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Palaeogeography, Palaeoclimatology, Palaeoecology

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The oldest known paleosol profiles on Earth: 3.46 Ga Panorama Formation,Western Australia

Gregory J. RetallackDepartment of Geological Sciences, University of Oregon, Eugene, Oregon 97403-1272, USA

A R T I C L E I N F O

Keywords:ArcheanPilbaraBariteAcid sulfate weatheringAtmospheric oxygen

A B S T R A C T

Subaerial volcanic flank and floodplain facies of the 3.46 Ga Panorama Formation have been recognized on thebasis of trough cross-bedded sandstones, lapilli tuffs, and abundant nodularized barite sand crystals, like thosenoted in other Archean non-marine facies. These barite-nodular layers are interpreted as alluvial paleosols forthe following reasons. They show different degree of bedding disruption scaled to nodule size beneath a sharpupper boundary, like desert soil profiles. They show multiple generations of cracking, clay skins, and up-profiledestruction of feldspar and rock fragments, compatible with weathering of labile constituents of the parentmaterial. Loss of alkali and alkaline earth elements and phosphorus up profile are also features of chemicalweathering. Geochemical mass balance calculations (tau analysis) shows that the profiles lost mass and labileelements, unlike unaltered sediments and tuffs. Rare earth element analysis also shows light rare earth retentionas in weathering, as opposed to sedimentation or hydrothermal alteration. Barite of nodules in the paleosols ismobilized to concentration under very acidic conditions (pH < 3) and is very stable under less acidic condi-tions. These Archean alluvial paleosols may have formed by acid sulfate weathering by sulfuric acid, rather thanthe currently more common hydrolysis by carbonic acid. This archaic system of widespread acid sulfateweathering has since been marginalized to a few playa lakes, deep water tables, and sulfur springs.

1. Introduction

Archean paleosols at major geological unconformities have beenstudied for some time (Buick et al., 1995; Rye and Holland, 1998), butArchean alluvial paleosols have eluded detection until recently(Retallack, 2014a; Nabhan et al., 2016; Retallack et al., 2016). Withoutclear indications like root traces, Precambrian alluvial paleosols can bedifficult to distinguish from ordinary sedimentary beds (Retallack,2011; Maslov and Grazhdankin, 2011; Kolesnikov et al., 2012, 2015).Field indications of Precambrian paleosols include vertical threads oflikely soil crust organisms (Retallack, 2011, 2013a), calcareous nodules(Retallack, 2011), and gypsum sand crystals (Retallack, 2012, 2013a),but confident identification of Precambrian alluvial paleosols requires avariety of petrographic and geochemical analyses (Retallack, 2011,2012, 2013a). A variety of lines of evidence for Precambrian alluvialpaleosols was reviewed by Retallack (2016): oxidation rinds, loesstexture, tsunamites, playa minerals, boron content, boron isotopes,carbon/sulfur ratios, total/reactive iron ratios, lapilli and crystal tuffs,ice deformation, birefringence fabrics, evaporitic sand crystals, geo-chemical tau analysis, and carbon‑oxygen isotopic covariance. Of these,Tarhan et al. (2015) found that the most convincing were evaporitesand crystals and nodules, of the kind commonly called desert roses

(Tarr, 1933) or Wüstenrosen (Dietrich, 1975). Such features are turningout to be abundant in Archean alluvial paleosols in both WesternAustralia (Retallack, 2014a; Retallack et al., 2016), and South Africa(Nabhan et al., 2016). This paper documents additional examples fromthe Panorama Formation of the Pilbara Craton of Western Australia, asevidence for the idea that acid-sulfate weathering (sulfuricization ofSoil Survey Staff, 2014) may have been more widespread on landduring the Archean than it has been subsequently. The paleosols de-scribed here are currently the oldest known paleosols on Earth (Ohmotoet al., 2014; Retallack, 2014a; Nabhan et al., 2016; Retallack et al.,2016), although paleosols as old as 3.7 Ga have been recognized onMars, and those also have sulfate nodules (Retallack, 2014c). Two ofthe Martian paleosols may appear controversial because described asplaya lake deposits with early diagenetic sulfates (Schieber et al.,2017), otherwise known as Solonchak soils (Food and AgricultureOrganization, 1974), or acid sulfate soils (Fitzpatrick et al., 2008).

1.1. Geological setting

The principal locality studied for this work is the PanoramaFormation in the area here informally called “Panorama Portal”, wherethe road to Dresser Mine leaves the floodplain of the North Shaw River,

http://dx.doi.org/10.1016/j.palaeo.2017.10.013Received 4 April 2017; Received in revised form 4 October 2017; Accepted 9 October 2017

E-mail address: [email protected].

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0031-0182/ © 2017 Elsevier B.V. All rights reserved.

Please cite this article as: Retallack, G.J., Palaeogeography, Palaeoclimatology, Palaeoecology (2017), http://dx.doi.org/10.1016/j.palaeo.2017.10.013

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http://dx.doi.org/10.1016/j.palaeo.2017.10.013


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and enters hills 30 km south of the Port Hedland-Marble Bar highway,and 10 km northwest of Panorama homestead (abandoned), in eastPilbara Shire, Western Australia (Figs. 1–2). Coordinates of the centralridge of this locality (Fig. 2A) are 21.01580°S 119.38178°E.

The geological age of the Panorama Formation is uncertain in thestudy area, but UePb ages from rhyolite zircons have been determinedin both Glen Herring Creek as 3430 ± 4 Ma (Nelson, 2002), and onPanorama Ridge as 3458 ± 1.9 Ma (Thorpe et al., 1992). This latterdate is equivalent to the UePb age of 3459 ± 18 Ma of the intrusiveNorth Pole Monzogranite (Thorpe et al., 1992), which may have been asubvolcanic intrusion to the active Panorama paleovolcano (vanKranendonk, 2000; Bolhar et al., 2005).

The study area exposes an unusually thick sequence of thePanorama Formation (Fig. 2A), regarded as flank facies of a paleo-volcanic stratocone (van Kranendonk, 2000; Bolhar et al., 2005). Thisvolcanic edifice remained a paleotopographic high because it is over-lain by Euro Basalt in the study area, whereas 28 m of Strelley PoolFormation unconformably overlies the Panorama Formation elsewhere(Fig. 1). The Panorama paleovolcano erupted onto pillow basalts of theMt. Ada Basalt (van Kranendonk, 2000; Bolhar et al., 2005). An areaidentified as an ancient vent of the paleovolcano by van Kranendonk(2000) and Bolhar et al. (2005) includes cross-cutting breccia pipes,with large blocks of banded iron formation of a caldera lake (Fig. 2B),and of carbonate-altered Mt. Ada Basalt. The ancient stratovolcanoflank facies is at least 1 km of agglomerates, rhyolite flows, and tuffs(Fig. 3–4). These tuffs include silicate accretionary lapilli (vanKranendonk, 2000), which disperse in water, and thus are evidence thatthe volcano flanks were subaerial (Retallack, 2014b). Central subaerialvolcanoes have also been inferred from a pattern of paleocurrents influvial facies of the Panorama Formation radiating from regional cen-ters (DiMarco and Lowe, 1989).

The original nature of the Panorama paleovolcano is best inferredfrom trace elements, because of widespread silicification and meta-morphic alteration (Brown et al., 2006). Some Panorama Formationtuffaceous siltstones have high Sc and Cr values, and low rare earth

elements and Th values, similar to local tholeiites and komatiites. Incontrast, most silicified rhyolitic tuffs of the Panorama Formation havelight rare earth enrichment and lack negative Eu anomalies, consistentwith melting of mantle eclogite, unlike other local Archaean rhyolites ofthe Duffer Formation (Cullers et al., 1993). Unsilicified calcareous tuffsof the Panorama Formation are similar in rare earth element compo-sition to rhyolitic tuffs, but have greater light rare earth enrichment,and high Ba and Sr (Fig. 5). These calcareous tuffs are not as enriched inlight rare earth elements as carbonatite tuffs (Woolley and Kempe,1989), including Archean (3.01 Ga) carbonatite from Greenland(Larsen and Rex, 1992; Bizzarro et al., 2002).

2. Materials and methods

The primary activity of this fieldwork was measuring sections of thePanorama Formation (Fig. 3), regarded as a subaerial volcanic edificeby Konhauser et al. (2002). Also documented were beds consideredlikely paleosols (Figs. 2C–I, 4), because of field characters of sharp topand gradational change below, dilational cracking patterns, and eva-poritic sand crystals. Potential paleosol beds were all named in thefield, within four distinct categories of Jurl, Ngumpu, Nganga andWanta, using descriptive words from Nyamal local aboriginal languages(Burgman, 2007). Each potential pedotype is a bed with a distinctivesequence of horizons, and the diagnostic features of each are given inTable 1. These are field categories to be tested as paleosol profiles bylaboratory data.

Oriented thin sections prepared vertical to bedding were pointcounted using a Swift automated stage and Hacker electronic countingbox (Supplementary material Tables S1–S2). Point counting was ad-justed for pervasive silicification and recrystallization to crystallitesminimally 5 μm across (Lepot et al., 2013), so that 5 μm grain size andsmaller were counted as clay, rather than the standard soil-clay cutoff at2 μm of soil science. Rock fragments were mainly sedimentary, withrare metamorphic and basaltic clasts. Evaporite minerals are nowquartz pseudomorphs, and identified from past research on polished

Fig. 1. Geological map of a paleovolcano in the 3.46 Ga Panorama Formation, North Pole Dome, Pilbara Craton, Western Australia (after van Kranendonk, 2000; Bolhar et al., 2005),showing location of measured sections of paleosols.

G.J. Retallack Palaeogeography, Palaeoclimatology, Palaeoecology xxx (xxxx) xxx–xxx

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slabs, large thin sections and electron microprobe analyses for com-parable local Archean cherts (Lambert et al., 1978; Brasier et al., 2002;Sugitani et al., 2003; Pinti et al., 2009). Original evaporite mineralogycould not be determined in thin section while point counting, but thepeudomorphs have a distinctive chalcedonic fabric. Opaque materialcounted includes both heavy minerals and organic matter. Two separate500-point counts were made for grain size and for mineral compositionof each sample, with errors (2σ) of± 2% for common (> 8%) com-ponents (Murphy, 1983). Chemical analyses of selected samples wereperformed by ALS Chemex of Vancouver, British Columbia, using XRFon glass discs for major elements, ICP on glass discs for trace elements,and FeO by Pratt titration (Supplementary material Tables S4–S6). Thestandard for the analyses was Canadian Certified Reference MaterialsProject standard SY-4, diorite gneiss from near Bancroft, Ontario. Errors(2σ) are from 89 replicate analyses of the standard in the same la-boratory. Titania values were below detection (< 0.01 wt%) for theJurl paleosol (R3775-R3779), but these were re-analyzed by ion mi-croprobe using an aggregation of multiple spectrometer intensities and

a refined blank correction in the CAMCOR instrument facility of theUniversity of Oregon (Donovan et al., 2011). Rare earth elements werenormalized to chondrite of Anders and Grevesse (1989). Bulk densitieswere determined from 20 to 40 g samples using paraffin at the Uni-versity of Oregon (Retallack, 1997a), with errors from 10 replicatedensity determinations of one of the Panorama Formation chert speci-mens (R4309). Thin sections and their source rocks are archived in theCondon Collection of the Museum of Natural and Cultural History of theUniversity of Oregon, in Eugene, Oregon.

3. Burial alteration

Soil formation is a form of early diagenesis, but reconstructing pa-leosols requires a thorough understanding of subsequent alterations,particularly metamorphism and silicification of individual beds.

Fig. 2. Field photographs of Panorama paleovolcano, banded iron formation, and paleosols and polished slabs of paleosol surface horizons covered with laminites, and of barite sandcrystals in Panorama Formation east Pilbara Craton, Western Australia. (A) ridge west of Panorama Portal (289–294 m in Fig. 3:21.02865°S 119.38087°E); (B) jasper-siderite banded ironwest of Panorama Portal (10-20 m in Fig. 3:21.03722°S 119.37359°E); (C) Nganga paleosol west of Panorama Portal (648 m in Fig. 3:21.03365°S 119.37915°E); (D) Jurl paleosol west ofPanorama Portal (293 m in Fig. 3:21.03229°S 119.97799°E); (E), Jurl paleosol southeast of Panorama Portal (311 m in Fig. 3:21.02554°S 119.38371°E), (F) Jurl paleosol north ofPanorama Portal (610 m in Fig. 3:21.01605°S 119.37734°E); (G) A horizon Wanta paleosol (R3774) west of Panorama Portal (63 m in Fig. 3: 21.03182°S 119.37310°E), (H) A horizon ofJurl paleosol (R3777) west of Panorama Portal (62 m in Fig. 3: 21.03182°S 119.37310°E); (I) By horizon of Jurl paleosol (R3755) west of Panorama Portal (294 m in Fig. 3: 21.03229°S119.97799°E). Sand crystals are at arrows.


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3.1. Metamorphism

Panorama Formation paleosols, metasediments, rhyolites and tuffssuffered chlorite-calcite-epidote lower greenschist facies burial meta-morphism (ca. 350 °C) like the overlying Strelley Pool Formation andEuro Basalt (Cullers et al., 1993; van Kranendonk and Pirajno, 2004;Allwood et al., 2006). Clayey rocks of the Panorama Formation werealso influenced by contact hydrothermal alteration to pyrophyllite (alsoca. 350 °C) by intrusion of the North Pole Monzogranite (Brown et al.,2006). The cherts studied here however were not altered by

hydrothermal fluids, because they lack large positive europiumanomalies due to high temperature plagioclase leaching and Cl-complexformation (Fig. 5), found in known hydrothermal chert of the Pilbararegion (Bolhar et al., 2005; Sugahara et al., 2010).

Greenschist facies regional metamorphism and pyrophyllitic hy-drothermal alteration are compatible with likely burial of the PanoramaFormation to depths of 14–17 km, and deep burial alterations of thepaleosols must be disentangled from ancient soil formation. AroundNorth Pole Dome, the overlying Strelley Pool Formation is 28 m thick,Euro Basalt 9400 m, Sulfur Springs Group 4941 m (van Kranendonk,2000), Gorge Creek Group 2610 m (van Kranendonk et al., 2006), for atotal of 16,980 m of overburden. Near Strelley Pool however, thethickness of Euro Basalt is only 1000 m (Buick et al., 1995), so over-burden beneath the Lalla Rookh and Fortescue Groups would have been

Fig. 3. Geological section of Panorama Formation in the paleovolcano flank facies, eastPilbara Craton, Western Australia (with paleoenvironmental interpretations from vanKranendonk, 2000).

Fig. 4. Detailed sections of Panorama Formation in the paleovolcano flank facies, eastPilbara Craton, Western Australia. Blue boxes to right of column show thickness (of A-Bk)and degree of development of paleosols (after Retallack, 1997a), white grains of barite areshown as volume percent, and hue is from a Munsell chart. (For interpretation of thereferences to colour in this figure legend, the reader is referred to the web version of thisarticle.)


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8580 m. The Lalla Rookh Formation of the overlying De Gray Groupwas confined to fault basins (Krapež, 1984), so its 3000 m did notoverlie the studied localities. Fortescue Group basalts and sedimentsnorthwest of Marble Bar are only 5240 m thick (Thorne and Trendall,2001; van Kranendonk, 2010), so toward the east total burial depthmay have been only 13,826 m.

3.2. Late diagenesis

Illitization is a common deep burial enrichment of potassium(Nesbitt and Young, 1989; Fedo et al., 1995; Novoselov and de SouzaFilho, 2015). Nevertheless, geochemical data on the putative paleosolschosen for detailed analysis shows weathering rather than illitizationtrends (Fig. 6A). Illitization has affected red laminites (part of Wantaprofile) and weathered upper portions of calcareous tuffs (Ngangaprofile) of the Panorama Formation with marked potash enrichments(Fig. 6B). In contrast the cherty paleosols (Jurl profiles) show minordepletion, not enrichment in potash, probably because of protection byearly silicification, as documented for the overlying Strelley Pool For-mation (Buick and Barnes, 1984; Hickman, 2008; Allwood et al., 2010).

Silicification from dissolved silica released by illitization of claysand pressure solution of sand grains is a common late diagenetic al-teration of sedimentary rocks (Siever, 1962), but unlikely for extensivesilicification of the Panorama Formation. Not only is there little evi-dence of a trend toward illite (Fig. 6A) or a potash enrichment in theserocks (Fig. 6B), but they were relatively poor in clay (Fig. 7). Parentmaterial in these plots is below the feldspar line (50% Al2O3) in Fig. 6Aand at origin in Fig. 6B. Cherts of the Panorama Formation do showpervasive crystal domains up to 5 μm in size, larger than usual forgeologically younger cherts with “pinpoint birefringence”, and perhapsdue to diagenetic (neomorphic) recystallization, as for cherts of theoverlying Strelley Pool Formation (Lepot et al., 2013).

3.3. Early diagenetic silicification

Most tuffs in the Panorama Formation have been extensively sili-cified like stromatolitic dolostones, sandstones and conglomerate of theoverlying Strelley Pool Formation (Allwood et al., 2007). Clues to sili-cification of Panorama Formation paleosols come from organic micro-fossils in comparable cherts of the Strelley Pool Formation (Sugitaniet al., 2010; Wacey et al., 2010, 2011), which Hickman (2008) con-siders equivalent to the microfossiliferous “Kitty's Gap Chert” (Westallet al., 2006). Microfossils are evidence of silicification before the cellscould decay: which takes about 160 days for cyanobacteria under an-oxic conditions (Harvey et al., 1995). Microfossils of the Strelley PoolFormation are mainly robust large spindles, which are widely acceptedas genuine for reasons of carbon isotopic composition, Raman spec-troscopy, and three-dimensional imaging (Sugitani et al., 2010; Waceyet al., 2010, 2011; Schopf et al., 2007; Brasier et al., 2015). These ob-servations contradict the idea that cherts of the Strelley Pool and Pa-norama Formations are silcretes formed in Archean outcrops duringCenozoic weathering (van Kranendonk, 2000). Australian Cenozoicsilcretes are chemically distinct, with much more titania (0.7–48.0 wt%: Wopfner, 1978; Taylor and Eggleton, 2001) than cherts of the Pa-norama Formation (0.006–0.14 wt%: Supplementary material TableS3).

The following paragraphs consider three plausible mechanisms ofearly diagenetic silicification of paleosols of the Panorama Formation:(1) hydrothermal, (2) biogenic, and (3) evaporitic. Silicification in hotsprings would be compatible with chert dikes and pyrophyllitic al-teration associated with intrusion of the North Pole Monzogranite, andnearby vent facies of the paleovolcano in the Panorama Formation(Kato and Nakamura, 2003; van Kranendonk, 2006). Furthermore, redcherty banded iron formation of the basal Panorama Formation shows astrong positive europium anomaly and heavy rare earth enrichmentcharacteristic of hydrothermal alternation, probably within waters of a

Fig. 5. Rare earth element analyses of putative paleosols of the Panorama Formation of Western Australia normalized to chondritic values of Anders and Grevesse (1989): (A) Wanta andNganga pedotypes; (B) Jurl and Ngumpu pedotypes.

Table 1Summary of Panorama Formation paleosol definition and classification.

Pedotype Nyamal meaning(Burgman, 2007)

Diagnosis US taxonomy (SoilSurvey Staff, 2014)

FAO map unit (Food andAgriculture Organization, 1974)

Australian classification(Isbell, 1996)

Jurl Salt Black massive cherty surface (A horizon) overgypsum crystal pseudomorphs in gray chert (Byhorizon)

Haplogypsid Orthic Solonchak Gray Sodosol

Nganga Sand Brown sandstone surface (A) horizon over whitecalcareous sandstone (C) horizon

Psamment Calcaric Regosol Arenic Rudosol

Ngumpu Narrow Black massive cherty surface (A horizon) overlaminated gray chert (C horizon).

Fluvent Eutric Fluvisol Stratic Rudosol

Wanta Crazy Red cracked and laminated surface (A horizon)over barite mottle pseudomorphs (By horizon)

Halaquept Gleyic Solonchak Gray-orthic Tenosol


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volcanic summit caldera (Bolhar et al., 2005). However, cherts of thepaleovolcano flank facies have a different and flatter rare earth elementpattern (Fig. 6), comparable with patterns in the Strelley Pool Forma-tion considered evidence for low temperature silicification (vanKranendonk, 2000; van den Boorn et al., 2010; Allwood et al., 2010).

An alternative source of silica in low temperature sedimentary en-vironments of many Precambrian microfossiliferous cherts is waters ofhigher silica concentrations than today, because predating evolution ofsiliceous microbes such as diatoms and radiolarians (Eriksson andWarren, 1983; Maliva et al., 2005). This view is undermined by de-monstration that silica spicules are secreted by sulfate-reducing bac-teria (Birnbaum and Wireman, 1985; Birnbaum et al., 1989), whichhave independently been inferred from microfossils and sulfur isotopicstudies of the overlying 3.42 Ga Strelley Pool Formation (Wacey et al.,2010, 2011; Bontognali et al., 2012), and underlying 3.48 Ga DresserFormation (Shen et al., 2001; Ueno et al., 2008). Recrystallization ofbiogenic spicules may have contributed to silicification of the PanoramaFormation.

The most likely mechanism of early silicification is formation asplaya cherts, comparable with magadiite (Hay, 1968; Eugster, 1969).Minerals such as gypsum (CaSO4.2H2O) are often precipitated at highpH (> 9), which mobilizes silica, and such alkaline solutions dissolvebiogenic opaline or volcanic silica in preference to quartz, creating si-lica gels in playas and intertidal flats (Eugster, 1969; Smith, 1979;Larsen, 2008). This model of caustic rather than neutral solutions,perhaps aided by sulfate-reducing bacteria is appealing for silica per-mineralization of many fossiliferous Precambrian intertidal to supra-tidal cherts (Maliva et al., 2005).

4. Testing likely paleosols

A variety of beds were identified in the field as potential paleosolslargely on the basis of destruction of primary sedimentary features, andthese beds were sampled for a battery of geochemical and petrographictests to determine whether they were paleosols, or unaltered tuffs orsedimentary beds. Most of these formed the top 30–40 cm of troughcross bedded sandstone (angular discordance to bed top in Fig. 2D–E).

4.1. Field observations

Many cherty beds of the Panorama Formation have common earlydiagenetic pseudomorphs of evaporite minerals which range in formfrom elongate to short crystals (Fig. 2D–F at arrows), variously cor-roded and or extended as rounded nodules (Fig. 2I at arrows). Thesevary in abundance and size from bed to bed, from small masses withlittle disrupting effect (Fig. 2D) on primary bedding to abundant no-dules up to 2 cm in diameter which totally obscure bedding (Fig. 2F).Furthermore, these nodular and crystalline masses are neither at thevery top nor base of their respective beds: they are sparse in the upper10 cm or so of a bed then increase in abundance below that and taperoff in abundance some distance above the plane or cross-bedded base ofthe bed (Fig. 2E–F). In all these ways, paleosols of the Panorama For-mation are superficially similar to salt-rich (By) horizons of soils(Retallack and Huang, 2010; Amundson et al., 2012), and paleosols ofthe 3.0 Ga Farrel Quartzite of Western Australia (Retallack et al., 2016),and 3.2 Ga Moodies Group of South Africa (Nabhan et al., 2016).

The bedding disruption fades downward from truncated uppersurfaces of the paleosols, including those without evaporites(Fig. 2C–F), and this may represent the upper surface of the originalsoil. These upper surfaces have organic carbon crusts with sharp uppersurface and wispy downward extensions (Fig. 2H), perhaps from mi-crobial soil crusts (Retallack, 2014a). Some upper surfaces to beds withnumerous evaporate pseudomorphs have a few laminations of hematite,including contorted layers (Fig. 2G), comparable with the cumulicsurface horizons of lowland soils (Retallack, 1997a). Most soils havesharp upper surfaces, but lowland soils have increments of additionalsediment before burial by a terminating tuff or river sand.

4.2. Petrographic observations

Bedding disruption is especially obvious in thin sections of theupper surface of putative paleosols which show lamination above asharp contact to near-vertical, chaotic fabric in thin sections cut verticalto regional bedding (Fig. 7A–E). This near-vertical fabric reveals mul-tiple generations of bedding disruption by cracking which havehomogenized original bedding, but there also are cases where lamina-tion is only mildly disrupted by vertical structures (Fig. 7F).

Other indications of soil formation in thin section are clay skins

Fig. 6. Illitization of paleosols in the Panorama Formation in the paleovolcano flank facies, east Pilbara Craton, Western Australia, showing (A) weathering trends toward alumina pole ofNesbitt and Young (1989) on a truncated triangular diagram, and (B) potash enrichment as tau values following Novoselov and de Souza Filho (2015).


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rimming weathered grains, which are embayed and dilated into frag-ments (Figs. 7G–H). Such dilation indicates alteration within a soilunder little confining pressure, unlike burial diagenetic alterationswhich result in overall reduction in volume. Furthermore, the clay skinshave high birefringence (Fig. 7G), a feature of soils and paleosols calledsepic plasmic fabric (Retallack, 1997a).

Evaporite pseudomorphs of the Panorama Formation in thin sectionshow abundant included grains characteristic of desert roses (Tarr,1933; Dietrich, 1975), as revealed by common sedimentary inclusions,and corroded and dusty rims (Fig. 8C–E). Some of these pseudomorphsare radially arranged like desert roses (Fig. 8C), and others are heavilynodularized (Fig. 8B). Graded beds overlying the paleosols have re-deposited crystal pseudomorphs (Fig. 8A), demonstrating that they hadformed and were eroded during breaks in deposition. Comparable ob-servations of sand crystals and redeposition reveal paleosols in the3.2 Ga Moodies Group of South Africa (Nabhan et al., 2016).

Sand crystals are distinct from crystals formed in open water orsaturated sediments, which are clean of inclusions because of displacive

growth or simple precipitation in marine evaporites (Warren, 2006;Ziegenbalg et al., 2010) and playa lakes (Hay, 1968; Renaut andTiercelin, 1994). Also different are void-filling, inclusion-free, barites ofhydrothermal vugs and veins, such as those of the 3.48 Ga DresserFormation (Shen et al., 2001; Ueno et al., 2008). Different again arespring tufas with cavity-filling pendant and erect barite fans, and la-minated to stromatolite-like isopachous barite coatings, also inclusion-free (Donovan et al., 1988; Bonny and Jones, 2008a, 2008b).

The original mineral of Archean chert pseudomorphs of evaporiteminerals in the Panorama Formation may have been barite (BaSO4),despite low Ba (8.2–3670 ppm) remaining after thorough early diage-netic silicification (Supplementary materials Table 3). Barite inclusionswithin other Archean pseudomorphs are near stoichiometric, with 59 to65 wt% BaO (Sugitani et al., 2003; Nabhan et al., 2016). The crystalswhen best preserved have the squat form of barite (Fig. 8E), unlikehexagonal needles of nahcolite (NaHCO3), fishtail twins of selenite-gypsum (CaSO4.2H2O), or cubes of anhydrite (CaSO4). Barite is themost common of pseudomorphed evaporite minerals reported from

Fig. 7. Petrographic thin sections of paleosols in the Panorama Formation of Western Australia, under crossed nicols (A–C, F–H) and plane light (D–E): showing filamentous verticaldisruptions of surface horizons of type Wanta silt loam (A), of type Jurl clay (B,D), of type Nganga loam (C) and of type Ngumpu clay loam (E–F); G, clay skins in surface horizon of typeJurl clay; H, rock fragments weathered to clay in A horizon of type Nganga clay loam. All thin sections were cut vertical to regional bedding and are curated in the Condon CollectionUniversity of Oregon: A, R3774; B, D, G, R3756; C, R3765; E–F, R3761; H, R3766.


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other Archean cherts: 3.0 Ga Farrel Quartzite (Sugitani et al., 2003),3.4 Ga Kromberg Formation (Lowe and Worrell, 1999), 3.4 Ga StrelleyPool Formation (Lowe, 1983; Allwood et al., 2007), 3.5 Ga Apex Chert(Pinti et al., 2009) and 3.5 Ga Dresser Formation (Lambert et al., 1978;Buick and Dunlop, 1990). In the Phanerozoic, gypsum-bearing paleo-sols are common (Retallack and Huang, 2010), but barite-bearing pa-leosols are rare and restricted to a few stratigraphic horizons (Tarr,1933; Beattie and Haldane, 1958; Olson, 1967; Skinner and Taylor,1967; Keyser, 1968; Laznicka, 1976; Bown and Kraus, 1988; Khalaf andEl Sayed, 1989; de Brodtkorb, 1989; Speczik and El Sayed, 1991; Garcesand Aguilar, 1994; Kreuser, 1995; Dietrich, 1975; Schrank andMahmoud, 2000; Retallack and Kirby, 2007; Retallack, 2009; Neuhaus,2010; Buck et al., 2010; Brock-Hon et al., 2012; McCarthy and Plint,2013; Jennings et al., 2015).

4.3. Petrographic trends

Point counting within individual beds also reveals changing pro-portions of minerals compatible with increased intensity of weatheringfrom bottom to top: increased proportion of clay, quartz, and opaqueminerals toward the surface, and declining abundance of feldspar androck fragments (Fig. 9B). Degree of weathering revealed by this metricvaries from minimal (Ngumpu, Nganga) to moderate (Wanta, Jurl) inproportion with bedding disruption by crystal pseudomorphs. Someminerals such as quartz and opaque oxides are resistant to weathering,but feldspar and rock fragments weather to clay in soils (Retallack,1997a). Furthermore, the extent of these weathering trends vary withthe degree to which original sedimentary bedding has been homo-genized by growth of nodules and cracking (Fig. 9A), revealing deeperchemical weathering associated with physical homogenization.

Barite pseudomorphs were found only in the cherty beds, not in thehighly calcareous bed considered an alkaline tuff from its unusualabundance of light rare earth elements (Fig. 5). That thick calcareousbed (Nganga profile) shows muted changes in proportions of

weatherable minerals until the very top, and so is grouped with theweakly developed profiles (Ngumpu).

4.4. Major element geochemical trends

Individual beds sampled as likely paleosols also show plausibletrends in major element geochemical composition, expressed as mole-cular (molar) weathering ratios (Fig. 9C). Parent material in these thinalluvial paleosols is the material directly below (labelled C for inter-preted sediment, or R for interpreted rock). Furthermore these thinprofiles are unusually deeply weathered with high base loss (aluminaover bases) for soils with evaporites (Amundson et al., 2012). Aluminaover bases shows little weathering for the one profile (Nganga), butexpected down-profile decline for others, especially marked in twosuperposed profiles (Wanta on Jurl profile). Another pronounced fea-ture is the high ratio of ferrous over ferric iron, like that of a water-logged (gleyed) soils in the modern world (Vepraskas and Sprecher,1997). Unlike the little-weathered calcareous tuff (Nganga profile), thecherty profiles have local enrichments of ferric iron, very low alkalies,alkaline earths and alumina, and modest salinization. These are featuresof acidic modern soil-forming processes, such as podzolization(Retallack, 1997a). The combination of podzolization-gleization withevaporite minerals is a distinctive feature of these and other Archeanpaleosols (Nabhan et al., 2016; Retallack et al., 2016), and somemodern soils (Benison and Bowen, 2015; Jennings et al., 2015).

4.5. Tau analysis

Geochemical mass balance or “tau analysis” is a method for calcu-lating both the mobilization of particular elements (mass transfer) andthe change in volume (strain) of soil formation (Brimhall et al., 1992).Mass transfer (τw,j in mass fraction) is derived from chemical con-centration of the element in soils (Cj,w in weight %) compared withparent material (Cp,w in weight %), correcting for bulk density of the

Fig. 8. Petrographic thin sections of evaporate pseudomorphs in the Panorama Formation of Western Australia, under plane light (A–C, E) and crossed nicols (D), showing redepositedbarite at base of graded bed (A), nodularized barite (B), crystal rosette of barite (C), corroded (left) and sand crystal (right) of barite (D), and sand crystal of barite (E). Paleosols are typeJurl clay (A–B, E), Jurl clay loam (C) and type Wanta silt loam (D). Sand crystals are within dashed outlines. Specimens in Condon Collection, University of Oregon are R3756A (A),R3756B (B), R4309 (C), R3775 (D) and R3757 (E).


8

soil (ρw in g cm−3) and parent material (ρp in g cm−3). Strain (εi,w as amass fraction) is estimated from an immobile element in soil (such as Tiused here). The relevant Eqs. (1) and (2) (below) are the basis forcalculating divergence from parent material composition.

= ⎡

⎣⎢

⋅⋅

⎤

⎦⎥ + −τ

ρ Cρ C

ε[ 1] 1j ww j w

p j pi w,

,

,,

(1)

= ⎡⎣⎢

⋅⋅

⎤⎦⎥ −ε

ρ Cρ C

1i wp j p

w j w,

,

, (2)

Soils and paleosols have a distinctive location in this style of plot-ting geochemical data: in lower left quadrant of volume collapse andelemental loss when plotted as mass transfer with strain (Fig. 10A), andwith depletion up-profile when plotted as mass transfer with depth(Fig. 10B). This is where the Panorama beds plot with a few exceptions,primarily of the Nganga profile, which is a little weathered tuff(Fig. 9C), and so plots in the dilate-and-gain quadrant of mass transferwith strain like a sedimentary or tuffaceous bed (Fig. 10A). TheNgumpu profile also shows little chemical differentiation, and may alsohave been a very weakly developed soil. Tight clustering of silica,alumina, alkalies and alkaline earth elements in this kind of analysis

Fig. 9. Mineral composition and molecular weathering ratios of paleosols in the Panorama Formation at Panorama Portal, Western Australia.


9

(Fig. 10A) over distances of only a few tens of centimeters (Fig. 10B) isevidence of chemical weathering.

4.6. Trace element trends

Chemical differences over short depths within the profiles are alsoseen in the abundance of rare earth elements, which can vary by anorder of magnitude within a single profile of the Panorama Formation(Fig. 5). The light rare earth enrichment of the Nganga bed is attributedto alkaline tuff parent material different from that of other beds, butdepleted in rare earth elements in the subsurface (C) horizons (Fig. 5A).Profiles with evaporite pseudomorphs (Wanta and Jurl) also have de-pletion in lower levels (Fig. 5A). A profile lacking evaporites and withclear bedding (Ngumpu profile) shows little change in rare earth ele-ment composition at different levels, and a flatter pattern than the otherprofiles (Fig. 5B). These depletions within profiles over only a fewcentimeters and heavy-rare-earth depletion (Jurl and Wanta versusNgumpu and Nganga) are evidence of acidic humid weathering at thelevel of the evaporate pseudomorphs (Nesbitt, 1979). These weatheringtrends can be contrasted with sedimentary redistribution and marinediagenesis of sediments. Sedimentary processes separate rare earthelements into clay rather than sand (Morey and Setterholm, 1997), butthe Panorama beds are poor in clay and do not show large changes ingrain size (Fig. 9). Marine diagenesis of sediments (halmyrolysis) doesnot change rare earth concentrations significantly (Clauer et al., 1990;Setti et al., 2004), unlike the relationships seen in the Panorama For-mation (Fig. 5).

5. Interpretation of paleosols

The named beds of the Panorama Formation appear to be paleosols,which are of interest because they yield information not only about thediversity of landforms of the past, but about a variety of paleoenvir-onmental conditions ranging from time for formation to climate andatmospheric composition (Retallack, 1997a). Each of the profilesnamed on the basis of their field appearance (Table 1) represents adifferent combination of formative influences, which are interpreted inthe following paragraphs (Table 2).

5.1. Classification of panorama formation paleosols

Field observations, together with geochemical and petrographicinformation (Fig. 9) on the paleosols allow interpretation within clas-sifications of modern soils of the various field pedotypes recognized inthe Panorama Formation (Table 1). Point count data (Fig. 9) also givethe option of naming particular paleosols from soil textural categoriesof their surface horizons, such as Jurl silt loam versus Jurl clay loam.Persistence of bedding, thin profiles, and muted chemical variationmark the Nganga and Ngumpu profiles as Entisols in the U.S Taxonomy(Soil Survey Staff, 2014). Surface lamination like banded iron forma-tion, an aquatic facies of the Panorama Formation (Bolhar et al., 2005),may be evidence that Wanta paleosols were intermittently waterlogged,as well as developing small barite nodules and crystals below whenbetter drained, so gleyed Inceptisols. Although banded iron formationshave been considered deep water, this does not seem to be the case for

Fig. 10. Tau analysis of paleosols in the Panorama Formation of Western Australia, showing (A) mass transfer with strain and (B) mass transfer with depth, and calculated by method ofBrimhall et al. (1992).


10

banded iron formation in the Panorama Formation (Bolhar et al., 2005),and likely photoferrotrophic origin is not possible in deep water (Croweet al., 2008). There are no baritic soils in the US taxonomy, althoughsome Ultisols have abundant barite at depth, and baritic paleosols arelocally abundant (Table 3). Nevertheless, Gypsids (with gypsum ratherthan barite in Soil Survey Staff, 2014) are a match for Jurl paleosols inprofile form and evaporite density. Comparable considerations can beused to classify the paleosols in classifications of Australia (Isbell, 1996)and of the Food and Agriculture Organization (1974). Classification ofthe Jurl paleosols in the FAO system as a Solonchak (among others inTable 1) opens the possibility of finding modern soil map units com-parable with the Panorama paleosol assemblage, which can be de-scribed by the FAO code as Zo with inclusions of Zg, Jd, and Rc (OrthicSolonchak, with Gleyic Solonchak, Dystric Fluvisol and Calcaric Re-gosol). The closest match to this in all mapped soil units around theworld is map unit Zo36-2a, with associated Zg, and inclusions of Rc, Rd,Xk and Yh (Calcaric Regosol, Dystric Regosol, Haplic Xerosol, andHaplic Yermosol) on alluvial sediments widespread in Western Aus-tralia, but most extensive near Lakes Disappointment and Carnegie(Food and Agriculture, Organization, 1978). Such soils are found inregions of arid grasslands and shrublands, but extend into the formerly

wooded “wheat belt”, for a range of mean annual precipitation of260–340 mm (Benison and Bowen, 2013, 2015). This region has meanannual temperature of 18–20 °C (Bureau of Meteorology, 2016). Theseparts of Western Australia have a mix of desert soils developed ondeeply weathered lateritic residuum from thick Ultisol and Oxisol pa-leosols of the Miocene (McKenzie et al., 2004). These Solonchak soilsare also called “acid saline lakes” (Benison et al., 2007; Benison andBowen, 2013), even though dry more often than flooded. A distinctivefeature of these lakes despite abundant salts is very acidic waters(pH 1.4–3.9), due to oxidation of sulfides unbuffered by local carbonatein soils or bedrock (Benison and Bowen, 2015). Halite and gypsum arethe principal evaporite minerals of the acidic lakes (Benison andBowen, 2013), but barite is also recorded in Western Australian acidsulfate soils with pH < 3 (Fitzpatrick et al., 2008). Another analog iscold sulfur springs where barite is precipitated at pH 7.3–8.4 fromoxidation and neutralization of anoxic and acidic brines from the sub-surface (Donovan et al., 1988; Bonny and Jones, 2008a, 2008b).

5.2. Parent material

Two distinct parent materials for paleosols of the Panorama

Table 2Summary of Panorama Formation paleosol interpretations.

Pedotype Climate Organisms Topography Parent material Soil duration

Jurl Humid (1154–1343 ± 182 mm MAP) temperate (8.7–10.3 ± 0.5 °C MAT) Microbial earth Floodplain Quartz-lithic silt 2408–70,151 yearsNganga Not diagnostic for climate Uncertain Coastal dunes Calcareous sand

and gravel10–1000 years

Ngumpu Not diagnostic for climate Microbial earth Streamside bar Quartz-lithic sand 10–1000 yearsWanta Not diagnostic for climate Microbial earth dominated

by iron-oxidizing bacteriaFloodplainswale

Quartz-lithic sand 1571–69,505 years

Note: MAP is mean annual precipitation and MAT is mean annual temperature.

Table 3Occurrences of pedogenic barite.

Ma Geological age Formation Location Reference

0.015 ± 0.0011 Pleistocene Lufkin loam soil College Station, Texas Jennings et al., 20150.015 ± 0.0011 Pleistocene Falba fine sandy loam soil Huntsville, Texas Carson et al., 19820.014 ± 0.0010 Pleistocene Bluford silt loam soil Bluford, Illinois Darmody et al., 19890.100 ± 0.05 Pleistocene High terrace Ultisol Bellavista, Peru Stoops and Zavaleta, 19782 ± 2 Pleistocene High terrace calcrete Mormon Mesa, Nevada Brock-Hon et al., 20125 ± 2 Pliocene Brucedale Parna Wagga Wagga, New South Wales Beattie and Haldane, 195810 ± 6 Mio-Pliocene Kuwait Group Kuwait Khalaf and El Sayed, 198910 ± 1 Miocene Edelény variegated clay Rudabánya, Hungary de Brodtkorb, 198916 ± 1 Barstovian Cucaracha Formation Panama Retallack and Kirby, 200716 ± 1 Barstovian Fort Randall Formation Bijou Hills, South Dakota Skinner and Taylor, 196726 ± 0.2 Chattian Rockenberg Formation Rockenberg, Germany Neuhaus, 201027 ± 2 Arikareean Monroe Creek Formation Ovid, Colorado de Brodtkorb, 198935 ± 5 Chadronian Carroza Formation La Papa, Mexico Buck et al., 201036 ± 1 Priabonian Qasr el Sagha Formation Fayum Depression, Egypt Bown and Kraus, 198850 ± 5 Ypresian El Naab Formation Bahariya Oasis, Egypt Speczik and El Sayed, 199190 ± 10 Cenomanian Dunvegan Formation Clayhurst, British Columbia McCarthy and Plint, 201390 ± 10 Cenomanian Sabaya Formation Kharga Oasis, Egypt Schrank and Mahmoud, 200090 ± 10 Cenomanian Bahariya Formation Bahariya Oasis, Egypt Speczik and El Sayed, 1991135 ± 5 Berriasian Morrison Formation Como Bluff, Wyoming Retallack, 2009151 ± 2 Tithonian Morrison Formation Thermopolis, Wyoming Jennings et al., 2015200 ± 5 Hettangian Elliott Formation Bethleham, South Africa Keyser, 1968235 ± 2 Carnian Cacheuta Formation Cacheuta, Argentina de Brodtkorb, 1989250 ± 3 Griesbachian Upper Beaufort Group Thaba Ncha, South Africa Keyser, 1968260 ± 2 Guadalupian Unidad Roja Superior Jaca, Spain Garces and Aguilar, 1994270 ± 2 Kungurian Mbuyura Formation Ruhuhu, Tanzania Kreuser, 1995270 ± 2 Kungurian Garber Sandstone Norman, Oklahoma Tarr, 1933, Olson, 1967270 ± 2 Kungurian Rotliegend Rothenberg, Gembeck, Germany Dietrich, 1975359 ± 10 Late Devonian Horn River Formation Twitya River, NWT, Canada Laznicka, 19762765 ± 10 Archean Mt Roe Basalt Whim Creek, Western Australia Teitler et al., 20153016 ± 13 Archean Farrel Quartzite Mt Grant Western Australia Retallack, 2014a3220 ± 10 Archean Moodies Group Barberton, South Africa Nabhan et al., 20163458 ± 1.9 Archean Panorama Formation North Pole Dome, Western Australia Herein


11

Formation are revealed by rare earth element analyses (Fig. 5) andpoint counting (Fig. 9), and these are now calcareous tuff and chertytuff. The calcareous tuff is an unsilicified lithic tuff with only thin andweakly developed paleosols (Nganga pedotype). The cherty tuffs werealso largely a lithic tuff, although acicular shapes are suggestive of avitric component, which may have aided silicification. Interbeddedflows and chemical composition of the Panorama Formation has beentaken as evidence that these silicified tuffs were originally rhyodaciticin composition (Cullers et al., 1993). These were parent materials tomost of the paleosols recognized: Ngumpu, Jurl and Wanta pedotypes.

5.3. Paleotopographic position

Evidence that the parent tuffs formed a well-drained volcanic aproncomes from facies analysis using observations such as accretionary la-pilli (van Kranendonk, 2000), and measurements such as paleocurrentsand stratigraphic thickness variations (Cullers et al., 1993). Calcareoussands capped by Nganga paleosols have cross bedding 1.5 m thick(Fig. 2C), which may have been coastal eolian dunes with as muchrelief. Then there is the evidence of abundant barite sand crystals inJurl and Wanta paleosols, which requires desiccation and precipitationin a porous medium (Tarr, 1933; Dietrich, 1975). The distribution ofbarite, scarce at the surface and peaking in abundance a short distancebelow (Figs. 2D–F, 9), is comparable with the distribution of gypsum inexcessively drained modern desert soils (Retallack and Huang, 2010;Amundson et al., 2012). Barite in modern soils forms near the watertable, because it requires neutralization, evaporation and oxidationfrom highly acidic and chemically reducing groundwater (Jennings andDriese, 2014; Jennings et al., 2015). Barite nodules in Jurl and Wantapaleosols do not extend much below 30 cm from the top of the profile,and this may have been the level of water table in dry playas, com-parable with Western Australian saline lakes (Benison and Bowen,2015).

Another indication of high water table may be laminae of hematitein the surface horizons of Wanta paleosols (Fig. 2G). These are most likebanded iron formations elsewhere in the Panorama Formation (Bolharet al., 2005), thought to have formed in an aqueous medium by mi-crobial photoferrolysis (Crowe et al., 2008). Although widely con-sidered marine, banded iron formations also have been recorded fromfreshwater crater lakes (Bolhar et al., 2005), and fluviodeltaic facies(Pufahl et al., 2013). This situation is the reverse of the situation inmodern soils, in which waterlogged soils are gray-green and welldrained soils red-brown (Vepraskas and Sprecher, 1997). Such upside-down redox relationships have been predicted as a consequence ofanoxic Archean atmosphere (Walker, 1987; Domagal-Goldman, 2014).

5.4. Time for formation

Paleosols within the Panorama Formation show different degrees ofdestruction of bedding and ripple marks as if they were moderately(Jurl, Wanta) to very weakly developed (Ngumpu, Nganga) in the fieldscale of Retallack (1997a). The application of this scale to these pa-leosols is problematic because their nodules are barite, rather thangypsum (Retallack and Huang, 2010), or low‑magnesium calcite ofmodern soils (Retallack, 2005).There are no modern chronosequencesfor barite in soils (Jennings and Driese, 2014; Jennings et al., 2015).Barite is rare under acidic (< pH 3) conditions of saline lakes withgypsum and halite (Fitzpatrick et al., 2008), but barite also precipitatesfrom waters of pH 7.3–8.4 in cold sulfur springs (Donovan et al., 1988;Bonny and Jones, 2008a, 2008b).

Despite these caveats, nodules in soils take millennia to coalesce(Retallack, 1997a). Two transfer function for estimating paleosol ages(A in years) come from diameter of low magnesium calcite nodules (S incm) in radiocarbon-dated soils of New Mexico (Retallack, 2005), andabundance of gypsum crystals (G as % surface area) of the Negev andAtacama Deserts (Retallack, 2013a), according to the following

formulae:

=A S3920· 0.34 (3)

with r2 = 0.57, s.e. = ± 1800 yrs

= +A G3987· 5774 4

with r2 = 0.95, s.e. = ± 15,000 yrsIf the sand crystals and nodules of the Panorama Formation were

calcite, their size in 13 Jurl paleosols would be evidence of4208 ± 1800 yr, and in 7 Wanta paleosols, of 3371 ± 1800 yr(Supplementary information Table S6). If on the other hand, they weregypsum crystals and had densities measured in the field, the duration of13 Jurl paleosols average 55,151 ± 15,000 years, and 7 Wanta pa-leosols average 44,505 ± 15,000 years. These estimates vary by anorder of magnitude, but provide an envelope of possibilities. A 130Xeisotopic estimate of exposure time before burial of< 40,000 years wasobtained for bedded (non-pedogenic) barite of the 3.48 Ga DresserFormation (Pujol et al., 2009).

5.5. Paleoclimate

Chemical index of alteration (I) is a general ratio of alumina overbases, calculated as follows

=+ + +

I mAl OmAl O mCaO mNa O mK O

100( )[( ) ( ) ( ) ( )]

2 3

2 3 2 2 (5)

This ratio is< 65 in glacial-frigid climates and> 80 in tropicalclimates, but between these limits in temperate climates (Nesbitt andYoung, 1989). Burial illitization would reduce this value toward coolerclimates, but was not a discernible problem for most of the early-sili-cified paleosols (Fig. 6). The 19 available analyses of the cherty pa-leosols (Ngumpu, Jurl, Wanta), and of their parent materials as a guideto sedimentary source terranes, have a chemical index of alteration of67 ± 12, which suggests cool temperate to periglacial climate.

More specific pedogenic paleothermometers are based on soils ofmodern woody vegetation (Sheldon et al., 2002; Gallagher andSheldon, 2013), not applicable to likely microbial earths of Archeanpaleosols (Retallack, 2014a). A paleothermometer based on modernsoils under lichen-shrub tundra vegetation of Iceland (Óskarsson et al.,2012) may be the best available option, predicting temperature (T in°C) from chemical index of weathering (W), another base depletionmetric, as in the following formula

=+ +

W mAl OmAl O mCaO mNa O

100( )[( ) ( ) ( )]

2 3

2 3 2 (6)

=T W0.21 –8.93 (7)

with r2 = 0.81, s.e. = ± 0.5 °C.These calculations give the following temperate mean annual pa-

leotemperatures for the lower A horizons of moderately developedpaleosols: 10.3 ± 0.5 °C for type Jurl clay, 9.1 ± 0.5 °C for Jurl siltloam, and 8.7 ± 0.5 °C for Jurl clay loam. These results are consistentwith “temperate paleoclimate” (< 40 °C) interpreted for the 3.4 GaBuck Reef Chert of South Africa, from combined oxygen and hydrogenisotopes (Hren et al., 2009), revising earlier estimates of 55-85 °C fromoxygen isotopes alone (Knauth and Lowe, 2003). Temperatures< 75 °Care indicated by 6–24 vol% original quartz in the Panorama Formationpaleosols (Fig. 9): most quartz would have dissolved at higher tem-peratures (Sleep and Hessler, 2006). One counterargument is that ac-tive volcanism may have buried quartz before it dissolved (Lowe,2007), but quartz was exposed for millennia given the degree of che-mical weathering of the paleosols (Figs. 5 and 10). Temperate paleo-temperatures are reasonable considering likely mid-paleolatitude posi-tion of the Panorama Formation: between 20.5 ± 5o for the ca.3.46 Ga upper Apex Chert (Suganuma et al., 2006), and 59.0 ± 8.8o

for the 2.86 ± 0.2 Ga Millindinna Complex of the West Pilbara(Schmidt and Embleton, 1985).


12

A widely used indicator of mean annual precipitation (mm) forpaleosols (Sheldon et al., 2002) uses chemical index of alterationwithout potash (A of Eq. (5) without mK2O), in the following re-lationship.

=P e221 W0.0197 (8)

with r2 = 0.72, s.e. = ± 182 mm.This paleohyetometer gives humid mean annual precipitation:

1343 ± 182 mm for the type Jurl clay, 1154 ± 182 mm for Jurl clayloam, 1200 ± 182 mm for Jurl silt loam. These estimates are based onsoils of vascular plants unlike Archean vegetation, if any, so must betreated with suspicion. However, other indications of weathering suchas light rare earth enrichment of the Panorama Formation (Fig. 5) arecomparable with modern weathering under forests (Nesbitt, 1979;Morey and Setterholm, 1997).

Abundant evaporite pseudomorphs in paleosols within thePanorama Formation give the superficial appearance of soils of aridpaleoclimates (Retallack and Huang, 2010). Unlike gypsum reportedfrom some Archean paleosols (Nabhan et al., 2016), crystals and no-dules pseudomorphed by chert in the Panorama Formation may havebeen barite (BaSO4), as indicated by elevated Ba (8.2–3670 ppm) andcrystal form (Fig. 8D–E). Modern pedogenic barite crystals and spher-ulites are mostly microscopic (Kohut and Durdas, 1993; Shahid andJenkins, 1994; Howari et al., 2002), but some soils have barite nodulesup to 3.8 cm in diameter (Jennings and Driese, 2014; Jennings et al.,2015) like those of some paleosols (Retallack and Kirby, 2007; Jenningset al., 2015). Pedogenic barite is best known in humid forested soilswith 1000–1200 mm mean annual precipitation, and is controlled moreby microbially induced precipitation, and pH or redox changes at thewater table, than by paleoclimate (Retallack and Kirby, 2007; Jenningset al., 2015). In cold sulfur springs barite is precipitated from oxidationand neutralization of saline anoxic brines from the subsurface (Donovanet al., 1988; Bonny and Jones, 2008a, 2008b). Barite is favored by acidsulfate weathering at low pH (< 3) and is insoluble at higher pH, butgypsum forms and also is soluble at higher pH (4–9: Carson et al.,1982). There is no clear relationship between depth of barite and meanannual precipitation as there is for gypsum (Retallack and Huang,2010).

5.6. Paleoatmospheric carbon dioxide

Paleosols of the Panorama Formation are thin but deeply weathered(Fig. 9) and could be used to calculate soil CO2 using the paleoba-rometer of Sheldon (2006), if barite did not introduce the probability ofacid sulfate weathering found in modern baritic soils (Jennings andDriese, 2014). The model of Sheldon (2006) is predicated on alkali andalkaline earth depletion by carbonic acid from soil CO2, but thesebaritic paleosols may have been weathered by sulfuric acid as well. Thestrong chemical weathering observed is compatible with high ArcheanCO2 levels of 30,000 to 300,000 ppmv (100–1000 PAL) inferred byLowe and Tice (2004) from nahcolite formation at assumed Archeantemperatures of 70 °C, but such high temperatures are not supported bychemical composition of Panorama paleosols (section 6.5), oxygen andhydrogen isotopic studies (Hren et al., 2009), or preservation of quartz(Sleep and Hessler, 2006). A minimum level of 2500 ppmv (8.9 PAL)CO2 has been inferred from 3.2 Ga ferrous‑carbonate weathering rinds(Hessler et al., 2004), but such fluvial pebbles may have been isolatedfrom the atmosphere by groundwater. Estimates of 1500–9000 ppmv(5–32 PAL) CO2 for 3.0 Ga come from chemical weathering of theJerico Dam paleosol of South Africa (Grandstaff et al., 1986). This muchCO2 would give acid rain (pH 4.0–4.5 according to Ohmoto et al.,2014), and soil-microbial CO2 and H2SO4 could drop soil water pH to 3favoring acid sulfate weathering (Carson et al., 1982).

Even the most extreme of these estimates is short of the amountneeded for a greenhouse capable of maintaining likely temperateArchean paleotemperatures (Hren et al., 2009) given the faint young

sun. Other greenhouse gases are needed, including water vapor, CH4,C2H6, SO2, and COS (carbonyl sulfide: Ohmoto et al., 2014; Haqq-Misraet al., 2008; Ueno et al., 2009; Rosing et al., 2010; Kasting andKirschvink, 2012). Methane is allowed by low (160 ppm) H2 valueslikely for the Archean (Kasting and Kirschvink, 2012). An atmospherewith three times the current mass of N2 and a H2 mixing ratio of 0.1,would also have created an adequate greenhouse (Wordsworth andPierrehumbert, 2013), but N2 in the atmosphere was limited to 1.1 to0.5 bars judging from nitrogen and argon isotopic ratios in fluid in-clusions of the 3.5 Ga Dresser Formation of Western Australia (Martyet al., 2013). The gas SO2 proposed theoretically for early Mars (Halevyet al., 2007) and early Earth (Claire et al., 2014) is compatible with thepaleovolcanic setting and abundance of sulfate in paleosols of the Pa-norama Formation.

5.7. Paleoatmospheric oxygen

Very low levels of atmospheric oxygen are indicated by high ferrousto ferric iron ratios of paleosols in the Panorama Formation (Fig. 9),comparable with swamp soils today exhausted of oxygen by microbialrespiration (Vepraskas and Sprecher, 1997). However, paleosols of thePanorama Formation were not waterlogged, as indicated by sandcrystals and nodules of barite (Fig. 8), like those of other Archean welldrained paleosols (Nabhan et al., 2016). Laminations of hematite andgoethite are a feature of the uppermost surface of Wanta paleosols(Fig. 2G), comparable with aquatic banded iron formation in the Pa-norama Formation (Bolhar et al., 2005). The 3.5 Ga Marble Bar Chert isalso an aquatic banded‑iron formation with original hematite (Hoashiet al., 2009), best explained by chemolithotrophy of iron-oxidizingbacteria under reducing conditions (Konhauser et al., 2002; Croweet al., 2008). In most cases Archean and Cenozoic hematite can bedistinguished by Munsell hue, because oxidation of the modern outcropis brownish red (10YR-7.5YR), whereas Archean hematite is fire enginered (5R to 10R). Exceptions are Archean rocks locally stained byoverlying Mesozoic and Cenozoic paleosols (McLoughlin, 1996; Morrisand Ramanaidou, 2007), or local oxidation along deep fissures (Katoet al., 2009; Li et al., 2012).

Deeply weathered Panorama paleosols are a guide to soil exposureto O2, and thus partial pressure of atmospheric O2 (pO2 in atmo-spheres). This can be done by modifying the method of Sheldon (2006),using integrated whole profile oxidation of iron and manganese insteadof base loss, by Eq. 9 below.

=⎢⎣

+ ⎥⎦

pO F

A κK P D αL

2

1000O O2 2

(9)

∫∑==

=M ρ

Cτ δZ2

100pj p

Z

Z Dj w z

,

0 , ( )j w,

(10)

= −−

B 0.511

e0.49

k0.27 (11)

Variables needed for these calculations are F (mol O2/cm2)= summed molar mass transfer gains of Fe2O3 and MnO through theprofile using Eq. (10); ρp (g.cm−3) = bulk density of parent material;Cj,p (weight %) = chemical concentration of an element (j) in parentmaterial (p); τ,j,w (mass fraction) = mass transfer of a specified (j)element in a soil horizon (w); Z (cm) = depth in soil represented byanalysis; A (years) = duration of soil formation calculated here usingEq. (3); KO2

(mol./kg.bar) = Henry's Law constant for O2 (=0.00125,range 0.0012–0.0013 from Aachib et al., 2004); P (cm) = mean annualprecipitation calculated here using Eq. (8); κ (s cm3/mol year) = sec-onds per year divided by volume per mole of gas at standard tem-perature and pressure (=1409); DO2

(cm2/s) = diffusion constant for O2

in air (=0.203 at 20 °C, range from 0 to 40 °C is 0.179–0.227 fromDenny, 1993); α (fraction) = ratio of diffusion constant for O2 in soil


13

divided by diffusion constant for O2 in air (=0.2, range 0.09–0.32 fromAachib et al., 2004); L (cm) = original depth to water table (after de-compaction using fractional compaction B); B (fraction) = compactionof inceptisol due to burial (using Eq. (11) after Sheldon and Retallack,2001); k (km) = depth of burial (=13.826 km following vanKranendonk, 2000; van Kranendonk et al., 2006). Each one of thesevariables has its own error, and Gaussian error propagation has beencalculated for O2 following Hughes and Hase (2010) and equationsgiven in Supplementary information Table S7.

The results for Archean O2 are 711 ± 131 ppm(0.0034 ± 0.00063 times preindustrial atmospheric level or PAL) forthe Jurl clay, 75 ± 191 ppm (0.00026 ± 0.00091 PAL) for the Jurlclay loam, 88 ± 222 ppm (0.00042 ± 0.0011 PAL) for the Jurl siltloam, and 57 ± 208 ppm (0.00027 ± 0.00099 PAL) for the Wantasilt loam. These are maximal levels of O2 because based on the short soilformation times inferred from Eq. (3), rather than Eq. (4) which wouldgive minimal levels of O2 an order of magnitude lower. These estimatesof atmospheric O2 based on Panorama paleosols are within the range of20–1000 ppm (0.00095–0.0047 PAL) O2 estimated by a differentmethod of calculating oxygen demand of the 3.0 Ga Jerico Dam pa-leosol of South Africa and associated uraniferous conglomerates(Grandstaff et al., 1986). Other comparable estimates of 12–63 ppm(0.0006–0.003 PAL) O2 come from lack of oxidative Cr and U recyclingin 2.9 Ga Nsuze paleosols of South Africa (Crowe et al., 2013), and of20–200 ppm (0.00095–0.0095 PAL) O2 from cerium anomalies in the3.0–3.3 Ga Keonjhar paleosol deep weathering remnant of India(Mukhopadhyay et al., 2014). Low atmospheric O2 is also indicated byredox-sensitive minerals, including pyrite (FeS), uraninite (UO2), side-rite (FeCO3) and gersdorfffite (NiAsS), common (57–85%) in heavymineral separations from fluvial siliciclastic sediments some 3.2–2.7 Gain age in the Pilbara region (Rasmussen and Buick, 1999). Consequentlack of an ozone layer in the stratosphere is also indicated by mass-independent fractionation of sulfur isotopes in Archean pyrites andbarites of the Pilbara region (Ueno et al., 2008; Shen et al., 2009). Incontrast, Ohmoto et al. (2014) has argued for 100,000 ppm (0.48 PAL)atmospheric O2 based on the 3.42 Ga pre-Strelley paleosols of WesternAustralia (Ohmoto et al., 2014), and considers the chemically reducedcondition of the Jerico Dam paleosol due to hydrothermal fluids orburial gleization by organic acids after formation in an oxidizing at-mosphere of at least 3000 ppm (0.014 PAL) O2 (Ohmoto, 1996). Myown field inspection of the pre-Strelley “laterites” of Ohmoto et al.(2014) found that they were instead banded iron formations unrelatedto the paleosol. There is no clear hydrothermal enrichment of heavyrare earth elements in the Jericho Dam profile (Kimberly andGrandstaff, 1986) and burial gleization to 3.3 m depth in that paleosolis unlikely considering an order of magnitude less gleization in paleo-sols of productive Triassic forest ecosystems (Retallack, 1997b).

5.8. Organisms

Direct evidence of life in paleosols of the Panorama Formation re-mains elusive, although Jurl paleosols in the 3.0 Ga Farrel Quartziteincluded microfossils interpreted as photosynthetic sulfur bacteria,methanogens and actinobacteria (Retallack et al., 2016). Microbes mayexplain how paleosols in the Panorama Formation have such oxidizedminerals as barite (Wanta and Jurl) and hematite (Wanta) in the nearlyanoxic atmosphere inferred for these paleosols (section 5.7). Photo-synthetic sulfur bacteria (Chromatiaceae and Chlorobiaceae; Youssefet al., 2010) play a major role in precipitation of barite in modern coldsulfur springs within 12 m of the source, as demonstrated by dielfluctuations in aqueous sulfate and incubations, and stable isotopiccomparisons at Zodletone Spring, Oklahoma (Senko et al., 2004). He-matite in banded iron formations is thought to have formed in an anoxicaqueous medium by microbial photoferrolysis (Konhauser et al., 2002;Crowe et al., 2008), and a comparable origin is likely for banded ironformation within the Panorama Formation (Bolhar et al., 2005).

Modern acid-sulfate soils are not inimical to life but have a distinctivearray of acidophilic microbes, including Actinobacteria, Proteobacteriaand Bacteroidetes in acidic Gypsids (Solonchaks) of Western Australia(Benison et al., 2008; Mormile et al., 2009).

A line of evidence commonly used to infer life in Precambrian pa-leosols does not work for the Panorama Formation. Significant molardepletion of phosphorus (Fig. 10) requires organic ligands at pH > 5.5and stable aluminum (Neaman et al., 2005). Aluminum may appearerratic or retained compared with other elements (Figs. 6,9), but tauanalysis shows that aluminum was mildly leached from the paleosolsexcept for ferruginous top of Wanta (Fig. 10). This and abundant bariteare evidence of pH < 3 (Carson et al., 1982; Fitzpatrick et al., 2008).Thus apatite may have been dissolved abiotically by sulfuric acid inthese paleosols.

There is evidence of life in intertidal stromatolites before and afterdeposition of the Panorama Formation from permineralized micro-fossils in the overlying Strelley Pool Formation (Sugitani et al., 2010,2013; Wacey et al., 2010, 2011; Brasier et al., 2015; Duda et al., 2016)and underlying Mt. Ada Basalt (Awramik et al., 1983, 1988; vanKranendonk, 2006; Hickman, 2013) and Dresser Formation (vanKranendonk et al., 2008). Even older claims of aquatic life in stroma-tolites and banded iron formations come from 3.7-Ga rocks westGreenland (Nutman et al., 2016) and 4.2-Ga rocks of northern Quebec(Dodd et al., 2017).

6. Archean acid sulfate weathering

Baritic paleosols are common and widespread in silicified tuffs ofthe 3.46 Ga Panorama Formation. Although exposure of the PanoramaFormation is imperfect (Fig. 2A), detailed sections show an average of1.4 baritic paleosols per meter, and with 90 m of suitable facies in theformation (Fig. 3), the formation may contain at least 126 successivebaritic paleosols. Furthermore, baritic paleosol intervals can be tracedover 4 km laterally from the fault separated volcanic vent facies towardthe northern margin of outcrop (Fig. 1). Individual paleosols can bewalked 2 km laterally along the most resistant ridges (Fig. 2A). Thesepaleosols formed on a small dacitic stratovolcano (Fig. 11), but bariticpaleosols in the 3.2 Ga Moodies Group of South Africa formed in quartzsandstones of braided streams in a stable cratonic back-arc basin withreworked rhyodacitic tuffs (Nabhan et al., 2016). In South Africa,baritic paleosols are common within 220 m of section traceable re-gionally for 80 km laterally. Baritic nodular paleosols were thus wide-spread globally during the Archean, and this contrasts with the rarity ofbarite in modern and fossil soils (Table 3).

Baritic paleosols represent a distinctive style of acid-sulfate weath-ering rare in three distinct settings of modern soils. First are thin pro-files of Gypsids (Solonchaks, Sodosols) in and around acidic playa lakesin subhumid to semiarid regions (McKenzie et al., 2004; Fitzpatricket al., 2008; Benison and Bowen, 2013, 2015). Second are at depths of ameter or so near the water table in acidic Ultisols (Podzoluvisols) ofhumid uplands (Stoops and Zavaleta, 1978; Darmody et al., 1989;Jennings and Driese, 2014; Jennings et al., 2015). Third are thick ter-races of Entisols (Solonchaks) around cold sulfur springs of humid cli-mates (Donovan et al., 1988; Bonny and Jones, 2008a, 2008b). Theplaya option may be the best analog to Archean paleosols consideringtheir profile form, but barite is rare in such settings (Fitzpatrick et al.,2008), and not included in the list of minerals by Benison and Bowen(2013). The deep water table option sometimes has common and largenodules (Jennings et al., 2015), and is like barite nodules in the 2.65-GaMt. Roe paleosol (Teitler et al., 2015), but is not like the shallow bariticpaleosols of Panorama Formation (Fig. 9), Farrel Quartzite (Retallacket al., 2016), or Moodies Group (Nabhan et al., 2016). The barite de-posits of cold sulfur springs match the abundance of barite in Archeanpaleosols, but are limited in lateral distribution and have very differenttextures, including cavity-filling barite fans and stromatolite-like iso-pachous barite coatings (Donovan et al., 1988; Bonny and Jones,


14

2008a, 2008b).Baritic paleosols are rare through the Phanerozoic (Table 3), when

they are found mainly at times of CO2 greenhouse crises (Retallack andKirby, 2007). Gypsum-bearing paleosols are much more common atother times in the fossil record (Retallack and Huang, 2010). At severalpoints during Permian-Triassic transition, for example, CO2 estimatedfrom stomatal index reached an astounding 7832 ppm (Retallack,2009), with consequent decline in both pH and Eh of paleosols(Retallack et al., 2011; Retallack, 2013b). Such transient atmosphericcrises can be regarded as a reversion of life history toward Precambrianbiotic and atmospheric conditions, enabling a resurgence of bariticpaleosols.

A common thread in all occurrences of barite in soils is extremefluctuation of pH and redox (Carson et al., 1982; Jennings et al., 2015).Barium and barite is remarkably insoluble at pH > 3, which promotesits persistence and precipitation, but can only be transported and con-centrated in solution at pH < 3 (Tarr, 1933; Carson et al., 1982). Inacid sulfate soils such low pH is created by sulfuric acid from the oxi-dation of pyrite, and the pyrite in turn requires an anoxic environment(Carson et al., 1982). Thus unlike post Archean weathering dominatedby the weak acid, carbonic acid, Archean weathering may have beenlargely produced by the strong acid, sulfuric acid. This difference mayhave contributed to the surprising degree of weathering of Archeanshales and paleosols weathered to bauxites (Serdyuchenko, 1968; Dashet al., 1987) considering likely low productivity microbial earth com-munities (Retallack, 2014a), or perhaps abiotic weathering sometimesassumed (Rye and Holland, 1998). This whole archaic system ofwidespread acid sulfate weathering has been driven underground andinto marginal environments. Data on baritic paleosols of the Proter-ozoic and early Paleozoic are lacking (Table 3), so it remains a mysterywhen and how that marginalization occurred.

7. Conclusions

Baritic paleosols are common and widespread, not only in volcanicapron facies of the 3.46-Ga Panorama Formation, but in thick paleosolsbetween flows of 2.76-Ga Mt. Roe Basalt of Western Australia, onfloodplain facies of the 3.2-Ga Moodies Group of South Africa, andcoastal plain facies of the 3.0-Ga Farrel Quartzite of Western Australia(Table. 3). At 3.46 Ga, baritic paleosols of the Panorama Formation arecurrently the oldest known on Earth. Baritic nodular paleosols of acidsulfate weathering may have been widespread during the Archean, incontrast with the rarity of barite in modern soils and Phanerozoic pa-leosols. Barite in soils requires extreme fluctuation of pH and redox, andcan only be transported and concentrated in solution at pH < 3. In acidsulfate modern soils, such low pH is created by sulfuric acid from theoxidation of pyrite in an oxidizing atmosphere. Archean paleosols suchas those of the Panorama Formation have utilization of iron and man-ganese indicative of anoxic atmosphere, so that oxidation may havebeen provided by photosynthetic sulfur-oxidizing bacteria. Unlike postArchean weathering dominated by fungi and weak acid (carbonic acid),Archean weathering may have been largely produced by sulfur bacteriaand strong acid (sulfuric acid). Archaic and widespread surficial acid-sulfate weathering has been driven underground and into marginalenvironments by Palaeoproterozoic atmospheric oxidation, into playalakes, cold sulfur springs, and deep water tables within soils.

Acknowledgments

Jeremy Retallack helped during fieldwork. Also helpful were dis-cussions with Abigail Allwood, Bruce Runnegar, Debra Jennings, StevenDriese, Martin Van Kranendonk, Malcolm Walter, James Farquhar, andRoger Buick.

Fig. 11. Graphical abstract of reconstructed soils in the3.46 Ga Panorama Formation of Western Australia.


15

Appendix A. Supplementary data

Supplementary data to this article can be found online at https://doi.org/10.1016/j.palaeo.2017.10.013.

References

Aachib, M., Mbonimpa, M., Aubertin, M., 2004. Measurement and prediction of theoxygen diffusion coefficient in unsaturated media with applications to soil covers.Water Air Soil Pollut. 156, 163–193.

Allwood, A.C., Walter, M.R., Marshall, C.P., 2006. Raman spectroscopy reveals thermalpalaeoenvironments of c.3.5 billion-year-old organic matter. Vib. Spectrosc. 41,190–197.

Allwood, A.C., Walter, M.R., Burch, I.W., Kamber, B.S., 2007. 3.42 billion-year-oldstromatolite reef from the Pilbara Craton of Western Australia: ecosystem-scale in-sights to early life on Earth. Precambrian Res. 158, 198–227.

Allwood, A.C., Kamber, B.S., Walter, M.R., Burch, I.W., Kanik, I., 2010. Trace elementsrecord depositional history of an Early Archean stromatolitic carbonate platform.Chem. Geol. 270, 148–163.

Amundson, R., Dietrich, W., Bellugi, D., Ewing, S., Nishiizumi, K., Chong, G., Owen, J.,Finkel, R., Heimsath, A., Stewart, B., Caffee, M., 2012. Geomorphological evidencefor the late Pliocene onset of hyperaridity in the Atacama Desert. Geol. Soc. Am. Bull.124, 1048–1070.

Anders, E., Grevesse, N., 1989. Abundances of the elements: meteoritic and solar.Geochim. Cosmochim. Acta 53, 197–214.

Awramik, S.M., Schopf, J.W., Walter, M.R., 1983. Filamentous fossil bacteria from theArchean of Western Australia. Precambrian Res. 20, 357–374.

Awramik, S.M., Schopf, J.W., Walter, M.R., 1988. Carbonaceous filaments from NorthPole, Western Australia: are they bacteria in Archean stromatolites? Precambrian Res.39, 303–309.

Beattie, J.A., Haldane, A.D., 1958. The occurrence of palygorskite and barytes in certainparna soils of the Murrumbidgee region, New South Wales. Austral. J. Sci. 20,274–275.

Benison, K.C., Bowen, B.B., 2013. Extreme sulfur-cycling in acid brine lake environmentsof Western Australia. Chem. Geol. 351, 154–167.

Benison, K.C., Bowen, B.B., 2015. The evolution of end-member continental waters: theorigin of acidity in southern Western Australia. GSA Today 25 (6), 1–10.

Benison, K.C., Bowen, B.B., Oboh-Ikuenobe, F.E., Jagniecki, E.A., LaClair, D.A., Story,S.L., Mormile, M.R., Hong, B.-Y., 2007. Sedimentology of acid saline lakes in southernWestern Australia: newly described processes and products of an extreme environ-ment. J. Sediment. Res. 77, 366–388.

Benison, K.C., Jagniecki, E.A., Edwards, T.B., Mormile, M.R., Storrie-Lombardi, M.C.,2008. “Hairy blobs:” microbial suspects preserved in modern and ancient extremelyacid lake evaporites. Astrobiology 8, 1–15.

Birnbaum, S.J., Wireman, J.W., 1985. Sulfate-reducing bacteria and silica solubility: apossible mechanism for evaporite diagenesis and silica precipitation in banded ironformations. Can. J. Earth Sci. 22, 1905–1909.

Birnbaum, S.J., Wireman, J.W., Borowski, R., 1989. Silica precipitation induced by theanaerobic sulfate reducing bacterium Desulfovibrio desulfuricans: effects upon cellmorphology and implications for preservation. In: Crick, R.E. (Ed.), Origin,Evolution, and Modern Aspects of Biomineralization in Plants and Animals. Springer,New York, pp. 507–516.

Bizzarro, M., Simonetti, A., Stevenson, R.K., David, J., 2002. Hf isotope evidence for ahidden mantle reservoir. Geology 30, 771–774.

Bolhar, R., Van Kranendonk, M.J., Kamber, B.S., 2005. A trace element study of siderite-jasper banded iron formation in the 3.45 Ga Warrawoona Group, Pilbara Craton:Formation from hydrothermal fluids and shallow seawater. Precambrian Res. 137,93–114.

Bonny, S.M., Jones, B., 2008a. Petrography and textural development of inorganic andbiogenic lithotypes in a relict barite tufa deposit at Flybye Springs, NT, Canada.Sedimentology 55, 275–303.

Bonny, S.M., Jones, B., 2008b. Experimental precipitation of barite (BaSO4) amongstreamers of sulfur-oxidizing bacteria. J. Sediment. Res. 78, 357–365.

Bontognali, T.R.R., Sessions, A.L., Allwood, A.C., Fischer, W.W., Grotzinger, J.P.,Summons, R.E., Eiler, J.M., 2012. Sulfur isotopes of organic matter preserved in 3.45-billion-year-old stromatolites reveal microbial metabolism. U.S. Nat. Acad. Sci. Proc.109, 15146–15151.

Bown, T.M., Kraus, M.J., 1988. Geology and paleoenvironment of the Oligocene JebelQatrani Formation and adjacent rocks, Fayum depression, Egypt. US Geol. Surv. Prof.Pap. 1452, 1–60.

Brasier, M.D., Green, O.R., Jephcoat, A.P., Kleppe, A.K., van Kranendonk, M.J., Lindsay,J.V., Stewart, A., Grassineau, N.V., 2002. Questioning the evidence for Earth's oldestfossils. Nature 416, 76–81.

Brasier, M.D., Antcliffe, J., Saunders, M., Wacey, D., 2015. Changing the picture of Earth'searliest fossils (3.5–1.9 Ga) with new approaches and new discoveries. U.S. Nat.Acad. Sci. Proc. 112, 4859–4864.

Brimhall, G.H., Chadwick, O.A., Lewis, C.J., Compston, W., Williams, I.S., Danti, K.J.,Dietrich, W.E., Power, M.E., Hendricks, D., Bratt, J., 1992. Deformational masstransport and invasive processes in soil evolution. Science 255, 695–702.

Brock-Hon, A.L., Robins, C.R., Buck, B.J., 2012. Micromorphological investigation ofpedogenic barite in Mormon Mesa petrocalcic horizons, Nevada USA: implication forgenesis. Geoderma 179, 1–8.

de Brodtkorb, M.K., 1989. Barite nodules and concretions. In: de Brodtkorb, M.K. (Ed.),Non-Metalliferous Stratabound Ore Fields. Van Nostrand Reinhold, New York, pp.

233–254.Brown, A.J., Cudahy, T.J., Walter, M.R., 2006. Hydrothermal alteration at the Panorama

Formation, North Pole Dome, Pilbara Craton, Western Australia. Precambrian Res.151, 211–233.

Buck, B.J., Lawton, T.F., Brock, A.L., 2010. Evaporitic paleosols in continental strata ofthe Carroza Formation, La Popa Basin, Mexico: record of Paleogene climate and salttectonics. Geol. Soc. Am. Bull. 122, 1011–1026.

Buick, R., Barnes, K.R., 1984. Cherts in the Warrawoona Group: Early Archean silicifiedsediments deposited in shallow-water environments. In: Muhling, J., Groves, D.I.,Blake, T.S. (Eds.), Archean and Proterozoic Basins of the Pilbara, Western Australia—Evolution and Mineralization Potential. Univ. Western Australia Geol. Dept. Univ.Extens. Publ. 9, pp. 37–53.

Buick, R., Dunlop, J.S.R., 1990. Evaporitic sediments of early Archean age from theWarrawoona, North Pole, Western Australia. Sedimentology 37, 247–277.

Buick, R., Thronetee, J.R., McHaughton, N.J., Smith, J.B., Barley, M.E., Savage, M., 1995.Record of emergent continental crust ~3.5 billion years ago in the Pilbara craton ofAustralia. Nature 375, 574–576.

Bureau of Meteorology, 2016. Australian Government Website. http://www.bom.gov.au/climate/data/ (accessed April 7, 2016).

Burgman, A., 2007. Nyamal Dictionary, English-Nyamal Finderlist and Topical Wordlist.Wangka Maya, Port Hedland (27 p).

Carson, C.D., Fanning, D.S., Dixon, J.B., 1982. Alfisols and ultisols with acid sulfateweathering in Texas. In: Kittrick, J.A., Fanning, D.S., Hossner, L.R. (Eds.), AcidSulfate Weathering. Soil Sci. Soc. Amer. Spec. Publ. 10pp. 127–146.

Claire, M.W., Kasting, J.F., Domagal-Goldman, S.D., Stuëken, E., Buick, R., Meadows,V.S., 2014. Modeling the signature of sulfur mass-independent fractionation pro-duced in the Archean atmosphere. Geochim. Cosmochim. Acta 141, 365–380.

Clauer, N., O'Neil, J.R., Bonnnot-Courtois, C., Holtzapfel, T., 1990. Morphological, che-mical, and isotopic evidence for an early diagenetic evolution of detrital smectite inmarine sediments. Clay Clay Miner. 38, 33–46.

Crowe, S.A., Jones, C., Katsev, S., Magen, C., O'Neill, A.H., Sturm, A., Canfield, D.E.,Haffner, G.D., Mucci, A., Sundby, B., Fowle, D.A., 2008. Photoferrotrophs thrive in anArchean Ocean analogue. U.S. Nat. Acad. Sci. Proc. 105, 15938–15943.

Crowe, S.A., Døssing, L.N., Beukes, N.J., Bau, M., Kruger, S.J., Frei, R., Canfield, D.E.,2013. Atmospheric oxygenation three billion years ago. Nature 501, 535–538.

Cullers, R.L., DiMarco, M.J., Lowe, D.R., Stone, J., 1993. Geochemistry of silicified, felsic,volcaniclastic sediments from the early Archean panorama formation, Pilbara block,Western Australia: an evaluation of depositional and post-depositional processes withspecial emphasis on rare earth elements. Precambrian Res. 60, 99–116.

Darmody, R.G., Harding, S.D., Hassett, J.J., 1989. Barite authigenesis in surficial soils ofmid-continental United States. In: Miles, D.L. (Ed.), Water-Rock Interaction. A.A.Balkema, Rotterdam, pp. 183–186.

Dash, B., Sahni, K.N., Bowes, D.A., 1987. Geochemistry and original nature ofPrecambrian knondalites in the Eastern Ghats, Orissa, India. Trans. Roy. Soc.Edinburgh. Earth Sci. 78, 115–127.

Denny, M.W., 1993. Air and Water: The Biology and Physics of Life's Media. PrincetonUniv. Press, Princeton (343 p).

Dietrich, M., 1975. Wüstenrose. Verein. Freunde Mineral. Geol. 26. pp. 217–219.DiMarco, M.J., Lowe, D.R., 1989. Stratigraphy and sedimentology of an early Archean

felsic volcanic sequence, Eastern Pilbara block, Western Australia, with special re-ference to the duffer formation and implications for crustal evolution. PrecambrianRes. 44, 147–169.

Dodd, M.S., Papineau, D., Grenne, T., Slack, J.F., Rittner, M., Pirajno, F., O'Neil, J., Little,C.T., 2017. Evidence for early life in Earth's oldest hydrothermal vent precipitates.Nature 543, 60–64.

Domagal-Goldman, S., 2014. The upside-down biosphere: evidence for the partiallyoxygenated oceans during the Archean Eon. In: Shaw, G.H. (Ed.), Earth's EarlyAtmosphere and Surface Environment. Geol. Soc. Amer. Spec. Pap. 504. pp. 97–99.

Donovan, R.N., Younger, P., Ditzell, C., 1988. Some aspects of the geology of ZodletoneMountain, Southwestern Oklahoma. In: Hayward, O.T. (Ed.), Geological Society ofAmerica Centennial Field Guide Southeast Section. Geological Society of America,Boulder, pp. 99–102.

Donovan, J.J., Lowers, H.A., Rusk, B.G., 2011. Improved electron probe microanalysis oftrace elements in quartz. Am. Mineral. 96, 274–281.

Duda, J.P., Van Kranendonk, M.J., Thiel, V., Ionescu, D., Strauss, H., Schäfer, N., Reitner,J., 2016. A rare glimpse of paleoarchean life: geobiology of an exceptionally pre-served microbial mat facies from the 3.4 Ga Strelley Pool Formation, WesternAustralia. PLoS One 11, e0147629.

Eriksson, K.A., Warren, J.K., 1983. A paleohydrologic model for Early Proterozoic dolo-mitization and silicification. Precambrian Res. 21, 299–321.

Eugster, H.P., 1969. Inorganic bedded cherts from the Magadi area, Kenya. Contrib.Mineral. Petrol. 22, 1–31.

Fedo, C.M., Nesbitt, H.W., Young, G.M., 1995. Unraveling the effects of potassium me-tasomatism in sedimentary rocks and paleosols, with implications for paleo-weathering conditions and provenance. Geology 23, 921–924.

Fitzpatrick, R., Degens, B., Baker, A., Raven, M., Shand, P., Smith, M., Rogers, S., George,R., 2008. Avon Basin, WA Wheatbelt: Acid sulfate soils and salt efflorescences in opendrains and receiving environments. In: Fitzpatrick, R., Shand, P. (Eds.), Inland AcidSulfate Soil Systems across Australia. CRC LEME Open File Report 249, pp. 189–204.

Food and Agriculture Organization, 1974. Soil Map of the World, 1:500,000. I. Legend.UNESCO, Paris (59 p).

Food and Agriculture Organization, 1978. Soil Map of the World, 1:500,000. X.Australasia. UNESCO, Paris (221 p).

Gallagher, T.M., Sheldon, N.D., 2013. A new paleothermometer for forest paleosols andits implications for Cenozoic climate. Geology 41, 647–651.

Garces, B.L.V., Aguilar, J.G., 1994. Permian saline lakes in the Aragón-Béarn Basin,


16

https://doi.org/10.1016/j.palaeo.2017.10.013

https://doi.org/10.1016/j.palaeo.2017.10.013

http://refhub.elsevier.com/S0031-0182(17)30361-9/rf0005






























































































http://www.bom.gov.au/climate/data/

http://www.bom.gov.au/climate/data/




































































Western Pyrenees. In: Renaut, R.W., Last, W.M. (Eds.), Sedimentology andGeochemistry of Modern and Ancient Saline Lakes. Soc.Sediment. Geol. Spec. Publ.50pp. 267–290.

Grandstaff, D.E., Edelman, M.J., Foster, R.W., Zbinden, E.M., Kimberley, M.M., 1986.Chemistry and mineralogy of Precambrian paleosols at the base of the Dominion andPongola Groups (Transvaal, South Africa). Precambrian Res. 32, 97–131.

Halevy, I., Zuber, M.T., Schrag, D.P., 2007. A sulfur dioxide climate feedback on earlyMars. Science 318, 1903–1907.

Haqq-Misra, J.D., Domagal-Goldman, S.D., Kasting, P.J., Kasting, J.F., 2008. A revisedhazy methane greenhouse for the Archean Earth. Astrobiology 8, 1127–1137.

Harvey, H.R., Tuttle, J.H., Bell, J.T., 1995. Kinetics of phytoplankton decay during si-mulated sedimentation: changes in biochemical composition and microbial activityunder oxic and anoxic conditions. Geochim. Cosmochim. Acta 59, 3367–3377.

Hay, R.L., 1968. Chert and its sodium-silicate precursors in sodium-carbonate lakes ofEast Africa. Contrib. Mineral. Petrol. 17, 255–274.

Hessler, A.M., Lowe, D.R., Jones, R.L., Bird, D.K., 2004. A lower limit for atmosphericcarbon dioxide levels 3.2 billion years ago. Nature 428, 736–738.

Hickman, A.H., 2008. Regional review of the 3424–3350 Ma Strelley Pool Formation,Pilbara Craton, Western Australia. West. Austral. Geol. Surv. Rec. 2008/15, 1–61.

Hickman, A.H., 2013. Geology of the North Shaw 1:100 000 Sheet. Western AustraliaGeological Survey, Perth (1p).

Hoashi, M., Bevacqua, D.C., Otake, T., Watanabe, Y., Hickman, A.H., Utsunomiya, S.,Ohmoto, H., 2009. Primary hematite formation in an oxygenated sea 3.46 billionyears ago. Nat. Geosci. 2, 301–306.

Howari, F.M., Goodell, P.C., Miyamoto, S., 2002. Spectral properties of salt crusts formedon saline soils. J. Environ. Qual. 31, 1453–1461.

Hren, M.T., Tice, M.M., Chamberlain, C.P., 2009. Oxygen and hydrogen isotope evidencefor a temperate climate 3.42 billion years ago. Nature 462, 205–208.

Hughes, I., Hase, I.P.A., 2010. Measurements and their Uncertainties: A Practical Guide toModern Error Analysis. Oxford Univ. Press, Oxford (136 p).

Isbell, R.F., 1996. The Australian Soil Classification. C.S.I.R.O, Collingwood (152 p).Jennings, D.S., Driese, S.G., 2014. Understanding barite and gypsum precipitation in

upland acid-sulfate soils: an example from a Lufkin series toposequence, South-cen-tral Texas, U.S.A. Sediment. Geol. 299, 106–118.

Jennings, D.S., Driese, S.G., Dworkin, S.I., 2015. Comparison of modern and ancientbarite-bearing acid-sulfate soils using micromorphology, geochemistry and field re-lationships. Sedimentology 62, 1078–1099.

Kasting, J., Kirschvink, J., 2012. Evolution of a habitable planet. In: Impey, C., Lunine, J.,Funes, J. (Eds.), Frontiers of Astrobiology. Cambridge Univ. Press, New York, pp.115–131.

Kato, Y., Nakamura, K., 2003. Origin and global tectonic significance of Early Archeancherts from the Marble Bar Greenstone Belt, Pilbara Craton, Western Australia.Precambrian Res. 125, 191–243.

Kato, Y., Suzuki, K., Nakamura, K., Hickman, A.H., Nedachie, M., Kusakabe, M.,Bevacqua, D.C., Ohmoto, H., 2009. Hematite formation by oxygenated groundwatermore than 2.76 billion years ago. Earth Planet. Sci. Lett. 278, 40–44.

Keyser, A.W., 1968. Some indications of arid climate during the deposition of the BeaufortSeries. Geol. Surv. South Africa Annals 5, 77–78.

Khalaf, F.I., El Sayed, M.I., 1989. Occurrence of barite-bearing mud balls within theMio–Pliocene terrestrial sand (Kuwait group) in southern Kuwait, Arabian gulf.Sediment. Geol. 64, 197–202.

Kimberly, M.M., Grandstaff, D.E., 1986. Profiles of elemental concentrations inPrecambrian paleosols in basaltic and granitic parent materials. Precambrian Res. 32,133–154.

Knauth, L.P., Lowe, D.R., 2003. High Archean climatic temperature inferred from oxygenisotope geochemistry of cherts in the 3.5 Ga Swaziland Supergroup, South Africa.Geol. Soc. Am. Bull. 115, 566–580.

Kohut, C.K., Durdas, M.J., 1993. Evaporite mineralogy and trace-element content of salt-affected soils in Alberta. Can. J. Soil Sci. 73, 399–409.

Kolesnikov, A.V., Grazhdankin, D.V., Maslov, A.V., 2012. Arumberia-type structures in theupper Vendian of the Urals. Dokl. Earth Sci. 447, 1233–1239.

Kolesnikov, A.V., Marusina, V.V., Nagovitsin, K.E., Maslov, A.V., Grazhdankin, D.V.,2015. Ediacaran biota in the aftermath of the Kotlinian crisis: Asha group of the SouthUrals. Precambrian Res. 263, 59–78.

Konhauser, K.O., Hamade, T., Raiswell, R., Morris, R.C., Ferris, F.G., Southam, G.,Canfield, D.E., 2002. Could bacteria have formed the Precambrian banded iron for-mations? Geology 30, 1079–1082.

Krapež, B., 1984. Sedimentation in a small, fault-bounded basin — Lalla RookhSandstone, East Pilbara Block. In: Muhling, J., Groves, D.I., Blake, T.S. (Eds.),Archean and Proterozoic Basins of the Pilbara, Western Australia — Evolution andMineralization Potential. Univ. Western Austral. Geol. Dept. Univ. Extens. Publ. 9pp.89–110.

Kreuser, T., 1995. Tectonic and climatic controls of lacustrine sedimentation in preriftand rift settings in the Permian–Triassic of East Africa. J. Paleolimnol. 13, 3–19.

Lambert, I.B., Donnelly, T.H., Dunlop, J.S.R., Groves, D.I., 1978. Stable isotopic compo-sitions of early Archean sulfate deposits of probably evaporitic and volcanogenicorigins. Nature 276, 808–811.

Larsen, D., 2008. Revisiting silicate authigenesis in the Pliocene-Pleistocene Lake Tecopabeds, southeastern California: depositional and hydrological controls. Geosphere 4,612–639.

Larsen, L.M., Rex, D.C., 1992. A review of the 2500 Ma span of alkaline-ultramafic, po-tassic and carbonatitic magmatism in West Greenland. Lithos 28, 367–402.

Laznicka, P., 1976. Barite nodules of possibly late diagenetic origin from Twitya Riverarea, Mackenzie Mountains, Northwest Territories. Can. J. Earth Sci. 13, 1446–1455.

Lepot, K., Williford, K.H., Ushikubo, T., Sugitani, K., Mimura, K., Spicuzza, M.J., Valley,J.W., 2013. Texture-specific isotopic compositions in 3.4 Gyr old organic matter

support selective preservation in cell-like structures. Geochim. Cosmochim. Acta 112,66–86.

Li, W., Johnson, C.M., Beard, B.L., 2012. U–Th–Pb isotope data indicate Phanerozoic agefor oxidation of the 3.4 Ga Apex Basalt. Earth Planet. Sci. Lett. 319-320, 197–206.

Lowe, D.R., 2007. A comment on weathering of quartz as an Archean climatic indicatorby N.H. Sleep and A.M. Hessler. Earth Planet. Sci. Lett. 253, 530–533.

Lowe, D.R., 1983. Restricted shallow-water sedimentation of early Archean stromatoliticand evaporitic strata of the Strelley Pool Chert, Pilbara Block, Western Australia.Precambrian Res. 19, 239–283.

Lowe, D.R., Tice, M.M., 2004. Geologic evidence for Archean atmospheric and climaticevolution: fluctuating levels of CO2, CH4, and O2 with an overriding tectonic control.Geology 32, 493–496.

Lowe, D.R., Worrell, G.F., 1999. Sedimentology, mineralogy, and implications of silicifiedevaporites in the Kromberg Formation, Barberton Greenstone Belt, South Africa. In:Lowe, D.R., Byerly, G.R. (Eds.), Geologic Evolution of the Barberton Greenstone Belt,South Africa. Geol. Soc. Amer. Spec. Publ. 329pp. 167–188.

Maliva, R.G., Knoll, A.H., Simonson, B.M., 2005. Secular change in the Precambrian silicacycle: insights from chert petrology. Geol. Soc. Am. Bull. 117, 835–845.

Marty, B., Zimmermann, L., Pujol, M., Burgess, R., Philippot, P., 2013. Nitrogen isotopecomposition and density of the Archean atmosphere. Science 342, 101–104.

Maslov, A.V., Grazhdankin, D.V., 2011. Petrologiya, geokhimiya, litokhimicheskeoso'ennosti porod verkhnego benda s prizhnakami vliyaniya proshessov pedogenezaOkrestnosti r Ust'-Katav (petrology, geochemistry and lithochemical features of upperVendian rocks with signs of pedogenic processes) near the city of Ust-katau. TrudyInst. Geol. Geophys. Ural Branch Russian Acad. Sci. 158, 113–118.

McCarthy, P.J., Plint, A.G., 2013. A pedostratigraphic approach to nonmarine sequencestratigraphy: a three-dimensional paleosol-landscape model from the Cretaceous(Cenomanian) Dunvegan Formation, Alberta and British Columbia, Canada. In:Driese, S.G. (Ed.), New Frontiers in Paleopedology and Terrestrial Paleoclimatology.Soc. Econ. Paleont. Mineral. Spec. Pap. 44pp. 159–177.

McKenzie, N., Jacquier, D., Isbell, R., Brown, K., 2004. Australian Soils and Landscapes.CSIRO, Melbourne (416 p).

McLoughlin, S., 1996. Early Cretaceous floras of Western Australia. West. Austral. Mus.Rec. 18, 19–65.

Morey, G.B., Setterholm, D.R., 1997. Rare earth elements in weathering profiles and se-diments of Minnesota: implications for provenance studies. J. Sediment. Res. 67,105–115.

Mormile, M.R., Hong, B.-Y., Benison, K.C., 2009. Molecular analysis of the microbialcommunities of Mars analog lakes in Western Australia. Astrobiology 9, 920–930.

Morris, R.C., Ramanaidou, E.R., 2007. Genesis of the channel iron deposits (CID) of thePilbara region Western Australia. Austral. J. Earth Sci. 54, 733–756.

Mukhopadhyay, J., Crowley, Q.G., Ghosh, S., Ghosh, G., Charkrabarti, K., Misra, B.,Heron, K., Bose, S., 2014. Oxygenation of the Archean atmosphere: new paleosolconstraints from eastern India. Geology 42, 923–926.

Murphy, C.P., 1983. Counting pores and illuvial clay in thin section. Geoderma 31,133–150.

Nabhan, S., Luber, T., Scheffler, T., Heubeck, C., 2016. Climatic and geochemical im-plications of Archean pedogenic gypsum in the Moodies Group (∼3.2 Ga), BarbertonGreenstone Belt, South Africa. Precambrian Res. 275, 119–134.

Neaman, A., Chorover, J., Brantley, S.L., 2005. Element mobility patterns record organicligands in soils on early Earth. Geology 33, 117–120.

Nelson, D.R., 2002. Tuffaceous rhyolite, shark well (sample 160221). Western AustraliaGeological Survey Geochronology Dataset 238, 106–108.

Nesbitt, H.W., 1979. Mobility and fractionation of rare earth elements during weatheringof a granodiorite. Nature 279, 206–210.

Nesbitt, H.W., Young, G.M., 1989. Formation and diagenesis of weathering profiles. J.Geol. 97, 129–147.

Neuhaus, S., 2010. New age for Paleogene/Neogene clastics at the northern terminationof the Upper Rhine Graben (Hesse, Germany). Zeit. Deutsche Gesell. Geowiss. 161,303–322.

Novoselov, A.A., de Souza Filho, C.R., 2015. Potassium metasomatism of Precambrianpaleosols. Precambrian Res. 262, 67–83.

Nutman, A.P., Bennett, V.C., Friend, C.R., Van Kranendonk, M.J., Chivas, A.R., 2016.Rapid emergence of life shown by discovery of 3,700-million-year-old microbialstructures. Nature 537, 535–538.

Ohmoto, H., 1996. Evidence in pre-2.2 Ga paleosols for the early evolution of atmosphericoxygen and terrestrial biota. Geology 24, 1135–1138.

Ohmoto, H., Watanabe, Y., Lasaga, A.C., Naraoka, H., Johnson, I., Brainard, J., Chorney,A., 2014. Oxygen, iron and sulfur geochemical cycles on the early Earth: problemsand contradictions. In: Shaw, G.H. (Ed.), Earth's Early Atmosphere and SurfaceEnvironment. Geol. Soc. Amer. Spec. Pap. 504pp. 55–95.

Olson, E.C., 1967. Early Permian vertebrates of Oklahoma. Oklahoma Geol. Surv. Circ.74, 1–111.

Óskarsson, B.V., Riishuus, M.S., Arnalds, O., 2012. Climate-dependent chemical weath-ering of volcanic soils in Iceland. Geoderma 189–190, 635–651.

Pinti, D.L., Mineau, R., Clement, V., 2009. Hydrothermal alteration and microfossil ar-tefacts of the 3,465-million-year-old apex chert. Nat. Geosci. 2, 640–643.

Pufahl, P.K., Pirajno, F., Hiatt, E.E., 2013. Riverine mixing and fluvial iron formation: anew type of Precambrian biochemical sediment. Geology 41, 1235–1238.

Pujol, M., Marty, B., Bumard, P., Philippot, P., 2009. Xenon in Archean barite: weak decayof 130Ba, mass dependent isotopic fractionation and implication for barite formation.Geochim. Cosmochim. Acta 73, 6834–6846.

Rasmussen, B., Buick, R., 1999. Redox state of the Archean atmosphere: evidence fromdetrital heavy minerals in ca. 3250–2750 Ma sandstones from the Pilbara Craton,Australia. Geology 27, 115–118.

Renaut, R.W., Tiercelin, J.-J., 1994. Lake Bogoria, Kenya Rift Valley— a sedimentological


17











































































































































































overview. In: Renaut, R.W., Last, W.M. (Eds.), Sedimentology and Geochemistry ofModern and Ancient Saline Lakes. Soc. Sedim. Geol. Spec. Publ. 50pp. 101–124.

Retallack, G.J., 1997a. A Colour Guide to Paleosols. Wiley, Chichester (346 p).Retallack, G.J., 1997b. Palaeosols in the upper Narrabeen Group of New South Wales as

evidence of Early Triassic palaeoenvironments without exact modern analogues.Austral. J. Earth Sci. 44, 185–201.

Retallack, G.J., 2005. Pedogenic carbonate proxies for amount and seasonality of pre-cipitation in paleosols. Geology 33, 333–336.

Retallack, G.J., 2009. Greenhouse crises of the past 300 million years. Geol. Soc. Am. Bull.121, 1441–1455.

Retallack, G.J., 2011. Neoproterozoic glacial loess and limits to snowball earth. Geol. Soc.London J. 168, 1–19.

Retallack, G.J., 2012. Were Ediacaran siliciclastics of South Australia coastal or deepmarine? Sedimentology 59, 1208–1236.

Retallack, G.J., 2013a. Ediacaran life on land. Nature 493, 89–92.Retallack, G.J., 2013b. Permian and Triassic greenhouse crises. Gondwana Res. 24,

90–103.Retallack, G.J., 2014a. Precambrian life on land. Palaeobotanist 63, 1–15.Retallack, G.J., 2014b. Volcanosedimentary paleoenvironments of Ediacaran fossils in

Newfound land. Geol. Soc. Am. Bull. 126, 619–638.Retallack, G.J., 2014c. Paleosols and paleoenvironments of early Mars. Geology 42,

755–758.Retallack, G.J., 2016. Field and laboratory tests for recognition of Ediacaran paleosols.

Gondwana Res. 36, 94–110.Retallack, G.J., Huang, C.-M., 2010. Depth to gypsic horizon as a proxy for paleopreci-

pitation in paleosols of sedimentary environments. Geology 38, 403–406.Retallack, G.J., Kirby, M.X., 2007. Middle Miocene global change and paleogeography of

Panama. Palaios 22, 667–679.Retallack, G.K., Sheldon, N.D., Carr, P.F., Fanning, M., Thompson, C.A., Williams, M.L.,

Jones, B.G., Hutton, A., 2011. Multiple Early Triassic greenhouse crises impededrecovery from Late Permian mass extinction. Palaeogeogr. Palaeoclimatol.Palaeoecol. 308, 233–251.

Retallack, G.J., Krinsley, D.H., Fischer, R., Razink, J.J., Langworthy, K., 2016. Archeancoastal-plain paleosols and life on land. Gondwana Res. 40, 1–20.

Rosing, M.T., Bird, D.K., Sleep, N.H., Bjerrum, C.J., 2010. No climate paradox under thefaint early sun. Nature 464, 744–747.

Rye, R., Holland, H.D., 1998. Paleosols and the evolution of atmospheric oxygen: a cri-tical review. Amer. J. Sci. 298, 621–672.

Schieber, J., Bish, D., Coleman, M., Reed, M., Hausrath, E.M., Cosgrove, J., Gupta, S.,Minitti, M.E., Edgett, K.S., Malin, M., 2017. Encounters with an unearthly mudstone:understanding the first mudstone found on Mars. Sedimentology 64, 311–358.

Schmidt, P.W., Embleton, B.J.J., 1985. Prefolding and overprint magnetic signatures inPrecambrian (2.9–2.7 Ga) igneous rocks from the Pilbara Craton and HamersleyBasin, NW Australia. J. Geophys. Res. 90 (B4), 2967–2984.

Schopf, J.W., Kudryavtsev, A.B., Czaja, A.D., Tripathi, A.B., 2007. Evidence of Archeanlife: stromatolites and microfossils. Precambrian Res. 158, 141–155.

Schrank, E., Mahmoud, M.S., 2000. New taxa of angiosperm pollen and associated pa-lynomorphs from the early late Cretaceous of Egypt (Mahgrabi Formation, KhargaOasis). Rev. Palaeobot. Palynol. 112, 167–188.

Senko, J.M., Campbell, B.S., Henriksen, J.R., Elshahed, M.S., Dewers, T.A., Krumholz,L.R., 2004. Barite deposition resulting from phototrophic sulphide-oxidizing bacterialactivity. Geochim. Cosmochim. Acta 68, 773–780.

Serdyuchenko, D.P., 1968. Metamorphosed weathering rusts of the Precambrian: theirmetallogenis and petrogaphic fabric. In: Hejtman, B. (Ed.), Precambrian Geology.Proc. 13th Int. Geol. Congr. Prague 4pp. 37–42.

Setti, M., Marinoni, L., López-Galindo, L., 2004. Mineralogical and geochemical char-acteristics (major, minor, trace elements and REE) of detrital and authigenic clayminerals in a Cenozoic sequence from Ross Sea, Antarctica. Clay Miner. 39, 405–421.

Shahid, S.A., Jenkins, D.A., 1994. Mineralogy and micromorphology of salt crusts fromthe Punjab, Pakistan. In: Ringrose-Voase, A.J., Humphreys, G.S. (Eds.), SoilMicromorphology: Studies in Management and Genesis. Elsevier, Amsterdam, pp.799–800.

Sheldon, N.D., 2006. Precambrian paleosols and atmospheric CO2 levels. PrecambrianRes. 147, 148–155.

Sheldon, N.D., Retallack, G.J., 2001. Equation for compaction of paleosols due to burial.Geology 29, 247–250.

Sheldon, N.D., Retallack, G.J., Tanaka, S., 2002. Geochemical climofunctions from NorthAmerican soils and application to paleosols across the Eocene-Oligocene boundary inOregon. J. Geol. 110, 687–696.

Shen, Y., Buick, R., Canfield, D.E., 2001. Isotopic evidence for microbial sulphate re-duction in the early Archaean era. Nature 410, 77–81.

Shen, Y., Farquhar, J., Masterson, A., Kaufman, A.J., Buick, R., 2009. Evaluating the roleof microbial sulfate reduction in the early Archean using quadruple isotope sys-tematics. Earth Planet. Sci. Lett. 279, 383–391.

Siever, R., 1962. Silica solubility, 0°–200 °C, and the diagenesis of siliceous sediments. J.Geol. 70, 127–150.

Skinner, M.F., Taylor, B.E., 1967. A revision of the geology and paleontology of the BijouHills, South Dakota. Am. Mus. Novit. 300, 1–53.

Sleep, N.H., Hessler, A.M., 2006. Weathering of quartz as an Archean climatic indicator.Earth Planet. Sci. Lett. 241, 594–602.

Smith, G.I., 1979. Subsurface stratigraphy and geochemistry of late quaternary evapor-ites, Searles Lake, California. U.S. Geol. Surv. Prof. Pap. 1043, 1–130.

Soil Survey Staff, 2014. Keys to Soil Taxonomy. Natural Resources Conservation Service,Washington DC (600 p).

Speczik, S., El Sayed, A.A.Y., 1991. Fluid inclusion studies of Bahariya barites (WesternDesert, Egypt). Acta Geol. Pol. 41, 109–116.

Stoops, G.J., Zavaleta, A., 1978. Micromorphological evidence of barite neoformation insoils. Geoderma 20, 63–70.

Sugahara, H., Sugitani, K., Mimura, K., Yamashita, F., Yamamoto, K., 2010. A systematicrare-earth elements and yttrium study of Archean cherts at the Mount Goldsworthygreenstone belt in the Pilbara Craton: implications for the origin of microfossil-bearing black cherts. Precambrian Res. 177, 73–87.

Suganuma, Y., Hamano, Y., Niitsuma, S., Hoashi, M., Hisamitsu, T., Niitsuma, N.,Kodama, K., Nedachi, M., 2006. Paleomagnetism of the marble bar Chert member,Western Australia: implications for apparent polar wander path for Pilbara Cratonduring Archaean time. Earth Planet. Sci. Lett. 252, 360–371.

Sugitani, K., Mimura, K., Suzuki, K., Nagamine, K., Sugisaki, R., 2003. Stratigraphy andsedimentary petrology of an Archean volcanic/sedimentary succession at Mt.Goldsworthy in the Pilbara block, Western Australia: implications of evaporite(nahcolite) and barite deposition. Precambrian Res. 120, 55–79.

Sugitani, K., Lepot, K., Nagaoka, T., Mimura, K., van Kranendonk, M.J., Oehler, D.Z.,Walter, M.R., 2010. Biogenicity of morphologically diverse carbonaceous micro-structures from the ca 3400 Ma Strelley Pool Formation in the Pilbara Craton,Western Australia. Astrobiology 10, 899–920.

Sugitani, K., Mimura, K., Nagaoka, T., Lepot, K., Takeuchi, M., 2013. Microfossil as-semblage from the 3400 Ma Strelley pool formation in the Pilbara Craton, WesternAustralia: results from a new locality. Precambrian Res. 226, 59–74.

Tarhan, L.G., Droser, M.L., Gehling, J.G., 2015. Depositional and preservational en-vironments of the Ediacara Member, Rawnsley Quartzite (South Australia): assess-ment of paleoenvironmental proxies and the timing of ‘ferruginization. PrecambrianRes. 257, 124–136.

Tarr, W.A., 1933. The origin of the sand barites of the lower Permian of Oklahoma. Am.Mineral. 18, 260–272.

Taylor, G., Eggleton, R.A., 2001. Regolith Geology and Geomorphology. Wiley,Chichester (375 p).

Teitler, A., Philippot, P., Gérard, M., Le Hira, G., Fluteau, F., Ader, M., 2015. Ubiquitousoccurrence of basaltic-derived paleosols in the Late Archean Fortescue group,Western Australia. Precambrian Res. 267, 1–27.

Thorne, A.M., Trendall, A.F., 2001. Geology of the Fortescue Group, Pilbara Basin,Western Australia. Geol. Surv. West. Austral. Bull. 144, 1–249.

Thorpe, R.I., Hickman, A.H., Davis, D.W., Mortensen, J.K., Trendall, A.F., 1992.Constraints to models for Archean lead evolution from precise zircon U-Pb geo-chronology for the Marble Bar region, Pilbara Craton, Western Australia. In: Glover,J.E., Ho, S.E. (Eds.), The Archean: Terrains, Processes and Metallogeny. Univ. West.Austral. Geol. Extens. Publ. 22, pp. 395–407.

Ueno, Y., Ono, S., Rumble, D., Maruyama, S., 2008. Quadruple sulfur isotope analysis ofca. 3.5 Ga dresser formation: new evidence for microbial sulfate reduction in theearly Archean. Geochim. Cosmochim. Acta 72, 5675–5691.

Ueno, Y., Johnson, M.S., Danielache, S.O., Eskeberg, C., Pandey, A., Yoshida, N., 2009.Geological sulfur isotopes indicate elevated OCS in the Archean atmosphere, solvingfaint young sun paradox. U.S. Nat. Acad. Sci. Proc. 106, 14784–14789.

van den Boorn, S.H.-J.M., van Bergen, M.J., Vroon, P.Z., de Vries, S.T., Nijman, W., 2010.Silicon isotope and trace element constraints on the origin of ~3.5 Ga cherts: im-plications for early Archean marine environments. Geochim. Cosmochim. Acta 74,1077–1103.

van Kranendonk, M.J., 2000. Geology of the North Shaw 1:100,000 sheet. ExplanatoryNotes Geological Survey of Western Australia 1–86.

van Kranendonk, M.J., 2006. Volcanic degassing, hydrothermal circulation and theflourishing of early life on Earth: a review of the evidence from c. 3490–3420 Marocks of the Pilbara Supergroup, Pilbara Craton, Western Australia. Earth Sci. Rev.74, 197–240.

van Kranendonk, M.J., 2010. Geology of the 1:100,000 Coongan Sheet. Geol. Surv. West.Australia Explan. Notespp. 1–67.

van Kranendonk, M.J., Pirajno, F., 2004. Geochemistry of metabasalts and hydrothermalalteration zones associated with c. 3450 Ga chert and barite deposits: implications forthe geological setting of the Warrawoona group, Pilbara Craton, Australia. Geochem.Explor. Environmental Anal. 4, 253–278.

van Kranendonk, M.J., Hickman, A.H., Smithies, R.H., Williams, I.R., Bagas, L., Farrell,T.R., 2006. Revised lithostratigraphy of Archaean supracrustal and intrusive rocks inthe northern Pilbara Craton, Western Australia. Geol. Surv. West. Austral. Rec. 2006(/15), 1–55.

van Kranendonk, M.J., Philippot, P., Lepot, K., Bodorkos, S., Pirajno, F., 2008. Geologicalsetting of the Earth's oldest fossils in the ca. 3.5 Ga dresser formation, Pilbara Craton,Western Australia. Precambrian Res. 167, 93–124.

Vepraskas, M.J., Sprecher, S.W., 1997. Summary. In: Vepraskas, M.J., Sprecher, S.W.(Eds.), Aquic Conditions and Hydric Soils: The Problem Soils. Soil Sci. Soc. Amer.Spec. Publ. 150pp. 153–156.

Wacey, D., McLoughlin, N., Whitehouse, M.J., Kilburn, M.R., 2010. Two coexisting sulfurmetabolisms in a ca. 3400 Ma sandstone. Geology 38, 1115–1118.

Wacey, D., Kilburn, M.R., Saunders, M., Cliff, J., Brasier, M.D., 2011. Microfossils ofsulphur-metabolizing cells in 3.4-billion-year-old rock of Western Australia. Nat.Geosci. 4, 698–702.

Walker, J.C.G., 1987. Was the Archaean biosphere upside down? Nature 329, 710–712.Warren, J.K., 2006. Evaporites; Sediments, Resources and Hydrocarbons. Springer, Berlin

(1035 p).Westall, F., de Vries, S.T., Nijman, J.N., Marchesini, D., Severine, A., 2006. The 3.446 Ga

“Kittys Gap Chert”, an Early Archean Microbial Ecosystem. In: Reimold, W.U.,Gibson, R.L. (Eds.), Processes on the Early. Earth. Geol. Soc. Amer. Spec. Pap. 405pp.105–131.

Woolley, A.R., Kempe, D.R.C., 1989. Carbonatites: nomenclature, average chemicalcompositions and element distribution. In: Bell, K. (Ed.), Carbonatites. UnwinHyman, London, pp. 1–14.


18













































































































































































Wopfner, H., 1978. Silcretes of Northern South Australia and adjacent regions. In:Langford-Smith, T. (Ed.), Silcrete in Australia. Dept. Geogr. Univ., New England,Armidale, pp. 93–141.

Wordsworth, R., Pierrehumbert, R., 2013. Hydrogen-nitrogen greenhouse warming inEarth's early atmosphere. Science 339, 64–67.

Youssef, N.H., Couger, M.B., Elshahed, M.S., 2010. Fine-scale bacterial beta diversity

within a complex ecosystem (Zodletone Spring, OK, USA): the role of the rare bio-sphere. Plos One 5, e12414-1241.

Ziegenbalg, S.B., Brunner, B., Rouchy, J.M., Birgel, D., Pierred, C., Böttcher, M.E., Caruso,A., Immenhauser, A., Peckmann, J., 2010. Formation of secondary carbonates andnative sulfur in sulfate-rich Messinian strata, Sicily. Sediment. Geol. 227, 37–50.


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