 and introduced into the plasma torch using a Cetac Aridus II or ESI Apex Omega desolvating nebulizer system with no auxiliary N2 flow. On-peak zero was determined at the start of each measurement sequence by analyzing a clean aliquot of the same HNO3 in which samples were measured. Sample and standard concentrations between 5 ppm and 30 ppm were used in different measurement sessions depending on sensitivity and the entrance slit being used, but most analyses making use of 10-12 ppm iron sample and standard solutions. Measurements were made with the use of bracketing standards matched to sample concentrations within ± 5 %. Measurements of 56 Fe were made on a 10 10 Ω amplifier resistor because signal intensities were generally higher than 50 V, and 10 11 Ω amplifier resistors were used for measurement of 54 Cr, 54 Fe, 57 Fe, 58 Fe, and 60 Ni.

For experimentally synthesized samples, the mass-dependent Fe isotopic fractionation (δ 56 Fe) was also determined by standard Fe isotopic analytical methods. A quartz cyclonic spray chamber was used to introduce 1 ppm solutions into the Neptune operating in medium resolution mode, resulting in a signal of ~7 V. Isotopic compositions were determined by standard-sample bracketing. All Fe isotopic analyses of experimentally synthesized samples (both triple isotopic and conventional) were normalized to IRMM-524 during analyses, but subsequently renormalized to the measured isotopic compositions of the starting batches of FeSm material used in respective experiments. All errors were combined in quadrature as appropriate. The average fractionation factor we determined for the pyrite precipitation reaction was αFeS-pyrite = 1.0023 ± 0.0003 (95% C.I.), consistent with the results of Guilbaud et al. (2011) [START_REF] Guilbaud | Abiotic Pyrite Formation Produces a Large Fe Isotope Fractionation[END_REF]. This fractionation factor was determined by calculating the average difference between the FeSm and pyrite splits from each serum bottle experiment. The degree of pyritization (DOPthe fraction of the total Fe in the pyrite pool) was calculated from the total iron masses in each split indicated by concentration measurements and known dilution factors. The low DOP values (max ~14 %) obtained in our experiments were not conducive to fitting the data to a Rayleigh distillation trend, however the distance between linear trends plotted through δ 56 Fe vs. degree of pyritization for the FeSm and pyrite data also gave an average fractionation factor of αFeS-pyrite = 1.0023 (Fig. S2).

We saw no systematic difference between short and long column purification techniques in triple Fe isotopic data for IF sample JD-C165A, and pyrite sample SF-1 599.8 Py, which were each processed multiple times using either column procedure to check the reproducibility of our measurements in the absence of geostandard materials, which have been analyzed to this level of precision (Fig. S2). In all cases, the individual pyrite analyses for this repeat sample have ϵ 56 Fe values that significantly more positive than the value of the anticipated for IF with the same δ 57 Fe value (Fig. S2). In addition, as a check for possible matrix effects in the preparation of IF and pyrite samples, we performed a matrix test with IRMM-524 standard iron solution. Briefly, aliquots of an IF sample (REX 187.5) and a pyrite sample (SF-1 623.6 Py) were passed through the short column purification procedure and the eluted matrix from each was collected. These matrix cuts were further purified by being passed through this column chemistry procedure again.

These matrix cuts were mixed with a solution of IRMM-524 back to the original iron concentration, and the iron was purified with two passes on short columns in the same manner as other samples. The ϵ 56 Fe values of both matrix-adjusted solutions and a pure solution of IRMM-524 were within error of zero and all identical within error (Fig. S2), suggesting that sample matrices did not systematically affect our ϵ 56 Fe analyses.

Modeling methods -Calculation of oxic and sulfidic sink sizes

Triple Fe isotopic systematics allow the isotopic composition of any given isotopically depleted pyrite to be deconvolved into contributions from KIE during pyrite precipitation, and isotopic depletion in dissolved Fe 2+ resulting from the removal of isotopically heavy Fe 3+

(oxyhydr)oxides. Rough estimates for those contributions can be solved graphically (Fig. 2A), however the uncertainty on those estimates requires propagation of all uncertainties on both the KIE and oxidation MFLs, in addition to uncertainty of individual pyrite measurements. This error propagation was performed using a Monte Carlo method in a MATLAB script.

For each pyrite and IF datapoint, an array of simulated datapoints was generated by randomly sampling 1000 times from a normal distribution defined by the 95 % C.I. of the measured ϵ' 56 Fe and δ' 57 Fe values and standard deviations. From the 1000 sets of randomly generated IF datapoints, Fe 2+ evolutions in ϵ' 56 Fe vs. δ' 57 Fe space that could be driven by Fe 3+ (oxyhydr)oxide removal. An array of 1000 values for the kinetic slope was generated by randomly sampling 1000 times from a normal distribution defined by the 95 % C.I. of the experimentally determined slope. The intercept of a line with the kinetic slope passing through the pyrite ϵ' 56 Fe vs. δ' 57 Fe, with the MFL fitted to the IF data, was found by solution of simultaneous equations (Fig. S4). The IF line, and the line defining the KIE trajectory which connects pyrite datapoints to the IF line, are given respectively by the equations: ϵ' 56 Fe = a ox δ' 57 Fe + b ox ϵ' 56 Fe = a KIE δ' 57 Fe + b KIE where a is the empirically determined slope of the MFL, and b is the intercept with the vertical axis. For known ϵ' 56 Fe and δ' 57 Fe values of a given pyrite, the second equation can be rearranged as:

b KIE = ϵ' 56 Fe py -a KIE δ' 57 Fe py

The intercept with the IF MFL is then found by equating the line equations, giving (after substitution):

δ' 57 Fe int = (ϵ' 56 Fe py -a KIE δ' 57 Fe py -b ox ) (a ox -a KIE )

where the subscript int refers to the value at the intercept (note in the main text this is denoted with the subscript wo) (Fig. S4). The δ' 57 Fe value at the intercept with the IF MFL (δ' 57 Feint) was taken as the isotopic composition of the pyrite-forming water mass that had already undergone iron oxidation, and the magnitude of the KIE during pyritization was calculated as the difference between the intercept value and the δ' 57 Fe of pyrite. This exercise was repeated 1000 times for each depleted pyrite sample using the randomly generated values of all parameters described. We note that in the Main Text, the discussion focusses on δ' 56 Fe values as these are more commonly reported in the literature, but here the exact procedure is detailed as performed, and conversion between δ' 57 Fe and δ' 56 Fe values is done by dividing by the values of θ 56/57 presented in this study.

The extent of Fe 2+ oxidation (Fox) to give a certain δ' 57 Feint was calculated using a Rayleigh distillation model: δ' 57 Fe int = δ' 57 Fe i + 1000(α -1) ln(1 -F ox )

where α is the fractionation factor during in Fe 2+ oxidation and precipitation that gives fractionation 1000×(α-1) = 0.5 ‰/amu during Fe 3+ (oxyhydr)oxide removal [START_REF] Dauphas | Iron Isotope Systematics[END_REF], the subscript i denotes the starting δ' 57 Fe for a hydrothermal Fe 2+ source of ~-0.3 ‰ [START_REF] Rouxel | Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF][START_REF] Beard | Iron isotope constraints on Fe cycling and mass balance in oxygenated Earth oceans[END_REF]. Iterations where the inferred δ' 57 Fe of seawater was more positive than this value were rejected as they would imply a negative degree of Fe 2+ oxidation.

The fraction of pyrite precipitation (Fpy in Main Text) was also determined in a Rayleigh distillation model. We calculate fpy by assuming that bulk pyrite nodules represent the cumulative product of pyrite precipitated from the initial Fe 2+ pool. This contrasts with in situ measurement of individual nodule layers that better approximate instantaneous precipitate compositions during the growth of pyrite grains [START_REF] Marin-Carbonne | Coupled Fe and S isotope variations in pyrite nodules from Archean shale[END_REF][START_REF] Marin-Carbonne | Situ Fe and S isotope analyses in pyrite from the 3.2 Ga Mendon Formation (Barberton Greenstone Belt, South Africa): Evidence for early microbial iron reduction[END_REF]. The magnitude of the KIE that was expressed in the product was calculated as the difference between the product and initial reactant: for each pyrite sample spanned a large range, but these variations were strongly correlated (Fig. S5). This is because a more negative estimate for the isotopic composition of seawater (which implies a larger Fox), gives a smaller estimate for the fractionation during the precipitation of pyrite from the oceanic iron pool (which implies a larger Fpy). These two effects have an opposite impact on the estimate of Fss, therefore, Fss estimates vary less than Fox or Fpy.

Triple-Fe-isotope-derived Fox and Fpy from pyrites can also be used to estimate the relative fluxes of Fe 2+ and S (as SO4 [START_REF] Moorbath | Early Archaean Age for the Isua Iron Formation, West Greenland[END_REF], and the geologic setting and Fe isotopic systematics of these IF have been discussed in the literature elsewhere [START_REF] Tsikos | Iron isotopes constrain biogeochemical redox cycling of iron and manganese in a Palaeoproterozoic stratified basin[END_REF][START_REF] Thibon | Dynamics of oceanic iron prior to the Great Oxygenation Event[END_REF][START_REF] Dauphas | Clues from Fe Isotope Variations on the Origin of Early Archean BIFs from Greenland[END_REF].

Pyrite grains from organic rich shales, a few bulk organic rich shales, and bulk IF samples were selected for this study. Pyrite grains and bulk shales were sampled from a subset of drill core

FiguresFig. 1 .

 1 Figures

Fig. 2 .

 2 Fig. 2.Interpreting triple-Fe-isotopic signatures of isotopically light pyrites. A) Schematic of two-step process involved in pyrite formation. Iron with starting compositions resembling hydrothermal fluids (gold circle) is oxidized, driving residual Fe 2+ to light compositions along the Fe 2+ oxidation MFL. Partial pyrite precipitation from this Fe 2+ subsequently causes fractionations along the kinetic pyrite formation MFL. The approach is detailed in FigureS4and equations in[START_REF]Materials and methods are available as supplementary materials on Science Online[END_REF]. B) Pyrite data and contours for Fox and fpy in triple Fe isotope space. Fpy, the fraction of total upwelled Fe deposited in pyrite, is calculated as Fpy = fpy × (1 -Fox). C) Pyrite data and contours of Fox/Fpy (relative size of oxic and pyrite sedimentary Fe sinks). Bold contours at 4 and 10 indicate thresholds for net O2 source vs. sink behavior for volcanic H2S/SO2 inputs ratios of 1[START_REF] Gaillard | Atmospheric oxygenation caused by a change in volcanic degassing pressure[END_REF] and 0 (8), respectively.

Fig. 3 D

 3 Fig. 3 D) Fox and Fpy, and molar O2 yield estimates, from Monte Carlo error propagation. Violin plots for Fox and Fpy probability densities for results determined from propagation of errors on the ϵ' 56 Fe and δ' 57 Fe values for each measured pyrite and the errors on the slopes of the MFLs (24). Blue filled (open) squares: Estimated molar O2 yields per mole of pyrite buried for individual samples

11 M

 11 et al.(2011) [START_REF] Guilbaud | Abiotic Pyrite Formation Produces a Large Fe Isotope Fractionation[END_REF] and references therein, which have consistently been demonstrated to produce pyrite precipitates from an FeSm (mackinawite) reactant. All reagents were of analytical grade, and solutions were prepared using 18 MΩcm deionized water and sparged for 30 min with O2-free grade N2 before use. Solutions were prepared and solid FeSm was synthesized in N2-filled recirculating Saffron alpha anoxic chamber under O2-free conditions. FeSm was precipitated by mixing 100 mL of 0.6 M iron (Fe 2+ ) solution prepared with Mohr's salt ((NH4)2Fe(SO4)2•6H2O; Sigma Aldrich) with 100 mL of 0.6 M sulfide solution made with Na2S•9H2O (Sigma Aldrich).This reaction produced a black precipitate. The precipitate was filtered using a Buchner filter with Whatman™ No. 1 filter paper, resuspended in sparged water and the filtration was repeated three times. The freshly precipitated FeSm was freeze-dried overnight on a Mini-Lyotrap (LTE) freezedryer then transferred back to the anoxic chamber and stored under O2-free conditions until use.The low-metal complexing MOPS (3-(N-morpholino)propanesulfonic acid) buffer (3-Nmorpholinopropanesulfonic acid, pKa = 7.31, Fisher) was made by dissolution of its sodium salt and in sparged water, buffered to pH 6 by NaOH titration. Ti 3+ citrate was prepared by adding 5 mL 15% TiCl3 to 50 mL 0.2 M Na citrate and buffered to pH 7 with Na2CO3. The solutions were stored in the glove box under O2-free conditions until use.The pyrite precipitation experiments were prepared in the glove box. Approximately 300 mg of the freeze-dried FeSm was weighed into serum bottles, 10 mL of 0.05 M MOPS buffer solution was added, and the bottles were sealed with rubber stoppers and aluminum crimper seals. The sealed bottles were attached to a gas transfer manifold via a hypodermic needle inserted through the rubber stopper, and the manifold and reaction bottle were flushed with O2-free grade N2 and pumped down to -14 PSI (-97 kPa, at full vacuum) three times. A sealed serum bottle containing 800 mg solid Na2S•9H2O was attached to the manifold via a hypodermic needle and flushed and pumped three times. A syringe was used to inject 2 mL of sparged, 50 vol% H2SO4 into the Na2S•9H2O-containing bottle to generate H2S. After H2S transfer into the FeSm-containing serum bottle, the pressure was adjusted to slight under-pressure (~ -2.5 PSI or ~ -17 kPa) by N2 addition.The needle holes in the serum bottle septa were covered with silicone sealant and transferred to a 40°C oven to allow the pyrite precipitation reaction to take place. After different, pre-determined reaction durations, the serum bottle reaction vessels were removed from the oven and frozen to stop the reaction.Once frozen, the serum bottles were unsealed under flushing N2 and excess H2S in the headspace was removed. The bottles were then re-stoppered, the stoppers pierced with a hypodermic needle under flushing N2, and the bottles left in the freeze-dryer for a day. The freezedried serum bottles were transferred to the anoxic chamber, and 2 mL sparged water and a few drops of the Ti 3+ citrate were added to poise the Eh at low negative values to prevent FeSm oxidation and ensure full dissolution of FeSm in a preferential dissolution protocol developed by[START_REF] Rickard | The composition of nanoparticulate mackinawite, tetragonal iron(II) monosulfide[END_REF] [START_REF] Rickard | The composition of nanoparticulate mackinawite, tetragonal iron(II) monosulfide[END_REF] and modified and calibrated by[START_REF] Guilbaud | Abiotic Pyrite Formation Produces a Large Fe Isotope Fractionation[END_REF] [START_REF] Heard | Constraints on the coevolution of oxic and sulfidic ocean iron sinks from Archean-Paleoproterozoic iron isotope records[END_REF]. The serum bottles were resealed and placed to a fume hood for preferential dissolution. In the fume hood, 20 mL of sparged 1.2 M HCl was injected into the serum bottle via hypodermic syringe to fully dissolve only FeSm and MOPS salt. Remaining solids, essentially pyrite, were separated by filtering on a 0.45 µm Millipore filter, and rinsing with sparged water. The FeSm in HCl solutions were adjusted to 50 mL by addition of water and a 10 mL (20%) cut was dried down in clean Savillex Teflon beakers for transport and isotopic analysis. Pyrite was dissolved with drops of concentrated HNO3, solutions were adjusted to 50 mL by addition of water and a 10 mL (20%) cut was dried down in clean Savillex Teflon beakers for transport and isotopic analysis. A 20 mL cut was taken for pyrite samples SB5 Py and SB6 Py, which were produced in short (4.66 hours) duration experiments and for which low pyrite iron yields were anticipated. In the Origins Laboratory at the University of Chicago, samples were dissolved in Aqua Regia with drops of HClO4 at 140°C, and dried down twice, then treated three times with 2 mL of H2O2 to remove organic carbon salts left in the FeSm solutions by MOPS. The solutions were then re-dissolved in 5 mL 6M HCl ready for iron purification. A small (5 µL) aliquot of each solution was dried down and redissolved in 0.3 M HNO3 to check the iron concentration of these solutions using MC-ICP-MS and determine the correct amount of volume of each sample solution to be passed through iron purification.

δ' 57

 57 Fe py -δ'57 Fe int = (F py -1) × 1000(α -1) ln(1 -F py ) F py where we assumed the maximum fractionation during pyrite precipitation (via FeSm) from Fe 2+ of 1000×(α-1) = -1.55 ‰/amu. Iterations requiring a larger instantaneous fractionation were rejected. The value of Fpy was determined by solving this equation numerically. The cumulative product composition cannot evolve past the initial composition of the reactant reservoir, when the reactant reservoir has been completely consumed. Therefore, iterations requiring a positive offset of the pyrite datapoint from δ' 57 Fe of seawater were rejected. Once the 1000 iterations for calculation of Fox, and Fpy, had been completed, the fractional size of the sulfidic sink (Fss in Main Text) was calculated as Fpy × (1 -Fox) for each pyrite with a confidence interval dictated by the range of values generated in the 1000 iterations. Monte Carlo simulation estimates of Fpy and Fox

  samples previously studied for Fe isotopic variations by[START_REF] Rouxel | Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF] 2006) [START_REF] Rouxel | Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF][START_REF] Rouxel | Response to Comment on "Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF], New pyrite grains were picked at the University of Hawaii. The nature of these grains was described in detail by[START_REF] Rouxel | Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF] [START_REF] Kump | Increased subaerial volcanism and the rise of atmospheric oxygen 2.5 billion years ago[END_REF]. Pyrite grains in organic-rich shales that were subsampled in our study occurred as nodules ~ 1 mm to 1 cm in diameter, with C-rich inclusions in variable amounts.The nodular pyrite either had no internal structure, or was composed of concentrically laminated, fine-grained pyrite or bladed pyrite crystals. Euhedral pyrite crystals commonly overgrew the outer part of the nodules. Shale lamina typically bends around pyrite nodules, which supports interpretations of their origins as being formed early on during diagenesis. Pyrite nodules often display complex features such as multiple-growth bands or composite nodules formed by coalescence of several nodules. Dissolution and reprecipitation of primary sulfide nodules could have happened in some samples and likely resulted in formation of massive pyrite, often characterized by euhedral grains free of C-rich inclusions.Localized dissolution-reprecipitation is unlikely to have affected Fe-isotopic compositions of pyrites. In the large numbers of samples analyzed per formation by Rouxel et al. (2006) (9) strongly negative δ56 Fe values were a consistent feature and no link between Fe isotopic composition and the nature of individual pyrite grains was reported, which supports the notion that these are primary sedimentary signatures and not the results of later alteration of the host rocks. Heard and Dauphas (2020)[START_REF] Heard | Constraints on the coevolution of oxic and sulfidic ocean iron sinks from Archean-Paleoproterozoic iron isotope records[END_REF] recently discussed the fidelity of the pyrite Fe isotope record as an archive of primary sedimentary signatures. In brief, the resistance of this system to metamorphic overprinting due: to the high abundance of Fe, low solubility of pyrite and small size of Fe isotopic fractionations at metamorphic temperatures all make it unlikely that primary sedimentary Fe isotopic signatures have been compromised by secondary processes that may nonetheless have affected the texture of pyrite grains. In practice, Marin-Carbonne et al. (2020)[START_REF] Marin-Carbonne | Situ Fe and S isotope analyses in pyrite from the 3.2 Ga Mendon Formation (Barberton Greenstone Belt, South Africa): Evidence for early microbial iron reduction[END_REF] recently demonstrated through in situ work that Archean pyrites that experienced late fluid circulation that led to partial recrystallization and alteration of S isotopic systematics, did not modify the Fe isotopic composition, in line with our expectations outlined above.

  Resolving the Role of Ocean Iron Sinks in Early Atmospheric Oxygenation

	Andrey Bekker
	Correspondence to: andyheard@uchicago.edu
	This PDF file includes:
	Materials and Methods
	Supplementary Text
	Figs. S1 to S7
	Tables S1 to S4

Andy W. Heard, Nicolas Dauphas, Romain Guilbaud, Olivier J. Rouxel, Ian B. Butler, Nicole X. Nie,

  Cr and 60 Ni interferences were corrected for using the exponential law. All of the Fe isotopes have molecular interferences with argide ions ( 40 Ar 14 N + ,40 Ar 16 O + ,40 Ar16 O 1 H + , and40 Ar 18 O + ), which present a significant hindrance to obtaining the requisite precision to resolve mass-dependent fractionation laws. Therefore, measurements were made on the flat-topped peak shoulder in high-resolution mode using a standard Neptune entrance slit. A few sample analyses made use of 'ultra-high-resolution mode' utilizing the highresolution mode of a Thermo Element 2 entrance slit. Results were consistent with those obtained using the standard HR method, but offered no improvement in precision whilst requiring higher iron concentrations to obtain the same signal. Nickel or aluminum sampler and H skimmer cones were used. Standard-sample bracketing was used to correct isotopic ratio measurements for instrumental mass fractionation, and Fe isotopic ratios of samples are reported relative to the average isotopic ratios of the bracketing standard solutions of IRMM-524, which has an identical isotopic composition to IRMM-014. The exponential law was initially used to calculate ϵ values by fixing57 Fe/ 54 Fe to 0.362549, the value of IRMM-014. The δ and ϵ values of samples are given

	by
					δ = [( Fe i	⁄	Fe 54 ) sample ( Fe i ⁄	⁄	54	Fe ) std	-1] ×10 3 ,
					ϵ= [( Fe i	⁄	Fe 54	) sample *	( Fe i ⁄	⁄	Fe 54 ) std *	-1] ×10 4 ,
	where i = 56, 57, or 58 and the * indicates ratios corrected for mass fractionation by internal
	normalization to a fixed reference 57 Fe/ 54 Fe ratio using the exponential law(17, 20),
	ln( Fe i	⁄	Fe 54	) sample, corrected *	= ln( Fe i	⁄	Fe 54	) sample, measured -ln	( Fe 57 ( Fe ⁄ 57 ⁄ 54 Fe Fe ) sample, measured 54 ) fixed reference	ln (m Fe i ln (m Fe ⁄ m Fe 54 57 m Fe ) 54

  2-, H2S, and S8) to shelf environments in the Neoarchean and pre-GOE Paleoproterozoic. Assuming an FeS2 stoichiometry for sedimentary sulfide, the Fe/S ratio in the shelf sediments is given by (Fe/S)sed = [Fpy + Fox]/[2 × Fpy], where Fpy and Fox are calculated based on the triple-Fe-isotopic measurements presented above. We only consider the relative magnitude they indicate that locally developed euxinic conditions[START_REF] Scott | Late Archean euxinic conditions before the rise of atmospheric oxygen[END_REF][START_REF] Reinhard | A Late Archean Sulfidic Sea Stimulated by Early Oxidative Weathering of the Continents[END_REF][START_REF] Olson | Volcanically modulated pyrite burial and ocean-atmosphere oxidation[END_REF] are not representative of the global ocean iron cycle where the burial of Fe 3+ -(oxyhydr)oxides was the dominant iron sink.For ca. 2.5 Ga pyrites, an additional source of sulfate from oxidative weathering of continental sulfides (4) might explain why (Fe/S)sed values are lower relative to those in the older Jeerinah and Lokammona formations (TableS4).Sample MaterialsGeological setting and age constraints for shale-hosted pyrite samples are given by[START_REF] Rouxel | Iron Isotope Constraints on the Archean and Paleoproterozoic Ocean Redox State[END_REF] [START_REF] Kump | Increased subaerial volcanism and the rise of atmospheric oxygen 2.5 billion years ago[END_REF]. Ages and geological units for all samples in this study are provided in TableS3. References for age constraints for pyrite and shale samples, and most IF samples, are provided in Rouxel et al. (2005; 2006) (6, 23), and Planavsky et al. (2012) (7), respectively. Age data for samples from the Hotazel Formation and Isua Supracrustal Belt IFs are provided by Gumsley et al. (2017) (2), and Moorbath et al. (1973)

	from continental weathering (0.4 Tmol/yr) and hydrothermal venting at mid-ocean ridges (0.18
	Tmol/yr) (42), to subaerial volcanic SO2 emissions (0.4-1.4 Tmol/yr) (3), continental weathering
	supply of SO4 2-(11 Tmol/yr) (42), and hydrothermal H2S supply (0.17 Tmol/yr) (42), is
	approximately 0.05. The Neoarchean-Paleoproterozoic pyrite record thus suggests a ratio of Fe to
	S supply to the oceans 14-40 times greater than modern. Archean hydrothermal and continental
	Fe 2+ fluxes exceeded modern levels, due to higher mantle heat flow (42), low sulfate content in
	the ocean (43), and continental weathering in contact with anoxic atmosphere. On another hand,
	the weathered SO4 2-flux and hydrothermal H2S flux (43) were lower throughout the Archean. Our
	calculated (Fe/S)sed values do not necessarily require weaker Archean volcanic SO2 fluxes than
	now, but

of oxide and sulfide iron sinks (e.g., iron silicates or carbonates are not considered) because they are the only ones that have large known isotopic effects, and that can influence redox balance in the Archean. Central estimates of (Fe/S)sed have a range of 0.7 to 2.0 across all samples (Table

S4

).

The Jeerinah and Lokammona formation pyrites have (Fe/S)sed ≥ 1.3, while younger pyrites deposited shortly before the GOE have lower ratios. For comparison, the ratio of modern Fe input

Acknowledgments: Discussions with Stephanie Olson and Jake Waldbauer were appreciated.

Nigel Brauser and Anne Davis are thanked for comments on the clarity of the manuscript. This work was funded by NASA grants 80NSSC17K0744 to ND and AWH (Habitable Worlds), 359NNX17AE86G to ND (LARS), and NNX17AE87G to ND (Emerging Worlds). Participation by AB was supported by NSERC Discovery and Acceleration grants.

List of Supplementary Materials:

Materials and Methods

Supplementary Text

Figs. S1 to S7

Tables S1 to S4

Supplementary Text

Background on mass fractionation laws for Fe isotopes

Instantaneous fractionations

Numerous reaction pathways have been proposed to create the >4 ‰ δ 56 Fe range in Archean IFs, shales, and pyrites. The extent of Fe isotopic fractionation is insufficient to discriminate between different scenarios for sedimentary iron cycling in the Archean oceans, because many fractionation processes can generate a large and indistinguishable range in delta values.

Considering two isotopic ratios could resolve this ambiguity, because the slope of δ 56 Fe vs. δ 57 Fe reflects the mass fractionation law (MFL) governing a process, and this law may be distinct for the fractionation processes at the root of various scenarios proposed for the Archean Fe isotopic record. Mass-dependent fractionation is described with a power law

θ 56 57 ⁄ where x αA/B are fractionation factors for isotope x between reservoirs A and B, and 56/57 θ is the mass dependent exponent or slope in triple Fe isotope space [START_REF] Dauphas | Mass Fractionation Laws, Mass-Independent Effects, and Isotopic Anomalies[END_REF]. Natural processes imparting different slopes of MFLs in three-isotope diagrams have been identified for O [START_REF] Angert | Effects of photorespiration, the cytochrome pathway, and the alternative pathway on the triple isotopic composition of atmospheric O2[END_REF][START_REF] Cao | Equilibrium mass-dependent fractionation relationships for triple oxygen isotopes[END_REF][START_REF] Young | On the Δ17O budget of atmospheric O2[END_REF][START_REF] Levin | Triple oxygen isotope variations in sedimentary rocks[END_REF], Mg [START_REF] Young | Kinetic and equilibrium mass-dependent isotope fractionation laws in nature and their geochemical and cosmochemical significance[END_REF][START_REF] Davis | Isotopic mass fractionation laws for magnesium and their effects on 26Al-26Mg systematics in solar system materials[END_REF], S [START_REF] Farquhar | Multiple sulphur isotopic interpretations of biosynthetic pathways: implications for biological signatures in the sulphur isotope record[END_REF][START_REF] Ono | Mass-dependent fractionation of quadruple stable sulfur isotope system as a new tracer of sulfur biogeochemical cycles[END_REF][START_REF] Johnston | Sulphur isotopes and the search for life: strategies for identifying sulphur metabolisms in the rock record and beyond[END_REF], Ca [START_REF] Zhang | Calcium and titanium isotopic fractionations during evaporation[END_REF], Ti [START_REF] Zhang | Calcium and titanium isotopic fractionations during evaporation[END_REF], and Fe [START_REF] Mccoy-West | The Fe and Zn isotope composition of deep mantle source regions: Insights from Baffin Island picrites[END_REF][START_REF] Nie | Iron and oxygen isotope fractionation during iron UV photo-oxidation: Implications for early Earth and Mars[END_REF], but this has yet to be investigated for Fe isotopes in sedimentary rocks with sufficient precision to resolve distinct slopes. The slope θ 56/57 for the triple-Fe-isotope diagram is given by θ 56 , where δ' is related to the standard δ notation by δ' = 1000×ln((δ 1000 ⁄ )+1).

Slopes vary only subtly between MFLs and therefore it is convenient for the purposes of visualization to express one isotopic ratio in terms of its deviation from an arbitrary reference law in parts per 10,000 by using ϵ' notation [START_REF] Dauphas | Mass Fractionation Laws, Mass-Independent Effects, and Isotopic Anomalies[END_REF][START_REF] Nie | Iron and oxygen isotope fractionation during iron UV photo-oxidation: Implications for early Earth and Mars[END_REF] where ϵ' 56 Fe = [δ' 56 Fe-0.672×δ' 57 Fe]×10.

In ϵ' 56 Fe vs. δ' 57 Fe diagrams MFLs are straight lines, and when the high-temperature equilibrium limit law with θ 56/57 = 0.678 is used as the reference law, ϵ' 56 Fe values are 0 if fractionation follows the exponential law. Other MFLs have positive (negative) slopes if θ 56/57 is larger (smaller) than 0.678.

These laws describe mass-dependent fractionation in a single step process. This approach is an oversimplification in cases where isotopes have been fractionated via several geochemical pathways, via Rayleigh distillation, or where samples reflect mixtures of two (or more) isotopically distinct pools of the element of interest. These concerns are well-documented in the more mature field of triple O and S isotope study [START_REF] Angert | Effects of photorespiration, the cytochrome pathway, and the alternative pathway on the triple isotopic composition of atmospheric O2[END_REF][START_REF] Cao | Equilibrium mass-dependent fractionation relationships for triple oxygen isotopes[END_REF][START_REF] Young | On the Δ17O budget of atmospheric O2[END_REF][START_REF] Farquhar | Multiple sulphur isotopic interpretations of biosynthetic pathways: implications for biological signatures in the sulphur isotope record[END_REF][START_REF] Ono | Mass-dependent fractionation of quadruple stable sulfur isotope system as a new tracer of sulfur biogeochemical cycles[END_REF][START_REF] Johnston | Sulphur isotopes and the search for life: strategies for identifying sulphur metabolisms in the rock record and beyond[END_REF], but we show here that these concerns are of diminished importance in application to low-temperature Fe isotope systematics. This is because both Rayleigh distillation and mixing produce curves in ϵ' 56 Fe vs. δ' 57 Fe space that are practically indistinguishable from instantaneous MFLs over the natural range of Fe isotopic variations.

Rayleigh distillation

In the case of Rayleigh distillation, Ono et al. [START_REF] Ono | Mass-dependent fractionation of quadruple stable sulfur isotope system as a new tracer of sulfur biogeochemical cycles[END_REF] have shown that the closed-system evolution of a reactant reservoir (A) during formation of a product (B) results in an observed slope in three-isotope space for A that is distinct from the intrinsic slope of the instantaneous fractionation process (Fig. S3A). In this study, the evolution of the reactant reservoir corresponds to the generation of an isotopically light Fe 2+ pool through the removal of an isotopically heavy Fe 3+ oxide product. The evolution of the reactant, written in δ' notation, is

where δ' x FeA,i is the initial isotopic composition of the reactant, δ' x FeA is the isotopic composition of the reactant when a fraction 54 fA of the reactant A remains, and x α is the isotopic fractionation factor for isotope x in the reaction of A to form product B. In three isotope space, the isotopic composition of the reactant will evolve with an effective slope, θ 56/57 eff θ 56 57 ⁄ eff = δ' 56 Fe A -δ' 56 Fe A,i δ' 57 Fe A -δ' 57 Fe A,i = ( α-1

56

)×ln(

where θ 56/57 inst is the intrinsic slope for the instantaneous reaction. The distinctions between θinst and θeff are significant in the O and S isotope systems, where fractionations and relative isotopic mass differences are large. In the case of Fe isotopes, specifically Rayleigh distillation of aqueous Fe 2+ driven by oxidation and removal of Fe 3+ minerals, the relevant values for θ 56/57 inst and 57 α are 0.678 and on the order of 1.0015 respectively [START_REF] Guilbaud | Abiotic Pyrite Formation Produces a Large Fe Isotope Fractionation[END_REF]. Inserting these values results in θ 56/57 eff ≈ 0.6778, which is smaller than the intrinsic slope by only 0.0002 and not resolvable from the intrinsic slope for any naturally occurring range of fractionations (Fig. S3A). Nie et al. [START_REF] Nie | Iron and oxygen isotope fractionation during iron UV photo-oxidation: Implications for early Earth and Mars[END_REF] provided a slightly different relation between the effective and intrinsic slopes for UV photooxidation, with 56 α = 1.0012, and determined the same result, that the effective slope for the evolving reactant reservoir was smaller than the instantaneous slope by just 0.0002 and thus the two slopes were indistinguishable within current measurement uncertainties. These calculations imply that discussing theoretical, single-step MFLs is an appropriate approximation for the evolution in triple Fe isotope space of an Fe 2+ reservoir affected by oxidation and removal of Fe 3+ products.

Mixing

Unlike fractionation processes, mixing lines are non-linear in logarithmic coordinates for isotopic ratios (Fig. S3B), so another pertinent question for this study is the degree to which ϵ' 56 Fe vs. δ' 57 Fe data may be explained by mixing systematics rather than distinct intrinsic MFLs. We define two isotopic reservoirs A and B as endmembers in a mixing array, and for the sake of simplicity (though this is not required), describe the linear tie-line between these two points as an instantaneous mass fractionation with the intrinsic slope and fractionation factors θ 56/57 inst and αA/B, respectively. Following Ono et al. [START_REF] Ono | Mass-dependent fractionation of quadruple stable sulfur isotope system as a new tracer of sulfur biogeochemical cycles[END_REF] the magnitude of the effect of mixing on ϵ' 56 Fe values is given as ϵ' 56 Fe mix -ϵ' 56 Fe MFL = [ln (

where ϵ' 56 Femix and ϵ' 56 FeMFL refer to the points on the mixing curve and the MFL connecting A and B, respectively, and 54 fA is the fraction of 54 Fe from A in the mixture. For ϵ' 56 Fe vs. δ' 57 Fe coordinates, this mixing curve is concave upward (Fig. S3B), analogous to the familiar example of Δ 33 S vs. δ 34 S. When the most highly fractionated value in the Archean dataset, δ' 57 Fe ≈ -5 ‰, is used as the endmember mixing with another endmember at δ' 57 Fe = 0 ‰, the maximum difference between ϵ' 56 Femix and ϵ' 56 FeMFL is 0.0069 (Fig. S3B), which is an order of magnitude smaller than typical analytical precision for measurement of these values and hence they are unresolvable. Therefore, while mixing lines in these coordinates are not strictly linear, in practice they are very well approximated by straight lines and our data cannot be explained through nonlinear mixing of endmembers lying on a single MFL.

Iron isotope MFLs

Mass-dependent triple-Fe-isotopic systematics have been explored in few publications to date, and only once previously in the context of low temperature aqueous geochemistry. [START_REF] Tsikos | Iron isotopes constrain biogeochemical redox cycling of iron and manganese in a Palaeoproterozoic stratified basin[END_REF], are all within error of one another suggests that fractionations of Fe isotopes driven by Fe 2+ -Fe 2+ equilibration may follow the equilibrium MFL regardless of the oxidation process involved. It will be important for future studies to constrain the value of θ 56/57 for the remaining proposed oxidation pathway for IF, anoxygenic photoferrotrophy [START_REF] Kappler | Deposition of banded iron formations by anoxygenic phototrophic Fe(II)-oxidizing bacteria[END_REF][START_REF] Konhauser | Iron formations: A global record of Neoarchaean to Palaeoproterozoic environmental history[END_REF], however our results to date suggest it is unlikely that triple Fe isotopic systematics will be able to identify the oxidation pathway for IF due to the tendency of Fe 2+ and Fe 3+ to rapidly isotopically equilibrate.

Precipitation of pyrite is a kinetically controlled process associated with a large kinetic isotope effect (KIE) that enriches early precipitates in the light isotopes of Fe [START_REF] Guilbaud | Abiotic Pyrite Formation Produces a Large Fe Isotope Fractionation[END_REF][START_REF] Rolison | Iron isotope fractionation during pyrite formation in a sulfidic Precambrian ocean analogue[END_REF]. The slope θ 56/57 is anticipated to be shallower than that of equilibrium isotope fractionation [START_REF] Young | Kinetic and equilibrium mass-dependent isotope fractionation laws in nature and their geochemical and cosmochemical significance[END_REF][START_REF] Dauphas | Mass Fractionation Laws, Mass-Independent Effects, and Isotopic Anomalies[END_REF], an expectation borne out in triple Fe isotopic measurements of olivine phenocrysts affected by iron diffusion [START_REF] Mccoy-West | The Fe and Zn isotope composition of deep mantle source regions: Insights from Baffin Island picrites[END_REF]. However, the value of θ 56/57 relevant to pyrite precipitation did not have an empirical constraint prior to our study. Our pyrite precipitation experiments resulted in maximum DOP of ~14% (Fig. S2). By mass balance most iron was always left in the FeSm pool and large fractionations from the starting composition of the experiment were observed in the pyrite pool, which provides leverage to determine the slope of the instantaneous MFL associated with pyrite precipitation. Our triple Fe isotopic analysis was consistent with a single kinetic MFL, with a slope of 0.0233 ± 0.0057 in ϵ' 56 Fe vs. δ' 57 Fe space, which translates to a slope θ 56/57 KIE = 0.6743 ± 0.0006. This is much shallower slope than that of the equilibrium limit law associated with redox equilibrium, but steeper than the exponential kinetic law that describes pure atomic transport [START_REF] Young | Kinetic and equilibrium mass-dependent isotope fractionation laws in nature and their geochemical and cosmochemical significance[END_REF][START_REF] Dauphas | Mass Fractionation Laws, Mass-Independent Effects, and Isotopic Anomalies[END_REF].

Figures

Fig. S1.

Triple-Fe-isotope systematics for IFs, pyrites, and black shales in ϵ' 56 Fe vs. δ' 57 Fe space, normalized to the exponential law. Error bars and envelopes are 95% confidence intervals. The slopes of endmember MFLs associated with iron-redox processes, and KIEs during pyritization, are constrained through analysis of isotopically light Mn-rich IFs and laboratory pyrite precipitation via the H2S pathway [START_REF] Guilbaud | Abiotic Pyrite Formation Produces a Large Fe Isotope Fractionation[END_REF][START_REF]Materials and methods are available as supplementary materials on Science Online[END_REF][START_REF] Rickard | Chemistry of Iron Sulfides[END_REF], respectively. The slope of the IF MFL agrees with the theoretical equilibrium law, implying control by Fe 2+ -Fe 3+ equilibrium [START_REF] Dauphas | Mass Fractionation Laws, Mass-Independent Effects, and Isotopic Anomalies[END_REF][START_REF] Nie | Iron and oxygen isotope fractionation during iron UV photo-oxidation: Implications for early Earth and Mars[END_REF]. Pyrite synthesis defines a kinetic MFL for pyrite precipitation. Pre-GOE pyrites fall in an intermediate space between redox equilibrium and kinetic endmembers. 57 Feint and δ' 57 Fepy gives the expression of the KIE during pyritization, which is used to determine the degree of pyritization (fpy) of the pre-pyritization water mass assuming that the pyrite is a cumulative product of all precipitated pyrite. δ' 57 Feint is assumed to be the δ' 57 Fe value of the pre-pyritization water mass, and its isotopic composition is assumed to reflect the degree of isotopically heavy Fe 3+ -(oxyhydr)oxide removal (Fox) that took place prior to the formation of pyrite. Uncertainties are 95 % confidence intervals from Monte Carlo simulations. O2 yields assume volcanic H2S/SO2 input ratio of 0. For volcanic H2S/SO2 input ratio of 1, O2 yields are smaller by a factor of 2.5


K Isoird

M Lazar

M L Locatelli

C Raynaud

D Planson

J P Chante

Study of 4H-SiC High Voltage Bipolar Diodes Under Reverse Biases Using Electrical and OBIC Characterizations

Keywords: OBIC, JTE, Bipolar diode, Breakdown Voltage, Surface charge

4H-SiC P + NN + structures have been fabricated following Medici TM software simulation in order to block voltages as high as 6 kV. In particular, these diodes are realized by surrounding the emitter by a Aluminum-implanted ring called Junction Termination Extension (JTE). Electrical characterizations under reverse bias at room temperature and in various environments (air, SF6) show a premature breakdown of the diodes. This breakdown is localized at the emitter periphery. OBIC (Optical Beam Induced Current) measurements show a peak of photocurrent at the emitter junction edge, indicating the presence of a high electric field. These results involve an effectiveness of 60 % of the JTE. This is probably related to a low electrical activation of the implanted aluminum during the post-implantation annealing and to the presence of positive charges at the surface of the devices.

Introduction.

Silicon carbide presents interesting electrical properties for many applications and especially for the high voltage devices. To fully benefit the advantages of SiC and avoid any premature breakdown of the high voltage diodes, it is necessary to realize efficient junction terminations. The purpose of this work is to estimate the Junction Termination Extension (JTE) influence on the electrical behavior of 4H-SiC P + NN + diodes under reverse biases. In this paper we present current-voltage (I-V) characteristics measured in two different gaseous ambients (air, SF6). Analyses of these results are completed by OBIC (Optical Beam Induced Current) measurements. This technique allows to investigate the electric field distribution in the structure and especially at the junction periphery.

Experiments

Device structure P + NN + diodes are formed by multiple aluminum implantations at room temperature in n-type 4H-SiC epilayer (thickness = 40 µm, doping level = 1.1´10 15 cm -3 confirmed by C(V) measurement over a few µm depth from the surface). The JTE is created by five successive aluminum implantations at room temperature. Main parameters of the junction termination (depth = 0.5 µm, doping level = 2´10 17 cm -3 , length = 250 µm) are optimized with Medici TM software simulations (the ionization coefficients used are those given by Konstantinov et al. [1]) in order to obtain blocking voltages as high as 6 kV. Then the samples are annealed at 1700 °C during 30 minutes [2], and finally metallized for both anode and cathode ohmic contacts. No passivation layer covers the semiconductor surface. Two types of diodes are tested : diodes without JTE (D1), and diodes with JTE region (D2).

OBIC setup and method

The OBIC method has been previously used to investigate the influence of JTE on the electric field distribution in bipolar diodes. For photo-excitation a UV light is used at different wavelengths (ranging from 363.8 nm down to 300 nm) which is focused at the surface of the sample and modulated at 70 Hz by a chopper. The curves are obtained by scanning a diameter of the diode (400 µm diameter) with the laser beam (3.3 µm diameter). The detailed descriptions of the OBIC method and setup are given in Reference [3].

Results.

Electrical characterizations.

The I-V characteristics and tests of breakdown voltages were realized under DC conditions at room temperature in two different ambients (air and SF6 gas). The demonstrations of the breakdown were observed by means of a CCD camera and a video monitor. The forward I-V characteristics present good rectifying behaviors: the typical forward drop voltage at 300 K is 3 V at 10 A cm -². Under reverse bias in air ambient at 300 K, the curve presents two distinct zones (Fig. 1). In the first part of the initial characteristic, the current remains weak up to a reverse bias VR of -200 V. Beyond this voltage, the current rises quickly until VR = -400 V and tends to saturate in the range [10 -4 -10 -3 ] A/cm². When VR < -500 V, electric arcs begin to appear between the anode contact and some close etched or metallized devices. This I-V curve is not reproducible. During the second measurement in the first part, we observe an increase of the current. For the second part of the characteristic, we note that for a same voltage the measured current is generally weaker. Those results are completed by I-V measurements performed in SF6 ambient under small overpressure (Fig. 2). In these conditions we observe the same behavior than in the air ambient. However, luminous white points at the emitter periphery accompany the electric arcs. Towards 2000 V the arcs lead to the diode breakdown. This illustrates a localized breakdown of the diode induced by a high electric field at the P + N junction edge. For the diodes without junction termination, the maximum breakdown value reached is in the range [700-1000] V, these results are in good agreement with those obtained by MEDICI simulations. Fig. 1: Reverse electrical characteristics of diodes with JTE (D2), measured at 300 K in the air. Two bias sweeps are presented for each diode. The order of appearance in the legend corresponds to the chronological order of I-V measurement performed on each diode. Fig. 2: Reverse electrical characteristics of diodes with JTE, measured at 300 K in the SF6 ambient using a Tektronix 370A curve tracer. Three bias sweeps are presented for this diode and are numbered in the order of measurement.

OBIC measurements

Fig. 3 shows OBIC measurements performed on a D1 diode along a diameter line using 363.8 nm wavelength corresponding to a penetration length in 4H-SiC higher than 40 µm [4]. We observe an OBIC signal when the spot light illuminates the emitter non-covered by the metal. The photocurrent increases with the reverse voltage due to the space charge region (SCR) spreading inside the epilayer. At VR = -300 V, a peak of photocurrent emerges at the emitter edge, which indicates carriers multiplication due to a localized high electric field. The OBIC signal measured on a D2 diode with JTE (Fig. 4) presents the same behavior. No photocurrent is detected over JTE position but we recognize the signal due to the P + N junction which increases with the reverse voltage. However at VR = -300 V we detected a weak signal on the part of JTE close to the emitter. In order to observe the effect of JTE on electric field distribution it is necessary to investigate a lower depth of the epilayer by measuring OBIC signal using a shorter wavelength.

Fig. 5: OBIC measurements performed on a D2 diode with JTE at l = 300 nm with Popt = 2 for different reverse voltages.

Fig. 5 shows OBIC measurements on a D2 diode illuminated at a wavelength of 300 nm. Contrary to the curves obtained at l = 363.8 nm, at VR = 0 V we detected an OBIC signal on part of JTE. This photocurrent increases with the voltage and we also note that when the bias increases, the area where the signal is detected extends laterally. Those results illustrate the lateral spreading of the SCR under JTE region [5] that indicates a total depletion of JTE.

Discussion.

The electrical results have shown a premature breakdown of diodes with junction termination. SIMS analyses at different places on a wafer reveal an aluminum implanted dose of JTE lower than the

N P + x [µm] Photocurrent [A] V R = -300 V V R = -200 V V R = -100 V V R = 0 V -2
W cm 0 100 200 300 400 500 600 700 800 900 0,0 5,0x10 -10 1,0x10 -9 1,5x10 -9 2,0x10 -9 2,5x10 -9

JTE JTE P + x [µm] Photocurrent [A] V R = -300 V V R = -200 V V R = -100 V V R = -50 V V R = 0 V -2
W cm 0 100 200 300 400 500 600 700 800 900 1000 10 -11 1x10 -10 1x10 -9 1x10 -8

Lateral spreading of SCR under JTE region

JTE JTE P + Contact x [µm] Photocurrent [A] V R = 0 V V R = -50 V V R = -100 V V R = -200 V V R = -300 V -2 W cm
desired one (1´10 13 cm -2 ). The simulations carried out with these effective JTE doping values give a theoretical breakdown voltage (3 kV) higher than experimental value. Measurements obtained with OBIC method show an premature depletion of the JTE area and a localized breakdown at the emitter edge. Those results tend to confirm a certain protection inefficiency probably due to an aluminum electrical activation1 lower than 100 % [3]. However the low aluminum electrical activation does not seem to be the only explanation of the poor JTE efficiency. Indeed OBIC measurements performed on several diodes have revealed the presence of a high electric field at the emitter edge under weak voltages. Those results and the instability of current during I-V measurements can be due to the presence of a surface charge. The different steps of the process flow are certainly at the origin of this surface charge. Indeed the presence of a positive charge on the SiC surface tends to tighten the equipotential lines at the junction periphery which induces a degradation of the JTE performance [6] and causes the appearance of high electric field at the emitter edge. This last assumption seems to be confirmed by simulation of D1 and D2 diodes with a positive charge density of 7´10 11 cm -2 . We note that the positive charge has no influence on the breakdown value of D1 diodes. On the contrary, by comparing the results of D1 diode simulation with or without surface charge, we observe for a same voltage an increase of the electric field at the emitter periphery. For the D2 diode we clearly observe an effect of surface charge on the JTE performance. By taking account of the positive surface charges, the efficiency of junction termination is reduced by about 12 %. The premature breakdown of the components can be induced by a combined action of these two phenomena (presence of a positive charge on the SiC surface and low electrical activation of aluminum dopant).

Conclusion.

The OBIC measurements performed on planar 4H-SiC bipolar diodes have allowed to investigate the JTE influence on the electric field distribution at the device periphery and the place of the breakdown. These results correlated with the electrical characterizations in different ambients tend to prove that the premature breakdown of our structures is due to partial electrical activation of the implanted aluminum dose in the JTE guard and to the presence of positive charge on SiC surface.

2 ]Fig. 3 :Fig. 4 :

 234 Fig. 3: OBIC measurements performed on a D1 diode without JTE at l = 363.8 nm with Popt = 2 for different reverse voltages.

The aluminum electrical activation is the ratio of Al concentration incorporated in SiC lattice sites on the implanted Al concentration.

Acknowledgements.

This work was supported by GIRCEP French group, among which DGA and Schneider Electric Research Center. We are also grateful to R. Nipoti from LAMEL-Bologne and her colleagues for the realization of the ohmic contacts.


F Nallet

P Godignon

D Planson

C Raynaud

J P Chante

Realization of a High Current and Low RON 600V Current Limiting Device

Keywords: 4H-SiC, MOSFET, Current Limiting Device, Simulation, I-V Characteristics, Forward Current Density, Saturation

The first experimental results of a 600 V 4H-SiC current limiting device are shown. This device limits the current as the bias voltage increases. The forward conduction is ensured by an N type implanted channel (doping species: nitrogen) on top of a P + implanted layer (doping species: aluminum). The prototypes reach a saturation current density of 900 Acm -2 , with a specific onresistance of 13 mWcm². The 4H-SiC current limiting devices belong to the best set of Accu-MOSFETs devices reported in the literature.

Introduction

Silicon carbide technologies for the manufacturing of power devices seem to be promising in the near future. The first SiC power devices commercially available are Schottky diodes proposed by Infineon technologies (ranging: 300 V / 10 A and 600 V / 4-6 A) and Microsemi. In the field of the SiC switches, the current trends are the Accu-MOSFET and JFET structures which exhibit the best specific on-resistance/breakdown voltage ratio experimentally obtained (Accu-MOSFET: 16 mWcm 2 /1600 V [1] ; JFET: 10 mWcm 2 /600 V [2], 14 mWcm 2 /1800 V [2]). The physical properties of SiC allow to develop new solid state devices for applications not covered by the silicon technologies. In this sense, a silicon carbide current limiting component should be a solid state structure which has to work under high voltage and high temperature conditions when a shortcircuit current appears. The function requires to limit a short-circuit overcurrent during a time long enough to allow the opening of the power line by a circuit breaker under reliable conditions. Such a 4H-SiC component corresponding to the End User specified requirements (600 V / 50 A) led to the definition of the device technological parameters assisted by the finite-element simulation software DESSIS ISE ® developed by ISE TM TCAD [START_REF]ISE Integrated System Engineering, DESSIS, Operator's manual[END_REF]. The electrical static behavior of the first process run, realised at CNM (Barcelona), are discussed and led us to change some technological parameters in order to reduce the specific on-resistance and increase the saturation current density.

Device structure, simulation and technology

The design of the component is based on an Accu-MOSFET device whose structure is presented in Fig. 1. The wafer is composed by a N ++ substrate layer (r =20 mW.cm / 430 µm) on which a N type epitaxial layer (4´10 15 cm -3 / 7 µm) was grown (purchased from CREE Research). A single Al implantation is used for the realization of the buried P + region and the guard ring protection. Then, a full wafer multiple N implantation is processed to implement the N-channel (the N implantation come over the Al implantation, so the N-channel is a high compensated N-layer). After that, N ++ implantation for ohmic source contact, SiC etching to reach the P + layer, dry oxidation and both side metallization are processed to manufacture the device without passivation layer. In case of positive drain to source voltage bias, the conduction current is carried by the electrons whose path inside the device crosses the entire thickness of the wafer, the 'intercell' region and the channel region successively. Fig. 1 shows the device structure with its guard ring protection. For a drain to source voltage VDS>0 V, the drain to source current IDS increases linearly with VDS.

When the device reached the saturation mode, the current can not increase any more because the reverse space charge region of the P + /NCHANNEL junction is wide enough to pinch-off the channel. Consequently, the device is normally-on and the current limiting behavior is unidirectional (for VDS>0 V). Electrical simulations, with DESSIS ISE ® [3], were performed [START_REF] Nallet | [END_REF] to optimize the structure characteristics : low on-resistance (10-20 mWcm 2 and high saturation current density (500-1000 Acm -2 ). The technological choices led to a 6 mask levels process using standard process steps commonly used in a Si clean room except for the high temperature post-implantation annealing which is realized in our specific annealing equipment. The device voltage handling capability is 700 V with the specified epitaxial layer and a non optimized guard ring protection. The expected I-V characteristics given by a DESSIS ISE ® simulation using implanted doping profiles simulated by a specific software developed by E. Morvan [5] are shown in Fig. 2 (P + : 350 keV/4´10 13 cm -2 ; NCHANNEL: 30-180 keV/1.3´10 13 cm -2 ).

Fig. 2: Device static I-V simulation results obtained by DESSIS ISE ® on a limiting device with the following parameters: (P + : 350 keV/4´10 13 cm -2 ; NCHANNEL: 30-180 keV/1.3´10 13 cm -2 ).

Experimental results

The prototypes are designed following a multiple hexagonal cells device, with different values of the channel length (4, 5 and 6 µm) and the 'intercell' width (4, 6 and 8 µm) (see Fig. 1). The active area (surface of the source contact) is about 0.05 mm 2 and the conduction perimeter is 2880 µm.

] V DS [V]

In the first batch prototypes, ohmic contacts are made with nickel and annealed at 850°C. The 'gate' oxide thickness is about 90 nm and the post-implantation annealing process used is 1700°C during 30 min in an inductive furnace. The electrical characteristics were measured using a Keithley K237 (PMAX=10 W that means a 100 mA/100 V range). The best devices show a normally-on behavior with a high specific on-resistance (RON-SPE) of 150 mWcm 2 and a saturation current density of 150 Acm -2 . We observe a very soft saturation region as shown in the I-V characteristics of Fig. 3. The specific on-resistances obtained are very high compared to the expected value (10-20 mWcm 2 ) (Fig. 2) and the saturation current region is not clearly demonstrated. SIMS (Secondary Ion Mass Spectroscopy) analysis indicate a channel depth of 0.1 µm with an average doping level of 2D10 17 cm -3 . TLM (Transmission Line Method) measurements combined with the doping profiles (extracted from the SIMS profiles) allow us to estimate a channel electron mobility of 100 cm 2 V -1 s -1 . The effective charge density at the SiO2/4H-SiC interface were extracted from C-V measurements and are found to be negative in the 1.5-2D10 12 cm -2 range. This explains the normally-off behavior of most devices (SiO2/4H-SiC interface charge density is close to the channel dose, 2D10 12 cm -2 ). To improve the device characteristics, some technological changes were used for the second prototypes realization. The static I-V characteristics (measured with a Tektronix 370 curve tracer) of the second run prototypes show a very good behavior. The effects of the channel length and the intercell width are clearly visible and indicate that the intercell width has a low effect on RON-SPE, but decreasing LCHANNEL lead to clearly decrease RON-SPE. The best specific on-resistances measured are in the 13-17 mW cm 2 range (Fig. 4) and the highest saturation current density is 900 A cm -2 (Fig. 5), obtained for the devices designed with a channel length of 4 µm and an intercell width of 8 µm. These results are not so far from the DESSIS ISE ® simulations performed with the second run projected technological data. Further analysis will be necessary to extract the channel mobility and the SiO2/4H-SiC interface charge density. Nevertheless, it seems that a high channel electron mobility (close to the simulation value, 400 cm 2 V -1 s -1 ) is obtained with a low level of SiO2/4H-SiC interface charge density, when comparing the simulated I-V characteristic (without SiO2/4H-SiC interface charge density specification) and the corresponding experimental results of Fig. 5. Fig. 4: Experimental Specific on-resistance plotted versus the intercell width.

Fig. 5: Simulation/experiment comparison for current limiting devices (channel length: 4 µm / intercell width: 8 µm). For the experimental results using the Tektronix 370 curve tracer, the temperature inside the device, particularly at the end of the channel, is unknown due to self heating (I-V characteristic was limited to 30 V to avoid the device destruction. The 600 V capability should require a pulsed method to avoid self heating and show the real I-V static characteristic).

Conclusion

In this paper we presented the first experimental results of a new normally-on Accu-MOSFET for current limitation applications. The fabricated devices show a very high saturation current density (900 Acm -2 ) and a low specific on-resistance (13 mWcm 2 ) without gate drive, close to the best characteristics reported for a JFET device [2] for the 600 V breakdown voltage range. The fabrication-technology used is based on an implanted channel and only 6 mask levels are necessary.

Fig. 1 :

 1 Fig. 1: Vertical cross section of the device and its guard ring peripheral protection (the channel region, LCHANNEL, and the 'intercell' region are noted).

Fig. 3 :

 3 Fig. 3: Sampling of different I-V static characteristics obtained on limiting devices from the first batch (25°C).

Acknowledgment

We acknowledge SCHNEIDER ELECTRIC S.A. for financial and technical supports, and CNM of Barcelona who manufactured the prototypes.