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Potential photosynthetic impact on phosphate stromatolite formation after the Marinoan glaciation: Paleoceanographic implications
Accepted Manuscript Potential photosynthetic impact on phosphate stromatolite formation after the Marinoan glaciation: Paleoceanographic implications
Fumito Shiraishi, Saki Ohnishi, Yasutaka Hayasaka, Yusaku Hanzawa, Chizuru Takashima, Tomoyo Okumura, Akihiro Kano PII: DOI: Reference:
S0037-0738(18)30267-7 https://doi.org/10.1016/j.sedgeo.2018.11.014 SEDGEO 5420
To appear in:
Sedimentary Geology
Received date: Revised date: Accepted date:
25 September 2018 29 November 2018 30 November 2018
Please cite this article as: Fumito Shiraishi, Saki Ohnishi, Yasutaka Hayasaka, Yusaku Hanzawa, Chizuru Takashima, Tomoyo Okumura, Akihiro Kano , Potential photosynthetic impact on phosphate stromatolite formation after the Marinoan glaciation: Paleoceanographic implications. Sedgeo (2018), https://doi.org/10.1016/ j.sedgeo.2018.11.014
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ACCEPTED MANUSCRIPT
Potential photosynthetic impact on phosphate stromatolite formation after the Marinoan glaciation: Paleoceanographic
a,*,
Saki Ohnishi a, Yasutaka Hayasaka a, Yusaku Hanzawa a,
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Fumito Shiraishi
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implications
a
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Chizuru Takashima b, Tomoyo Okumura c, Akihiro Kano d
Department of Earth and Planetary Systems Science, Hiroshima University, 1-3-1
Faculty of Culture and Education, Saga University, 1 Honjo-Machi, Saga 840-8502,
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b
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Kagamiyama, Higashihiroshima, Hiroshima 739-8526, Japan.
Center for Advanced Marine Core Research, Kochi University, Kochi 783-8502,
Japan.
Department of Earth and Planetary Science, The University of Tokyo, 7-3-1 Hongo,
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d
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c
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Japan.
Bunkyo-ku, Tokyo 113-0033, Japan.
* Corresponding author. E-mail address: [email protected] (F. Shiraishi).
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Abstract
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This study investigated the origin and the depositional age of phosphate
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stromatolites contained in the Neoproterozoic Salitre Formation, Brazil. The
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stromatolites exhibited columnar shapes and were intercalated with laminated
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dolostones. Their depositional environment was interpreted to have been an evaporitic ramp where erosion and reworking by waves prevailed. Laser
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ablation-inductively coupled plasma-mass spectrometry measurements of the
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stromatolites yielded a U–Pb age of 616 ± 32 Ma, suggesting that they were
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deposited after the Marinoan glaciation. Several types of filamentous structures
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were recognized in the stromatolites; many of them were “pseudofossil” ambient inclusion trails, but some were possibly microfossils. To consider the general
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presence of photosynthetic microorganisms on the surface of shallow-water stromatolites,
the
involvement
of
microbial
photosynthesis
in
phosphate
stromatolite formation was expected. Numerical calculations determined that photosynthesis could have induced phosphate mineral precipitation at the time of deposition if the phosphorus concentration was above ca. 5 M. To achieve such high
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concentration at the shallow ocean, globally elevated phosphorus concentration was considered to be the ultimate factor in addition to the local process(es) including
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upwelling and/or evaporation. Therefore, the phosphate stromatolites in the Salitre
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Formation are evidence of the development of a phosphorus-rich ocean after the
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Marinoan glaciation.
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Keywords: Brazil, Ediacaran, phosphate stromatolite, phosphorus concentration,
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photosynthesis, U–Pb dating
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1. Introduction
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The late Neoproterozoic was an era of significant environmental change related with severe glaciations, which occurred at least two times: the Sturtian glaciation
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occurred ca. 720–700 Ma and the Marinoan glaciation occurred ca. 650–635 Ma (Fairchild and Kennedy, 2007; Hoffman and Li, 2009; Macdonald et al., 2010). Since the main source of phosphorus in the ocean is from continental weathering (Froelich et al., 1982; Föllmi, 1996), incensed chemical weathering after these severe glaciations resulted in unusually high [P] (brackets denote concentration) in the
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ocean (Planavsky et al., 2010). High [P] enhanced primary production, organic carbon burial, and consequent atmospheric oxygenation, which might have
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triggered the evolution and radiation of metazoans (Donnelly et al., 1990; Lenton
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and Watson, 2004; Papineau, 2010; Planavsky et al., 2010). Concomitantly, global
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phosphogenic events occurred through the Ediacaran and Cambrian periods (Cook
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and Shergold, 1984, 1986; Papineau, 2010; Drummond et al., 2015).
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In general, sedimentary phosphorite is considered to be formed by early diagenesis after deposition: microbial degradation of buried organic matter
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liberates PO43− in pore water, inducing precipitation of and/or replacement by
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phosphate minerals, which are sometimes further concentrated by physical
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processes such as winnowing and reworking (Sheldon, 1981; Froelich et al., 1982;
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Van Cappellen and Berner, 1991; Glenn et al., 1994; Jarvis et al., 1994; Krajewski et al., 1994; Föllmi, 1996; Filippelli, 2011). Another process potentially involved in the
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phosphorite formation is Fe-redox cycling (alternatively, the Fe-P shuttle): phosphorus adsorbed on Fe-oxides is released during early diagenesis due to reduction, and PO43− liberated into pore water causes phosphate mineral precipitation (Heggie et al., 1990; Krajewski et al., 1994; Glenn et al., 1994; Jarvis et al., 1994). Recently, there has been interest in examining the contribution of
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chemolithoautotrophic sulfide-oxidizing bacteria. They store polyphosphate under oxic conditions, but when exposed to anoxia and/or sulfide they hydrolyze
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polyphosphate to liberate PO43− into pore water causing precipitation of phosphate
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minerals (Schulz and Schulz, 2005; Goldhammer et al., 2010; Brock and
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Schulz-Vogt, 2011; Crosby and Bailey, 2012; Bailey et al., 2013; Crosby et al., 2014).
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Most Phanerozoic sedimentary phosphorites were formed in the outer shelf or
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upper slope where nutrient supply via upwelling leads to a high rate of organic matter burial (O’Brien and Veeh, 1980; Bremner and Rogers, 1990; Ruttenberg and
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Berner, 1993; Schwennicke et al., 2000). In contrast, Precambrian sedimentary
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phosphorites were mainly formed within very shallow epicontinental seas (Cook
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and Shergold, 1984; Krajewski et al., 1994; Drummond et al., 2015). These
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phosphorites sometimes exhibit unique stromatolitic forms (Chauhan, 1979; Southgate, 1980; Misi and Kyle, 1994; Bertrand-Sarfati et al., 1997; Shen et al.,
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2000), suggesting a depositional process different from that occurred in the Phanerozoic. In particular, the presence of phosphate stromatolites containing the occasional microfossil suggests that the benthic microbial community play a crucial role in phosphate mineral precipitation (Banerjee, 1971; Chauhan, 1979; Krajewski et al., 1994; Föllmi, 1996; Schwennicke et al., 2000; Papineau, 2010; Papineau et al.,
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2016); however, the causal relationship is still unclear. Late Neoproterozoic phosphate stromatolites are found in the Salitre Formation,
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in the upper part of the Una Group in São Francisco Craton, Brazil (Fig. 1). The
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Salitre Formation is composed mainly of carbonate and overlies glacial diamictites
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of the Bebedouro Formation. Based on research on Phanerozoic sedimentary
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phosphorites, previous studies have considered these phosphate stromatolites to
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have been formed by early diagenesis in an anoxic region of the deposits (Misi and Kyle, 1994; Caird et al., 2017). However, the observations in this study are
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incompatible with that interpretation as discussed below, suggesting that the
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depositional process needs to be reinterpreted.
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In addition, the depositional age of the Salitre Formation is still controversial; it
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is unclear whether it is formed after the Sturtian or Marinoan glaciations. An U–Pb age of ca. 875 Ma was obtained from detrital zircon in the underlying Bebedouro
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Formation (Babinski et al., 2007), but this did not provide a constraint for the controversy above. Although a Pb–Pb age of 514 ± 33 Ma was obtained from stromatolitic carbonates of the Salitre Formation, it recorded a fluid percolation event (Trindade et al., 2004). Using Sr isotope stratigraphy, Misi and Veizer (1998) suggested the depositional age of the Salitre Formation to 670–600 Ma. However,
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reliable age data is still lacking. The depositional age of the Bambuí Group (Fig. 1A) correlative to the Una Group
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in the São Francisco Craton is also controversial. Previously, the Bambuí Group was
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considered to be post-Sturtian deposits, because a U–Pb age of 880 Ma was obtained
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from detrital zircon in the Jequitai Formation (equivalent to the Bebedouro
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Formation), and a Pb–Pb age of 740 ± 22 Ma was obtained from basal limestone of
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the Sete Lagoas Formation (equivalent to the Salitre Formation) (Babinski et al., 2007; Vieira et al., 2007; Misi et al., 2007, 2011; Sial et al., 2010). However, a U–Pb
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age of ca. 610 Ma was reported for detrital zircon from the middle section of the Sete
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Lagoas Formation, which necessitated the reconsideration of depositional age
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(Caxito et al., 2012). Caxito et al. (2012) compiled litho- and isotope stratigraphy
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and proposed that the Sete Lagoas Formation was deposited after the Marinoan glaciation (ca. 635–610 Ma). By following this compilation, recent studies
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considered that the correlative Salitre Formation could have formed after the Marinoan glaciation (Caird et al., 2017; and references therein). However, Paula-Santos et al. (2015) argued that the Sete Lagoas Formation was much younger and did not relate to either the Sturtian or the Marinoan glaciation. Thus, the age of Neoproterozoic successions in the São Francisco Craton are not
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well constrained, mainly due to an extensive fluid percolation event that hinders reliable isotopic age determination (Babinski et al., 1999, 2007; D’Agrella-Filho et
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al., 2000; Trindade et al., 2004; Guimarães et al., 2011). However, it is possible to
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age date using apatite from phosphate stromatolites, because apatite is much more
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resistant to post-depositional alteration than calcite (the closure temperature of the
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apatite U–Pb system is ca. 500–600°C) and has provided reliable ages of geological
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material (Sano and Terada, 1999; Barfod et al., 2002). This study therefore applied laser ablation-inductively coupled plasma-mass spectrometry (LA-ICP-MS) to
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directly determine the U–Pb age of the phosphate stromatolites.
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This study aims to elucidate the origin and depositional age of phosphate
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stromatolites in the Salitre Formation, the results of which will provide insight into
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understanding paleoceanography after the Neoproterozoic severe glaciation.
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2. Geological settings
2.1. Neoproterozoic successions of the São Francisco Craton
The major Neoproterozoic successions of the São Francisco Craton are the Una
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Group and the correlative Bambuí Group (Fig. 1A). These groups are considered to have been deposited at a shallow-marine epeiric ramp (Misi and Kyle, 1994; Misi
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and Veizer, 1998; Sial et al., 2010; Misi et al., 2011; Drummond et al., 2015; Caird et
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al., 2017). The Una Group is distributed mainly at the Irecê Basin, and
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unconformably overlies either Paleoproterozoic basement rocks or sandstones of the
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Mesoproterozoic Chapada Diamantina Group. The Una Group consists of glacial
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diamictites of the Bebedouro Formation and overlying limestones, dolostones, and siliciclastics of the Salitre Formation (Fig. 1B; Misi et al., 2011; Guimarães et al.,
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2011; Caxito et al., 2012). Caird et al. (2017) subdivided the Salitre Formation into
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Nova America, Gabriel, Jussara, and Irecê members in ascending order.
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Phosphorites are restricted to the Nova America member (Misi and Kyle, 1994; Sial
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et al., 2010; Caxito et al., 2012). Comparative successions of Nova America and Gabriel members are lacking in the Bambuí Group (Caird et al., 2017).
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Phosphate stromatolites of the Salitre Formation exhibit several forms, among which the most distinctive type is columnar one previously assigned as Jurussania
krilov (Misi and Kyle, 1994; Kyle and Misi, 1997; and references therein). Phosphate stromatolites are present in beds with thicknesses ranging from centimeters to meters and are interbedded with dolostones (Misi and Kyle, 1994).
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Phosphate stromatolites occasionally have cross-bedding and current-ripple bedding which suggests they are deposited in a subtidal to lower intertidal
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environment (Misi and Kyle, 1994).
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Some areas, including Três Irmãs (Fig. 1B), contain a Zn-Pb-Ag sulfide-rich zone
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at the stratigraphic position slightly above the phosphate stromatolites. This zone
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has been interpreted to have been formed by the extensive percolation of
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metal-bearing fluids within sediments during the late stage (Kyle and Misi, 1997).
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2.2. Investigated outcrop
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The outcrop investigated for this study is located ca. 5 km east of Irecê (Fig. 1B)
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and composed of alternating of phosphorites and dolostones (Fig. 1C). While the strata are relatively undeformed at the northern part of the outcrop, slump folds are
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prominent at the southern part. Dolostones are more significantly distorted by slump folding than the phosphorites. Columnar phosphate stromatolites of ca. 2 cm diameter are crowded in the phosphorites and sometimes bifurcate upward (Fig. 2A, B). The matrix between the columnar stromatolites is rich in dolomite. Cross sections of columnar stromatolites are usually circular (Fig. 2C), but occasionally
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exhibit elongation (Fig. 2D). Dolostones, on the other hand, are composed of fine dolomite crystals (Fig. 2E), and sometimes deformed by the slump folding (Fig. 2F).
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Boundaries between phosphorites and dolostones are transitional at the relatively
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undeformed northern part (Fig. 2G), but abrupt at the southern part where slump
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undeformed northern part (Fig. 1C; S2-1 to S2-11).
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folds are prominent (Fig. 2H). This study collected 11 samples from the relatively
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3.1. Petrographic analyses
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3. Methods
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Polished thin sections were observed under a polarized microscope (Eclipse LV100 POL, Nikon) and a microscope (Microphoto-FXA, Nikon) equipped with a
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cathodoluminescence (CL) system (ELM-3R, Premier American Technologies). For comparing with microfossils that was found (see Section 4.2), a modern cyanobacterial culture (Scytonema sp.; National Institute for Environmental Studies, Japan) was also observed. Thin sections of S2-2 and S2-7 were coated with carbon and used for the
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elemental mapping of P, Ca, Mg, F, Fe, and Si using an electron-probe microanalyzer (EPMA; JXA-8200, JEOL). For S2-7, EPMA analyses were conducted
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after LA-ICP-MS analyses described below. Each elemental map was output as a
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grayscale image and used to generate RGB composites in Adobe Photoshop (CS6,
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Adobe Systems).
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3.2. Bulk analyses of mineralogy, constituent elements, and stable isotopes
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Powdered samples of phosphorites and dolostones were prepared. For phosphorite,
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stromatolite and matrix parts were powdered separately (suffixes of s and m were
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added to the names of stromatolite and matrix samples, respectively).
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Aliquots from each powdered sample were used for analyzing mineralogy with a powder X-ray diffractometer with Cu K radiation (40 kV, 40 mA) and a graphite
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monochromator (MultiFlex, Rigaku). Mineral percentages were evaluated using the RIR (Reference Intensity Ratio) method of the ICDD qualitative analysis software. Another aliquot of each powdered sample was used to analyze constituent elements. About 100 mg of powder was collected in a centrifugal tube and dissolved by dropping 1 mL of 1M HCl followed by adding 0.25 mL of 35% HCl. After
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centrifugation, the supernatant was diluted with 2% HNO3, and Ca, P, Mg, Al, Si, Na, S, Sr, Ba, K, Fe, Mn, and Zn contents were measured using an inductively
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coupled plasma optical emission spectroscope (ICP-OES; iCAP6300, Thermo Fisher
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Scientific). The resulting data was used to calculate the elemental content of
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carbonate and phosphate minerals.
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A third aliquot was used for analyzing oxygen and carbon stable isotopes
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following the method of Furuyama et al. (2017). About 0.3 mg of powder was put in an airtight glass vial. After replacing the air in the vial with purified He gas, CO2
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was generated by reacting a drop (up to 0.02 mL) of anhydrous phosphoric acid at
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60°C for more than 6 h. The CO2 was separated from other gases using a gas
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preparation system (GasBench II, Thermo Fisher Scientific), and the oxygen and
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carbon isotopic compositions were measured using a mass spectrometer (Delta Plus, Thermo Fisher Scientific). The isotopic values were reported relative to the Vienna
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Peedee Belemnite (VPDB) standard, and the standard deviations of repeated measurements were better than 0.2‰. Oxygen isotopic values were also reported relative to the Vienna Standard Mean Ocean Water (VSMOW) standard for a reference. Correlation coefficients between each elemental content and isotopic value were
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calculated by the CORREL function of Excel (Microsoft) for evaluating how strongly
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two variables were related to each other.
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3.3. U–Pb dating of apatite
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U–Pb isotopes were measured following the methods of Katsube et al. (2012).
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Polished thin sections of four phosphorite samples (S2-4, S2-7, S2-9, and S2-11) were ultrasonicated in Milli-Q water for ca. 10 min and air-dried before
(UP-213,
New
Wave
Research)
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Laser
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measurement. The LA-ICP-MS systems used for this study were a 213 nm Nd-YAG and
an
ICP-MS
(7500,
Agilent).
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Measurements were performed using a mixed He–N2–Ar carrier gas system
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equipped with a small volume ablation cell, sample aerosol stabilizer (buffering chamber), and charcoal filter attachment. The sample surface was first
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decontaminated by a pre-abrasion with a laser spot diameter of 65 m, followed by the measurement with a spot diameter of 55 µm. Measurement areas were selected from the margin of the columnar stromatolites where the phosphate minerals were relatively pure (see below). 4–7 areas were selected from each sample, and 4–8 spots were measured in each area. Every time after measuring 10 spots of samples, NIST
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SRM 610 glass standard was measured for calibrating
206Pb/238U
and
207Pb/206Pb
ratios. The data reduction program Pepi-AGE (Dunkl et al., 2008) was used to
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process the raw data, including common Pb correction. Final plots were generated
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using Isoplot/Ex (Version 3.23; Ludwig, 2003). The isotopic ratios and ages were
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quoted at 2σ, and the weighted mean was given with a 95% confidence interval.
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The calibration using NIST SRM 610 glass standard described above was
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previously applied for determining U–Pb age of zircons by LA-ICP-MS, and demonstrated that the matrix effect was smaller than the analytical uncertainties
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(Orihashi et al., 2008). In addition, least ablation energies of NIST SRM 610 glass
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standard and apatite minerals are comparable, which is advantageous for unifying
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the fractionations of U and Pb during a deepening of ablation pit (e.g., Hirata and
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Nesbitt, 1995; Eggins et al., 1998). To evaluate the validity of the measured ages, the apatite crystals within Duluth Complex gabbro FC1 (Paces and Millar, 1993)
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were measured.
3.4. Evaluation of photosynthetic influence on francolite and calcite precipitation
As discussed later in this study, photosynthetic microorganisms were potentially
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involved in the formation of the phosphate stromatolites in the Salitre Formation. Therefore, it was necessary to determine the influence of photosynthesis on
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phosphate mineral precipitation; this was evaluated using numerical calculations.
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In general, marine sedimentary phosphate precipitates as francolite, a carbonate
Föllmi,
1996),
with
an
average
stoichiometry
of
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1994;
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fluorapatite mineral (Jahnke, 1984; Van Cappellen and Berner, 1991; Jarvis et al.,
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Ca9.54Na0.33Mg0.13(PO4)4.8(CO3)1.2F2.48 (Jahnke et al., 1983; Ruttenberg and Berner, 1993; Schenau et al., 2000). Therefore, the saturation state of francolite (fra) is
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defined as
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fra = {Ca2+}9.54{Na+}0.33{Mg2+}0.13{PO43−}4.8{CO32−}1.2{F−}2.48 / KSP_fra, (1)
where braces denote the activity of each ion and KSP_fra represents the solubility
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product of francolite. KSP_fra depends on {CO32−}, which was estimated using the methods of Jahnke et al. (1983). Photosynthetic influence on francolite precipitation was evaluated by calculations like those of Shiraishi (2012) who evaluated the photosynthetic influence on CaCO3 mineral precipitation: 200 mol L−1 of CO2 was removed from the given water using PHREEQC, a computer program for
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geochemical calculations, and the saturation state of francolite was compared before
fra = fra_aft − fra_bef
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(fra_bef) and after (fra_aft) CO2 removal as follows:
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(2)
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This value corresponds to a change expected at the sediment-water interface.
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Similar calculations were performed for calcite precipitation (cal_bef, cal_aft, and cal).
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The seawater chemistry used for the calculations was based on a major ion
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composition estimated from a mid ocean ridge brine and river water mixing model
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(Hardie, 2003). This study used the composition of 600 Ma ([Ca2+] = 12.7 mM, [Mg2+]
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= 48.6 mM, [K+] = 13.1 mM, [Na+] = 466.2 mM, and [SO42−] = 25.8 mM; Hardie, 2003), which was close to the depositional age of the Salitre Formation (616 ± 32
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Ma) as discussed below. This is almost comparable to the composition of modern standard seawater ([Ca2+] = 10.28 mM, [Mg2+] = 52.82 mM, [K+] = 10.21 mM, [Na+] = 469.06 mM, and [SO42−] = 28.24 mM; Zeebe and Wolf-Gladrow, 2001). The value used for water temperature was 25 °C. Modern standard seawater values were used for [Cl−] and [F−] ([Cl−] = 545.86 mM and [F−] = 0.07 mM; Zeebe and Wolf-Gladrow,
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2001). [P] is depleted in the modern surface ocean and is ca. 0.5–2.0 M even at the upwelling areas (Garcia et al., 2010). However, [P] in the aftermath of
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Neoproterozoic “Snowball Earth” glaciations was estimated to be 5 to 10 times
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used a value for [P] ranging from 0.1 to 20 M.
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higher than that of Phanerozoic ocean (Planavsky et al., 2010). Therefore, this study
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The ranges of pH and [DIC] (dissolved inorganic carbon) in the Neoproterozoic
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ocean are not well constrained, although they strongly affect the composition of phosphate chemical species (see below) and the magnitude of photosynthetic
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influence on water chemistry (Arp et al., 2001; Shiraishi, 2012). Indeed, the results
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based on model calculations are still controversial: Bartley and Kah (2004)
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considered that the Neoproterozoic [DIC] may have been approaching the modern
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level (ca. 2 mM), while Bristow and Kennedy (2008) considered that [DIC] in the Ediacaran ocean was higher than that in the modern ocean. However, chemical
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proxies can provide rough ranges for pH and [DIC]. Estimation based on boron isotopes suggested that the pH value was ca. 7.0–8.5 after the Marinoan glaciation (Kasemann et al., 2010). In addition, estimation based on triple oxygen isotope suggested that pCO2 (equilibrium CO2 partial pressure) values were ca. 6,000– 26,000 ppm in the immediate aftermath of Marinoan glaciation and ca. 2,000–7,000
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ppm around the Neoproterozoic/Cambrian boundary (Bao et al., 2008). To consider the depositional age of the Salitre Formation (616 ± 32 Ma) discussed below, a pCO2
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value of 2,000–26,000 ppm (i.e., 10−2.7–10−1.6 atm) is the maximum estimate for the
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Salitre Formation. These pH and pCO2 values constrained the [DIC] range to ca.
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0.56–177.83 mM (see Results), which was used for the calculations in this study.
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below had the following three constraints:
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In addition, the depositional characteristics of phosphate stromatolites described
1) The presence of primary phosphate stromatolites indicated a lack of substantial
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francolite dissolution, so fra_bef > 1.
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2) The presence of redeposited dolomite within the phosphate stromatolites
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dolomite) >1.
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indicated a lack of substantial dolomite dissolution, so dol (saturation state of
3) Scarcity of primary calcite within phosphate stromatolites indicated the lack of
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substantial calcite precipitation even inorganically, so cal_bef < 20 (≈101.3) (Shiraishi, 2012). However, phosphorus inhibits carbonate precipitation: for example, in the case where [P] is 10 M, ca. 0.4–0.6 higher calcite saturation in the logarithmic scale is required to achieving the precipitation rate comparable to the case without phosphorus (Burton and Walter, 1990). Therefore, this study,
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employed cal_bef < 80 (≈10(1.3
+ 0.6))
for the condition precluding substantial
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inorganic calcite precipitation.
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4. Results
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4.1. Petrographic characteristics
Dolostones exhibited parallel laminations ca. 1–15 mm thick, which represented
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repeating light-gray and dark-gray laminas (Fig. 3A). They were composed mainly
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of fine dolomite crystals, and the lamination represented different crystal sizes:
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light-gray was ca. 50 m and dark-gray was ca. 10 m (Fig. 3B–C). Although there
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were some rhombic euhedral crystals (Planar-e), many dolomite crystals showed irregular abraded shapes (Bone et al., 1992) rather than anhedral (also see below).
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Dolomite crystals showed similar CL characteristics regardless of their size and shape: the core part of each crystal exhibited dull luminescence without distinct zonation, while the limpid rim part (ca. 5 m thick), representing syntaxial overgrowth, exhibited brighter luminescence (Fig. 3C). Light-gray laminas sometimes contained angular to rounded phosphate intraclasts of ca. 1 mm
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diameter, with the long axis parallel to the bedding plane (Fig. 3D, F). Interstitial spaces of laminas were sometimes infilled by calcite cement exhibiting bright
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luminescence with zonation (Fig. 3C, D). Calcite cement rarely surrounded the
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Siliciclastics and iron minerals were scarce (Fig. 3E).
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dolomite (Fig. 3B, D), which might represent a rosette (Papineau et al., 2016).
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Phosphorites were composed of stromatolite part and infilling matrix part, and
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boundaries between them were sharp. However, some laminas were traceable across the boundary (Fig. 4A), suggesting that the protrusion of stromatolites was
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less than ca. 1 cm at the time of deposition. Phosphate stromatolites mostly
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exhibited columnar shape, while laminar stromatolites sometimes bridged the
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columns (Fig. 4B). Margins of the columnar stromatolites (ca. 1 mm) were relatively
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pure phosphate, while the inner part contained dolomite laminas of less than ca. 500 m thick as well as isolated fine dolomite crystals (Fig. 4B, C, G). The matrix
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part was dominated by rounded and sorted dolomite intraclasts of ca. 100–200 m in diameter, around which isopachous phosphate cortices were common (Fig. 4C, D). Dolomite intraclasts sometimes composed stromatolite laminas. Phosphate intraclasts of less than ca. 1 cm diameter were frequently contained in the matrix part (Fig. 4A). Although the appearance of dolomite in the phosphorites was
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different from that of dolostones, characteristics of individual dolomite crystals were similar: dull luminescence in the core and brighter luminescence in the rim,
T
and sometimes represented an abraded shape (Fig. 4D, G). The rim part exhibiting
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brighter luminescence was absent for the dolomite crystals directly contacting to
AN
US
they were infilled by calcite containing Fe and Si.
CR
the phosphate minerals (Fig. 4G). Cracks were common in the phosphorites, and
M
4.2. Filamentous structures
ED
Several types of filamentous structures were recognized in the investigated
PT
phosphate stromatolites (Fig. 5A). A major type was the structure known as
CE
ambient inclusion trails (AITs; Wacey et al., 2008a; She et al., 2016). AITs exhibited straight, curved, and spiral tubes of ca. 1–25 m diameter, and occasionally had
AC
terminal Fe-oxide minerals exhibiting equivalent diameter of the tubes (Fig. 5B, D). AITs exhibited polygonal cross sections, continuous longitudinal striations on the walls, calcite spar infillings (Fig. 5A, C), and ”starburst pattern” (Fig. 5D). Filamentous structures rarely exhibited septation-like structure (Fig. 5E) and orange-colored filament resembling iron-oxidizing bacteria (Fig. 5F), although they
22
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had terminal Fe-oxide minerals like AITs. The “starburst pattern” composed of micritic tubes of ca. 1 m diameter (Fig. 5G) apparently lacked terminal mineral
T
and arranged along the cracks as well as the border between phosphate and
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dolomite, implying the possibility of endolithic contamination e.g. by fungi in
CR
younger ages (Wacey et al., 2008a).
US
Microborings formed by ancient endolithic microorganisms were also rarely
AN
recognized, which appeared as micritic filaments of ca. 1 m diameter extending inward from the surface of isopachous phosphate cortex (Fig. 5H). Similar
M
microborings were reported from other phosphorites (Krajewski et al., 1994; and
ED
references therein). Isopachous phosphate cortex lacked microfossils except for
PT
microborings, which was different from the situation of granular phosphorites in
CE
the Neoproterozoic Doushantuo Formation (She et al., 2014). In addition to the diagenetic/destructive structures above, two possible
AC
microfossils were recognized from phosphate stromatolites. The more frequently observed structure was wavy filaments of ca. 8–12 m diameter. Although such features were also common to AITs, these were distinguished from AITs by the absence of terminal minerals and longitudinal striations, circular cross section with relatively uniform diameter, relatively uniform sinuous wave lengths (ca. 40–60 m),
23
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and micrite/phosphate infilling that resembled a trichome with EPS (extracellular polymeric substances) sheath (Fig. 5I, K). Indeed, the “sheath” was sometimes torn
T
(Fig. 5I, red arrow), twisted (Fig. 5K, green arrow), and showed a hormogonia-like
IP
appearance (Fig. 5I, blue arrow) that was similar to modern cyanobacteria (Fig. 5J).
CR
Less abundant microfossils were straight filaments of ca. 3 m diameter with
US
external microcrystalline carbonates that formed microbial mat-like structures
AN
within the laminas of phosphate stromatolites (Fig. 5L).
ED
M
4.3. Characteristics of bulk mineralogy, constituent elements, and stable isotopes
PT
XRD analysis indicated that the major constituent minerals of phosphorites and
CE
dolostones were fluorapatite and dolomite, respectively (Fig. 6). Fluorapatite should be practically described as francolite because of the ion substitutions described
AC
below, although the diffraction data for francolite is not available in the database used. Phosphorites also contained a certain amount of dolomite and the dolomite percentage was found to be higher in the matrix part. Minor mineral components were calcite (less than ca. 30%) and clay minerals (illite and montmorillonite; less than ca. 5%).
24
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Results of constituent element and stable isotope analyses are presented in Table 1 and Fig. 7 (results of analysis for Si and Zn content as well as those for 13C and
T
18O are also presented in Fig. 6). Based mainly on correlation coefficients (Fig. 7A),
IP
the measured elements and isotopes were categorized into three groups.
CR
1) P, Na, Sr, Ba, and S: P is the constituent element of francolite, and other
US
elements tend to be substituted by francolite (Jarvis et al., 1994).
AN
2) Mg, Ca, 18O, and 13C: Ca and Mg are the constituent elements of dolomite. Although Ca is also contained in francolite, its content is higher in dolomite (ca.
M
13% for dolomite, compared to ca. 3% for francolite). Positive correlation
ED
between Mg content, 18O, and 13C (Fig. 7A) largely reflects the isotopic
PT
fractionation related to dolomite as discussed below (Fig. 7B).
CE
3) Si, Al, K, Zn, Fe, and Mn: These elements are mainly associated with clay minerals, diagenesis, and metal-bearing fluids, although K and Si may partially
AC
be substituted in francolite (Jarvis et al., 1994). Si and Zn content (and Al and K content) are found to be elevated around S2-9 (Fig. 6), suggesting the influence of metal-bearing fluids that form Zn-Pb-Ag sulfide. The positive correlation between Fe and Mn content suggests their incorporation under reductive condition. Especially, dolomite-rich parts (the dolostones and the matrix part of
25
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phosphorites) exhibit higher Fe and Mn contents (Fig. 7C), supporting the
T
formation of dolomite under somewhat reductive condition (see Section 5.5).
CR
IP
4.4. U–Pb age of the phosphate stromatolites
US
The results of LA-ICP-MS measurements are presented in Fig. 8, and their
AN
background data are listed in Appendix A and B. Data for S2-4, S2-7, and S2-11 exhibited a well constrained linear array in the Terra–Wasserburg Concordia plot
M
(Fig. 8A, B, D), indicating variable contribution of common Pb and radiogenic Pb. In
ED
contrast, the data for S2-9 was concentrated in a relatively small domain (Fig. 8C),
PT
possibly due to the metal-bearing fluids percolation in this horizon as described
CE
above. This sample was considered an anomalous outlier and excluded from later discussion. For the remaining samples, some areas showed a large dispersion due to
AC
impurities and/or narrow variation in common Pb/radiogenic Pb ratio. Therefore, the areas with error less than 50% relative to the mean age values were accepted (areas 4 and 7 for S2-4; areas 1, 2, 4, and 6 for S2-7; areas 2, 4, 6, and 7 for S2-11), and their weighted mean age was 616 ± 32 Ma (Fig. 8E). This age supported the hypothesis that the phosphate stromatolites in the Salitre Formation were formed
26
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after the Marinoan glaciation. The measured age of the apatite crystals in the FC1 standard (1,045 ± 15 Ma;
T
Appendix A) was comparable to the reported values for apatite (1,150 ± 59 Ma;
IP
Norman and Nemchin, 2012) and zircon (1,099.0 ± 0.1 Ma; Paces and Millar, 1993).
CR
Taking into consideration the lower closure temperature of apatite, the result of this
US
study was more conformable to the age of zircon. In any case, this result supported
AN
the validity of the age obtained above and the application of the glass standard
M
calibration for apatite minerals.
PT
ED
4.5. Results of numerical calculations
CE
Table 2 shows the photosynthetic influence on the saturation states of francolite and calcite, calculated at pH 7.5 with [DIC] of 2, 5, and 10 mM. The saturation state
AC
of francolite was significantly increased by photosynthetic CO2 removal when [P] was higher than ca. 3–5 M. However, the photosynthetic influence was buffered at higher [DIC] levels: the increase in francolite saturation state was ca. 100-fold when [DIC] = 2 mM, while it was suppressed to ca. 4-fold at [DIC] = 10 mM. The saturation state of calcite was also increased by photosynthesis; however, its degree
27
ACCEPTED MANUSCRIPT
was much smaller than that of francolite (ca. 4-fold at [DIC] = 2 mM while ca. 2-fold at [DIC] = 10 mM).
T
Calculation results at wider ranges of pH and [DIC] are shown in Fig. 9.
IP
Phosphorus-independent constraints (i.e., pCO2, dol, and cal_bef) assumed for the
CR
Salitre Formation further restricted the ranges into a rhomboidal area of pH = ca.
US
7.0–8.4 and [DIC] = ca. 1–50 mM. This rhomboidal area was overlapped by the
AN
range restriction from phosphorus-dependent constraint (i.e., fra_bef) when [P] was above ca. 3 M (the areas without hatches in Fig. 9). On the other hand, the area of
M
high fra value was enlarged as [P] increased. The increase of fra value was
ED
suppressed at the condition of low pH and high [DIC] due to CO2 buffering
PT
(Shiraishi, 2012). In addition, fra value at low [DIC] was significantly high
CE
because the low DIC effect (Shiraishi, 2012) does not affect the phosphate chemical species, unlike the case of CaCO3 precipitation. Within the range satisfying all
AC
constraints, high fra values were achieved when [P] was higher than ca. 5 M. Similar result was also obtained for fra_aft.
5. Discussion
28
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5.1. Interpretation of filamentous structures
T
Some coccoid and filamentous microfossils interpreted as cyanobacteria have
IP
been reported from ancient phosphorites, the latter of which mostly exhibited
CR
unbranching tubular structures ca. 2–30 m diameter (Chauhan, 1979; Soudry and
US
Champetier, 1983; Krajewski et al., 1994; Zhang et al., 1998; Schwennicke et al.,
AN
2000; She et al., 2014). Such morphological characteristics also apply to the filamentous structures found in the phosphate stromatolites investigated in this
M
study; however, many of the structures found are “pseudofossil” AITs (as described
ED
earlier in this article). AITs have been also reported from other Neoproterozoic
PT
phosphorites (Xiao and Knoll, 1999; Wacey et al., 2008a; She et al., 2013, 2016). The
CE
formation process of AITs is generally as follows: 1) organic matter attached to metal-rich grains (typically pyrite) undergo decomposition during compaction and
AC
heating in the shallow burial stage, 2) the metal-rich grains are propelled by pressure solution due to the generated gas pressure, and 3) pyrites are sometimes changed to hematite pseudomorphs (Knoll and Barghoorn, 1974; Wacey et al., 2008a,b). AITs found in this study possessed reddish Fe-oxide grains as terminal minerals (e.g., Fig. 5B, D), probably representing hematite pseudomorphs
29
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generated by the oxidation of pyrite grains like the case of Ediacaran Doushantuo Formation (She et al., 2016). Therefore, the presence of AITs suggests the
T
development of an anoxic environment and a Fe/S redox cycle during diagenesis.
IP
Nonetheless, the contribution of such diagenetic processes on phosphate
US
constituents in the investigated outcrop (Table 1).
CR
stromatolite formation is considered to be minor because Fe and S are not the major
AN
On the other hand, two filamentous structures exhibiting characteristics similar to cyanobacteria are likely to be microfossils (Fig. 5I, K, L). Indeed, their
M
morphological features (waved filament of ca. 8–12 m wide with “sheaths” and
ED
“hormogonia”; Fig. 5I, K) resemble those of modern cyanobacteria including
PT
Phormidium (straight, coiled, or waved trichome of ca. 2–12 m wide with sheaths
CE
and hormogonia; Komárek and Hauer, 2013) and Arthrospira (open helix trichome of ca. 2.5–16 m wide with sheaths and hormogonia; Sili et al., 2012). In addition,
AC
thinner filaments with microcrystalline carbonates (Fig. 5L) resemble the encrusted sheath of cyanobacteria that is formed by high nucleation rates on acidic EPS sheaths (Shiraishi et al., 2017). If this is the case, this microfossil type represents one of rare Precambrian calcified cyanobacteria. Another potential microfossil is the filamentous structure with “septation” (Fig.
30
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5E). This resembles the sulfide-oxidizing bacteria related to phosphogenesis, which is characterized by septate and unbranching filaments with variable diameters
T
(ranging from 1–200 m; Bailey et al., 2013). However, this filamentous structure
IP
appears to have a terminal mineral grain like AITs, and the “septation” may
CR
represent traces of stepwise movement reported from AITs (Wacey et al., 2008a,b).
US
Orange-colored filament (Fig. 5F) may also represent a microfossil because it
AN
resembles possible filamentous iron-oxidizing bacteria having terminal knobs (Little et al., 2004). Microfossils of filamentous iron-oxidizing bacteria are also
M
reported from phosphate stromatolites of Paleoproterozoic Aravalli Group (Crosby
ED
et al., 2014), which may be a common feature of phosphate stromatolites. However,
PT
the orange-colored filament of the Salitre Formation lacks stalk-like structure, and
CE
similar orange-colored tubes are also formed by AITs (She et al., 2016). Thus, some of filamentous structures in the Salitre Formation may represent
AC
microfossils of cyanobacteria, sulfide-oxidizing bacteria, and iron-oxidizing bacteria; however, definitive evidence is lacking and detail investigation is necessary in future studies.
5.2. Interpretation of stable isotopes
31
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18O and 13C values of the investigated outcrop showed a positive correlation
T
(Fig. 7A, B). Such a trend in carbonate rocks is generally interpreted as the
IP
influence of diagenesis, along which both 18O and 13C values decrease (Banner
CR
and Hanson, 1990; Jacobson and Kaufman, 1999; Knauth and Kennedy, 2009). A
US
similar trend is reported for francolite alteration (McArthur et al., 1986; Jarvis et
AN
al., 1994). However, the Mn/Sr ratios for all samples were much lower than 2 (Table 1), suggesting the absence of substantial diagenetic influence at least for carbonate
M
rocks (Jacobson and Kaufman, 1999). In addition, Na contents were negatively
ED
correlated with 18O and 13C values (Fig. 7A), which was the opposite of the
PT
general trend of francolite diagenesis (McArthur, 1985). The positive correlation
CE
between 18O and 13C values is, instead, considered to reflect the constituent minerals: francolite and dolomite have relatively low and high isotopic values,
AC
respectively. Indeed,18O and 13C values positively correlated with the content of dolomite-related elements (i.e., Mg), while they negatively correlated with the content of francolite-related elements (i.e., P, Na, Sr, Ba, and S; Fig. 7A). Francolite contains oxygen as PO4 and CO3. The phosphoric acid treatment applied in this study liberates only the oxygen of CO3 (Kolodny and Kaplan, 1970;
32
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Shemesh et al., 1983), which has a 18O value similar to that of the coexisting calcite (Shemesh et al., 1988). The 18O value of dolomite, on the other hand, is ca. 3‰
T
heavier than that of coexisting calcite (Land, 1980; Vasconcelos et al., 2005). Indeed,
IP
18O values of dolostones in this study were heavier than that of the stromatolite
CR
part of the phosphorites, and their maximum difference was ca. 3‰. This suggests
US
that the observed difference in 18O values primarily reflects the difference in
AN
constituent minerals, and that both francolite and dolomite are precipitated from water having similar oxygen isotope composition (ca. −6‰). Such light 18O value is
M
the general trend of rocks formed after the Marinoan glaciation (Shields and Veizer,
ED
2002).
PT
For 13C, the isotopic fractionation during francolite precipitation is expected to
CE
be minimal, like in CaCO3 mineral precipitation (Jarvis et al., 1994). The 13C value of dolomite, on the other hand, has a ca. 2.4‰ heavier value than that of the
AC
coexisting calcite at 25°C (Horita, 2014). Indeed, the maximum difference of 13C values between the dolostones and the stromatolite part of the phosphorites was close to 2.4‰. This suggests that the difference in 13C values also primarily reflects the difference in constituent minerals, and that both francolite and dolomite are precipitated from water having similar carbon isotopic compositions (ca. −3‰). Such
33
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light 13C value is the general trend of rocks formed after the Marinoan glaciation (Halverson et al., 2010).
T
From the discussion above, it is concluded that the positive correlation between
IP
18O and 13C values in the investigated outcrop primarily reflects the isotopic
CR
fractionation related to dolomite. Indeed, the data fit this trend well (Fig. 7B).
US
Similar explanation may be applicable to the phosphate stromatolites of the
AN
Paleoproterozoic Aravalli Group, although the maximum differences between the stromatolitic column and intercolumn are somewhat larger than the expectation
ED
M
(4.7‰ for 18O and 4.0‰ for13C; Papineau et al., 2016).
CE
PT
5.3. Origin of phosphate: endogenous factors
Previous studies interpreted the origin of phosphate stromatolites in the Salitre
AC
Formation by relating them to early diagenesis in an anoxic environment. For example, Misi and Kyle (1994) considered that bacterial degradation of organic matter in the anoxic environment of the deposit caused localized phosphate enrichment in the pore water, which resulted in the concentration of carbonate fluorapatite by direct precipitation or replacement of calcium carbonate to form
34
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phosphate stromatolites. A similar interpretation was employed by Caird et al. (2017); they additionally suggested the possible involvement of a Fe-redox cycle and
T
chemosynthetic sulfur bacteria. In addition, Drummond et al. (2015) investigated
IP
the phosphatic siltstones of the Sete Lagoas Formation and interpreted them as
CR
follows: aeolian Fe-oxyhydroxides and clays delivered phosphorus to promote
US
cyanobacterial mat development, during which the Fe-redox cycle and activity of
AN
sulfide-oxidizing bacteria at the redox boundary caused phosphogenesis. Such diagenetic phosphogenesis model has been widely applied to ancient phosphorites
M
including those of the Proterozoic and Cambrian, many of which actually have
ED
intimate relationships with substantial amounts of black shale and pyrite (Cook
PT
and Shergold, 1984; Krajewski et al., 1994; Schwennicke et al., 2000; She et al.,
CE
2014; Hiatt et al., 2015; Cui et al., 2016), and occasionally contain filamentous microfossils resembling modern sulfide-oxidizing bacteria (Bailey et al., 2013).
AC
However, the diagenetic phosphogenesis model is incompatible with the features of phosphate stromatolites in the Salitre Formation with regard to following points. 1) The phosphate stromatolites of the Salitre Formation show sharp borders with surrounding deposits; this is commonly seen in other examples of phosphate stromatolites (Chauhan, 1979; Southgate, 1980; Banerjee et al., 1986; Krajewski
35
ACCEPTED MANUSCRIPT
et al., 1994; Kyle and Misi, 1997; Crosby et al., 2014). In addition, particles of the surrounding deposits are scarce in the investigated stromatolites except for
T
in laminas (Fig. 4A–C). Such fabrics support synsedimentary francolite
IP
precipitation, rather than early diagenetic cementation of particles by francolite.
CR
2) Previous studies proposed that the phosphate stromatolites could have been
US
formed by the replacement of carbonate stromatolites with francolite during
AN
early diagenesis (Föllmi, 1996; Misi and Kyle, 1994; Kyle and Misi, 1997). However, dolomite crystals are incorporated into phosphate stromatolites as
M
laminas and isolated crystals (Fig. 4B, C, E, G), despite that they may be also
ED
replaced by phosphate minerals (Jarvis et al., 1994). Such fabrics support
PT
synsedimentary francolite precipitation with incorporated dolomite crystals,
CE
rather than the replacement of precursor carbonate minerals (Schwennicke et al., 2000).
AC
3) Phosphate stromatolites of the Salitre Formation lack substantial amounts of black shale (i.e., organic matter) and pyrite, although AITs imply the presence of some organic matter and pyrite. This does not support the idea of significant organic matter decomposition, sulfate reduction, and sulfide oxidation contributing to the formation of the phosphate stromatolites. Although the
36
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sulfide-rich zone of Três Irmãs (Fig. 1B) contains abundant pyrites, they are formed after the phosphogenesis and do not have a genetic relationship (Misi
T
and Kyle, 1994; Kyle and Misi, 1997).
IP
4) A previous study interpreted that the depleted 13C values of dolomite around
CR
the phosphate stromatolites suggested the involvement of sulfate reduction
US
within organic-rich sediments, and that further depleted 13C values for
AN
carbonate fluorapatite suggested its precipitation in an anoxic environment (Misi and Kyle, 1994). However, the difference of 13C values between dolomite
M
and francolite can be better explained by the isotopic fractionation of dolomite,
ED
and the “depleted” value of francolite is well within the range of the global trend,
PT
as discussed above. Conversely, 13C values much lower than the investigated
CE
outcrop (ca. −3 to −1‰) have been reported from Ediacaran phosphorites of the Doushantuo Formation that are intimately related with substantial amount of
AC
black shales and pyrites (−8.5 to −6.3‰; She et al., 2014) and from modern dolomite formed in anoxic environments inhabited by sulfate reducing bacteria (ca. −10‰; Sánchez-Román et al., 2009). 5) Phosphorite-bearing sediments related with a Fe-redox cycle exhibit strong positive correlation between Fe and P contents (Jarvis et al., 1994). Conversely,
37
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the deposits in the investigated outcrop exhibited a weak negative correlation (Fig. 7A), which did not support the theory that a Fe-redox cycle made a
T
significant contribution.
IP
6) Mn is usually dissolved in pore water under reductive conditions (e.g., Hiatt and
CR
Pufahl, 2014), and potentially incorporated into francolite (Jarvis et al., 1994).
US
However, the content of Mn is negatively correlated with those of elements
AN
related with francolite (P, Na, Sr, Ba, and S; Fig. 7A). This trend suggests that the francolite composing the stromatolites is precipitated under oxic conditions
M
rather than an anoxic environment.
ED
One of the reasons why the previous studies employed the diagenetic
PT
phosphogenesis model to the Salitre Formation would be due to the absence of an
CE
alternative model. Behavior of phosphorus is often considered to be redox-sensitive because it tends to be adsorbed by redox-sensitive Fe-oxides; however, it should be
AC
noted that the phosphate chemical species themselves are pH-sensitive. This is demonstrated by using known constants (Fig. 10A): the relative abundance of PO 43− logarithmically increases with pH. It is well known that photosynthetic CO2/HCO3− assimilation causes a pH increase by shifting carbonate equilibrium, and that the concomitant increase of
38
ACCEPTED MANUSCRIPT
CO32− relative abundance (Fig. 10B) results in carbonate stromatolite formation due to an elevated CaCO3 saturation state (Shiraishi, 2012). Therefore, it is likely that a
T
photosynthetic pH increase also shifts the phosphate equilibrium causing an
IP
increase in the relative abundance of PO43−, elevating the saturation state of
CR
francolite (Eq. 1), and leading to phosphate stromatolite formation. Indeed, some
US
cyanobacteria are known to produce carbonate and polyphosphate inclusions
AN
intracellularly (Couradeau et al., 2012; Benzerara et al., 2014). Although definitive evidence is lacking, some of microfossils found in the investigated phosphate
M
stromatolites resemble cyanobacteria (Fig. 5I, K, L). Even if these microfossils are
ED
not cyanobacteria, it is reasonable to assume the presence of photosynthetic
PT
microorganisms on the stromatolite surface as this is generally considered to be the
Indeed,
CE
case for stromatolite-forming microbial mats (e.g., Stal, 2000). the
numerical
calculations
demonstrated
that
photosynthesis
AC
significantly increased the saturation state for francolite (Fig. 9; Table 2). Considering that the magnitude of increase in mineral saturation state is generally proportional to the precipitation quantity (Stumm and Morgan, 1996), the calculation results suggest that photosynthesis can induce substantial precipitation of francolite, instead of calcite, when [P] is above ca. 5 M under the constraints
39
ACCEPTED MANUSCRIPT
assumed for the Salitre Formation. Such high [P] also inhibits CaCO3 mineral precipitation, which might further contribute to the formation of phosphate, not
T
carbonate, stromatolites (Cook and Shergold, 1984). At [P] of ca. 10–20 M, however,
IP
significantly highfra_bef values suggest substantial inorganic precipitation of
CR
francolite, which is incompatible with the francolite preferentially associated with
US
stromatolites. From the discussion above, it is concluded that the francolite
AN
composing the investigated phosphate stromatolites is syndepositional and precipitated by microbial photosynthesis under the oxidative condition of deposit
M
surface. This interpretation solves questions that arose from the diagenetic
ED
phosphogenesis model. Although cyanobacterial EPS might also contribute to the
PT
phosphate stromatolite formation by providing nucleation sites (Schwennicke et al.,
CE
2000; She et al., 2013, 2014), it would define the locus of francolite precipitation (i.e., fabric) rather than the precipitation quantity (i.e., presence or absence of francolite)
AC
as is the case with carbonate stromatolites (Shiraishi et al., 2017).
5.4. Origin of phosphate: exogenous factors
In the scenario where the studied phosphate stromatolites originated from
40
ACCEPTED MANUSCRIPT
microbial photosynthesis, high [P] would need to be achieved in the shallow ocean environment where microbial mats, including cyanobacteria, developed. During the
T
Neoproterozoic era, unusually high [P] even in the shallow ocean was the global
IP
event (Planavsky et al., 2010) and stromatolites were relatively abundant (Riding,
CR
2006). Therefore, it seems likely that an environment favorable for phosphate
US
stromatolite formation is widespread at that time. However, the occurrence of
AN
well-developed phosphate stromatolite is rather restricted among the abundant Neoproterozoic phosphorites. This suggests the presence of an additional factor
ED
locally increased [P].
M
determining the locations where phosphate stromatolites form, which would be
PT
One factor that could potentially cause locally elevated [P] would be upwelling.
CE
For example, [P] is depleted at the surface of the modern ocean due to its consumption by primary producers, while it is as high as several M at depths
AC
greater than ca. 500 m (Karl et al., 2001). Therefore, [P] at upwelling zones is elevated up to ca. 0.5–2.0 M, even at the surface seawater near the coast (Garcia et al., 2010). Considering that [P] in the aftermath of the Neoproterozoic “Snowball Earth” glaciations was 5- to 10-times higher than that of Phanerozoic ocean (Planavsky et al., 2010), [P] at upwelling zones during that time might have
41
ACCEPTED MANUSCRIPT
achieved sufficient levels for photosynthetic francolite precipitation. Indeed, it was suggested that post-glacial oceanic overturn could drive the upwelling of
T
phosphorus-rich bottom water during the Neoproterozoic era (Cook and Shergold,
IP
1984; Donnelly et al., 1990; Shen et al., 2000; Papineau, 2010; She et al., 2014).
CR
Another potential factor is evaporation. The São Francisco Craton is considered to
US
have been in the subtropical high-pressure belt in the early Ediacaran (Drummond
AN
et al., 2015; and references therein), and evaporitic features are common in the Salitre Formation (Misi and Kyle, 1994). Therefore, the evaporitic seawater
M
condensation might further increase [P] to cause phosphogenesis, like the case of
ED
other evaporitic settings (Soudry and Champetier, 1983; and references therein).
PT
From the discussion above, it is considered that the necessary conditions for
CE
forming phosphate stromatolites in the Salitre Formation are 1) the process(es) locally elevating [P] in shallow oceans (e.g., upwelling and/or evaporation) in
AC
addition to 2) the globally high [P] and 3) a shallow ocean environment dominated by a benthic microbial community. Among these, the second condition is specific to the Neoproterozoic era. This is because the globally high [P] in the Neoproterozoic ocean is considered to be caused by the incensed chemical weathering during the severe glaciations and following extreme greenhouse conditions, leading to
42
ACCEPTED MANUSCRIPT
unprecedented continental phosphorus fluxes (Planavsky et al., 2010; She et al., 2014)．In addition, the scarcity of land vegetation (Brasier, 1992) and the increased
T
amount of felsic crusts (Caird et al., 2017) might also contribute to the increase of
IP
continental phosphorus fluxes during this time period. In any case, the
CR
Neoproterozoic severe glaciations that caused globally high [P] via chemical
US
weathering were considered to be the ultimate factor causing phosphate
AN
stromatolite formation. On the other hand, phosphate stromatolites of the Aravalli Group (ca. 1.7 Ga; Banerjee, 1971; Chauhan, 1979; Banerjee et al., 1986; Crosby et
M
al., 2014) have a significant lag time from the severe glaciation occurred around
ED
2.3–2.2 Ga (Kopp et al., 2005). This leaves open the possibility that a different
CE
PT
process might be involved, and further research is necessary.
AC
5.5. Origin of dolomite
The dolostones in the investigated outcrop were composed mainly of fine dolomite crystals of less than ca. 50 m in diameter (Figs. 3A–C; 4B, C, G) and dolomite intraclasts of ca. 100–200 m in diameter (Fig. 4C, D). Notwithstanding the differences in their appearance, similar CL characteristics suggest their common
43
ACCEPTED MANUSCRIPT
origin. Previous
studies
interpreted
that
the
dolomite
around
the
phosphate
T
stromatolites was formed by a near-surface seepage-reflux dolomitization process:
IP
dense brine was formed via high evaporation rate on the surface of supratidal flats,
CR
which moved to the subsurface and pervasively dolomitized the underlying
US
sediments (Misi and Kyle, 1994; Kyle and Misi, 1997). Dolomite crystals formed by
AN
such a process generally exhibit various CL characteristics depending on the content of Mn, an activator, and Fe, a quencher, namely the redox condition of the
M
pore water. For example, dolomite crystals of the Permian Phosphoria Rock
ED
Complex in the USA contain ca. 700 ppm of Mn and ca. 100 ppm of Fe, and exhibit
PT
bright luminescence (Hiatt and Pufahl, 2014). Those of the Permian Changxing
CE
Formation and the Triassic Feixianguan Formation contain ca. 20–250 ppm of Mn and ca. 30–630 ppm of Fe, and exhibit non to dull luminescence (Jiang et al., 2013).
AC
The dolostones and the matrix part of the phosphorites in the Salitre Formation contain ca. 20–60 ppm of Mn and ca. 150–800 ppm of Fe (Table 1), which plots in the area of dull luminescence (Fig. 7C). Indeed, the core part of these dolomite crystals exhibit dull luminescence (Figs. 3C, 4D, G), which is also consistent with the somewhat reduced condition expected from the near-surface seepage-reflux
44
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dolomitization. In contrast, the syntaxial limpid dolomite cement at the rim part exhibit bright luminescence, which is characteristic of a more reduced shallow
T
burial condition (Choquette and Hiatt, 2008). The absence of such syntaxial cement
IP
around the dolomite crystals that are isolated within and contacted to the francolite
CR
(Fig. 4G) also supports the idea that the syntaxial overgrowth occurs after the
US
emplacement of the dolomite crystals.
AN
In general, dolomites formed in evaporitic settings, including those of the seepage-reflux dolomitization, tend to have heavier 18O values (Warren, 2000),
M
although the values of the Salitre Formation are comparable to the general trend of
ED
coeval rocks as discussed above. This apparent conflict may be explained by the
PT
offset related to high temperatures during evaporation (Haas et al., 2017). For
CE
example, even if the 18O value of seawater is increased from ca. −6‰ to ca. −1‰ due to evaporation, the value of precipitated dolomite is the same at ca. −3‰ if the
AC
water temperature is increased from ca. 20 °C to ca. 40 °C (Land, 1983). Dolomite formation is also known from modern evaporitic settings, which is often intimately related with microorganisms. Such dolomite is characterized by the crystal shapes of spheroidal and dumbbell (Sánchez-Román et al., 2009; and references therein), while that of this study is by irregular abraded shape (see
45
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Section 4.1). This difference suggests that the microbial mediation is insignificant for the dolomite formation of the Salitre Formation.
T
Thus, the results of this study are consistent with the interpretation of previous
IP
studies that the dolomite around the phosphate stromatolites is formed by
CR
seepage-reflux dolomitization (or alternatively, evaporative pumping; Hsü and
US
Siegenthaler, 1969; McKenzie et al., 1980). However, previous studies considered
AN
the dolomitization to have occurred after the phosphogenesis due to the presence of 1) dolomite rhombs cutting through phosphate-rich laminas, and 2) reworked
M
phosphate intraclasts in dolomite matrix (Misi and Kyle, 1994; Kyle and Misi, 1997).
ED
However, the former fabric was not clearly indicated by previous studies, nor
PT
recognized by this study. In addition, the dolomite crystals also exhibit reworked
CE
features including an abraded shape (Fig. 4G) and intraclasts (Fig. 4C, D) (Bone et al., 1992). Moreover, the dolomites are sometimes surrounded by phosphate
AC
minerals, which is exemplified by the dolomite intraclasts with isopachous phosphate cortices (Fig. 4D, 5H) and the fine dolomite crystals incorporated into the phosphate stromatolites (Fig. 4B, C, E, G). These observations indicated that the dolomitization must have preceded the phosphogenesis, and that the dolomite crystals were reworked and redeposited after
46
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the seepage-reflux dolomitization. Interestingly, a similar clastic interpretation was applied
to
dolomite
accompanied
with
phosphate
stromatolites
in
the
T
Paleoproterozoic Aravalli Group (Banerjee, 1971). In addition, rosette-like fabric
IP
(Fig. 5B, D) in the dolostone was also reported from the Aravalli Group (Papineau et
CR
al., 2016), although their composition was different (a calcite “rosette” with dolomite
AN
US
core in this study, while apatite rosettes with chert cores in the Aravalli Group).
M
5.6. Depositional model of the investigated outcrop
ED
From the discussion above, the depositional model of phosphorites and dolostones
PT
in the investigated outcrop is established (Fig. 11).
CE
The Salitre Formation as well as the comparable Sete Lagoas Formation were formed at an evaporitic ramp (Misi and Kyle, 1994; Misi and Veizer, 1998; Sial et al.,
AC
2010; Misi et al., 2011; Drummond et al., 2015; Caird et al., 2017), and the dolomite crystals
were originally
formed on
a supratidal
flat
via
seepage-reflux
dolomitization. Somewhat consolidated dolomite crystals were eroded and reworked e.g., by waves, and the sand-sized dolomite intraclasts were accumulated and rounded at an intertidal zone under wave influence while the finer abraded
47
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dolomite crystals were redeposited mainly below the fair-weather wave base (FWWB;
Fig.
11A).
These
redeposited
dolomite
grains
were
relatively
T
unconsolidated in this stage. Under such depositional condition, [P] in the subtidal
IP
to lower intertidal zone was occasionally increased by the upwelling of
CR
phosphorus-rich bottom water and/or evaporitic seawater condensation. Increased
US
[P] promoted the development of microbial mats including cyanobacteria, and the
AN
resulting photosynthesis-induced francolite precipitation formed the consolidated phosphate stromatolites. Stromatolite morphology such as bifurcate-branching
M
columnar and elongated shapes were formed by wave and tidal currents (Hoffman,
ED
1976; Grotzinger, 1986; Caird et al., 2017) and/or difference in microbial
PT
communities (e.g., Papineau et al., 2005; Goh et al., 2009). Wave and tidal currents
CE
also fragmented the stromatolites to form phosphate intraclasts (Misi and Kyle, 1994). Ooid-like grains (rounded dolomite intraclasts with isopachous phosphate
AC
cortices; Figs. 4D, 5H) might indicate high energy setting supersaturated with francolite, although She et al. (2013) interpreted similar grains as a consequence of microbially-mediated accretional growth. Alternatively, these grains may represent diagenetic chemically-oscillating reactions proposed for granular iron formation and cherts (Papineau et al., 2017; Dodd et al., 2018). Regular alternation of dolomite and
48
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phosphate laminas observed in the inner part of the stromatolites may be related to cyclic (seasonal?) changes of the ecosystem (Krajewski et al., 1994).
T
High-order fluctuations in the relative sea level (Caird et al., 2017) formed
IP
alternating phosphate-rich and dolomite-rich layers (Fig. 11B). Rapid deposition in
CR
a steepened ramp setting triggered the sliding that lead to the formation of slump
US
folding (Vieira et al., 2007), although the conformable relationship (i.e., the
AN
transitional boundary between phosphorites and dolostones) was partially retained (Fig. 11C). During this sliding, relatively unconsolidated dolomite-rich layers were
M
significantly deformed, while consolidated phosphate-rich layers instead became
ED
fragmented. At the shallow burial stage (i.e., before substantial compaction), the
PT
syntaxial limpid dolomite cement overgrew under reductive condition, and AITs
CE
were formed within the phosphate stromatolites. After that, calcite spar cements exhibiting bright luminescence were precipitated, perhaps by a freshwater
AC
diagenesis during a period of lower sea level (Misi and Kyle, 1994; Kyle and Misi, 1997; Drummond et al., 2015). After consolidation of both layers, cracks were formed and the fluid percolated through them to form veins enriched in Fe, Mn, and Si.
49
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6. Conclusions
T
Petrographic and geochemical investigations of this study indicated that the
IP
phosphate stromatolites of the Salitre Formation were syndepositionally formed
CR
under the oxidative condition. Although “pseudofossil” AITs were common, the
US
phosphate stromatolites contained several filamentous structures resembled
AN
cyanobacteria, sulfide-oxidizing bacteria, and iron-oxidizing bacteria. U–Pb age of the phosphate stromatolites directly indicated that they deposited after the
stromatolites
were formed by photosynthesis-induced francolite
ED
phosphate
M
Marinoan glaciation. Numerical calculations demonstrated the possibility that the
PT
precipitation. To enable this process after the Marinoan glaciation, [P] higher than
CE
ca. 5 M was required in a shallow ocean setting where microbial mats including cyanobacteria developed. Although the local process(es) including upwelling and/or
AC
evaporation would be required, the unusually high [P] in the ocean during this time is the ultimate cause of phosphate stromatolite formation. Therefore, the phosphate stromatolite in the Salitre Formation is the consequence of the phosphorus-rich ocean after the Marinoan glaciation.
50
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Acknowledgements
T
We acknowledge Dr. Nilo S. Matsuda for supporting the field works, Mr. Yasuhiro
IP
Shibata for supporting EPMA analyses, Dr. Kazuya Tanaka for supporting
CR
ICP-OES measurements, and Mr. Takeru Omori for providing cyanobacterial
US
culture. We thank Dr. James Hein for valuable discussion. We express condolences
AN
on Dr. Augusto J. Pedreira who kindly supported our first fieldwork in Irecê. Dr. Dominic Papineau and an anonymous reviewer gave valuable and helpful
M
suggestions and comments. This research was supported by a JSPS Grants-in-Aid
ED
for Scientific Research Nos. 09J00250, 25800280, and 16H06022 to F.S, and
AC
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Figure and table captions
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Fig. 1. (A) Geological map of the São Francisco Craton. Inset indicates the location
CR
of the craton in Brazil. After Sial et al. (2010), Guimarães et al. (2011), and Caxito et
US
al. (2012). (B) Geological map of area indicated by red dashed line box in (A). Red
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and black circles represent the locations of the investigated outcrop and Três Irmãs, respectively. Modified after Misi and Kyle (1994) with revised stratigraphy of Caird
M
et al. (2017). (C) Detail map of the investigated outcrop. Locations of field pictures
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shown in Fig. 2 are also indicated.
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Fig. 2. Field pictures of the investigated outcrop. Length of pictured hammer is 33 cm. (A) Columnar phosphate stromatolites crowded in the phosphorite. (B) Closeup
AC
of the phosphate stromatolites. Spaces between the stromatolites are infilled by laminated dolomite-rich matrix. Arrows represent small gaps presumably formed by the slump folding. (C, D) Cross sections of phosphate stromatolites exhibiting circular (C) and elongated (D) shapes. (E, F) Laminated dolostones. Planar lamination (E) is sometimes deformed by the slump folding (F). Whitish parts are
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the remnant of carbonate veins presumably related to fluid percolation. (G, H) Borders between phosphorite and dolostone exhibiting transitional (G) and abrupt
IP
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(H) changes. Whitish parts are the remnants of carbonate veins.
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Fig. 3. Petrographic characteristics of dolostone. (A) Scanned picture of a thin
US
section (S2-2) showing planar laminations composed of light-gray and dark-gray
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laminas. (B) Magnified region from (A) (plane-polarized light: PPL). Interstices between the dolomite laminas are filled with calcite cement (e.g., red arrows).
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Calcite cements rarely surround the dolomite (white arrow). (C) Magnified region
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from (B) showing the same microscopic field of view (left: PPL, right: CL). Note that
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the lamination represents the difference in crystal size of dolomite, and that each
CE
dolomite crystal exhibits the core part with dull luminescence and the rim part with bright luminescence regardless of crystal size. Calcite cement in the lower part
AC
exhibits bright luminescence with zonation. (D) P, Ca, and Mg elemental map of the region indicated in (B). White square represents the area of picture shown in (C). Light blue, green, and yellow colors roughly represent dolomite, calcite, and francolite, respectively. Note that calcite infilled the spaces between the dolomite laminas, and that the light-gray lamina contains phosphate intraclasts. White
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arrow represents calcite cements surrounding the dolomite. (E) Fe, Si, and F elemental map of the region indicated in (B). White square represents the area of
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picture shown in (C). (F) Accumulated phosphate intraclasts occasionally observed
CR
IP
in the dolostones (S2-2; PPL).
US
Fig. 4. Petrographic characteristics of phosphorite. (A) Scanned picture of a thin
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section (S2-3), in which whitish and grayish parts are the phosphate stromatolite and the matrix, respectively. Red dashed lines represent some of the laminas. (B)
M
Scanned picture of a thin section (S2-7). Many dolomite laminas of grayish color are
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recognizable within whitish phosphate stromatolite. (C) Magnified region from (B)
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(PPL). Red dashed circles represent the traces of LA-ICP-MS measurements. (D)
CE
Magnified region from (C) showing the same microscopic field of view (left: PPL, middle: XPL (cross-polarized light), right: CL). Note isopachous phosphate cortices
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around the rounded dolomite intraclasts (e.g., arrows). (E) P, Ca, and Mg elemental map of the region indicated in (C). The white square represents the area of the picture shown in (D). Light blue, green, and yellow colors roughly represent dolomite, calcite, and francolite, respectively. (F) Fe, Si, and F elemental map of the region indicated in (C). The white square represents the area of the picture shown
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in (D). (G) Magnified view of dolomite lamina of the phosphate stromatolite (S2-4) showing the same microscopic field of view (left: PPL, middle: XPL, right: CL).
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Relatively large isolated dolomite crystal within phosphate is partly fractured
CR
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(yellow arrows).
US
Fig. 5. Filamentous structures observed in the phosphate stromatolites. (A)
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Marginal part of the phosphate stromatolite (S2-8) showing the same microscopic field of view (left: PPL, right: XPL). Three types of filamentous structures are
M
recognized: gently coiled with micritic infillings (red arrows; possible microfossil),
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thin and relatively straight with spiritic infillings (light blue arrows; AITs), and
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thick spiral with spiritic infillings (green arrows; AITs). (B) Magnified region from
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(A). Red arrow represents terminal Fe-oxide mineral. (C) Magnified region from (A). Polygonal cross section (e.g., red arrows) and longitudinal striation (e.g., blue
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arrows) are recognized. (D) ”Starburst pattern” of AITs formed by propagating the tubes in several directions (S2-9; PPL). Terminal Fe-oxide minerals exhibiting equivalent diameter of the tubes are recognized. (E) Filamentous structure exhibiting septation-like structure (black arrows) and the terminal Fe-oxide mineral (red arrow) (S2-4; PPL). (F) Filamentous structure exhibiting the coating of
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tube wall with orange-colored Fe-oxide minerals and the terminal Fe-oxide mineral (red arrow) (S2-4; PPL). (G) ”Starburst pattern” composed of micritic filaments of ca.
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1 m diameter, distributing along the border between phosphate and dolomite
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(S2-11; PPL). Terminal mineral is not recognized. (H) Microborings formed at the
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surface of isopachous phosphate cortex (S2-8; PPL). (I) Filamentous structure
US
representing possible microfossil (S2-9; PPL). Structures similar to hormogonia
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(blue arrow) and teared sheath (red arrow) are recognized. Two direction arrow represents approximate wave length. (J) Structures seen in modern filamentous
M
cyanobacteria (PPL): torn sheath (red arrow), hormogonia (blue arrow), and twisted
ED
sheath (green arrow). (K) Filamentous structure representing possible microfossil
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(S2-9; PPL). A structure similar to the twisted sheath (green arrow) is seen. Two
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direction arrow represents approximate wave length. (L) Filamentous structure representing possible microfossil (S2-4; PPL) that forms microbial mat-like
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structures within the lamina of phosphate stromatolite. Straight filaments are surrounded by microcrystalline carbonates.
Fig. 6. Columnar section of the investigated outcrop presented with the mineral composition, the Si and Zn contents, and the oxygen and carbon isotopes. A
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dolostone lens occurred around 24 m represents an intercalation by the slump folding. Si and Zn contents tend to be higher around S2-9 (gray horizon), suggesting
IP
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the influence of metal-bearing fluids that formed Zn-Pb-Ag sulfide.
CR
Fig. 7．Summary of constituent elements and stable isotopes. (A) Correlation
US
coefficients between each parameter. +1 (red) and −1 (blue) indicate perfect positive
AN
and negative correlations, respectively. (B) Cross plot of oxygen and carbon isotopes. Gray dashed line represents the trend of isotopic fractionation related with dolomite
M
(a slope of 3 / 2.4). (C) Cross plot of Fe and Mn contents. Gray dashed line represents
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the border between non and dull luminescence for CL of carbonate minerals
CE
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(Machel and Burton, 1991).
Fig. 8. Summary of LA-ICP-MS measurements. (A–D) Locations of measured areas
AC
within the thin sections are shown in the left, and corresponding Terra–Wasserburg Concordia plots are shown in the right. Areas used for calculating the weighted mean age are indicated by warm colors, while those excluded for calculation are indicated by cold colors. (E) Weighted mean of age yielded from 10 areas exhibiting the error less than 50% relative to the mean age values. Each error bar represents
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2.
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Fig. 9. Contour plots of fra at wide range of pH, [DIC], and [P]. Black solid line,
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black dashed line, gray solid line, and gray dashed line represent pCO2 = 10−1.6 atm,
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pCO2 = 10−2.7 atm, cal_bef = 80, and dol_bef = 1, respectively. White dashed lines
US
represent the contour of fra_bef. Hatched area does not satisfy the necessary
AN
conditions assumed for the Salitre Formation.
M
Fig. 10. Relative abundance of (A) phosphate and (B) carbonate species at various
ED
pH. Equilibrium constants of phosphate species are pK1 = 2.1, pK2 = 7.2, and pK3 =
CE
Morgan, 1996).
PT
12.3, and those of carbonate species are pK1 = 6.3 and pK2 = 10.3 (Stumm and
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Fig. 11. Depositional model of phosphorites and dolostones in the investigated outcrop. (A) Dolomite formed via seepage-reflux dolomitization was eroded and reworked.
[P]
increase
by
upwelling
and/or
evaporation
resulted
in
photosynthesis-induced francolite precipitation to form phosphate stromatolites. (B) Relative sea-level fluctuations formed the alternation of phosphate-rich and
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dolomite-rich layers. (C) Rapid deposition in a steepened ramp setting triggered the
T
sliding that lead to the formation of slump folding.
CR
US
(footnote) Standard deviation (s.d.) represents 1.
IP
Table 1. Constituent elements and stable isotopes.
AN
Table 2. Photosynthetic influence on the saturation states of francolite and calcite calculated at pH 7.5 with [DIC] of 2, 5, and 10 mM.
M
(footnote) bef and aft represent the values before and after the CO2 removal,
PT
ED
respectively.
CE
Appendix A. Detail data of LA-ICP-MS measurements shown in Fig. 8. r.s.e. represents the relative standard error. Lower intercept ages are presented only
AC
when the error is less than 50% relative to the mean age values. The measured data of the apatite crystals in the FC1 standard is also presented.
Appendix B. Isotopic values of NIST SRM 610 glass standard used for the calibration.
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Table 1 Sampl e
Ca %
S2-1 S2-2 S2-3s S2-3m S2-4s S2-4m S2-5 S2-6 S2-7s
27. 5 27. 5 23. 9 25. 7 23. 4 32. 0 27. 8 31. 3 21. 8
P
s.d .
0.2
2.6
0
3 7
16. 0.01 4
2 0.02 7
17. 0.00 0
2
10. 0.04 6
0.4 2.4 0.1 1.6 0.2
0.01
15. 0.03
0.1 7.6 0.3
Al
% s.d. % s.d. ppm
0.2 6.0 0.2
Mg
0
2.4 1.5 7.7 1.0 5.6
9
9
0.02 911 6
5
0.00 332 9
3
0.02 185 0
0
0.00 145 1
4
0.02 269 9
4
0.02 11. 0.01 100 3
5
7
0.00 12. 0.00 2
17. 0.00 4
8.9
0.00 469
4
1 0.7
2 0.00 3
2 906 360
Si
s.d .
ppm
Na
s.d .
12 602 27
1 838
163
5
1
16 496
4
8 213
1
3 315 11
166 1 263 1 101 7 264 0 174
s.d .
5
4
ppm
3 891 12 16 8 5
164 9 284 0 122 1 288
Sr
2
6
173
0
Ba
s.d pp s.d pp .
m 9 207
.
m 2
8
2 16
15 476
5 18
4 212
1 10
10 479
3 20
6 347
1 11
D E
2 435
8 129
1
5
5 105
4 349
4 334
7 102
1
4
1 203
1
194 0
20
190 4
12 463
pp
0.0
m
1
30
0.0 7
27
0.0 9
1 16
0.0 7
0.0 8 0.1 1 0.0 3 0.0 5 0.0 7 82
15 23
9 18 31 32 10
Fe
s.d. 0.08 3 0.05
Mn
pp s.d pp m
.
m
0.0 5
U N
0.17
0.5
1
0.07 3 8 0.03 0 0.19 2 0.32 8 0.36 2 0.00 1
251 0.6 54 182 0.8 14 519 5.1 59 113 0.4 12 371 3.7 54 186 1.4 26 148 1.5 28 100 0.4 11
pp
T P
I R
s.d.
C S
240 0.6 24
Mn/S
Zn
0.1
A
M
1 450
C A
K
s.d.
13 408
T P
E C
8 370 10 137
ppm
S
4
0.0 5 7 0.0 5 0.5 6 0.1 9 0.2 9 0.0 4
m
3 2 1 1 1 2 2 2 1
r
s.d.
0.0 2
0.0 1 0.0 2 0.0 2 0.0 2 0.0 2 0.0 5 0.0 4 0.0 1
13C ‰ VPDB
0.12
-1.2
0.13
-2.5
0.03
-2.8
0.28
-1.7
0.03
-2.6
0.16
-2.0
0.20
-0.8
0.28
-1.4
0.02
-2.6
18O s.d. 0.1 2 0.1 3 0.0 9 0.0 6 0.0 6 0.0 5 0.0 7 0.0 8 0.0 5
‰ VPDB -4.0 -5.3 -4.5 -3.3 -4.8 -4.3 -3.5 -3.0 -5.7
18O s.d. 0.0 8 0.2 1 0.1 1 0.0 9 0.0 4 0.1 1 0.0 6 0.0 5 0.0 4
‰ VSMOW 26.8 25.4 26.2 27.5 26.8 26.5 27.3 27.8 25.9
s.d. 0.0 8 0.2 1 0.1 1 0.0 9 0.0 4 0.1 1 0.0 6 0.0 5 0.0 4
ACCEPTED MANUSCRIPT
S2-7m S2-8s S2-8m S2-9s S2-9m
27. 2 23. 6 26. 5 28. 2 32. 1
S2-10 26. s
7
S2-10 25. m S2-11s S2-11 m
8 25. 8 29. 2
0.3 0.3 0.2 0.5 0.4 0.2 0.2 0.2 0.2
11. 0.01 2
7
17. 0.01 0
4
14. 0.03 6
5
18. 0.00 0
4
12. 0.03 5
3
17. 0.00 3
2
12. 0.05 7
3
17. 0.01 1
0
11. 0.01 4
3
5.1 1.0 2.8 0.3 4.3 0.8 4.1 1.0 5.0
0.01 509 2
0
0.01 342 0
3
0.02 453 5
8
0.00 701 3
2
0.02 640 4
5
0.00 160 1
5
0.03 402 8
3
0.00 281 7
8
0.00 476 9
5
22 635
4
7 676
5
9 14 25
102
6
9 178 1 214 1
10 14
7 682 4
118 8
9 13
10 605
3
160 4 220 2 201 7 265 5 194 5 256 3 188 6 223 8
22 11 8 14 9 18 10
C A
4
167
3
0 222 9 198 5 265 1 208 1 272 6 220 5
12 334
4
242
7
175
1
3
8
3 494
7 21
9 443
4 22
30 594
3 29
14 407
2 16
12 549
4 15
D E
12 421
T P
10
E C
7 613
158
0.0
24
4 0.1
55
9 0.0
70
3 0.5 0 0.0 3 0.1
180 227
2 11
23 594
5 26
12 371
4 14
0.0 8 0.2 9 0.1 4
83
4 0.50 6 0.08 9 1.17 2
120 230 39 76
421 1.9 45 147 1.5 12 320 0.9 29
1
0.60 0
1.52 5 0.08 1 0.79 0
0.2 3 0.1 2
9
743 0.1 48 148 1.0
9
411 2.3 29 76 0.3
9
231 2.5 29
2 1
0.0 3
7
0.0 7
0.0 5 0.0 5 0.0 9 0.0 4 0.2 4
2 3 4 1 1 1 2
0.14
T P
0.0
I R
0.0
C S
159 1.3
U N
0.41
A
M 1
0.11
1
0.0 2
0.0 3 0.0 3 0.0 2 0.0 2 0.0 3 0.0 3
-2.0
0.02
-2.7
0.07
-1.7
0.02
-3.4
0.12
-1.0
0.02
-2.1
0.07
-1.4
0.02
-3.4
0.08
-1.7
0.1 6 0.0 7 0.0 9 0.0 8 0.1 1 0.1 1 0.1 1 0.1 3 0.0 7
-4.0 -4.9 -4.6 -6.1 -4.1 -4.1 -3.2 -5.6 -4.3
0.0 4 0.0 5 0.1 0 0.0 5 0.1 0 0.1 1 0.0 9 0.0 9 0.1 5
26.8 26.8 26.2 25.5 26.7 27.6 27.6 25.2 26.5
0.0 4 0.0 5 0.1 0 0.0 5 0.1 0 0.1 1 0.0 9 0.0 9 0.1 5
ACCEPTED MANUSCRIPT
Table 2
fra_bef
[P] (M)
fra_aft
cal_bef
cal_aft
0.0
0.0
1.3
5.8
1
0.0
1.4
1.3
5.8
2
0.4
39.8
1.3
5.8
3
2.7
278.0
1.3
5.8
5
30.8
3197.2
1.3
5.6
10
854.6
87704.7
1.3
5.6
15
5977.2
604704.1
1.3
5.6
23849.9 2369767.4
1.3
5.6
20
CR
pH 7.5, [DIC] = 5 mM (pCO2 = 10-2.3 atm) 0.0
0.0
3.2
6.8
1
0.0
0.0
3.2
6.8
2
0.1
1.0
3.2
6.8
3
0.6
7.2
3.2
5
7.3
83.2
3.2
10
204.4
2306.7
15
1426.4
16049.4
20
5674.8
63549.3
AN
US
0.1
6.8 6.8 6.6
3.2
6.6
3.2
6.6
0.0
6.3
9.3
0.0
6.3
9.3
0.1
6.3
9.3
pH 7.5, [DIC] = 10 mM (pCO2 = 10
M
3.2
-2.0
atm)
0.0
1
0.0
2
0.0
3
0.2
0.7
6.3
9.3
5
2.4
8.6
6.3
9.3
10
66.4
239.1
6.3
9.3
464.3
1670.2
6.3
9.3
1852.5
6631.4
6.3
9.3
PT
CE
AC
20
ED
0.1
15
84
IP
0.1
T
pH = 7.5, [DIC] = 2 mM (pCO2 = 10-2.7 atm)
Figure 1
Figure 2
Figure 3
Figure 4
Figure 5
Figure 6
Figure 7
Figure 8
Figure 9
Figure 10
Figure 11
Fumito Shiraishi, Saki Ohnishi, Yasutaka Hayasaka, Yusaku Hanzawa, Chizuru Takashima, Tomoyo Okumura, Akihiro Kano PII: DOI: Reference:
S0037-0738(18)30267-7 https://doi.org/10.1016/j.sedgeo.2018.11.014 SEDGEO 5420
To appear in:
Sedimentary Geology
Received date: Revised date: Accepted date:
25 September 2018 29 November 2018 30 November 2018
Please cite this article as: Fumito Shiraishi, Saki Ohnishi, Yasutaka Hayasaka, Yusaku Hanzawa, Chizuru Takashima, Tomoyo Okumura, Akihiro Kano , Potential photosynthetic impact on phosphate stromatolite formation after the Marinoan glaciation: Paleoceanographic implications. Sedgeo (2018), https://doi.org/10.1016/ j.sedgeo.2018.11.014
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ACCEPTED MANUSCRIPT
Potential photosynthetic impact on phosphate stromatolite formation after the Marinoan glaciation: Paleoceanographic
a,*,
Saki Ohnishi a, Yasutaka Hayasaka a, Yusaku Hanzawa a,
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Fumito Shiraishi
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implications
a
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Chizuru Takashima b, Tomoyo Okumura c, Akihiro Kano d
Department of Earth and Planetary Systems Science, Hiroshima University, 1-3-1
Faculty of Culture and Education, Saga University, 1 Honjo-Machi, Saga 840-8502,
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b
M
Kagamiyama, Higashihiroshima, Hiroshima 739-8526, Japan.
Center for Advanced Marine Core Research, Kochi University, Kochi 783-8502,
Japan.
Department of Earth and Planetary Science, The University of Tokyo, 7-3-1 Hongo,
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d
CE
c
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Japan.
Bunkyo-ku, Tokyo 113-0033, Japan.
* Corresponding author. E-mail address: [email protected] (F. Shiraishi).
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Abstract
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This study investigated the origin and the depositional age of phosphate
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stromatolites contained in the Neoproterozoic Salitre Formation, Brazil. The
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stromatolites exhibited columnar shapes and were intercalated with laminated
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dolostones. Their depositional environment was interpreted to have been an evaporitic ramp where erosion and reworking by waves prevailed. Laser
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ablation-inductively coupled plasma-mass spectrometry measurements of the
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stromatolites yielded a U–Pb age of 616 ± 32 Ma, suggesting that they were
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deposited after the Marinoan glaciation. Several types of filamentous structures
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were recognized in the stromatolites; many of them were “pseudofossil” ambient inclusion trails, but some were possibly microfossils. To consider the general
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presence of photosynthetic microorganisms on the surface of shallow-water stromatolites,
the
involvement
of
microbial
photosynthesis
in
phosphate
stromatolite formation was expected. Numerical calculations determined that photosynthesis could have induced phosphate mineral precipitation at the time of deposition if the phosphorus concentration was above ca. 5 M. To achieve such high
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concentration at the shallow ocean, globally elevated phosphorus concentration was considered to be the ultimate factor in addition to the local process(es) including
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upwelling and/or evaporation. Therefore, the phosphate stromatolites in the Salitre
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Formation are evidence of the development of a phosphorus-rich ocean after the
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Marinoan glaciation.
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Keywords: Brazil, Ediacaran, phosphate stromatolite, phosphorus concentration,
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photosynthesis, U–Pb dating
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1. Introduction
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The late Neoproterozoic was an era of significant environmental change related with severe glaciations, which occurred at least two times: the Sturtian glaciation
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occurred ca. 720–700 Ma and the Marinoan glaciation occurred ca. 650–635 Ma (Fairchild and Kennedy, 2007; Hoffman and Li, 2009; Macdonald et al., 2010). Since the main source of phosphorus in the ocean is from continental weathering (Froelich et al., 1982; Föllmi, 1996), incensed chemical weathering after these severe glaciations resulted in unusually high [P] (brackets denote concentration) in the
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ocean (Planavsky et al., 2010). High [P] enhanced primary production, organic carbon burial, and consequent atmospheric oxygenation, which might have
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triggered the evolution and radiation of metazoans (Donnelly et al., 1990; Lenton
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and Watson, 2004; Papineau, 2010; Planavsky et al., 2010). Concomitantly, global
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phosphogenic events occurred through the Ediacaran and Cambrian periods (Cook
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and Shergold, 1984, 1986; Papineau, 2010; Drummond et al., 2015).
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In general, sedimentary phosphorite is considered to be formed by early diagenesis after deposition: microbial degradation of buried organic matter
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liberates PO43− in pore water, inducing precipitation of and/or replacement by
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phosphate minerals, which are sometimes further concentrated by physical
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processes such as winnowing and reworking (Sheldon, 1981; Froelich et al., 1982;
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Van Cappellen and Berner, 1991; Glenn et al., 1994; Jarvis et al., 1994; Krajewski et al., 1994; Föllmi, 1996; Filippelli, 2011). Another process potentially involved in the
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phosphorite formation is Fe-redox cycling (alternatively, the Fe-P shuttle): phosphorus adsorbed on Fe-oxides is released during early diagenesis due to reduction, and PO43− liberated into pore water causes phosphate mineral precipitation (Heggie et al., 1990; Krajewski et al., 1994; Glenn et al., 1994; Jarvis et al., 1994). Recently, there has been interest in examining the contribution of
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chemolithoautotrophic sulfide-oxidizing bacteria. They store polyphosphate under oxic conditions, but when exposed to anoxia and/or sulfide they hydrolyze
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polyphosphate to liberate PO43− into pore water causing precipitation of phosphate
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minerals (Schulz and Schulz, 2005; Goldhammer et al., 2010; Brock and
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Schulz-Vogt, 2011; Crosby and Bailey, 2012; Bailey et al., 2013; Crosby et al., 2014).
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Most Phanerozoic sedimentary phosphorites were formed in the outer shelf or
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upper slope where nutrient supply via upwelling leads to a high rate of organic matter burial (O’Brien and Veeh, 1980; Bremner and Rogers, 1990; Ruttenberg and
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Berner, 1993; Schwennicke et al., 2000). In contrast, Precambrian sedimentary
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phosphorites were mainly formed within very shallow epicontinental seas (Cook
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and Shergold, 1984; Krajewski et al., 1994; Drummond et al., 2015). These
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phosphorites sometimes exhibit unique stromatolitic forms (Chauhan, 1979; Southgate, 1980; Misi and Kyle, 1994; Bertrand-Sarfati et al., 1997; Shen et al.,
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2000), suggesting a depositional process different from that occurred in the Phanerozoic. In particular, the presence of phosphate stromatolites containing the occasional microfossil suggests that the benthic microbial community play a crucial role in phosphate mineral precipitation (Banerjee, 1971; Chauhan, 1979; Krajewski et al., 1994; Föllmi, 1996; Schwennicke et al., 2000; Papineau, 2010; Papineau et al.,
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2016); however, the causal relationship is still unclear. Late Neoproterozoic phosphate stromatolites are found in the Salitre Formation,
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in the upper part of the Una Group in São Francisco Craton, Brazil (Fig. 1). The
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Salitre Formation is composed mainly of carbonate and overlies glacial diamictites
CR
of the Bebedouro Formation. Based on research on Phanerozoic sedimentary
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phosphorites, previous studies have considered these phosphate stromatolites to
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have been formed by early diagenesis in an anoxic region of the deposits (Misi and Kyle, 1994; Caird et al., 2017). However, the observations in this study are
M
incompatible with that interpretation as discussed below, suggesting that the
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depositional process needs to be reinterpreted.
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In addition, the depositional age of the Salitre Formation is still controversial; it
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is unclear whether it is formed after the Sturtian or Marinoan glaciations. An U–Pb age of ca. 875 Ma was obtained from detrital zircon in the underlying Bebedouro
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Formation (Babinski et al., 2007), but this did not provide a constraint for the controversy above. Although a Pb–Pb age of 514 ± 33 Ma was obtained from stromatolitic carbonates of the Salitre Formation, it recorded a fluid percolation event (Trindade et al., 2004). Using Sr isotope stratigraphy, Misi and Veizer (1998) suggested the depositional age of the Salitre Formation to 670–600 Ma. However,
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reliable age data is still lacking. The depositional age of the Bambuí Group (Fig. 1A) correlative to the Una Group
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in the São Francisco Craton is also controversial. Previously, the Bambuí Group was
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considered to be post-Sturtian deposits, because a U–Pb age of 880 Ma was obtained
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from detrital zircon in the Jequitai Formation (equivalent to the Bebedouro
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Formation), and a Pb–Pb age of 740 ± 22 Ma was obtained from basal limestone of
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the Sete Lagoas Formation (equivalent to the Salitre Formation) (Babinski et al., 2007; Vieira et al., 2007; Misi et al., 2007, 2011; Sial et al., 2010). However, a U–Pb
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age of ca. 610 Ma was reported for detrital zircon from the middle section of the Sete
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Lagoas Formation, which necessitated the reconsideration of depositional age
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(Caxito et al., 2012). Caxito et al. (2012) compiled litho- and isotope stratigraphy
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and proposed that the Sete Lagoas Formation was deposited after the Marinoan glaciation (ca. 635–610 Ma). By following this compilation, recent studies
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considered that the correlative Salitre Formation could have formed after the Marinoan glaciation (Caird et al., 2017; and references therein). However, Paula-Santos et al. (2015) argued that the Sete Lagoas Formation was much younger and did not relate to either the Sturtian or the Marinoan glaciation. Thus, the age of Neoproterozoic successions in the São Francisco Craton are not
7
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well constrained, mainly due to an extensive fluid percolation event that hinders reliable isotopic age determination (Babinski et al., 1999, 2007; D’Agrella-Filho et
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al., 2000; Trindade et al., 2004; Guimarães et al., 2011). However, it is possible to
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age date using apatite from phosphate stromatolites, because apatite is much more
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resistant to post-depositional alteration than calcite (the closure temperature of the
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apatite U–Pb system is ca. 500–600°C) and has provided reliable ages of geological
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material (Sano and Terada, 1999; Barfod et al., 2002). This study therefore applied laser ablation-inductively coupled plasma-mass spectrometry (LA-ICP-MS) to
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directly determine the U–Pb age of the phosphate stromatolites.
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This study aims to elucidate the origin and depositional age of phosphate
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stromatolites in the Salitre Formation, the results of which will provide insight into
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understanding paleoceanography after the Neoproterozoic severe glaciation.
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2. Geological settings
2.1. Neoproterozoic successions of the São Francisco Craton
The major Neoproterozoic successions of the São Francisco Craton are the Una
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Group and the correlative Bambuí Group (Fig. 1A). These groups are considered to have been deposited at a shallow-marine epeiric ramp (Misi and Kyle, 1994; Misi
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and Veizer, 1998; Sial et al., 2010; Misi et al., 2011; Drummond et al., 2015; Caird et
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al., 2017). The Una Group is distributed mainly at the Irecê Basin, and
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unconformably overlies either Paleoproterozoic basement rocks or sandstones of the
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Mesoproterozoic Chapada Diamantina Group. The Una Group consists of glacial
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diamictites of the Bebedouro Formation and overlying limestones, dolostones, and siliciclastics of the Salitre Formation (Fig. 1B; Misi et al., 2011; Guimarães et al.,
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2011; Caxito et al., 2012). Caird et al. (2017) subdivided the Salitre Formation into
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Nova America, Gabriel, Jussara, and Irecê members in ascending order.
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Phosphorites are restricted to the Nova America member (Misi and Kyle, 1994; Sial
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et al., 2010; Caxito et al., 2012). Comparative successions of Nova America and Gabriel members are lacking in the Bambuí Group (Caird et al., 2017).
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Phosphate stromatolites of the Salitre Formation exhibit several forms, among which the most distinctive type is columnar one previously assigned as Jurussania
krilov (Misi and Kyle, 1994; Kyle and Misi, 1997; and references therein). Phosphate stromatolites are present in beds with thicknesses ranging from centimeters to meters and are interbedded with dolostones (Misi and Kyle, 1994).
9
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Phosphate stromatolites occasionally have cross-bedding and current-ripple bedding which suggests they are deposited in a subtidal to lower intertidal
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environment (Misi and Kyle, 1994).
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Some areas, including Três Irmãs (Fig. 1B), contain a Zn-Pb-Ag sulfide-rich zone
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at the stratigraphic position slightly above the phosphate stromatolites. This zone
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has been interpreted to have been formed by the extensive percolation of
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metal-bearing fluids within sediments during the late stage (Kyle and Misi, 1997).
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2.2. Investigated outcrop
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The outcrop investigated for this study is located ca. 5 km east of Irecê (Fig. 1B)
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and composed of alternating of phosphorites and dolostones (Fig. 1C). While the strata are relatively undeformed at the northern part of the outcrop, slump folds are
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prominent at the southern part. Dolostones are more significantly distorted by slump folding than the phosphorites. Columnar phosphate stromatolites of ca. 2 cm diameter are crowded in the phosphorites and sometimes bifurcate upward (Fig. 2A, B). The matrix between the columnar stromatolites is rich in dolomite. Cross sections of columnar stromatolites are usually circular (Fig. 2C), but occasionally
10
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exhibit elongation (Fig. 2D). Dolostones, on the other hand, are composed of fine dolomite crystals (Fig. 2E), and sometimes deformed by the slump folding (Fig. 2F).
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Boundaries between phosphorites and dolostones are transitional at the relatively
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undeformed northern part (Fig. 2G), but abrupt at the southern part where slump
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undeformed northern part (Fig. 1C; S2-1 to S2-11).
CR
folds are prominent (Fig. 2H). This study collected 11 samples from the relatively
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3.1. Petrographic analyses
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3. Methods
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Polished thin sections were observed under a polarized microscope (Eclipse LV100 POL, Nikon) and a microscope (Microphoto-FXA, Nikon) equipped with a
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cathodoluminescence (CL) system (ELM-3R, Premier American Technologies). For comparing with microfossils that was found (see Section 4.2), a modern cyanobacterial culture (Scytonema sp.; National Institute for Environmental Studies, Japan) was also observed. Thin sections of S2-2 and S2-7 were coated with carbon and used for the
11
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elemental mapping of P, Ca, Mg, F, Fe, and Si using an electron-probe microanalyzer (EPMA; JXA-8200, JEOL). For S2-7, EPMA analyses were conducted
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after LA-ICP-MS analyses described below. Each elemental map was output as a
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grayscale image and used to generate RGB composites in Adobe Photoshop (CS6,
US
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Adobe Systems).
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3.2. Bulk analyses of mineralogy, constituent elements, and stable isotopes
M
Powdered samples of phosphorites and dolostones were prepared. For phosphorite,
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stromatolite and matrix parts were powdered separately (suffixes of s and m were
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added to the names of stromatolite and matrix samples, respectively).
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Aliquots from each powdered sample were used for analyzing mineralogy with a powder X-ray diffractometer with Cu K radiation (40 kV, 40 mA) and a graphite
AC
monochromator (MultiFlex, Rigaku). Mineral percentages were evaluated using the RIR (Reference Intensity Ratio) method of the ICDD qualitative analysis software. Another aliquot of each powdered sample was used to analyze constituent elements. About 100 mg of powder was collected in a centrifugal tube and dissolved by dropping 1 mL of 1M HCl followed by adding 0.25 mL of 35% HCl. After
12
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centrifugation, the supernatant was diluted with 2% HNO3, and Ca, P, Mg, Al, Si, Na, S, Sr, Ba, K, Fe, Mn, and Zn contents were measured using an inductively
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coupled plasma optical emission spectroscope (ICP-OES; iCAP6300, Thermo Fisher
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Scientific). The resulting data was used to calculate the elemental content of
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carbonate and phosphate minerals.
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A third aliquot was used for analyzing oxygen and carbon stable isotopes
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following the method of Furuyama et al. (2017). About 0.3 mg of powder was put in an airtight glass vial. After replacing the air in the vial with purified He gas, CO2
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was generated by reacting a drop (up to 0.02 mL) of anhydrous phosphoric acid at
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60°C for more than 6 h. The CO2 was separated from other gases using a gas
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preparation system (GasBench II, Thermo Fisher Scientific), and the oxygen and
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carbon isotopic compositions were measured using a mass spectrometer (Delta Plus, Thermo Fisher Scientific). The isotopic values were reported relative to the Vienna
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Peedee Belemnite (VPDB) standard, and the standard deviations of repeated measurements were better than 0.2‰. Oxygen isotopic values were also reported relative to the Vienna Standard Mean Ocean Water (VSMOW) standard for a reference. Correlation coefficients between each elemental content and isotopic value were
13
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calculated by the CORREL function of Excel (Microsoft) for evaluating how strongly
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two variables were related to each other.
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3.3. U–Pb dating of apatite
US
U–Pb isotopes were measured following the methods of Katsube et al. (2012).
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Polished thin sections of four phosphorite samples (S2-4, S2-7, S2-9, and S2-11) were ultrasonicated in Milli-Q water for ca. 10 min and air-dried before
(UP-213,
New
Wave
Research)
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Laser
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measurement. The LA-ICP-MS systems used for this study were a 213 nm Nd-YAG and
an
ICP-MS
(7500,
Agilent).
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Measurements were performed using a mixed He–N2–Ar carrier gas system
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equipped with a small volume ablation cell, sample aerosol stabilizer (buffering chamber), and charcoal filter attachment. The sample surface was first
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decontaminated by a pre-abrasion with a laser spot diameter of 65 m, followed by the measurement with a spot diameter of 55 µm. Measurement areas were selected from the margin of the columnar stromatolites where the phosphate minerals were relatively pure (see below). 4–7 areas were selected from each sample, and 4–8 spots were measured in each area. Every time after measuring 10 spots of samples, NIST
14
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SRM 610 glass standard was measured for calibrating
206Pb/238U
and
207Pb/206Pb
ratios. The data reduction program Pepi-AGE (Dunkl et al., 2008) was used to
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process the raw data, including common Pb correction. Final plots were generated
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using Isoplot/Ex (Version 3.23; Ludwig, 2003). The isotopic ratios and ages were
CR
quoted at 2σ, and the weighted mean was given with a 95% confidence interval.
US
The calibration using NIST SRM 610 glass standard described above was
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previously applied for determining U–Pb age of zircons by LA-ICP-MS, and demonstrated that the matrix effect was smaller than the analytical uncertainties
M
(Orihashi et al., 2008). In addition, least ablation energies of NIST SRM 610 glass
ED
standard and apatite minerals are comparable, which is advantageous for unifying
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the fractionations of U and Pb during a deepening of ablation pit (e.g., Hirata and
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Nesbitt, 1995; Eggins et al., 1998). To evaluate the validity of the measured ages, the apatite crystals within Duluth Complex gabbro FC1 (Paces and Millar, 1993)
AC
were measured.
3.4. Evaluation of photosynthetic influence on francolite and calcite precipitation
As discussed later in this study, photosynthetic microorganisms were potentially
15
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involved in the formation of the phosphate stromatolites in the Salitre Formation. Therefore, it was necessary to determine the influence of photosynthesis on
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phosphate mineral precipitation; this was evaluated using numerical calculations.
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In general, marine sedimentary phosphate precipitates as francolite, a carbonate
Föllmi,
1996),
with
an
average
stoichiometry
of
US
1994;
CR
fluorapatite mineral (Jahnke, 1984; Van Cappellen and Berner, 1991; Jarvis et al.,
AN
Ca9.54Na0.33Mg0.13(PO4)4.8(CO3)1.2F2.48 (Jahnke et al., 1983; Ruttenberg and Berner, 1993; Schenau et al., 2000). Therefore, the saturation state of francolite (fra) is
ED
M
defined as
CE
PT
fra = {Ca2+}9.54{Na+}0.33{Mg2+}0.13{PO43−}4.8{CO32−}1.2{F−}2.48 / KSP_fra, (1)
where braces denote the activity of each ion and KSP_fra represents the solubility
AC
product of francolite. KSP_fra depends on {CO32−}, which was estimated using the methods of Jahnke et al. (1983). Photosynthetic influence on francolite precipitation was evaluated by calculations like those of Shiraishi (2012) who evaluated the photosynthetic influence on CaCO3 mineral precipitation: 200 mol L−1 of CO2 was removed from the given water using PHREEQC, a computer program for
16
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geochemical calculations, and the saturation state of francolite was compared before
fra = fra_aft − fra_bef
T
(fra_bef) and after (fra_aft) CO2 removal as follows:
CR
IP
(2)
US
This value corresponds to a change expected at the sediment-water interface.
AN
Similar calculations were performed for calcite precipitation (cal_bef, cal_aft, and cal).
M
The seawater chemistry used for the calculations was based on a major ion
ED
composition estimated from a mid ocean ridge brine and river water mixing model
PT
(Hardie, 2003). This study used the composition of 600 Ma ([Ca2+] = 12.7 mM, [Mg2+]
CE
= 48.6 mM, [K+] = 13.1 mM, [Na+] = 466.2 mM, and [SO42−] = 25.8 mM; Hardie, 2003), which was close to the depositional age of the Salitre Formation (616 ± 32
AC
Ma) as discussed below. This is almost comparable to the composition of modern standard seawater ([Ca2+] = 10.28 mM, [Mg2+] = 52.82 mM, [K+] = 10.21 mM, [Na+] = 469.06 mM, and [SO42−] = 28.24 mM; Zeebe and Wolf-Gladrow, 2001). The value used for water temperature was 25 °C. Modern standard seawater values were used for [Cl−] and [F−] ([Cl−] = 545.86 mM and [F−] = 0.07 mM; Zeebe and Wolf-Gladrow,
17
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2001). [P] is depleted in the modern surface ocean and is ca. 0.5–2.0 M even at the upwelling areas (Garcia et al., 2010). However, [P] in the aftermath of
T
Neoproterozoic “Snowball Earth” glaciations was estimated to be 5 to 10 times
CR
used a value for [P] ranging from 0.1 to 20 M.
IP
higher than that of Phanerozoic ocean (Planavsky et al., 2010). Therefore, this study
US
The ranges of pH and [DIC] (dissolved inorganic carbon) in the Neoproterozoic
AN
ocean are not well constrained, although they strongly affect the composition of phosphate chemical species (see below) and the magnitude of photosynthetic
M
influence on water chemistry (Arp et al., 2001; Shiraishi, 2012). Indeed, the results
ED
based on model calculations are still controversial: Bartley and Kah (2004)
PT
considered that the Neoproterozoic [DIC] may have been approaching the modern
CE
level (ca. 2 mM), while Bristow and Kennedy (2008) considered that [DIC] in the Ediacaran ocean was higher than that in the modern ocean. However, chemical
AC
proxies can provide rough ranges for pH and [DIC]. Estimation based on boron isotopes suggested that the pH value was ca. 7.0–8.5 after the Marinoan glaciation (Kasemann et al., 2010). In addition, estimation based on triple oxygen isotope suggested that pCO2 (equilibrium CO2 partial pressure) values were ca. 6,000– 26,000 ppm in the immediate aftermath of Marinoan glaciation and ca. 2,000–7,000
18
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ppm around the Neoproterozoic/Cambrian boundary (Bao et al., 2008). To consider the depositional age of the Salitre Formation (616 ± 32 Ma) discussed below, a pCO2
T
value of 2,000–26,000 ppm (i.e., 10−2.7–10−1.6 atm) is the maximum estimate for the
IP
Salitre Formation. These pH and pCO2 values constrained the [DIC] range to ca.
CR
0.56–177.83 mM (see Results), which was used for the calculations in this study.
AN
below had the following three constraints:
US
In addition, the depositional characteristics of phosphate stromatolites described
1) The presence of primary phosphate stromatolites indicated a lack of substantial
M
francolite dissolution, so fra_bef > 1.
ED
2) The presence of redeposited dolomite within the phosphate stromatolites
CE
dolomite) >1.
PT
indicated a lack of substantial dolomite dissolution, so dol (saturation state of
3) Scarcity of primary calcite within phosphate stromatolites indicated the lack of
AC
substantial calcite precipitation even inorganically, so cal_bef < 20 (≈101.3) (Shiraishi, 2012). However, phosphorus inhibits carbonate precipitation: for example, in the case where [P] is 10 M, ca. 0.4–0.6 higher calcite saturation in the logarithmic scale is required to achieving the precipitation rate comparable to the case without phosphorus (Burton and Walter, 1990). Therefore, this study,
19
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employed cal_bef < 80 (≈10(1.3
+ 0.6))
for the condition precluding substantial
T
inorganic calcite precipitation.
CR
IP
4. Results
AN
US
4.1. Petrographic characteristics
Dolostones exhibited parallel laminations ca. 1–15 mm thick, which represented
M
repeating light-gray and dark-gray laminas (Fig. 3A). They were composed mainly
ED
of fine dolomite crystals, and the lamination represented different crystal sizes:
PT
light-gray was ca. 50 m and dark-gray was ca. 10 m (Fig. 3B–C). Although there
CE
were some rhombic euhedral crystals (Planar-e), many dolomite crystals showed irregular abraded shapes (Bone et al., 1992) rather than anhedral (also see below).
AC
Dolomite crystals showed similar CL characteristics regardless of their size and shape: the core part of each crystal exhibited dull luminescence without distinct zonation, while the limpid rim part (ca. 5 m thick), representing syntaxial overgrowth, exhibited brighter luminescence (Fig. 3C). Light-gray laminas sometimes contained angular to rounded phosphate intraclasts of ca. 1 mm
20
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diameter, with the long axis parallel to the bedding plane (Fig. 3D, F). Interstitial spaces of laminas were sometimes infilled by calcite cement exhibiting bright
T
luminescence with zonation (Fig. 3C, D). Calcite cement rarely surrounded the
CR
Siliciclastics and iron minerals were scarce (Fig. 3E).
IP
dolomite (Fig. 3B, D), which might represent a rosette (Papineau et al., 2016).
US
Phosphorites were composed of stromatolite part and infilling matrix part, and
AN
boundaries between them were sharp. However, some laminas were traceable across the boundary (Fig. 4A), suggesting that the protrusion of stromatolites was
M
less than ca. 1 cm at the time of deposition. Phosphate stromatolites mostly
ED
exhibited columnar shape, while laminar stromatolites sometimes bridged the
PT
columns (Fig. 4B). Margins of the columnar stromatolites (ca. 1 mm) were relatively
CE
pure phosphate, while the inner part contained dolomite laminas of less than ca. 500 m thick as well as isolated fine dolomite crystals (Fig. 4B, C, G). The matrix
AC
part was dominated by rounded and sorted dolomite intraclasts of ca. 100–200 m in diameter, around which isopachous phosphate cortices were common (Fig. 4C, D). Dolomite intraclasts sometimes composed stromatolite laminas. Phosphate intraclasts of less than ca. 1 cm diameter were frequently contained in the matrix part (Fig. 4A). Although the appearance of dolomite in the phosphorites was
21
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different from that of dolostones, characteristics of individual dolomite crystals were similar: dull luminescence in the core and brighter luminescence in the rim,
T
and sometimes represented an abraded shape (Fig. 4D, G). The rim part exhibiting
IP
brighter luminescence was absent for the dolomite crystals directly contacting to
AN
US
they were infilled by calcite containing Fe and Si.
CR
the phosphate minerals (Fig. 4G). Cracks were common in the phosphorites, and
M
4.2. Filamentous structures
ED
Several types of filamentous structures were recognized in the investigated
PT
phosphate stromatolites (Fig. 5A). A major type was the structure known as
CE
ambient inclusion trails (AITs; Wacey et al., 2008a; She et al., 2016). AITs exhibited straight, curved, and spiral tubes of ca. 1–25 m diameter, and occasionally had
AC
terminal Fe-oxide minerals exhibiting equivalent diameter of the tubes (Fig. 5B, D). AITs exhibited polygonal cross sections, continuous longitudinal striations on the walls, calcite spar infillings (Fig. 5A, C), and ”starburst pattern” (Fig. 5D). Filamentous structures rarely exhibited septation-like structure (Fig. 5E) and orange-colored filament resembling iron-oxidizing bacteria (Fig. 5F), although they
22
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had terminal Fe-oxide minerals like AITs. The “starburst pattern” composed of micritic tubes of ca. 1 m diameter (Fig. 5G) apparently lacked terminal mineral
T
and arranged along the cracks as well as the border between phosphate and
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dolomite, implying the possibility of endolithic contamination e.g. by fungi in
CR
younger ages (Wacey et al., 2008a).
US
Microborings formed by ancient endolithic microorganisms were also rarely
AN
recognized, which appeared as micritic filaments of ca. 1 m diameter extending inward from the surface of isopachous phosphate cortex (Fig. 5H). Similar
M
microborings were reported from other phosphorites (Krajewski et al., 1994; and
ED
references therein). Isopachous phosphate cortex lacked microfossils except for
PT
microborings, which was different from the situation of granular phosphorites in
CE
the Neoproterozoic Doushantuo Formation (She et al., 2014). In addition to the diagenetic/destructive structures above, two possible
AC
microfossils were recognized from phosphate stromatolites. The more frequently observed structure was wavy filaments of ca. 8–12 m diameter. Although such features were also common to AITs, these were distinguished from AITs by the absence of terminal minerals and longitudinal striations, circular cross section with relatively uniform diameter, relatively uniform sinuous wave lengths (ca. 40–60 m),
23
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and micrite/phosphate infilling that resembled a trichome with EPS (extracellular polymeric substances) sheath (Fig. 5I, K). Indeed, the “sheath” was sometimes torn
T
(Fig. 5I, red arrow), twisted (Fig. 5K, green arrow), and showed a hormogonia-like
IP
appearance (Fig. 5I, blue arrow) that was similar to modern cyanobacteria (Fig. 5J).
CR
Less abundant microfossils were straight filaments of ca. 3 m diameter with
US
external microcrystalline carbonates that formed microbial mat-like structures
AN
within the laminas of phosphate stromatolites (Fig. 5L).
ED
M
4.3. Characteristics of bulk mineralogy, constituent elements, and stable isotopes
PT
XRD analysis indicated that the major constituent minerals of phosphorites and
CE
dolostones were fluorapatite and dolomite, respectively (Fig. 6). Fluorapatite should be practically described as francolite because of the ion substitutions described
AC
below, although the diffraction data for francolite is not available in the database used. Phosphorites also contained a certain amount of dolomite and the dolomite percentage was found to be higher in the matrix part. Minor mineral components were calcite (less than ca. 30%) and clay minerals (illite and montmorillonite; less than ca. 5%).
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Results of constituent element and stable isotope analyses are presented in Table 1 and Fig. 7 (results of analysis for Si and Zn content as well as those for 13C and
T
18O are also presented in Fig. 6). Based mainly on correlation coefficients (Fig. 7A),
IP
the measured elements and isotopes were categorized into three groups.
CR
1) P, Na, Sr, Ba, and S: P is the constituent element of francolite, and other
US
elements tend to be substituted by francolite (Jarvis et al., 1994).
AN
2) Mg, Ca, 18O, and 13C: Ca and Mg are the constituent elements of dolomite. Although Ca is also contained in francolite, its content is higher in dolomite (ca.
M
13% for dolomite, compared to ca. 3% for francolite). Positive correlation
ED
between Mg content, 18O, and 13C (Fig. 7A) largely reflects the isotopic
PT
fractionation related to dolomite as discussed below (Fig. 7B).
CE
3) Si, Al, K, Zn, Fe, and Mn: These elements are mainly associated with clay minerals, diagenesis, and metal-bearing fluids, although K and Si may partially
AC
be substituted in francolite (Jarvis et al., 1994). Si and Zn content (and Al and K content) are found to be elevated around S2-9 (Fig. 6), suggesting the influence of metal-bearing fluids that form Zn-Pb-Ag sulfide. The positive correlation between Fe and Mn content suggests their incorporation under reductive condition. Especially, dolomite-rich parts (the dolostones and the matrix part of
25
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phosphorites) exhibit higher Fe and Mn contents (Fig. 7C), supporting the
T
formation of dolomite under somewhat reductive condition (see Section 5.5).
CR
IP
4.4. U–Pb age of the phosphate stromatolites
US
The results of LA-ICP-MS measurements are presented in Fig. 8, and their
AN
background data are listed in Appendix A and B. Data for S2-4, S2-7, and S2-11 exhibited a well constrained linear array in the Terra–Wasserburg Concordia plot
M
(Fig. 8A, B, D), indicating variable contribution of common Pb and radiogenic Pb. In
ED
contrast, the data for S2-9 was concentrated in a relatively small domain (Fig. 8C),
PT
possibly due to the metal-bearing fluids percolation in this horizon as described
CE
above. This sample was considered an anomalous outlier and excluded from later discussion. For the remaining samples, some areas showed a large dispersion due to
AC
impurities and/or narrow variation in common Pb/radiogenic Pb ratio. Therefore, the areas with error less than 50% relative to the mean age values were accepted (areas 4 and 7 for S2-4; areas 1, 2, 4, and 6 for S2-7; areas 2, 4, 6, and 7 for S2-11), and their weighted mean age was 616 ± 32 Ma (Fig. 8E). This age supported the hypothesis that the phosphate stromatolites in the Salitre Formation were formed
26
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after the Marinoan glaciation. The measured age of the apatite crystals in the FC1 standard (1,045 ± 15 Ma;
T
Appendix A) was comparable to the reported values for apatite (1,150 ± 59 Ma;
IP
Norman and Nemchin, 2012) and zircon (1,099.0 ± 0.1 Ma; Paces and Millar, 1993).
CR
Taking into consideration the lower closure temperature of apatite, the result of this
US
study was more conformable to the age of zircon. In any case, this result supported
AN
the validity of the age obtained above and the application of the glass standard
M
calibration for apatite minerals.
PT
ED
4.5. Results of numerical calculations
CE
Table 2 shows the photosynthetic influence on the saturation states of francolite and calcite, calculated at pH 7.5 with [DIC] of 2, 5, and 10 mM. The saturation state
AC
of francolite was significantly increased by photosynthetic CO2 removal when [P] was higher than ca. 3–5 M. However, the photosynthetic influence was buffered at higher [DIC] levels: the increase in francolite saturation state was ca. 100-fold when [DIC] = 2 mM, while it was suppressed to ca. 4-fold at [DIC] = 10 mM. The saturation state of calcite was also increased by photosynthesis; however, its degree
27
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was much smaller than that of francolite (ca. 4-fold at [DIC] = 2 mM while ca. 2-fold at [DIC] = 10 mM).
T
Calculation results at wider ranges of pH and [DIC] are shown in Fig. 9.
IP
Phosphorus-independent constraints (i.e., pCO2, dol, and cal_bef) assumed for the
CR
Salitre Formation further restricted the ranges into a rhomboidal area of pH = ca.
US
7.0–8.4 and [DIC] = ca. 1–50 mM. This rhomboidal area was overlapped by the
AN
range restriction from phosphorus-dependent constraint (i.e., fra_bef) when [P] was above ca. 3 M (the areas without hatches in Fig. 9). On the other hand, the area of
M
high fra value was enlarged as [P] increased. The increase of fra value was
ED
suppressed at the condition of low pH and high [DIC] due to CO2 buffering
PT
(Shiraishi, 2012). In addition, fra value at low [DIC] was significantly high
CE
because the low DIC effect (Shiraishi, 2012) does not affect the phosphate chemical species, unlike the case of CaCO3 precipitation. Within the range satisfying all
AC
constraints, high fra values were achieved when [P] was higher than ca. 5 M. Similar result was also obtained for fra_aft.
5. Discussion
28
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5.1. Interpretation of filamentous structures
T
Some coccoid and filamentous microfossils interpreted as cyanobacteria have
IP
been reported from ancient phosphorites, the latter of which mostly exhibited
CR
unbranching tubular structures ca. 2–30 m diameter (Chauhan, 1979; Soudry and
US
Champetier, 1983; Krajewski et al., 1994; Zhang et al., 1998; Schwennicke et al.,
AN
2000; She et al., 2014). Such morphological characteristics also apply to the filamentous structures found in the phosphate stromatolites investigated in this
M
study; however, many of the structures found are “pseudofossil” AITs (as described
ED
earlier in this article). AITs have been also reported from other Neoproterozoic
PT
phosphorites (Xiao and Knoll, 1999; Wacey et al., 2008a; She et al., 2013, 2016). The
CE
formation process of AITs is generally as follows: 1) organic matter attached to metal-rich grains (typically pyrite) undergo decomposition during compaction and
AC
heating in the shallow burial stage, 2) the metal-rich grains are propelled by pressure solution due to the generated gas pressure, and 3) pyrites are sometimes changed to hematite pseudomorphs (Knoll and Barghoorn, 1974; Wacey et al., 2008a,b). AITs found in this study possessed reddish Fe-oxide grains as terminal minerals (e.g., Fig. 5B, D), probably representing hematite pseudomorphs
29
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generated by the oxidation of pyrite grains like the case of Ediacaran Doushantuo Formation (She et al., 2016). Therefore, the presence of AITs suggests the
T
development of an anoxic environment and a Fe/S redox cycle during diagenesis.
IP
Nonetheless, the contribution of such diagenetic processes on phosphate
US
constituents in the investigated outcrop (Table 1).
CR
stromatolite formation is considered to be minor because Fe and S are not the major
AN
On the other hand, two filamentous structures exhibiting characteristics similar to cyanobacteria are likely to be microfossils (Fig. 5I, K, L). Indeed, their
M
morphological features (waved filament of ca. 8–12 m wide with “sheaths” and
ED
“hormogonia”; Fig. 5I, K) resemble those of modern cyanobacteria including
PT
Phormidium (straight, coiled, or waved trichome of ca. 2–12 m wide with sheaths
CE
and hormogonia; Komárek and Hauer, 2013) and Arthrospira (open helix trichome of ca. 2.5–16 m wide with sheaths and hormogonia; Sili et al., 2012). In addition,
AC
thinner filaments with microcrystalline carbonates (Fig. 5L) resemble the encrusted sheath of cyanobacteria that is formed by high nucleation rates on acidic EPS sheaths (Shiraishi et al., 2017). If this is the case, this microfossil type represents one of rare Precambrian calcified cyanobacteria. Another potential microfossil is the filamentous structure with “septation” (Fig.
30
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5E). This resembles the sulfide-oxidizing bacteria related to phosphogenesis, which is characterized by septate and unbranching filaments with variable diameters
T
(ranging from 1–200 m; Bailey et al., 2013). However, this filamentous structure
IP
appears to have a terminal mineral grain like AITs, and the “septation” may
CR
represent traces of stepwise movement reported from AITs (Wacey et al., 2008a,b).
US
Orange-colored filament (Fig. 5F) may also represent a microfossil because it
AN
resembles possible filamentous iron-oxidizing bacteria having terminal knobs (Little et al., 2004). Microfossils of filamentous iron-oxidizing bacteria are also
M
reported from phosphate stromatolites of Paleoproterozoic Aravalli Group (Crosby
ED
et al., 2014), which may be a common feature of phosphate stromatolites. However,
PT
the orange-colored filament of the Salitre Formation lacks stalk-like structure, and
CE
similar orange-colored tubes are also formed by AITs (She et al., 2016). Thus, some of filamentous structures in the Salitre Formation may represent
AC
microfossils of cyanobacteria, sulfide-oxidizing bacteria, and iron-oxidizing bacteria; however, definitive evidence is lacking and detail investigation is necessary in future studies.
5.2. Interpretation of stable isotopes
31
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18O and 13C values of the investigated outcrop showed a positive correlation
T
(Fig. 7A, B). Such a trend in carbonate rocks is generally interpreted as the
IP
influence of diagenesis, along which both 18O and 13C values decrease (Banner
CR
and Hanson, 1990; Jacobson and Kaufman, 1999; Knauth and Kennedy, 2009). A
US
similar trend is reported for francolite alteration (McArthur et al., 1986; Jarvis et
AN
al., 1994). However, the Mn/Sr ratios for all samples were much lower than 2 (Table 1), suggesting the absence of substantial diagenetic influence at least for carbonate
M
rocks (Jacobson and Kaufman, 1999). In addition, Na contents were negatively
ED
correlated with 18O and 13C values (Fig. 7A), which was the opposite of the
PT
general trend of francolite diagenesis (McArthur, 1985). The positive correlation
CE
between 18O and 13C values is, instead, considered to reflect the constituent minerals: francolite and dolomite have relatively low and high isotopic values,
AC
respectively. Indeed,18O and 13C values positively correlated with the content of dolomite-related elements (i.e., Mg), while they negatively correlated with the content of francolite-related elements (i.e., P, Na, Sr, Ba, and S; Fig. 7A). Francolite contains oxygen as PO4 and CO3. The phosphoric acid treatment applied in this study liberates only the oxygen of CO3 (Kolodny and Kaplan, 1970;
32
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Shemesh et al., 1983), which has a 18O value similar to that of the coexisting calcite (Shemesh et al., 1988). The 18O value of dolomite, on the other hand, is ca. 3‰
T
heavier than that of coexisting calcite (Land, 1980; Vasconcelos et al., 2005). Indeed,
IP
18O values of dolostones in this study were heavier than that of the stromatolite
CR
part of the phosphorites, and their maximum difference was ca. 3‰. This suggests
US
that the observed difference in 18O values primarily reflects the difference in
AN
constituent minerals, and that both francolite and dolomite are precipitated from water having similar oxygen isotope composition (ca. −6‰). Such light 18O value is
M
the general trend of rocks formed after the Marinoan glaciation (Shields and Veizer,
ED
2002).
PT
For 13C, the isotopic fractionation during francolite precipitation is expected to
CE
be minimal, like in CaCO3 mineral precipitation (Jarvis et al., 1994). The 13C value of dolomite, on the other hand, has a ca. 2.4‰ heavier value than that of the
AC
coexisting calcite at 25°C (Horita, 2014). Indeed, the maximum difference of 13C values between the dolostones and the stromatolite part of the phosphorites was close to 2.4‰. This suggests that the difference in 13C values also primarily reflects the difference in constituent minerals, and that both francolite and dolomite are precipitated from water having similar carbon isotopic compositions (ca. −3‰). Such
33
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light 13C value is the general trend of rocks formed after the Marinoan glaciation (Halverson et al., 2010).
T
From the discussion above, it is concluded that the positive correlation between
IP
18O and 13C values in the investigated outcrop primarily reflects the isotopic
CR
fractionation related to dolomite. Indeed, the data fit this trend well (Fig. 7B).
US
Similar explanation may be applicable to the phosphate stromatolites of the
AN
Paleoproterozoic Aravalli Group, although the maximum differences between the stromatolitic column and intercolumn are somewhat larger than the expectation
ED
M
(4.7‰ for 18O and 4.0‰ for13C; Papineau et al., 2016).
CE
PT
5.3. Origin of phosphate: endogenous factors
Previous studies interpreted the origin of phosphate stromatolites in the Salitre
AC
Formation by relating them to early diagenesis in an anoxic environment. For example, Misi and Kyle (1994) considered that bacterial degradation of organic matter in the anoxic environment of the deposit caused localized phosphate enrichment in the pore water, which resulted in the concentration of carbonate fluorapatite by direct precipitation or replacement of calcium carbonate to form
34
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phosphate stromatolites. A similar interpretation was employed by Caird et al. (2017); they additionally suggested the possible involvement of a Fe-redox cycle and
T
chemosynthetic sulfur bacteria. In addition, Drummond et al. (2015) investigated
IP
the phosphatic siltstones of the Sete Lagoas Formation and interpreted them as
CR
follows: aeolian Fe-oxyhydroxides and clays delivered phosphorus to promote
US
cyanobacterial mat development, during which the Fe-redox cycle and activity of
AN
sulfide-oxidizing bacteria at the redox boundary caused phosphogenesis. Such diagenetic phosphogenesis model has been widely applied to ancient phosphorites
M
including those of the Proterozoic and Cambrian, many of which actually have
ED
intimate relationships with substantial amounts of black shale and pyrite (Cook
PT
and Shergold, 1984; Krajewski et al., 1994; Schwennicke et al., 2000; She et al.,
CE
2014; Hiatt et al., 2015; Cui et al., 2016), and occasionally contain filamentous microfossils resembling modern sulfide-oxidizing bacteria (Bailey et al., 2013).
AC
However, the diagenetic phosphogenesis model is incompatible with the features of phosphate stromatolites in the Salitre Formation with regard to following points. 1) The phosphate stromatolites of the Salitre Formation show sharp borders with surrounding deposits; this is commonly seen in other examples of phosphate stromatolites (Chauhan, 1979; Southgate, 1980; Banerjee et al., 1986; Krajewski
35
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et al., 1994; Kyle and Misi, 1997; Crosby et al., 2014). In addition, particles of the surrounding deposits are scarce in the investigated stromatolites except for
T
in laminas (Fig. 4A–C). Such fabrics support synsedimentary francolite
IP
precipitation, rather than early diagenetic cementation of particles by francolite.
CR
2) Previous studies proposed that the phosphate stromatolites could have been
US
formed by the replacement of carbonate stromatolites with francolite during
AN
early diagenesis (Föllmi, 1996; Misi and Kyle, 1994; Kyle and Misi, 1997). However, dolomite crystals are incorporated into phosphate stromatolites as
M
laminas and isolated crystals (Fig. 4B, C, E, G), despite that they may be also
ED
replaced by phosphate minerals (Jarvis et al., 1994). Such fabrics support
PT
synsedimentary francolite precipitation with incorporated dolomite crystals,
CE
rather than the replacement of precursor carbonate minerals (Schwennicke et al., 2000).
AC
3) Phosphate stromatolites of the Salitre Formation lack substantial amounts of black shale (i.e., organic matter) and pyrite, although AITs imply the presence of some organic matter and pyrite. This does not support the idea of significant organic matter decomposition, sulfate reduction, and sulfide oxidation contributing to the formation of the phosphate stromatolites. Although the
36
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sulfide-rich zone of Três Irmãs (Fig. 1B) contains abundant pyrites, they are formed after the phosphogenesis and do not have a genetic relationship (Misi
T
and Kyle, 1994; Kyle and Misi, 1997).
IP
4) A previous study interpreted that the depleted 13C values of dolomite around
CR
the phosphate stromatolites suggested the involvement of sulfate reduction
US
within organic-rich sediments, and that further depleted 13C values for
AN
carbonate fluorapatite suggested its precipitation in an anoxic environment (Misi and Kyle, 1994). However, the difference of 13C values between dolomite
M
and francolite can be better explained by the isotopic fractionation of dolomite,
ED
and the “depleted” value of francolite is well within the range of the global trend,
PT
as discussed above. Conversely, 13C values much lower than the investigated
CE
outcrop (ca. −3 to −1‰) have been reported from Ediacaran phosphorites of the Doushantuo Formation that are intimately related with substantial amount of
AC
black shales and pyrites (−8.5 to −6.3‰; She et al., 2014) and from modern dolomite formed in anoxic environments inhabited by sulfate reducing bacteria (ca. −10‰; Sánchez-Román et al., 2009). 5) Phosphorite-bearing sediments related with a Fe-redox cycle exhibit strong positive correlation between Fe and P contents (Jarvis et al., 1994). Conversely,
37
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the deposits in the investigated outcrop exhibited a weak negative correlation (Fig. 7A), which did not support the theory that a Fe-redox cycle made a
T
significant contribution.
IP
6) Mn is usually dissolved in pore water under reductive conditions (e.g., Hiatt and
CR
Pufahl, 2014), and potentially incorporated into francolite (Jarvis et al., 1994).
US
However, the content of Mn is negatively correlated with those of elements
AN
related with francolite (P, Na, Sr, Ba, and S; Fig. 7A). This trend suggests that the francolite composing the stromatolites is precipitated under oxic conditions
M
rather than an anoxic environment.
ED
One of the reasons why the previous studies employed the diagenetic
PT
phosphogenesis model to the Salitre Formation would be due to the absence of an
CE
alternative model. Behavior of phosphorus is often considered to be redox-sensitive because it tends to be adsorbed by redox-sensitive Fe-oxides; however, it should be
AC
noted that the phosphate chemical species themselves are pH-sensitive. This is demonstrated by using known constants (Fig. 10A): the relative abundance of PO 43− logarithmically increases with pH. It is well known that photosynthetic CO2/HCO3− assimilation causes a pH increase by shifting carbonate equilibrium, and that the concomitant increase of
38
ACCEPTED MANUSCRIPT
CO32− relative abundance (Fig. 10B) results in carbonate stromatolite formation due to an elevated CaCO3 saturation state (Shiraishi, 2012). Therefore, it is likely that a
T
photosynthetic pH increase also shifts the phosphate equilibrium causing an
IP
increase in the relative abundance of PO43−, elevating the saturation state of
CR
francolite (Eq. 1), and leading to phosphate stromatolite formation. Indeed, some
US
cyanobacteria are known to produce carbonate and polyphosphate inclusions
AN
intracellularly (Couradeau et al., 2012; Benzerara et al., 2014). Although definitive evidence is lacking, some of microfossils found in the investigated phosphate
M
stromatolites resemble cyanobacteria (Fig. 5I, K, L). Even if these microfossils are
ED
not cyanobacteria, it is reasonable to assume the presence of photosynthetic
PT
microorganisms on the stromatolite surface as this is generally considered to be the
Indeed,
CE
case for stromatolite-forming microbial mats (e.g., Stal, 2000). the
numerical
calculations
demonstrated
that
photosynthesis
AC
significantly increased the saturation state for francolite (Fig. 9; Table 2). Considering that the magnitude of increase in mineral saturation state is generally proportional to the precipitation quantity (Stumm and Morgan, 1996), the calculation results suggest that photosynthesis can induce substantial precipitation of francolite, instead of calcite, when [P] is above ca. 5 M under the constraints
39
ACCEPTED MANUSCRIPT
assumed for the Salitre Formation. Such high [P] also inhibits CaCO3 mineral precipitation, which might further contribute to the formation of phosphate, not
T
carbonate, stromatolites (Cook and Shergold, 1984). At [P] of ca. 10–20 M, however,
IP
significantly highfra_bef values suggest substantial inorganic precipitation of
CR
francolite, which is incompatible with the francolite preferentially associated with
US
stromatolites. From the discussion above, it is concluded that the francolite
AN
composing the investigated phosphate stromatolites is syndepositional and precipitated by microbial photosynthesis under the oxidative condition of deposit
M
surface. This interpretation solves questions that arose from the diagenetic
ED
phosphogenesis model. Although cyanobacterial EPS might also contribute to the
PT
phosphate stromatolite formation by providing nucleation sites (Schwennicke et al.,
CE
2000; She et al., 2013, 2014), it would define the locus of francolite precipitation (i.e., fabric) rather than the precipitation quantity (i.e., presence or absence of francolite)
AC
as is the case with carbonate stromatolites (Shiraishi et al., 2017).
5.4. Origin of phosphate: exogenous factors
In the scenario where the studied phosphate stromatolites originated from
40
ACCEPTED MANUSCRIPT
microbial photosynthesis, high [P] would need to be achieved in the shallow ocean environment where microbial mats, including cyanobacteria, developed. During the
T
Neoproterozoic era, unusually high [P] even in the shallow ocean was the global
IP
event (Planavsky et al., 2010) and stromatolites were relatively abundant (Riding,
CR
2006). Therefore, it seems likely that an environment favorable for phosphate
US
stromatolite formation is widespread at that time. However, the occurrence of
AN
well-developed phosphate stromatolite is rather restricted among the abundant Neoproterozoic phosphorites. This suggests the presence of an additional factor
ED
locally increased [P].
M
determining the locations where phosphate stromatolites form, which would be
PT
One factor that could potentially cause locally elevated [P] would be upwelling.
CE
For example, [P] is depleted at the surface of the modern ocean due to its consumption by primary producers, while it is as high as several M at depths
AC
greater than ca. 500 m (Karl et al., 2001). Therefore, [P] at upwelling zones is elevated up to ca. 0.5–2.0 M, even at the surface seawater near the coast (Garcia et al., 2010). Considering that [P] in the aftermath of the Neoproterozoic “Snowball Earth” glaciations was 5- to 10-times higher than that of Phanerozoic ocean (Planavsky et al., 2010), [P] at upwelling zones during that time might have
41
ACCEPTED MANUSCRIPT
achieved sufficient levels for photosynthetic francolite precipitation. Indeed, it was suggested that post-glacial oceanic overturn could drive the upwelling of
T
phosphorus-rich bottom water during the Neoproterozoic era (Cook and Shergold,
IP
1984; Donnelly et al., 1990; Shen et al., 2000; Papineau, 2010; She et al., 2014).
CR
Another potential factor is evaporation. The São Francisco Craton is considered to
US
have been in the subtropical high-pressure belt in the early Ediacaran (Drummond
AN
et al., 2015; and references therein), and evaporitic features are common in the Salitre Formation (Misi and Kyle, 1994). Therefore, the evaporitic seawater
M
condensation might further increase [P] to cause phosphogenesis, like the case of
ED
other evaporitic settings (Soudry and Champetier, 1983; and references therein).
PT
From the discussion above, it is considered that the necessary conditions for
CE
forming phosphate stromatolites in the Salitre Formation are 1) the process(es) locally elevating [P] in shallow oceans (e.g., upwelling and/or evaporation) in
AC
addition to 2) the globally high [P] and 3) a shallow ocean environment dominated by a benthic microbial community. Among these, the second condition is specific to the Neoproterozoic era. This is because the globally high [P] in the Neoproterozoic ocean is considered to be caused by the incensed chemical weathering during the severe glaciations and following extreme greenhouse conditions, leading to
42
ACCEPTED MANUSCRIPT
unprecedented continental phosphorus fluxes (Planavsky et al., 2010; She et al., 2014)．In addition, the scarcity of land vegetation (Brasier, 1992) and the increased
T
amount of felsic crusts (Caird et al., 2017) might also contribute to the increase of
IP
continental phosphorus fluxes during this time period. In any case, the
CR
Neoproterozoic severe glaciations that caused globally high [P] via chemical
US
weathering were considered to be the ultimate factor causing phosphate
AN
stromatolite formation. On the other hand, phosphate stromatolites of the Aravalli Group (ca. 1.7 Ga; Banerjee, 1971; Chauhan, 1979; Banerjee et al., 1986; Crosby et
M
al., 2014) have a significant lag time from the severe glaciation occurred around
ED
2.3–2.2 Ga (Kopp et al., 2005). This leaves open the possibility that a different
CE
PT
process might be involved, and further research is necessary.
AC
5.5. Origin of dolomite
The dolostones in the investigated outcrop were composed mainly of fine dolomite crystals of less than ca. 50 m in diameter (Figs. 3A–C; 4B, C, G) and dolomite intraclasts of ca. 100–200 m in diameter (Fig. 4C, D). Notwithstanding the differences in their appearance, similar CL characteristics suggest their common
43
ACCEPTED MANUSCRIPT
origin. Previous
studies
interpreted
that
the
dolomite
around
the
phosphate
T
stromatolites was formed by a near-surface seepage-reflux dolomitization process:
IP
dense brine was formed via high evaporation rate on the surface of supratidal flats,
CR
which moved to the subsurface and pervasively dolomitized the underlying
US
sediments (Misi and Kyle, 1994; Kyle and Misi, 1997). Dolomite crystals formed by
AN
such a process generally exhibit various CL characteristics depending on the content of Mn, an activator, and Fe, a quencher, namely the redox condition of the
M
pore water. For example, dolomite crystals of the Permian Phosphoria Rock
ED
Complex in the USA contain ca. 700 ppm of Mn and ca. 100 ppm of Fe, and exhibit
PT
bright luminescence (Hiatt and Pufahl, 2014). Those of the Permian Changxing
CE
Formation and the Triassic Feixianguan Formation contain ca. 20–250 ppm of Mn and ca. 30–630 ppm of Fe, and exhibit non to dull luminescence (Jiang et al., 2013).
AC
The dolostones and the matrix part of the phosphorites in the Salitre Formation contain ca. 20–60 ppm of Mn and ca. 150–800 ppm of Fe (Table 1), which plots in the area of dull luminescence (Fig. 7C). Indeed, the core part of these dolomite crystals exhibit dull luminescence (Figs. 3C, 4D, G), which is also consistent with the somewhat reduced condition expected from the near-surface seepage-reflux
44
ACCEPTED MANUSCRIPT
dolomitization. In contrast, the syntaxial limpid dolomite cement at the rim part exhibit bright luminescence, which is characteristic of a more reduced shallow
T
burial condition (Choquette and Hiatt, 2008). The absence of such syntaxial cement
IP
around the dolomite crystals that are isolated within and contacted to the francolite
CR
(Fig. 4G) also supports the idea that the syntaxial overgrowth occurs after the
US
emplacement of the dolomite crystals.
AN
In general, dolomites formed in evaporitic settings, including those of the seepage-reflux dolomitization, tend to have heavier 18O values (Warren, 2000),
M
although the values of the Salitre Formation are comparable to the general trend of
ED
coeval rocks as discussed above. This apparent conflict may be explained by the
PT
offset related to high temperatures during evaporation (Haas et al., 2017). For
CE
example, even if the 18O value of seawater is increased from ca. −6‰ to ca. −1‰ due to evaporation, the value of precipitated dolomite is the same at ca. −3‰ if the
AC
water temperature is increased from ca. 20 °C to ca. 40 °C (Land, 1983). Dolomite formation is also known from modern evaporitic settings, which is often intimately related with microorganisms. Such dolomite is characterized by the crystal shapes of spheroidal and dumbbell (Sánchez-Román et al., 2009; and references therein), while that of this study is by irregular abraded shape (see
45
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Section 4.1). This difference suggests that the microbial mediation is insignificant for the dolomite formation of the Salitre Formation.
T
Thus, the results of this study are consistent with the interpretation of previous
IP
studies that the dolomite around the phosphate stromatolites is formed by
CR
seepage-reflux dolomitization (or alternatively, evaporative pumping; Hsü and
US
Siegenthaler, 1969; McKenzie et al., 1980). However, previous studies considered
AN
the dolomitization to have occurred after the phosphogenesis due to the presence of 1) dolomite rhombs cutting through phosphate-rich laminas, and 2) reworked
M
phosphate intraclasts in dolomite matrix (Misi and Kyle, 1994; Kyle and Misi, 1997).
ED
However, the former fabric was not clearly indicated by previous studies, nor
PT
recognized by this study. In addition, the dolomite crystals also exhibit reworked
CE
features including an abraded shape (Fig. 4G) and intraclasts (Fig. 4C, D) (Bone et al., 1992). Moreover, the dolomites are sometimes surrounded by phosphate
AC
minerals, which is exemplified by the dolomite intraclasts with isopachous phosphate cortices (Fig. 4D, 5H) and the fine dolomite crystals incorporated into the phosphate stromatolites (Fig. 4B, C, E, G). These observations indicated that the dolomitization must have preceded the phosphogenesis, and that the dolomite crystals were reworked and redeposited after
46
ACCEPTED MANUSCRIPT
the seepage-reflux dolomitization. Interestingly, a similar clastic interpretation was applied
to
dolomite
accompanied
with
phosphate
stromatolites
in
the
T
Paleoproterozoic Aravalli Group (Banerjee, 1971). In addition, rosette-like fabric
IP
(Fig. 5B, D) in the dolostone was also reported from the Aravalli Group (Papineau et
CR
al., 2016), although their composition was different (a calcite “rosette” with dolomite
AN
US
core in this study, while apatite rosettes with chert cores in the Aravalli Group).
M
5.6. Depositional model of the investigated outcrop
ED
From the discussion above, the depositional model of phosphorites and dolostones
PT
in the investigated outcrop is established (Fig. 11).
CE
The Salitre Formation as well as the comparable Sete Lagoas Formation were formed at an evaporitic ramp (Misi and Kyle, 1994; Misi and Veizer, 1998; Sial et al.,
AC
2010; Misi et al., 2011; Drummond et al., 2015; Caird et al., 2017), and the dolomite crystals
were originally
formed on
a supratidal
flat
via
seepage-reflux
dolomitization. Somewhat consolidated dolomite crystals were eroded and reworked e.g., by waves, and the sand-sized dolomite intraclasts were accumulated and rounded at an intertidal zone under wave influence while the finer abraded
47
ACCEPTED MANUSCRIPT
dolomite crystals were redeposited mainly below the fair-weather wave base (FWWB;
Fig.
11A).
These
redeposited
dolomite
grains
were
relatively
T
unconsolidated in this stage. Under such depositional condition, [P] in the subtidal
IP
to lower intertidal zone was occasionally increased by the upwelling of
CR
phosphorus-rich bottom water and/or evaporitic seawater condensation. Increased
US
[P] promoted the development of microbial mats including cyanobacteria, and the
AN
resulting photosynthesis-induced francolite precipitation formed the consolidated phosphate stromatolites. Stromatolite morphology such as bifurcate-branching
M
columnar and elongated shapes were formed by wave and tidal currents (Hoffman,
ED
1976; Grotzinger, 1986; Caird et al., 2017) and/or difference in microbial
PT
communities (e.g., Papineau et al., 2005; Goh et al., 2009). Wave and tidal currents
CE
also fragmented the stromatolites to form phosphate intraclasts (Misi and Kyle, 1994). Ooid-like grains (rounded dolomite intraclasts with isopachous phosphate
AC
cortices; Figs. 4D, 5H) might indicate high energy setting supersaturated with francolite, although She et al. (2013) interpreted similar grains as a consequence of microbially-mediated accretional growth. Alternatively, these grains may represent diagenetic chemically-oscillating reactions proposed for granular iron formation and cherts (Papineau et al., 2017; Dodd et al., 2018). Regular alternation of dolomite and
48
ACCEPTED MANUSCRIPT
phosphate laminas observed in the inner part of the stromatolites may be related to cyclic (seasonal?) changes of the ecosystem (Krajewski et al., 1994).
T
High-order fluctuations in the relative sea level (Caird et al., 2017) formed
IP
alternating phosphate-rich and dolomite-rich layers (Fig. 11B). Rapid deposition in
CR
a steepened ramp setting triggered the sliding that lead to the formation of slump
US
folding (Vieira et al., 2007), although the conformable relationship (i.e., the
AN
transitional boundary between phosphorites and dolostones) was partially retained (Fig. 11C). During this sliding, relatively unconsolidated dolomite-rich layers were
M
significantly deformed, while consolidated phosphate-rich layers instead became
ED
fragmented. At the shallow burial stage (i.e., before substantial compaction), the
PT
syntaxial limpid dolomite cement overgrew under reductive condition, and AITs
CE
were formed within the phosphate stromatolites. After that, calcite spar cements exhibiting bright luminescence were precipitated, perhaps by a freshwater
AC
diagenesis during a period of lower sea level (Misi and Kyle, 1994; Kyle and Misi, 1997; Drummond et al., 2015). After consolidation of both layers, cracks were formed and the fluid percolated through them to form veins enriched in Fe, Mn, and Si.
49
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6. Conclusions
T
Petrographic and geochemical investigations of this study indicated that the
IP
phosphate stromatolites of the Salitre Formation were syndepositionally formed
CR
under the oxidative condition. Although “pseudofossil” AITs were common, the
US
phosphate stromatolites contained several filamentous structures resembled
AN
cyanobacteria, sulfide-oxidizing bacteria, and iron-oxidizing bacteria. U–Pb age of the phosphate stromatolites directly indicated that they deposited after the
stromatolites
were formed by photosynthesis-induced francolite
ED
phosphate
M
Marinoan glaciation. Numerical calculations demonstrated the possibility that the
PT
precipitation. To enable this process after the Marinoan glaciation, [P] higher than
CE
ca. 5 M was required in a shallow ocean setting where microbial mats including cyanobacteria developed. Although the local process(es) including upwelling and/or
AC
evaporation would be required, the unusually high [P] in the ocean during this time is the ultimate cause of phosphate stromatolite formation. Therefore, the phosphate stromatolite in the Salitre Formation is the consequence of the phosphorus-rich ocean after the Marinoan glaciation.
50
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Acknowledgements
T
We acknowledge Dr. Nilo S. Matsuda for supporting the field works, Mr. Yasuhiro
IP
Shibata for supporting EPMA analyses, Dr. Kazuya Tanaka for supporting
CR
ICP-OES measurements, and Mr. Takeru Omori for providing cyanobacterial
US
culture. We thank Dr. James Hein for valuable discussion. We express condolences
AN
on Dr. Augusto J. Pedreira who kindly supported our first fieldwork in Irecê. Dr. Dominic Papineau and an anonymous reviewer gave valuable and helpful
M
suggestions and comments. This research was supported by a JSPS Grants-in-Aid
ED
for Scientific Research Nos. 09J00250, 25800280, and 16H06022 to F.S, and
AC
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Figure and table captions
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Fig. 1. (A) Geological map of the São Francisco Craton. Inset indicates the location
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of the craton in Brazil. After Sial et al. (2010), Guimarães et al. (2011), and Caxito et
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al. (2012). (B) Geological map of area indicated by red dashed line box in (A). Red
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Fig. 2. Field pictures of the investigated outcrop. Length of pictured hammer is 33 cm. (A) Columnar phosphate stromatolites crowded in the phosphorite. (B) Closeup
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the remnant of carbonate veins presumably related to fluid percolation. (G, H) Borders between phosphorite and dolostone exhibiting transitional (G) and abrupt
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(H) changes. Whitish parts are the remnants of carbonate veins.
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Fig. 3. Petrographic characteristics of dolostone. (A) Scanned picture of a thin
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section (S2-2) showing planar laminations composed of light-gray and dark-gray
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laminas. (B) Magnified region from (A) (plane-polarized light: PPL). Interstices between the dolomite laminas are filled with calcite cement (e.g., red arrows).
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Calcite cements rarely surround the dolomite (white arrow). (C) Magnified region
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from (B) showing the same microscopic field of view (left: PPL, right: CL). Note that
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the lamination represents the difference in crystal size of dolomite, and that each
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dolomite crystal exhibits the core part with dull luminescence and the rim part with bright luminescence regardless of crystal size. Calcite cement in the lower part
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exhibits bright luminescence with zonation. (D) P, Ca, and Mg elemental map of the region indicated in (B). White square represents the area of picture shown in (C). Light blue, green, and yellow colors roughly represent dolomite, calcite, and francolite, respectively. Note that calcite infilled the spaces between the dolomite laminas, and that the light-gray lamina contains phosphate intraclasts. White
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arrow represents calcite cements surrounding the dolomite. (E) Fe, Si, and F elemental map of the region indicated in (B). White square represents the area of
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picture shown in (C). (F) Accumulated phosphate intraclasts occasionally observed
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in the dolostones (S2-2; PPL).
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Fig. 4. Petrographic characteristics of phosphorite. (A) Scanned picture of a thin
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section (S2-3), in which whitish and grayish parts are the phosphate stromatolite and the matrix, respectively. Red dashed lines represent some of the laminas. (B)
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Scanned picture of a thin section (S2-7). Many dolomite laminas of grayish color are
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recognizable within whitish phosphate stromatolite. (C) Magnified region from (B)
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(PPL). Red dashed circles represent the traces of LA-ICP-MS measurements. (D)
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Magnified region from (C) showing the same microscopic field of view (left: PPL, middle: XPL (cross-polarized light), right: CL). Note isopachous phosphate cortices
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around the rounded dolomite intraclasts (e.g., arrows). (E) P, Ca, and Mg elemental map of the region indicated in (C). The white square represents the area of the picture shown in (D). Light blue, green, and yellow colors roughly represent dolomite, calcite, and francolite, respectively. (F) Fe, Si, and F elemental map of the region indicated in (C). The white square represents the area of the picture shown
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in (D). (G) Magnified view of dolomite lamina of the phosphate stromatolite (S2-4) showing the same microscopic field of view (left: PPL, middle: XPL, right: CL).
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Relatively large isolated dolomite crystal within phosphate is partly fractured
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(yellow arrows).
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Fig. 5. Filamentous structures observed in the phosphate stromatolites. (A)
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Marginal part of the phosphate stromatolite (S2-8) showing the same microscopic field of view (left: PPL, right: XPL). Three types of filamentous structures are
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recognized: gently coiled with micritic infillings (red arrows; possible microfossil),
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thin and relatively straight with spiritic infillings (light blue arrows; AITs), and
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thick spiral with spiritic infillings (green arrows; AITs). (B) Magnified region from
CE
(A). Red arrow represents terminal Fe-oxide mineral. (C) Magnified region from (A). Polygonal cross section (e.g., red arrows) and longitudinal striation (e.g., blue
AC
arrows) are recognized. (D) ”Starburst pattern” of AITs formed by propagating the tubes in several directions (S2-9; PPL). Terminal Fe-oxide minerals exhibiting equivalent diameter of the tubes are recognized. (E) Filamentous structure exhibiting septation-like structure (black arrows) and the terminal Fe-oxide mineral (red arrow) (S2-4; PPL). (F) Filamentous structure exhibiting the coating of
77
ACCEPTED MANUSCRIPT
tube wall with orange-colored Fe-oxide minerals and the terminal Fe-oxide mineral (red arrow) (S2-4; PPL). (G) ”Starburst pattern” composed of micritic filaments of ca.
T
1 m diameter, distributing along the border between phosphate and dolomite
IP
(S2-11; PPL). Terminal mineral is not recognized. (H) Microborings formed at the
CR
surface of isopachous phosphate cortex (S2-8; PPL). (I) Filamentous structure
US
representing possible microfossil (S2-9; PPL). Structures similar to hormogonia
AN
(blue arrow) and teared sheath (red arrow) are recognized. Two direction arrow represents approximate wave length. (J) Structures seen in modern filamentous
M
cyanobacteria (PPL): torn sheath (red arrow), hormogonia (blue arrow), and twisted
ED
sheath (green arrow). (K) Filamentous structure representing possible microfossil
PT
(S2-9; PPL). A structure similar to the twisted sheath (green arrow) is seen. Two
CE
direction arrow represents approximate wave length. (L) Filamentous structure representing possible microfossil (S2-4; PPL) that forms microbial mat-like
AC
structures within the lamina of phosphate stromatolite. Straight filaments are surrounded by microcrystalline carbonates.
Fig. 6. Columnar section of the investigated outcrop presented with the mineral composition, the Si and Zn contents, and the oxygen and carbon isotopes. A
78
ACCEPTED MANUSCRIPT
dolostone lens occurred around 24 m represents an intercalation by the slump folding. Si and Zn contents tend to be higher around S2-9 (gray horizon), suggesting
IP
T
the influence of metal-bearing fluids that formed Zn-Pb-Ag sulfide.
CR
Fig. 7．Summary of constituent elements and stable isotopes. (A) Correlation
US
coefficients between each parameter. +1 (red) and −1 (blue) indicate perfect positive
AN
and negative correlations, respectively. (B) Cross plot of oxygen and carbon isotopes. Gray dashed line represents the trend of isotopic fractionation related with dolomite
M
(a slope of 3 / 2.4). (C) Cross plot of Fe and Mn contents. Gray dashed line represents
ED
the border between non and dull luminescence for CL of carbonate minerals
CE
PT
(Machel and Burton, 1991).
Fig. 8. Summary of LA-ICP-MS measurements. (A–D) Locations of measured areas
AC
within the thin sections are shown in the left, and corresponding Terra–Wasserburg Concordia plots are shown in the right. Areas used for calculating the weighted mean age are indicated by warm colors, while those excluded for calculation are indicated by cold colors. (E) Weighted mean of age yielded from 10 areas exhibiting the error less than 50% relative to the mean age values. Each error bar represents
79
ACCEPTED MANUSCRIPT
2.
T
Fig. 9. Contour plots of fra at wide range of pH, [DIC], and [P]. Black solid line,
IP
black dashed line, gray solid line, and gray dashed line represent pCO2 = 10−1.6 atm,
CR
pCO2 = 10−2.7 atm, cal_bef = 80, and dol_bef = 1, respectively. White dashed lines
US
represent the contour of fra_bef. Hatched area does not satisfy the necessary
AN
conditions assumed for the Salitre Formation.
M
Fig. 10. Relative abundance of (A) phosphate and (B) carbonate species at various
ED
pH. Equilibrium constants of phosphate species are pK1 = 2.1, pK2 = 7.2, and pK3 =
CE
Morgan, 1996).
PT
12.3, and those of carbonate species are pK1 = 6.3 and pK2 = 10.3 (Stumm and
AC
Fig. 11. Depositional model of phosphorites and dolostones in the investigated outcrop. (A) Dolomite formed via seepage-reflux dolomitization was eroded and reworked.
[P]
increase
by
upwelling
and/or
evaporation
resulted
in
photosynthesis-induced francolite precipitation to form phosphate stromatolites. (B) Relative sea-level fluctuations formed the alternation of phosphate-rich and
80
ACCEPTED MANUSCRIPT
dolomite-rich layers. (C) Rapid deposition in a steepened ramp setting triggered the
T
sliding that lead to the formation of slump folding.
CR
US
(footnote) Standard deviation (s.d.) represents 1.
IP
Table 1. Constituent elements and stable isotopes.
AN
Table 2. Photosynthetic influence on the saturation states of francolite and calcite calculated at pH 7.5 with [DIC] of 2, 5, and 10 mM.
M
(footnote) bef and aft represent the values before and after the CO2 removal,
PT
ED
respectively.
CE
Appendix A. Detail data of LA-ICP-MS measurements shown in Fig. 8. r.s.e. represents the relative standard error. Lower intercept ages are presented only
AC
when the error is less than 50% relative to the mean age values. The measured data of the apatite crystals in the FC1 standard is also presented.
Appendix B. Isotopic values of NIST SRM 610 glass standard used for the calibration.
81
ACCEPTED MANUSCRIPT
Table 1 Sampl e
Ca %
S2-1 S2-2 S2-3s S2-3m S2-4s S2-4m S2-5 S2-6 S2-7s
27. 5 27. 5 23. 9 25. 7 23. 4 32. 0 27. 8 31. 3 21. 8
P
s.d .
0.2
2.6
0
3 7
16. 0.01 4
2 0.02 7
17. 0.00 0
2
10. 0.04 6
0.4 2.4 0.1 1.6 0.2
0.01
15. 0.03
0.1 7.6 0.3
Al
% s.d. % s.d. ppm
0.2 6.0 0.2
Mg
0
2.4 1.5 7.7 1.0 5.6
9
9
0.02 911 6
5
0.00 332 9
3
0.02 185 0
0
0.00 145 1
4
0.02 269 9
4
0.02 11. 0.01 100 3
5
7
0.00 12. 0.00 2
17. 0.00 4
8.9
0.00 469
4
1 0.7
2 0.00 3
2 906 360
Si
s.d .
ppm
Na
s.d .
12 602 27
1 838
163
5
1
16 496
4
8 213
1
3 315 11
166 1 263 1 101 7 264 0 174
s.d .
5
4
ppm
3 891 12 16 8 5
164 9 284 0 122 1 288
Sr
2
6
173
0
Ba
s.d pp s.d pp .
m 9 207
.
m 2
8
2 16
15 476
5 18
4 212
1 10
10 479
3 20
6 347
1 11
D E
2 435
8 129
1
5
5 105
4 349
4 334
7 102
1
4
1 203
1
194 0
20
190 4
12 463
pp
0.0
m
1
30
0.0 7
27
0.0 9
1 16
0.0 7
0.0 8 0.1 1 0.0 3 0.0 5 0.0 7 82
15 23
9 18 31 32 10
Fe
s.d. 0.08 3 0.05
Mn
pp s.d pp m
.
m
0.0 5
U N
0.17
0.5
1
0.07 3 8 0.03 0 0.19 2 0.32 8 0.36 2 0.00 1
251 0.6 54 182 0.8 14 519 5.1 59 113 0.4 12 371 3.7 54 186 1.4 26 148 1.5 28 100 0.4 11
pp
T P
I R
s.d.
C S
240 0.6 24
Mn/S
Zn
0.1
A
M
1 450
C A
K
s.d.
13 408
T P
E C
8 370 10 137
ppm
S
4
0.0 5 7 0.0 5 0.5 6 0.1 9 0.2 9 0.0 4
m
3 2 1 1 1 2 2 2 1
r
s.d.
0.0 2
0.0 1 0.0 2 0.0 2 0.0 2 0.0 2 0.0 5 0.0 4 0.0 1
13C ‰ VPDB
0.12
-1.2
0.13
-2.5
0.03
-2.8
0.28
-1.7
0.03
-2.6
0.16
-2.0
0.20
-0.8
0.28
-1.4
0.02
-2.6
18O s.d. 0.1 2 0.1 3 0.0 9 0.0 6 0.0 6 0.0 5 0.0 7 0.0 8 0.0 5
‰ VPDB -4.0 -5.3 -4.5 -3.3 -4.8 -4.3 -3.5 -3.0 -5.7
18O s.d. 0.0 8 0.2 1 0.1 1 0.0 9 0.0 4 0.1 1 0.0 6 0.0 5 0.0 4
‰ VSMOW 26.8 25.4 26.2 27.5 26.8 26.5 27.3 27.8 25.9
s.d. 0.0 8 0.2 1 0.1 1 0.0 9 0.0 4 0.1 1 0.0 6 0.0 5 0.0 4
ACCEPTED MANUSCRIPT
S2-7m S2-8s S2-8m S2-9s S2-9m
27. 2 23. 6 26. 5 28. 2 32. 1
S2-10 26. s
7
S2-10 25. m S2-11s S2-11 m
8 25. 8 29. 2
0.3 0.3 0.2 0.5 0.4 0.2 0.2 0.2 0.2
11. 0.01 2
7
17. 0.01 0
4
14. 0.03 6
5
18. 0.00 0
4
12. 0.03 5
3
17. 0.00 3
2
12. 0.05 7
3
17. 0.01 1
0
11. 0.01 4
3
5.1 1.0 2.8 0.3 4.3 0.8 4.1 1.0 5.0
0.01 509 2
0
0.01 342 0
3
0.02 453 5
8
0.00 701 3
2
0.02 640 4
5
0.00 160 1
5
0.03 402 8
3
0.00 281 7
8
0.00 476 9
5
22 635
4
7 676
5
9 14 25
102
6
9 178 1 214 1
10 14
7 682 4
118 8
9 13
10 605
3
160 4 220 2 201 7 265 5 194 5 256 3 188 6 223 8
22 11 8 14 9 18 10
C A
4
167
3
0 222 9 198 5 265 1 208 1 272 6 220 5
12 334
4
242
7
175
1
3
8
3 494
7 21
9 443
4 22
30 594
3 29
14 407
2 16
12 549
4 15
D E
12 421
T P
10
E C
7 613
158
0.0
24
4 0.1
55
9 0.0
70
3 0.5 0 0.0 3 0.1
180 227
2 11
23 594
5 26
12 371
4 14
0.0 8 0.2 9 0.1 4
83
4 0.50 6 0.08 9 1.17 2
120 230 39 76
421 1.9 45 147 1.5 12 320 0.9 29
1
0.60 0
1.52 5 0.08 1 0.79 0
0.2 3 0.1 2
9
743 0.1 48 148 1.0
9
411 2.3 29 76 0.3
9
231 2.5 29
2 1
0.0 3
7
0.0 7
0.0 5 0.0 5 0.0 9 0.0 4 0.2 4
2 3 4 1 1 1 2
0.14
T P
0.0
I R
0.0
C S
159 1.3
U N
0.41
A
M 1
0.11
1
0.0 2
0.0 3 0.0 3 0.0 2 0.0 2 0.0 3 0.0 3
-2.0
0.02
-2.7
0.07
-1.7
0.02
-3.4
0.12
-1.0
0.02
-2.1
0.07
-1.4
0.02
-3.4
0.08
-1.7
0.1 6 0.0 7 0.0 9 0.0 8 0.1 1 0.1 1 0.1 1 0.1 3 0.0 7
-4.0 -4.9 -4.6 -6.1 -4.1 -4.1 -3.2 -5.6 -4.3
0.0 4 0.0 5 0.1 0 0.0 5 0.1 0 0.1 1 0.0 9 0.0 9 0.1 5
26.8 26.8 26.2 25.5 26.7 27.6 27.6 25.2 26.5
0.0 4 0.0 5 0.1 0 0.0 5 0.1 0 0.1 1 0.0 9 0.0 9 0.1 5
ACCEPTED MANUSCRIPT
Table 2
fra_bef
[P] (M)
fra_aft
cal_bef
cal_aft
0.0
0.0
1.3
5.8
1
0.0
1.4
1.3
5.8
2
0.4
39.8
1.3
5.8
3
2.7
278.0
1.3
5.8
5
30.8
3197.2
1.3
5.6
10
854.6
87704.7
1.3
5.6
15
5977.2
604704.1
1.3
5.6
23849.9 2369767.4
1.3
5.6
20
CR
pH 7.5, [DIC] = 5 mM (pCO2 = 10-2.3 atm) 0.0
0.0
3.2
6.8
1
0.0
0.0
3.2
6.8
2
0.1
1.0
3.2
6.8
3
0.6
7.2
3.2
5
7.3
83.2
3.2
10
204.4
2306.7
15
1426.4
16049.4
20
5674.8
63549.3
AN
US
0.1
6.8 6.8 6.6
3.2
6.6
3.2
6.6
0.0
6.3
9.3
0.0
6.3
9.3
0.1
6.3
9.3
pH 7.5, [DIC] = 10 mM (pCO2 = 10
M
3.2
-2.0
atm)
0.0
1
0.0
2
0.0
3
0.2
0.7
6.3
9.3
5
2.4
8.6
6.3
9.3
10
66.4
239.1
6.3
9.3
464.3
1670.2
6.3
9.3
1852.5
6631.4
6.3
9.3
PT
CE
AC
20
ED
0.1
15
84
IP
0.1
T
pH = 7.5, [DIC] = 2 mM (pCO2 = 10-2.7 atm)
Figure 1
Figure 2
Figure 3
Figure 4
Figure 5
Figure 6
Figure 7
Figure 8
Figure 9
Figure 10
Figure 11