# Sulfur isotope mass-independent fractionation in impact deposits of the 3.2 billion-year-old Mapepe Formation, Barberton Greenstone Belt, South Africa

Mark van Zuilen, Pascal Philippot, Martin Whitehouse, Aivo Lepland

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Mark van Zuilen, Pascal Philippot, Martin Whitehouse, Aivo Lepland. Sulfur isotope mass-independent fractionation in impact deposits of the 3.2 billion-year-old Mapepe Formation, Barberton Greenstone Belt, South Africa. Geochimica et Cosmochimica Acta, 2014, 142, pp.429-441. 10.1016/j.gca.2014.07.018. insu-03581092

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# Sulfur isotope mass-independent fractionation in impact deposits of the 3.2 billion-year-old Mapepe Formation, Barberton Greenstone Belt, South Africa

M.A. van Zuilen<sup>a,*</sup>, P. Philippot<sup>a</sup>, M.J. Whitehouse<sup>b</sup>, A. Lepland<sup>c,d,e</sup>

Theoretical and experimental studies have shown that atmospheric SO₂ isotopologue self-shielding effects in the 190–220 nm region of the solar spectrum are the likely cause for mass independent fractionation of sulfur isotopes (S-MIF). The main products of this photochemical reaction – SO₃ and S⁰ – typically define a compositional array of ca. Δ³³S/δ³⁴S = 0.06–0.14. This is at odds with the generally observed trend in Archean sulfides, which broadly defines an array of ca. Δ³³S/δ³⁴S = 0.9. Various explanations have been proposed, including a diminution of δ³⁴S caused by chemical and biogenic mass-dependent fractionation of sulfur isotopes (S-MDF), mixing with photolytic products produced during felsic volcanic events, or partial blocking of the low-wavelength part of the spectrum due to the presence of reduced atmospheric gases or an organic haze. Early in Earth history large meteorite impacts would have ejected dust and gas clouds into the atmosphere that shielded solar radiation and affected global climate. It is thus likely that at certain time intervals of high meteorite flux the atmosphere was significantly perturbed, having an effect on atmospheric photochemistry and possibly leaving anomalous sulfur isotopic signatures in the rock record. Here we describe the sulfur isotopic signatures in sulfides of spherule beds S2, S3 and S4 of the Barberton Greenstone Belt, South Africa. In particular, in spherule bed S3 – and to a lesser extent S4 – a trend of ca. Δ³³S/δ³⁴S = 0.23 is observed that closely follows the expected trend for SO₂-photolysis in the 190–220 nm spectral range. This suggests that an impact dust cloud (deposited as spherule beds), which sampled the higher region of the atmosphere, specifically incorporated products of SO₂ photolysis in the 190–220 nm range, and blocked photochemical reactions at higher wavelengths (250–330 nm band). By implication, the generally observed Archean trend appears to be the result of mixing of different MIF-S sources arising from a variety of photochemical reactions that took place in the lower part of the atmosphere.

## 1. INTRODUCTION

It is generally accepted that mass-independent fractionation of sulfur isotopes (S-MIF) – defined as (1)

$$ \Delta^{33}\text{S} = \delta^{33}\text{S} - 1000 * [(1 + \delta^{34}\text{S}/1000)^{0.515} - 1], \quad \text{with} \quad (2) \\ \delta^{3x}\text{S} = [(^{3x}\text{S}/^{32}\text{S})_{\text{sample}}/(^{3x}\text{S}/^{32}\text{S})_{\text{CDT}} - 1] * 1000 \quad \text{is caused by photochemical reactions of gaseous sulfur compounds in a reduced atmosphere (Farquhar et al., 2000). In the} \\ \text{>2.4 Ga Archean atmosphere the photolytic disproportionation of volcanic SO}_2 \text{ produced elemental sulfur (S}^0\text{) and sulfate (SO}_4\text{) aerosols with a positive (Δ}^{33}\text{S} > 0) and} $$

<footnote><sup>a</sup> Géobiosphère Actuelle & Primitive, Institut de Physique du Globe de Paris, Sorbonne Paris Cité, Univ Paris Diderot, UMR 7154 CNRS, F-75005 Paris, France</footnote>

<footnote><sup>b</sup> Swedish Museum of Natural History, Box 50007, SE104-05 Stockholm, Sweden</footnote>

<footnote><sup>c</sup> Geological Survey of Norway, Leiv Eirikssons vei 39, 7491 Trondheim, Norway</footnote>

<footnote><sup>d</sup> Tallinn University of Technology, Institute of Geology, 19086 Tallinn, Estonia</footnote>

<footnote><sup>e</sup> Centre for Arctic Gas Hydrate, Environment and Climate, University of Tromsø, 9037 Tromsø, Norway</footnote>

<footnote>* Corresponding author. Tel.: +33 (0)183957527.</footnote>

<footnote>E-mail address: vanzuilen@ipgp.fr (M.A. van Zuilen).</footnote>
negative (Δ³³S < 0) S-MIF signature, respectively. Although the actual mechanism for photolytic isotope fractionation is still debated, it is clear that sulfur S-MIF is largely caused by self-shielding of SO₂ in the spectral range 190–220 nm (Lyons, 2007, 2009; Ono et al., 2013). It was recently shown that photo-excitation of SO₂ in the spectral range 250–330 nm could also lead to S-MIF (Whitehill and Ono, 2012; Whitehill et al., 2013). Furthermore, mass-independent kinetic isotope effects in reactions such as SO photolysis or S + S₂ = S₃, can also contribute to the S-MIF signal (Masterson et al., 2011). Overall, once S-MIF is created it requires two or more atmospheric exit channels – such as e.g. elemental sulfur aerosols (S⁰) and sulphuric acid droplets (H₂SO₄) – for preservation in the rock record (Ono et al., 2003; Lyons, 2007). Since these exit channels can vary as a function of reducing conditions (presence of e.g. CH₄, hydrocarbon haze), S-MIF variation through time can provide insight in changing atmospheric compositions throughout the Archean Eon (Zerkle et al., 2012; Kurzweil et al., 2013).

Early in Earth history the relatively high frequency of large meteorite impacts would have caused ejection of dust and gas clouds into the atmosphere that shielded solar radiation and affected global climate (Toon et al., 1997). It is thus reasonable to assume that at certain time intervals of high meteorite flux the atmosphere was significantly perturbed, possibly causing anomalous sulfur isotopic signatures that could be preserved in the rock record. The actual induced changes to the atmosphere depend on the characteristics of the impact target; carbonate deposits release large amounts of CO₂ and CO, organic-rich deposits generate soot (a particularly strong absorber of short-wave radiation), evaporites release SO₂- or SO₃-gases, and ocean impacts cause the discharge of large amounts of water vapor (Pierazzo and Artemieva, 2012). A large meteorite impact could thus have caused a short-term shift in atmospheric composition, and therefore possibly have left a characteristic sulfur-MIF signature in Archean impact-related deposits. Directly after an impact a cloud of molten rock droplets is generated, which is quenched in the atmosphere and settles as spherules. Several of such impact-related spherule beds have been recognized in Early Archean terrains (Kyte et al., 2003; Lowe et al., 2003; Glikson, 2010). Their impact-related origin is substantiated by syn-sedimentary rip-up clasts (derived from high-energy erosion caused by tsunami waves), extraterrestrial Ir and Cr-isotope anomalies (Kyte et al., 2003), and the presence of Ni-spinel (Krull-Davatzes et al., 2010). Since such spherules were formed high in the atmosphere, they could have adsorbed sulfur aerosols and therefore have recorded specific photolytic isotope anomalies.

Here we present the variation of δ³⁴S and Δ³³S in sulfides of spherule beds in the 3.24 Ga Mapepe Formation, Fig Tree Group, Barberton Greenstone Belt, South Africa and compare this with other sediments in the Mapepe Formation and underlying Mendon Formation (Philippot et al., 2012). Furthermore, a comparison is made with the generally observed trend of Archean sulfides (Archean Reference Array, ARA), and an assessment is provided of the temporary influence of meteorite

impacts on the composition and shielding capacity of the Archean atmosphere.

## 2. GEOLOGY AND SAMPLE DESCRIPTION

The Onverwacht Group, including the Mendon Formation, consists primarily of komatiitic and basaltic volcanic rocks, alternating with cherty sedimentary layers (Fig. 1). The overlying Fig Tree Group, including the Mapepe Formation, primarily consists of terrigenous sedimentary rocks interbedded with felsic volcaniclastic rocks. The entire sequence has been affected by hydrothermal alteration and lower greenschist-facies metamorphism. This caused serpentinization of komatiites, and pervasive silicification and/or carbonation of all igneous and sedimentary sequences. For a description of the different spherule bed units within the Barberton Greenstone Belt (BGB) we adopt here the stratigraphy given in Lowe et al. (2003) (Fig. 1). Spherule bed S2 occurs in the southern part of the BGB at the base of the Fig Tree Group and has been dated at 3260 Ma (Byerly et al., 1996). It consists of a low proportion of spherules (predominantly <10%, and up to 50% at the type locality) mixed with silt to cobble-sized cherty detritus eroded from cherts and komatiitic volcanic units of the underlying Mendon Formation (Lowe et al., 2003). Two separate graded beds have been identified in S2 that have been interpreted as the result of surge-like current and/or wave events associated with an impact event (Lowe et al., 2003). In the southern part of the BGB, spherule bed S3 occurs within the Mapepe Formation about 50–150 m above the base of the Fig Tree Group (Fig. 1). In the northern part of the bed of the BGB, the Onverwacht-Fig Tree contact is dated at 3243 Ma (Kröner et al., 1991), and spherule bed S3 occurs directly at the base of the Fig Tree Group (Fig. 1). At its type locality in the southern part of the BGB, S3 occurs as a 70 cm layer within a silicified section of sandstone, mudstone and conglomerate, called Jay's Chert (Lowe and Nocita, 1999). S3 is the most widely outcropping spherule bed in the BGB, and has been studied in great detail (Krull-Davatzes et al., 2012). Several examples have been shown of S3 slabs that consist entirely of 0.1–4 mm size spherules (Lowe et al., 2003), and a variety of spherule types have been described in literature. In some sections that show little or no evidence of reworking by currents, a bimodal distribution is observed of large silica-rich spherules (microcrystalline quartz with relict fan-shaped growth textures) and smaller Cr-rich sericite spherules (Lowe and Byerly, 1986). Both types commonly have thin, opaque rutile-rich rims. Minor mineral phases in these spherule beds typically include chlorite, spinel, apatite, rutile, barite, feldspar, siderite, and sulfide minerals (primarily pyrite, and minor chalcopyrite). Diagenesis and metamorphism have caused pervasive alteration of all mineral phases, the only intact primary mineral phase being Ni-rich chromite (Krull-Davatzes et al., 2012). Spherule bed S4 is only known from one outcrop, stratigraphically 6 meters above S3, within Jay's Chert. It consists largely of spherules, irregular chlorite-rich grains, and fine chloritic matrix, and occurs interbedded with coarse, conglomeratic, highly energetic fan-delta deposits (Lowe et al., 2003).
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Fig. 1. Stratigraphy of the Onverwacht Group and Fig Tree Group, Barberton Greenstone Belt (BGB), South Africa (after Lowe et al., 2003). Spherule beds are indicated by S1–S4 throughout the sections of the Hooggenoeg Formation and the Mapepe Formation. The southern part of the BGB contains the Songimvelo Block (Lowe et al., 1999), south of the Granville Grove Fault. The northern part of the BGB contains areas north of the Inyoka Fault.

Rock samples were collected from spherule beds S2, S3, and S4 during the 2003 Astrobiology field excursion (organized by D. Lowe and G. Byerly) (Lowe and Byerly, 2003). Sample S2 (25°54.557'S, 31°01.009'E) was collected ca. 1 km west along strike from key locality SAF-374 (Lowe et al., 2003). Samples S3 (25°54.897'S, 31°01.134'E) and S4 (25°54.907'S, 31°01.135'E) were collected at the type locality SAF-179 (Lowe et al., 2003). Conventional 30 µm petrographic thin sections were prepared of each sample. After examination by petrographic microscope, areas of interest were cut from the thin sections by diamond wire saw, embedded in an Epotec 301 epoxy mount, and polished using a diamond paste. A total of seven polished slabs (S2: SPH-5, S3: SPH-1, SPH-6, AL-1, S4: SPH-2, SPH-4, SPH-8) were prepared for SIMS analysis and Raman spectroscopy. Samples of spherule bed S2 consist of silicified detritus (Fig. 2a). The individual clasts consist of black chert and volcanic rock material from the underlying Mendon Formation. The succession largely represents sedimentary deposition from large wave currents, and spherules occur only sporadically (Fig. 2b). The few spherules that were observed did not contain any pyrite, so only pyrite

grains were selected from a black chert clast (zone 1, Fig. 2c and d), and from the hydrothermal silica matrix (zone 2, Fig. 2e and f). In thin sections from spherule beds S3 and S4, many spherules are observed (Fig. 3), which consist largely of chlorite and quartz. In some cases pyrite grains occur within or line the outer parts of individual spherules (Figs. 3c, and 4a and b), but pyrite is also found scattered throughout the thin sections. In both S3 and S4 large euhedral pyrite grains are observed, which clearly represent later addition of a sulfur source during hydrothermal alteration (Fig. 5).

## 3. METHODS

### 3.1. Raman spectroscopy

A Renishaw In Via Raman spectrometer coupled to an Olympus BX61 confocal microscope was used to identify pyrite and other possible sulfide phases (for detailed description of instrument characteristics and analytical settings, see Sforna et al., 2014). Measurements were made through a 50× objective in static mode (fixed at
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Fig. 2. (a) Thin section image of sample SPH-5 of spherule bed S2. The sample consists of silicified detritus, with clasts representing black chert and volcanic deposits of the underlying Mendon Formation. Two zones were cut out, one representing a black chert clast (zone 1), and one representing the silicified matrix (zone 2). (b) Transmitted light image of a spherule containing an opaque rutile-rich rim. Spherules are rare in bed S2, and occur disseminated throughout the rock matrix. (c) Transmitted light image showing a black chert clast in zone 1, consisting of fluffy carbonaceous grains embedded in finely laminated quartz. Closely associated with the carbonaceous phase are two small pyrite grains and one larger aggregate of pyrite grains. (d) Reflected light image showing the pyrite grains and the locations of ion probe analyses. (e) Transmitted light image showing zone 2 consisting of a large aggregate of small pyrite grains within a silicified zone of the sample. (f) Transmitted light image showing the locations of ion probe analyses. In all panels the scale bar is 100 µm.

1150 cm⁻¹) for 2 × 20 s running time, using a 514 nm laser (on-sample intensity of ca. 0.4 mW). Pyrite was identified with 337, 372, and 425 cm⁻¹ Raman peaks by direct comparison with pyrite standard R070173 from the RRUFF online database. Raman analysis was performed within SIMS pits to ensure that each isotope analysis was actually made in a pyrite crystal (Fig. 4).

## 3.2. Secondary ion mass spectrometry (SIMS)

Polished rock thin section samples were mounted in 25 mm diameter epoxy blocks. Subsequently, a section was cut from these blocks and gold-coated before co-mounting with a set of epoxy-embedded sulfide standards. These standards include the S-MDF pyrites Balmat

(δ³⁴S = +15.1 ± 0.2‰) and Ruttan (δ³⁴S = +1.2 ± 0.1‰) (Crowe and Vaughan, 1996), as well as S-MIF pyrite 248474 (δ³⁴S = +1.99 ± 0.36‰, Δ³³S = 3.31 ± 0.38‰) from a ca. 3.7–3.8 Ga banded iron formation of the Isua Greenstone Belt, Southern West Greenland (Whitehouse, 2013). The δ³⁴S and δ³³S values of sample spots were calibrated relative to the Ruttan reference pyrite, since there is a discrepancy in literature values for Balmat (Crowe and Vaughan, 1996; Cabral et al., 2013; Whitehouse, 2013). Pyrite from banded iron formation GGU 248474 (Isua Greenstone Belt, SW Greenland) was used as a monitor for Δ³³S measurements. Measurements of ³²S, ³³S, and ³⁴S were made in multicollector mode using the Cameca IMS 1280 ion microprobe at the NordSIMS facility of the Swedish Museum of Natural History, Stockholm. In order
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Fig. 3. Transmitted light images of individual spherules: (a) Spherules in bed S3 consisting primarily of chlorite, with small circular cavities filled with quartz. (b) Spherule in bed S3 showing a quartz-filled center, and a larger chlorite layer. (c) Spherule in bed S3 consisting mainly of chlorite, containing some small pyrite grains. (d) Spherule from bed S4, showing a similar chlorite filling and a quartz-filled center. Scale bars in all panels are 100 µm.

to remove the gold coating on the sample surface a
25 × 25 µm rastered area was pre-sputtered for 90 s, fol-
lowed by analysis. Sulfur isotopes were analyzed as S⁻ ions
produced by bombardment with a 2.5 nA intensity critically
focused (Gaussian) primary Cs⁺ beam (operated at 10 kV).
Charge compensation during analysis was achieved with a
low energy electron gun. The analytical pits were 10 µm
in diameter. A 2500 µm field aperture was used correspond-
ing to a field of view on the sample of 30 × 30 µm (magni-
fication 90×), and analyses were made at a mass resolving
power (M/ΔM) of 4860, sufficient to resolve <sup>33</sup>S⁻ from
<sup>32</sup>S<sup>1</sup>H<sup>-</sup>. Typical ion intensities of 9 × 10<sup>8</sup> counts per second
 cps) were obtained on <sup>32</sup>S<sup>-</sup>. On small grains a single SIMS
spot analysis was obtained, while on some larger euhedral
sulfides several spot analyses were made. Sulfur isotopic
ratios are calculated using equations 1 and 2, and are pre-
sented in conventional delta notation in per mil (‰) units
relative to CDT. Standard bracketing was achieved before
and after each set of 6 sample spot analyses. The average
external reproducibility, as estimated from repeated mea-
surements of the sulfur standards, was generally below
0.3‰ for δ<sup>33</sup>S, 0.2‰ for δ<sup>34</sup>S, and 0.3‰ for Δ<sup>33</sup>S (2
std.dev.).

4. RESULTS

Pure pyrite grains that were identified in the thin sec-
tions were further divided into a few large euhedral crystals
(ca. 300–600 µm) and smaller crystals (<200 µm). The euhe-
dral pyrite crystals were used as a point of reference during
the SIMS analytical sessions, and many individual point

analyses were performed on them. For these large crystals
only the average values are reported (Fig. 5). Examples of
typical analysis spots are shown in Fig. 4a-c. In spherule
typical analysis spots are shown in Fig. 4a-c. In spherule
bed S3, a large zoned pyrite grain was identified (Fig. 4d),
and many SIMS analyses were carried out on the centre
床 ( zone A), first growth zone (zone B), and outer rim (zone
C). All SIMS spot analyses of spherule beds S2, S3, and
S4 are listed in Table S1. There is no correlation observed
between ion count rate (displayed as the ratio of ion counts
between ion count rate (displayed as the ratio of ion counts
on the sample and ion counts on a pyrite standard) and iso-
tope ratios ³³S/³²S and ³⁴S/³²S (Fig. S1). Measurements
with relative count rates against reference pyrite below 0.5
have been excluded from the dataset, assuming that such
values indicate that a non-pyrite or even a non-sulfide tar-
get was (partially) hit by the ion beam. Some pyrite grains
contain small impurities, such as chalcopyrite (Figs. 4e and
f and 5c) or anatase (Fig. 4g). The effect on isotope analysis
of chalcopyrite is shown in Fig. 5d. A difference in δ³⁴S of
ca. 4‰ is observed between chalcopyrite and pyrite. This
could represent a real isotopic difference between the two
sulfide phases, or it represents a mineral matrix effect that
is introduced during mass-dependent instrumental fraction-
ation. Raman spectroscopic analysis of SIMS pits (Fig. 4h)
was used to identify a chalcopyrite matrix, and these analyt-
ical points were removed from the dataset.

Spherule bed S2 contains two types of pyrite grains; one
type that only displays S-MIF (Δ³³S = 0 to +2.0‰) and the
other type that only displays S-MDF (δ³⁴S = -4.0 to 0.1‰)
(Fig. 6a). The observed sulfur isotope ratios thus appear to
confirm that the studied pyrite grains in S2 are all derived
from either erosional clasts of the Mendon Formation or
<figure><img src="image_5.png" /><figcaption>Fig. 4. (a) Example of SIMS analysis on a pyrite grain within a polished mount of spherule bed S3, sample SPH6. Transmitted light image showing a small round spherule (left) and the rim of a large spherule (right). Hydrothermal fluids caused permineralization of the spherules, leading to blue–grey euhedral quartz grains. Scale bar is 100 µm. (b) Reflected-light image showing the location of a pyrite grain with three ion probe pits. (c) Zoom-in on pyrite grain showing the ion probe pits of points 70, 71 and 72. For each pit a 25 µm pre-sputtering rectangle is visible, as well as the 10 µm area of analysis. Scale bar is 15 µm. (d) Zoned pyrite crystal in spherule bed S3, sample AL-01. Ion probe pits cover the entire surface from center to rim. Scale bar is 50 µm. (e) Reflected light image of sample SPH6 of spherule bed S3, showing a spherule structure containing several sulfide grains. Scale bar is 50 µm. (f) Zoom-in on upper sulfide grain, shown in (e). Raman spectroscopy in all ion probe pits confirmed that the grain is pyrite. However, one small region represents chalcopyrite. Scale bar is 15 µm. (g) Zoom-in on lower sulfide grains, shown in (e). Raman spectroscopy confirmed that these grains are chalcopyrite. The grain on the right also includes small crystals of anatase. Scale bar is 15 µm. (h) Raman spectra of pyrite, chalcopyrite, and anatase. Spectra are obtained in ion probe pits shown in (f) and (g). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>
<figure><img src="image_6.png" /><figcaption>Fig. 5. (a) Reflected light image of a large euhedral pyrite crystal in sample SPH-6 of spherule bed S3. Scale bar is 100 µm. (b) Electron microscope image (secondary electrons) of the same pyrite crystal, showing the exact locations of the ion probe pits. Scale bar is 100 µm. (c) Reflected light image of the red zone shown in (b). SIMS analyses 12 and 13 were obtained from a small chalcopyrite inclusion. Scale bar is 10 µm. (d) Sulfur isotope results of all analyses, showing a significant mass-dependent matrix effect for the two analyses on the chalcopyrite inclusion. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

from a sulfur source that was introduced at a later stage
during hydrothermal alteration. In this bed there are no
indications of an impact-related sulfur source. In contrast,
pyrite grains from spherule bed S3 appear to describe a
unique trend (Fig. 6c) that reflects mixing between a distinct
S-MIF end-member (δ³⁴S = +10‰, Δ³³S = +2‰) and the
juvenile mantle reservoir (δ³⁴S = 0‰, Δ³³S = 0‰). This
mixing trend is also observed in the zoned crystal shown
in Fig. 4d. The centre (zone A) of this crystal appears to
have preserved the S-MIF end-member, while the outer
zones B and C have mantle values (Fig. 6d). A similar trend
towards this distinct S-MIF end-member, although less well
defined, is observed in pyrite grains of spherule bed S4
(Fig. 6b). Overall, these trends suggest that spherule beds
S3 and S4 retained a primary sulfur component that may
be related to the impact-origin of these rocks.

5. DISCUSSION

**5.1. Comparison with other rocks of the Mapepe Formation**
**and underlying Mendon Formation**

Philippot et al. (2012) showed that pyrites in felsic ash layers in the Mapepe Formation of the Barberton Greenstone Belt show a negative Δ³³S/δ³⁴S trend (Fig. 7a). This trend, which is referred to as the Felsic Volcanic Array (FVA) has been attributed to an episode of intense volcanic SO₂ discharge into the atmosphere. Many pyrites in barite deposits of the Mapepe Formation show strongly negative δ³⁴S values, which can be attributed to mass-dependent

sulfur isotope fractionation caused by microbial and/or abi-
ologic sulfate-reduction (MDF-SR) (Shen et al., 2001, 2009;
Philippot et al., 2007; Ueno et al., 2008; Philippot et al.,
2012; Roerdink et al., 2013). Philippot et al. (2012) con-
cluded that the entire range of observed δ³⁴S and Δ³³S val-
ues of these barite-deposit pyrites can be explained by
different degrees of mixing between a photochemical S⁰
source related to the FVA and sulfides derived from
MDF-SR (Fig. 7b). In contrast, komatiitic volcanic units
and deep water carbonaceous cherts at the top of the under-
lying Mendon Formation do not shown any significant neg-
ative shift in δ³⁴S, and do not follow this FVA. Rather, they
overlap with the generally observed positive trend in
Archean sediments (Fig. 7a). The slope of this trend varies
throughout the Archean, and according to depositional
environment (Ono et al., 2003, 2009a,b; Kamber and
Whitehouse, 2007; Kaufman et al., 2007; Ueno et al.,
2008; Zerkle et al., 2012). Particularly in late Archean sed-
iments an overall slope of ca. Δ³³S/δ³⁴S = 0.9 is observed
(Ono et al., 2003, 2009a,b; Kamber and Whitehouse,
2007; Kaufman et al., 2007).

Recently, Roerdink et al. (2013) reported sulfur isotope
compositions of pyrites in barite deposits, volcanoclastic
deposits, conglomerates, cherts, and dolomites of the
Mapepe Formation (Fig. 7c). They observed a comparable
range in sulfur isotopic ratios as Philippot et al. (2012) with
the exception of the strongly anomalous FVA values in vol-
canic ash layers. They also noted a specific difference in iso-
topic trends between barite-rich sediments and barite-free
sediments, and concluded that most of the observed
<figure><img src="image_7.png" /><figcaption>Fig. 6. S-MDF (δ³⁴S) and S-MIF (Δ³³S) of all point analyses on pyrites in spherule bed samples. All data points shown in these figures are listed in Table S1. (a) Spherule bed S2; pyrite grains occurring in a black chert clast (red circles) and pyrite grains occurring in the silica matrix (green circles). (b) Spherule bed S4; pyrite grains (red circles). (c) Spherule bed S3; pyrite grains (red circles), large euhedral pyrite grains used as reference (blue squares), and a zoned pyrite crystal (yellow triangles). (d) Detailed analyses of the zoned pyrite grain from spherule bed S3, shown in Fig. 4d. A clear difference is observed between the center (zone A, black triangles), and the two outer rims (zones B, grey circles, and C, white squares). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

S-MDF and S-MIF values in the Mapepe Formation can be explained by shifts in environmental conditions in the depositional basin rather than mixing of products of primary atmospheric processes.

The important observation is made here that the S-isotope compositions in pyrites of spherule bed S3 (Fig. 7) – and to a lesser extent spherule bed S4 – show an anomalous trend that is completely different from the FVA trend (Philippot et al., 2012), cannot be explained by the mixing wedge proposed by Philippot et al. (2012), and cannot be explained by the environmental controls that were proposed by Roerdink et al. (2013). Only some large euhedral pyrite crystals (Fig. 5) and the outer rims of the zoned crystal (regions B and C) in spherule bed S3 (Fig. 4d) show values that are shifted to low δ³⁴S and Δ³³S, representing late stage mixing. For spherule bed S2 the pyrites found in the black chert clast (Fig. 2c and d) have δ³⁴S and Δ³³S values (Fig. 6a) directly overlapping with those of carbonaceous cherts of the underlying Mendon Formation (Fig. 7a, Philippot et al., 2012). Pyrites from the silica matrix (Fig. 2e and f) fall in the general mixing range defined by

Philippot et al. (2012) and Roerdink et al. (2013). It is concluded that pyrites of spherule beds S3 and S4 describe a unique trend that likely represents primary deposition of an impact dust cloud and carries an S⁰ component of atmospheric SO₂ photochemistry.

## 5.2. Possible sources of sulfur preserved in spherule beds

The observed S-MIF anomaly is only observed in a few small pyrite grains, or in the center of larger zoned pyrites. Clearly, the atmospheric photochemical S⁰ aerosol products that carried this anomalous isotopic trend, and that were incorporated into the dust cloud, only represent a small fraction of the total sulfide inventory in the spherule bed deposits. The spherule beds represent a giant impact. The bolide is estimated to have been ca. 30 km in diameter, causing a crater of ca. 500 km in diameter with a depth of ca. 16 km (Krull-Davatzes et al., 2014). On hitting the Earth's surface it would have released an enormous amount of water vapor, and would have excavated through sediments, the oceanic crust, into the upper mantle
<figure><img src="image_8.png" /><figcaption>Fig. 7. (a) $\delta^{34}\text{S}-\delta^{33}\text{S}$ plot of spherule bed pyrites (red circles), compared with pyrites from komatiites and black cherts of the Mendon Formation (black circles) and pyrites from the barite deposits and volcanoclastic ash layers of the Mapepe Formation (green circles) of the Barberton Greenstone Belt (data points from Philippot et al., 2012). Pyrites from the volcanic ash layers define a negative $\delta^{33}\text{S}/\delta^{34}\text{S}$ trend (Felsic Volcanic Array, FVA), while many of the pyrites in the barite deposits show a shift to strongly negative $\delta^{34}\text{S}$ caused by microbial and/or abiologic mass-dependent sulfate reduction (MDF-SR). The grey circles represent the generally observed Archean Reference Array (ARA). Sulfur isotope composition of sulfate is shown for barite deposits of the Mapepe Formation (purple crosses, data points from Bao et al. (2007) and Roerdink et al. (2012). (b) Zoom-in of Fig. 4a, showing the mixing wedge defined by the FVA and MDF-SR. (c) Sulfur isotopic composition data from Roerdink et al. (2013) for the Mapepe Formation. A clear difference is seen between pyrites in barite-rich sediments (blue circles) and barite-poor sediments (light green circles). A trend of MDF-SR is observed that is similar to that reported by Philippot et al. (2012). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

(Krull-Davatzes et al., 2014). The resulting dust cloud was globally distributed (Lowe et al., 2003), and consisted of rock melt particles and condensation particles, largely representing cherty sediments, basaltic rock and meteorite material. Only a small fraction of this dust cloud consisted of micron- and sub-micron-size particles that would have remained in the atmosphere for several months (Toon et al., 1997). Any atmospheric photochemical $S^0$ aerosols would thus mix with the sulfur component already present in the dust cloud. This dust-derived sulfur largely represents basaltic rock and meteorite material, and would thus carry a mantle-type sulfur isotopic signature. Indeed, it is seen that the mixing trend of the observed S-MIF anomaly crosses the origin in a $\delta^{34}\text{S}-\delta^{33}\text{S}$ plot (Fig. 6b–d). Another sulfur

source that would dilute the S-MIF anomaly of $S^0$ aerosol products is represented by the sulfide products of microbial and/or abiologic sulfate-reduction (MDF-SR). The Archean ocean episodically contained concentrations of sulfate that were high enough to cause precipitation of barite deposits (Huston and Logan, 2004; Philippot et al., 2012) and isotopic evidence of contemporaneous MDF-SR has been presented (Ohmoto et al., 1993; Roerdink et al., 2012, 2013; Shen et al., 2001, 2009; Ueno et al., 2008). One such short duration episode occurred 3.26–3.24 Ga ago, and overlapped with the deposition of spherule bed S3 (Fig. 1). Isotopic evidence for MDF-SR in sulfides of barite-rich rocks in the Mapepe Formation (Philippot et al., 2012; Roerdink et al., 2012, 2013)
(Fig. 7) further supports the possibility that a fraction of the sulfur inventory in spherule bed S3 could be derived from a seawater sulfate source. Indeed, some individual pyrite grains that were measured by SIMS showed a shift towards negative δ³⁴S (Fig. 6b and c). Overall, it can be concluded that sulfur from several sources was incorporated into the spherule beds, and aerosol products of high-altitude atmospheric photochemical reactions only formed a very small fraction of this total sulfur inventory. Nevertheless, in situ analysis of individual pyrite grains using SIMS showed that a specific anomaly could be identified and is preserved in the rock record.

### 5.3. Photochemistry of atmospheric SO₂ and the general Archean δ³⁴S-Δ³³S trend

The molecule SO₂ has two dominant adsorption bands in the UV-region, one at 190–220 nm and the other at 250–330 nm (Heicklen et al., 1980). In the region 190–220 nm, absorption is assigned to the electronic transition C¹B₂-X¹A₁, and photo-dissociation of SO₂ proceeds by the reaction: SO₂ + hv → *SO₂ → SO + O (where *SO₂ represents electronically excited SO₂). Subsequent photolysis of SO proceeds by the reaction: SO + hv (<230 nm) → S + O, while atomic oxygen reacts with residual SO₂ to form SO₃ (SO₂ + O → SO₃). Theoretical calculations and experimental studies have shown that SO₂ photolysis in the 190–220 nm range is associated with an SO₂ isotopologue self-shielding effect and leads to S⁰ products with positive δ³⁴S and Δ³³S that form an isotopic compositional array of ca. Δ³³S/δ³⁴S ≈ 0.06–0.14 (Lyons, 2007, 2009; Ono et al., 2013; Whitehill and Ono, 2012) (Fig. 8a). The δ³⁴S of the photolysis product S⁰ increases with increasing SO₂ pressure, and the slope Δ³³S/δ³⁴S increases with decreasing pressures of a nitrogen bath gas (Ono et al., 2013). At a certain maximum SO₂ pressure (at ca. 10 mbar SO₂) the self-shielding effect reaches saturation and δ³⁴S does not further increase. At even higher SO₂ pressures (>30 mbar SO₂), such as in the experiments of Masterson et al. (2011) using He as a bath gas, the SO₂ itself behaves as a bath gas and there is a diminution of the S-MIF signal.

In the region 250–330 nm of the spectrum absorption is attributed to the B¹B₁-X¹A₁ electronic transition, and photo-excitation of SO₂ proceeds by the reaction: SO₂ + hv → *SO₂. Photolysis does not occur, and instead photo-oxidation through self-reaction with ground state SO₂ takes place: *SO₂ + SO₂ → SO₃ + SO. The produced SO can react with other SO molecules (SO + SO → S + SO₂), or react with SO₃ (SO + SO₃ → 2SO₂). Photochemical experiments in the 250–330 nm region (Whitehill and Ono, 2012; Whitehill et al., 2013) showed that the product S⁰ has a relatively weak positive δ³⁴S, but strongly positive Δ³³S (Fig. 8a). The highest S-MIF signals are produced at relatively low SO₂ pressure, confirming that isotopologue self-shielding is not the dominant effect. Rather, it is suggested that differences in intersystem crossing (ISC) rates from an initially excited singlet state to a reactive or dissociative triplet state is responsible. The excited-state SO₂ which carries the S-MIF signature is difficult to preserve, and

<figure><img src="image_9.png" /></figure>

<figure><img src="image_10.png" /></figure>

Fig. 8. (a) Experimental data from Ono et al. (2013) showing δ³⁴S and Δ³³S values of S⁰ products from SO₂ photolysis in the 190–220 nm spectral range (circles), and experimental data from Whitehill et al. (2013) showing δ³⁴S and Δ³³S values of S⁰ products from SO₂ photoexcitation in the 250–330 nm spectral range (triangles). For Ono et al. (2013) experimental results at different bath gas (nitrogen) pressures are shown; purple circles at pN₂ = 1 bar, blue circles at pN₂ = 0.5 bar, and yellow circles at pN₂ = 0.25 bar. For Whitehill et al. (2013) experimental results at different total pressure are shown; purple triangles at P = 1013 mbar, blue triangles at P = 506 mbar, yellow triangles P = 253 mbar. Other trends shown in this figure are the Archean Reference Array (ARA) with a slope of Δ³³S = 0.9 δ³⁴S, and the Felsic Volcanic Array (FVA) with a slope of Δ³³S = -0.4 δ³⁴S. (b) The spherule bed S3 (red circles) defines a slope of Δ³³S = 0.23 δ³⁴S. This strongly deviates from the ARA trend but is rather close to the 0.06–0.14 slopes observed in the experiments of Ono et al. (2013) and Whitehill and Ono (2012). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

therefore unlikely to be recorded into the rock record. However, it has been shown (DeWitt et al., 2010) that UV radiation of SO₂ in the presence of methane can produce organosulfur aerosols, which could potentially act as an exit channel for this S-MIF signal. Also the excited-state SO₂ could react with small organic molecules such as C₂H₂ to form organosulfur aerosols (Whitehill and Ono, 2012).

The SO₂ photolysis reaction in the 190–220 nm adsorption range has been generally seen as the main source of atmospheric S-MIF in the Archean (Farquhar et al., 2000; Pavlov and Kasting, 2002; Lyons, 2007; Danielache et al., 2008; Masterson et al., 2011). However, in order to
explain the strongly positive $\Delta^{33}\text{S}$ values (>2.5‰) that are often observed, a very high atmospheric $\text{SO}_2$ column density should have been maintained throughout the Archean. Moreover, the observed positive $\Delta^{33}\text{S}$ values should be associated with very strongly positive $\delta^{34}\text{S}$ values. Clearly the ARA is at odds with theoretical and experimental studies of $\text{SO}_2$ photolysis in the 190–220 range (Fig. 8a). Lyons (2009) presented various explanations for a diminution of $\delta^{34}\text{S}$, including chemical and biogenic S-MDF effects during elemental sulfur formation in the atmosphere, or aqueous phase reactions of $\text{HS}^-$ with $\text{Fe}^{2+}$ and $\text{FeS}$ in the oceans. Whitehill and Ono (2012) suggested that a mixture of sulfur products from $\text{SO}_2$ photo-excitation in both the 190–220 nm and the 250–330 nm absorption bands could explain the specific trend that is observed in sulfide deposits throughout the Archean. This could be attained if the atmosphere contained methane, leading to the formation of organosulfur aerosols as an additional exit channel (DeWitt et al., 2010) and the enhanced preservation of the signature of $\text{SO}_2$-photo-excitation in the 250–330 nm range. Alternatively, the presence of an organic haze layer (Sagan and Chyba, 1997; Haqq-Misra et al., 2008) could have specifically blocked the lower part of the solar spectrum (Wolf and Toon, 2012), which would have reduced the rate of direct $\text{SO}_2$-photolysis in the 190–220 nm range and at the same time would have provided an environment rich in alkanes and $\text{C}_2\text{H}_2$ in which the signature of $\text{SO}_2$-photo-excitation in the 250–330 nm range could be preserved (Whitehill and Ono, 2012).

### 5.4. Impact-induced $\delta^{34}\text{S}$–$\Delta^{33}\text{S}$ trend in early Archean spherule beds

The overall trend $\Delta^{33}\text{S}/\delta^{34}\text{S} = 0.23$ that is observed in the pyrite grains in spherule bed S3, and to a lesser extent S4, deviates significantly from the ARA, but is relatively close to the theoretical and experimental trend for $\text{SO}_2$-photolysis in the 190–220 nm spectral range ($\Delta^{33}\text{S}/\delta^{34}\text{S} \approx 0.06–0.14$) (Lyons, 2007, 2009; Ono et al., 2013; Whitehill and Ono, 2012) (Fig. 8b). Since this trend is specifically observed in spherule beds, it is likely that it is related to an impact-induced disturbance in the atmosphere. The impact of a large bolide can cause the ejection of enormous amounts of water, gases, and rock dust, leading to an optically thick layer in the atmosphere (Toon et al., 1997) that blocks radiation in both the UV and visible range (Pierazzo and Artemieva, 2012). Under such circumstances any photochemical reactions would be largely prevented in the lower part of the atmosphere, and it is likely that the dust itself incorporated S-MIF products that were produced by a photolytic process higher up in the atmosphere.

In the modern atmosphere a specific S-MIF signature is typically observed at high altitude after episodes of large volcanic eruptions (Savarino et al., 2003; Baroni et al., 2007, 2008). Stratospheric sulfate aerosols (SSA) deposited in polar ice recorded an S-MIF signature that appears to be specifically linked to $\text{SO}_2$-photolysis in the 190–220 nm range (Ono et al., 2013). This effect is generally explained by an increased $\text{SO}_2$ column density in the stratosphere causing an isotopologue self-shielding effect. It is possible

that the Archean spherule beds recorded a similar upper-atmosphere S-MIF signature that was generated exclusively by $\text{SO}_2$-photolysis. At such high altitude in the Archean atmosphere methane and organic haze gases such as alkanes $\text{C}_2\text{H}_2$ were probably relatively scarce, and S-MIF products of $\text{SO}_2$ photoexcitation in the 240–350 nm range were not preserved. It was shown by Ono et al. (2013) that the ratio $\Delta^{33}\text{S}/\delta^{34}\text{S}$ of product S0 in a $\text{SO}_2$-photolysis reaction increases when the bath gas (nitrogen) pressure decreases (Fig. 8a). The observed $\Delta^{33}\text{S}/\delta^{34}\text{S} = 0.23$ in the spherule beds could thus indicate that $\text{SO}_2$-photolysis took place at high altitude under very low pressure conditions.

Based on $\delta^{34}\text{S}$ and $\Delta^{33}\text{S}$ measurements alone it is difficult to establish if these spherule beds indeed uniquely recorded a high-altitude $\text{SO}_2$-photolysis effect in the 190–220 nm range, or also involved other atmospheric photochemical processes and/or later mass-dependent chemical effects. For instance, the S-MIF signal that is recorded in sediments also depends on the relative photochemical exit channels of $\text{SO}_3$ and S0 (Zerkle et al., 2012; Kurzweil et al., 2013). An impact-induced shift in atmospheric redox conditions could have favored preservation of S0 over $\text{SO}_3$ and thus have diminished the positive S-MIF signal. Future measurements of $^{36}\text{S}$ in these spherule bed pyrites could possibly confirm an isotopologue self-shielding effect. The general Archean reference array has a $\Delta^{36}\text{S}/\Delta^{33}\text{S}$ ratio between –0.9 and –1.5 (Kaufman et al., 2007; Zerkle et al., 2012), $\text{SO}_2$-photolysis in the 190–220 nm range a ratio of –4.6 ± 1.3 (Ono et al., 2013), and $\text{SO}_2$-photoexcitation in the 240–350 nm range a ratio between 1.12 and 2.25 (Whitehill et al., 2013). Both photochemical processes show a dependence on both p$\text{SO}_2$ and bathing gas pressure. A detailed quadruple sulfur isotope study on these spherule beds pyrites could therefore potentially confirm or reject $\text{SO}_2$-photolysis as a source reaction.

## 6. CONCLUSIONS

The sulfur isotope composition was studied of pyrite grains from three spherule beds in the 3.24 Ga Mapepe Formation, Fig Tree Group, Barberton, South Africa. Each of these deposits represents a settling dust cloud after a large meteorite impact event. Since this material reached high altitudes in the atmosphere, it could have adsorbed sulfur aerosols and therefore have recorded specific photolytic isotope anomalies. The sample of spherule bed S2 contained only very few spherules, and pyrite grains in this rock either display a sulfur isotope signature similar to that found in black cherts in the underlying Mendon Formation, or represent a sulfur source that was introduced at a later stage during hydrothermal alteration. In contrast, the sulfur isotope compositions in pyrites of spherule bed S3 – and to a lesser extent spherule bed S4 – show an anomalous trend $\Delta^{33}\text{S}/\delta^{34}\text{S} = 0.23$, that reflects mixing between a distinct S-MIF end-member ($\delta^{34}\text{S} = +10\text{‰}$, $\Delta^{33}\text{S} = +2\text{‰}$) and the juvenile mantle reservoir ($\delta^{34}\text{S} = 0\text{‰}$, $\Delta^{33}\text{S} = 0\text{‰}$). This trend significantly differs from all other sediments in the Mendon and Mapepe Formations, and also differs from the general Archean array (ARA). It appears that this trend more closely overlaps with the theoretical and experimental
trend for SO₂-photolysis in the 190–220 nm spectral range (Δ³³S/δ³⁴S = 0.06–014). It is possible that impact dust particles sampled the high, upper layer of the atmosphere where photochemical reactions were dominated by SO₂ isotopologue self-shielding effects.

## ACKNOWLEDGMENTS

Samples for this study were collected during the 2003 NASA Astrobiology Institute Field Workshop “Archean Surface Processes”, in South Africa. M.A. van Zuilen thanks the organizers D. Lowe (Stanford University) and G. Byerly (Louisiana State University) for field guidance during this workshop. This work was supported by ANR grant eLIFE2 to P. Philippot and Labex UnivEarths to P. Philippot and M.A. van Zuilen. K. Lindén is thanked for technical support during sample preparation. We thank S. Ono, two anonymous reviewers, and Associate Editor F. Moynier for careful and constructive reviews. The Nordsim ion microprobe facility is operated under an agreement with the joint Nordic research cooperation agency, NOS-N. This is Nordsim contribution 371 and IPGP N° 3557.

## APPENDIX A. SUPPLEMENTARY DATA

Supplementary data associated with this article can be found, in the online version, at http://dx.doi.org/10.1016/j.gca.2014.07.018.

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