Redox conditions in the end-Early Triassic Panthalassa

Redox conditions in the end-Early Triassic Panthalassa

Palaeogeography, Palaeoclimatology, Palaeoecology 432 (2015) 15–28 Contents lists available at ScienceDirect Palaeogeography, Palaeoclimatology, Pal...

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Palaeogeography, Palaeoclimatology, Palaeoecology 432 (2015) 15–28

Contents lists available at ScienceDirect

Palaeogeography, Palaeoclimatology, Palaeoecology journal homepage: www.elsevier.com/locate/palaeo

Redox conditions in the end-Early Triassic Panthalassa Satoshi Takahashi a,⁎, Shin-ichi Yamasaki b, Kazuhiro Ogawa c, Kunio Kaiho c, Noriyoshi Tsuchiya b a b c

Department of Earth and Planetary Science, University of Tokyo, Hongo 7-3-1, Tokyo, Japan Graduate School of Environmental Studies, Tohoku University, Aramaki-Aza-Aoba 6-6-20, Sendai, Japan Graduate School of Science, Tohoku University, Aramaki-Aza-Aoba 6-3, Sendai, Japan

a r t i c l e

i n f o

Article history: Received 22 January 2015 Received in revised form 16 April 2015 Accepted 21 April 2015 Available online 12 May 2015 Keywords: Early Triassic Trace element Pyrite framboid Sulphur isotope Redox conditions Deep-sea

a b s t r a c t This study focuses on an upper Lower Triassic (Spathian) to lowermost Middle Triassic (Anisian) section representing the central Panthalassic deep sea. Analysed organic carbon isotope ratio (δ13Corg) records from the section demonstrate that lower values in the Spathian increase by up to 6‰ at the Spathian–Anisian transition. This trend accords with the carbonate carbon isotope (δ13Ccarb) record from shallow-water carbonate sections. Most horizons during late Spathian–early Anisian show features of redox conditions of not fully oxic but dysoxic conditions, inferred from low Mn, U, V, Mo and euhedral pyrite-dominated occurrences. Conversely, in the end-Spathian black-coloured beds and underlying siliceous claystone beds, relatively higher concentrations of redox-sensitive elements such as U, V, Mo and abundant pyrite framboids are detected. As enrichment factors of redox-sensitive elements are not much higher than the typical anoxic–sulphidic trend and large pyrite framboids are found, these trends suggest suboxic rather than strong anoxic conditions. These oxygen-poor conditions coincide with carbon isotope minimum values at the late Spathian. At the same time, reducing seawater conditions have been also reported in from continental sections. These coincidences imply global environmental perturbations that may have been related to the delayed recovery of life after the end-Permian mass extinction. © 2015 Elsevier B.V. All rights reserved.

1. Introduction After the Permian–Triassic mass extinction, marine ecosystems experienced several stressful environments through the Early Triassic, finally recovering during the Anisian in the early Middle Triassic (e.g., Chen and Benton, 2012; Algeo et al., 2013). The process that delayed recovery of marine animals in the Early Triassic is an important question that needs to be answered. The Early Triassic was characterised by carbon isotope perturbations (Fig. 1; Payne et al., 2004; Galfetti et al., 2007; Sun et al., 2012). The features of carbon isotopes combined with biostratigraphy provide a useful tool for global chronological correlation. Around the Early Triassic to Middle Triassic transition (Olenekian [Spathian] to Anisian), a significant swing from late Spathian minimum values to an early Anisian positive spike is observed in the carbonate carbon isotope ratio (δ13Ccarb) in shallowwater carbonate sections (Fig. 1). These carbon isotope features were labelled by Song et al. (2013, 2014), who defined N1–N4 after first initial negative and P1–P4 after first initial positive values (Fig. 1). The minimum values at late Spathian and following early Anisian maximum values are named N4 and P4, respectively. The recovery pattern after the mass extinction differs on the basis of the clades of conodonts, ammonites and possibly other nekton animals that diversified more rapidly than others in the Smithian. However, at a ⁎ Corresponding author. Tel.:+81 3 5841 4538. E-mail address: [email protected] (S. Takahashi).

http://dx.doi.org/10.1016/j.palaeo.2015.04.018 0031-0182/© 2015 Elsevier B.V. All rights reserved.

time similar to that of the late Spathian δ13Ccarb minimum values and the following increasing trend, the diversity of ammonites and conodonts decreased (Orchard, 2007; Stanley, 2009). Then, marine animals began to diversify after the early Anisian (e.g., as compiled by Chen and Benton [2012] and Sun et al. [2012]). Warm temperature conditions were implied by δ18O of conodont apatite, but this trend was not more significant than the increases in δ18O at the Griesbachian– Dienerian and Smithian–Spathian transitions (Romano et al., 2012; Sun et al., 2012). Stressful oceanic conditions during the Spathian have been reported in several regions. Oxygen-poor seawater conditions are one possible candidate for the cause of these temporal decreases and the delay of recovery because geochemical anoxic evidence, such as organic molecules and redox-sensitive elements, has been reported from some late Spathian sedimentary horizons from the Panthalassic ocean, Boreal deep sea and continental shelf carbonates (Takahashi et al., 2009a; Marenco et al., 2012; Grasby et al., 2012; Wignall et al., 2010; Song et al., 2012; Saito et al., 2014; Tian et al., 2014). Among these sedimentary settings, the Panthalassic deep ocean provides widely extended environmental information (Fig. 1). This study focuses on such pelagic deep-sea sediments from Japanese accretionary complexes (Matsuda and Isozaki, 1991; Ando et al., 2001; Fig. 1) and presents continuous stable carbon isotope ratio of organic matter (δ13Corg) curves across a Spathian–Anisian boundary section, namely the Momotaro–Jinja section (Mj section) in central Japan. This section preserves a lithological change from Spathian siliceous claystone to Anisian chert (Yamakita, 1987; Kakuwa, 1996; Isozaki, 1997). The

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Great Bank Guizhou 13

247

Middle Triassic

(eastern shallow Tethys, South China) -2

Lower Triassic

Ccarb

(VPDB ‰)

2

4

6

8

δ13Corg

Anisian -32

-30

-28

(VPDB ‰) Momotaro-Jinja section -26 (central Panthalassa, Japan)

~1.5 million years

P4

Spathian

Anisian

Middle Triassic

P3

Spathian

Lower Triassic

3m 2

N4

Smithian N3

1 0

P2 -32

-30

-28

-26

Dienerian N2

chert siliceous claystone black claystone black chert

P1

Griesbachian

N1 Upper Permian

251 Ma

0

Changhsingian

-2

0

2

4

6

8

Payne et al.(2004), Song et al.(2014)

Early-Middle Triassic

Fig. 1. Carbon isotopic profiles during the Early Triassic and Middle Triassic from a shallow-water carbonate section (Great Bank of Guizhou, South China; Payne et al., 2004) and a deepwater siliceous sedimentary section (Momotaro–Jinja section, central Japan; this study). A palaeogeographic map (by Ziegler et al., 1998) showing depositional areas of Lower–Middle Triassic sediment from South China and Japan is also shown. For δ13Corg of the Momotaro–Jinja section, a three-point running mean (blue line) and ±2 standard error (sky blue area) are shown.

deposition of these claystones is thought to reflect a decrease in biotic silica deposition and widespread oxygen-poor conditions during the Late Permian and Early Triassic (Isozaki, 1997; Algeo et al., 2010; Takahashi et al., 2009b; Algeo et al., 2011), and the recommencement of chert deposition is thought to reflect the process of biotic recovery, mainly silicic radiolarian productivity. Furthermore, Takahashi et al. (2009a) reported late Spathian development of somewhat-reducing water based on geochemical evidence, such as high dibenzothiophene (sulphur-bonded organic compounds) concentrations and high S/C ratios, and postulated that the development of oxygen-poor ocean was one possible reason for the delayed recovery in pelagic Panthalassa. Herein, we also present trace element compositions, sulphide mineral occurrences and their isotope ratios (δ34Ssulphide) for further discrimination of late Spathian redox conditions.

coloured siliceous claystone beds (1.1 m thick, Beds 1–17), a thick black chert bed (Bed 19) interbedded with black claystone, overlying dark greenish-grey-coloured siliceous claystone beds (Beds 21–24), muddy chert and chert beds (Beds 26–34), siliceous claystone beds (Beds 34–39) and muddy chert and chert beds (Beds 40–52). The upper limit of the section is the small fault. The base of this interval corresponds to the first occurrence of a chert bed interbedded in grey siliceous claystone beds (Bed group 7 of Takahashi et al., 2009a) overlying the red siliceous claystone bed of the lower part of the Mj section (Bed group 5 of Takahashi et al., 2009a). The base of the Anisian strata is placed between the black chert of Bed 19 and the muddy chert of Bed 26 (Yao and Kuwahara, 1997).

2. Geological setting and lithology

Samples were collected from the study interval in the Mj section. Weathered samples and samples with modern vegetation were avoided. Samples were sliced and polished on cut surfaces for observation. Cut surfaces were photographed using a flat-head scanner (Epson GTX970) with colour separation guides (Kodak Q-13). For sampling reproduction, rock colour information (Lab colour space) was measured by Adobe Photoshop software (Table 1). Significant white- and blackcoloured carbonaceous veins were observed in Beds 4, 20 and 53 (a typical example of Bed 20 is shown in Fig. 4). These samples are enriched in Si, Mn, Fe and Ca, which was confirmed by inductively coupled plasma mass spectrometry (ICP-MS) analysis (see underlined lists in Table 1) and X-ray microscope observations using a Horiba XGT2700. Because of these contaminations to the sedimentary chemical compositions, we did not include these samples in the following vertical plots.

The Mj section (first named by Yao and Kuwahara (1997) after the Momotaro–Jinja shrine), is located in the Mino Belt of central Japan and consists of Jurassic accretionary complexes (Fig. 2). The 7.5-mthick Mj section is composed of the following lithological facies in ascending order: 1.5-m-thick grey siliceous claystone, 1.5-m-thick red siliceous claystone (cut by a small fault at 0.25 m), 1.25-m-thick grey siliceous claystone interbedded with two chert beds, 0.16- and 0.24m-thick black chert beds and 1.5-m-thick alternating grey chert and siliceous claystone. The age of the section has been examined using radiolarian and conodont biostratigraphy (Yao and Kuwahara, 1997; Takahashi et al., 2009a), which indicates that the section ranges from Spathian (upper Olenekian) to lower Anisian. The Lower Triassic– Middle Triassic transition (equals Spathian–Anisian) is at the contact between the black chert and overlying grey chert, as indicated by the Parentactinia nakatsugawaensis/Hozmadia gifuensis zone boundary (the first occurrence of Triassocampe sp., which characterises the H. gifuensis assemblage; Figs. 2 and 3). This study focuses on the 3.3-m Spathian–Anisian transition interval, from 2 m below to ca. 1.3 m above the Olenekian–Anisian boundary (OAB) (Figs. 2 and 3). This interval is composed of dark greenish-grey-

3. Methods

3.1. Organic carbon isotope ratios The samples from the Mj section were powdered and then treated repeatedly with 6 N HCl to remove all carbonates and other acidsoluble minerals. Typically, 10–30 mg of sample powder was sealed in foil for the isotope analyses. Then, the carbon isotope ratios of the

S. Takahashi et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 432 (2015) 15–28

17

N Kiso-gawa River Inuyama

Photo of outcrop

30° 130°

PACIFIC OCEAN

135°

North

South Bed 1 (chert) Bed 19 (black chert) 0

1m

Fault

(upper limit of the section)

lower

upper

Momotaro Jinja Section

0

500m

Clastic rocks (middle Jurassic) Bedded chert (middle Triassic - lower Jurassic) Siliceous shale (lower Triassic) Fault

Fig. 2. Location and photograph of the Momotaro–Jinja section, central Japan. The geologic map around the section is based on Sugiyama (1997).

samples were analysed using a Finnigan MAT 252 mass spectrometer attached to an elemental analyser (Carlo Elba 1180) through an interface system (ConfloIII) at Tohoku University. The overall uncertainty of the δ13C values is within 0.2‰. The isotopic values are reported relative to the Vienna Pee Dee Belemnite (VPDB) standard.

by Takahashi et al. (2014). As a result (see supplemental information), the reliability of the EDXRF analysis was indicated by the good agreement between two analytical methods (regression line with a good relative coefficient). Based on comparisons of measured results, quantitated Mo values by EDXRF analysis are around 38% higher than the values by ICP-MS analysis from the outside.

3.2. EDXRF analysis 3.3. Normalisation Samples were sliced to a thickness of less than 3 mm and polished on the cut surfaces to remove metal contamination from the blade. The samples were then covered by several-fold polyvinyl chloride cloths and hammered into small pieces (~2 mm in diameter). Crushed samples were chosen to avoid veinlets and stains due to weathering. These pieces were washed in ion-exchange water, dried and then powdered in an agate planetary mill (PULVERISETTE 7, FRITSCH GmbH). Powdered samples were then pressed into pellets using a hydraulic press (Specac, Ltd). Major- and trace-element concentrations were determined using an energy-dispersive X-ray fluorescence spectrometer (EDXRF; PANanalytical Epsilon 5) at the Graduate School of Environmental Studies, Tohoku University. The instrument was calibrated using standard reference materials (listed in Matsunami et al. [2010] and Yamasaki et al. [2011]). Matsunami et al. (2010) and Yamasaki et al. (2011) did not report the quantitative performance of Mo. We therefore examined the reliability of Mo analysis by EDXRF, using pressed pellets from the deep-sea Permian–Triassic boundary (Takahashi et al., 2009b, 2010) previously quantified by ICP-MS analysis

To compare enrichments of elements, concentrations are given in the form of enrichment factors (XEF; Tribovillard et al., 2006), in which sample concentrations are normalised to the average value of upper continental crust (AUCC; McLennan, 2001), as shown in Table 1 and the vertical plots of Fig. 3:  XE F ¼ Xsample =Alsample =ðXAUCC =AlAUCC Þ;

ð1Þ

in which X and Al are the weight concentrations of elements X and Al, respectively. Using this normalised enrichment factor, we can assess how many times larger (or smaller) the detected elemental concentration is compared to crustal material before erosion and sedimentation. 3.4. Pyrite observation Thin sections were made from 15 selected horizons from the study section and observed under a transmitted-light microscope to determine

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Fig. 3. Lithologic column and vertical plots of carbon isotope ratio of organic matter (δ13Corg), concentration of Mn and its enrichment factor (MnEF), trace elemental concentrations of V, U and Mo and their enrichment factors (VEF, UEF and MoEF), and dibenzothiophene (DBT; reported by Takahashi et al., 2009a). For δ13Corg, a three-point running mean (blue line) and ±2 standard error (sky blue area) are shown. Elemental compositions are normalised by average values of upper continental crust (defined by McLennan, 2001) to form enrichment factors (EFs).

the area of euphedral pyrite and pyrite framboids. We also counted the number of pyrite framboids and their diameters. The thin sections were also observed using a backscattering electron microscope (JEOL JED2300F/SDD).

4. Results All results for the Mj section are shown in Tables 1 and 2 and Figs. 3, 5 and 6. 4.1. Organic carbon isotope ratio (δ13Corg)

3.5. Sulphide sulphur isotope ratios Sedimentary sulphide was extracted from 1 to 5 g of crushed sample from the Mj section based on the method of Canfield et al. (1986). Sulphide was removed from the sample by reaction with an excess of acidified 1 M CrCl2 solution. The mixture was boiled until all H2S evolved from the reaction of Cr2+ with sulphide. Then, the H2S was trapped as Ag2S by bubbling through an AgNO3 solution. The Ag2S was collected by filtration, dried and finally weighed. Stoichiometric calculations were used to convert the mass of Ag2S to the sulphide sulphur content of the sample. Isotopic determinations were carried out using the elemental analyser–isotope ratio mass spectrometer (EA-IRMS) method with a MAT-Finnigan 252 mass spectrometer. The δ34Ssulphide results were compared with the Canyon Diablo Troilite (CDT) standard. Errors for replicate samples are within ±0.5‰.

The δ13Corg values from the interval vary between − 31.5 and − 25.5‰ (Table 1; Fig. 3). The δ13Corg values in the lower part of the chert and siliceous claystone (Beds 1–8) vary within the relatively low range of − 31.4 to − 28.0‰. The values decrease from − 29.3 to − 31.1‰ between siliceous claystone Bed 11 and chert layer Bed 12 and remain low (−31.5 to −30.6‰) in the siliceous claystone of Beds 12 to 17 and the overlying black-coloured claystone and chert beds (Beds 18, 19). Then, δ13Corg increases to − 27.8‰ in the siliceous claystone of Bed 22 and increases further to − 26.3‰ in Bed 28, with small swings. Although δ13Corg decreases to − 29.7‰ in chert Bed 32, it then increases to −26.3‰ in the overlying siliceous claystone of Bed 34, before decreasing immediately to − 30.1‰ in Bed 35. The δ13Corg rises again towards the chert in Bed 42 (− 25.5‰). Then, the δ13Corg values range between − 29.7 and − 26.0‰, although the values

S. Takahashi et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 432 (2015) 15–28

fluctuate widely in this interval and decrease by about 1‰, with a 2‰ swing in the upper part of the chert and muddy chert of Beds 43–53. 4.2. Elemental compositions Selected elemental compositions in the study section are shown in Table 1. SiO2 is the major component (75.2–90.0%, equals 35.16–42.06% as Si) and Al2O3 is the second highest (0.39–14. 31%, equals 0.2–7.6% as Al) in most of the samples. Fe and Mn range from 0.59 to 6.5% and from 0.01 to 0.09%, respectively. A triangular diagram of these two elements and Al is shown in Fig. 5. In this plot, high Fe and Mn samples relative to Al are found in Beds 4, 20 Bed 53. Accepting the criteria by Adachi et al. (1986), these samples show significant hydrothermal influence. For this reason, their data were removed from the vertical plots for palaeoenvironmental reconstruction. Mn and MnEF show values of 0.01 to 0.18% and 0.36 to 2.63, respectively, or less than the detection limit. V and VEF range from 37.70 ppm to 484.12 ppm and from 0.70 to 5.38, respectively, being high in Beds 16, 18 and 19. Values of U and UEF are less than the detection limit to 4.98 ppm and 0.01 to 2.06, respectively (avoiding undetectable U samples). Slightly higher U and UEF overwhelming the AUCC value (Table 1) were detected from Bed 1 and Beds 14, 16, 18 and 19. Although most of the Mo concentrations vary within 0.45 to 3.06 ppm, with values lower than 1 ppm being less certain, significant high values of 16 to 25 ppm were detected in three horizons, Beds 12, 16 and 18 (Fig. 3). In these horizons, the Mo quantity is guaranteed with 41% error because an internal calibration curve was available (see Section 3.2). These high Mo concentrations accord with MoEF values of 14.2, 19.7 and 14.6, respectively. S (total S in Fig. 6) shows higher values (0.5–4.36%) in Beds 12–19 and peaks in Beds 9 and 16. In other horizons, S is low (b 0.3%) or beneath the detection limit (Fig. 6).

19

19. Significant swings by around 20‰ between −40.7 and −37.7‰ to − 25.3 and − 19.8‰, respectively, were detected in the overlying alternation of siliceous claystone and chert (Beds 22–34). Subsequently, δ34Ssulphide is relatively stable in Beds 36–46, with values of −24.0 to 20.1‰. Then, a general increase from −30.9 to − 20.0‰ occurs in the upper part of the chert and siliceous claystone of Beds 49 to 53. 5. Discussion 5.1. Correlation of the carbon isotopic variation

An example of the lower part of the study section (Bed 9) contains slight pyrites (areal 0.003% in the thin section; Fig. 6; Table 2). These pyrites are dominated by euhedral pyrite crystals; few pyrite framboids are present (only one grain per 16.4 mm2). Frequent pyrite framboid occurrences (more than 0.05 areal %; Fig. 7A) are observed in the uppermost Spathian black laminated horizons of Beds 9 and 16, and overlying black claystone bed (Bed 18), black chert (Bed 19) and siliceous claystone bed (Bed 22). Above these horizons, euhedral pyrite dominates, pyrite framboids are absent or few (less than 0.007 areal %; Fig. 7B) with the exception of Bed 46. Most pyrite framboid grains are recognised as a cluster of micro-crystals (Fig. 7C–D), occurrences of euhedral pyrite that overgrow around pyrite framboids are found (Fig. 7E). In some Anisian horizons of Beds 26 and 29, fossil-shaped pyrites, likely derived from spherical radiolarians, were recognised (Fig. 7F). Size distributions of pyrite framboids are drawn as figures for the samples with the framboids of more than 0.063 areal % in Fig. 6. As the size distribution of samples with less than 40 pyrite framboid grains in one thin section is difficult to discuss, their data are shown together in Table 2, for reference. Minimum diameters of most of pyrite framboid cross sections are 2 μm in most of the samples. Maximum diameter varies among the samples, reaching 40 μm in Beds 16 and 18 with abundant pyrite framboids. The calculated average diameters are 4.8–5.4 μm (Fig. 6). Standard deviations of diameters are 2.1–3.8 for Beds 12, 16, 18 and 19. A low standard deviation of 1.6 was detected in Beds 22 and 46, which exhibited relatively few framboid grains.

The δ13Corg of the Mj section tended to be low in the upper Spathian (−30‰) and then increased by 2 to 6‰ towards the lower Anisian horizons (Fig. 3). Above those horizons, the δ13Corg varied in the range of heavier values relative to those of the Spathian. Similar carbon isotopic variation has been reported for the carbonate carbon isotope ratio (δ13Ccarb) curve of global carbonate sections from, for example, South China (Fig. 1; Payne et al., 2004; Tong et al., 2007; Galfetti et al., 2007; Meyer et al., 2011; Sun et al., 2012; Song et al., 2013). In these studies, the δ13Ccarb curves had relatively low values in the middle part of the Spathian horizon, and an approximately 5‰ increase was found between the upper part of the Spathian and lower part of the Anisian horizon, labelled N4 and P4 by Song et al. (2014). The isotopic variations in δ13Ccarb from the shallow Tethys sections and in δ13Corg from the Mj section of the pelagic deep ocean are comparable in several ways (see comparison in Fig. 1). First, the decreasing trend of δ13Corg in the lower part of the study section (Beds 1–11) is comparable to that of δ13Ccarb from the lower part of the Spathian in the Palaeotethyan section. Second, the lowest values of δ13Corg from the horizons below the Spathian black chert in the Mj section (Beds 9–19) show simultaneous variation with the lower values of δ13Ccarb in the middle part of the Spathian Palaeotethyan section (N4; Fig. 1). Third, the subsequent increase in δ13Corg between Beds 19 and 28 at the Spathian–Anisian transition of the Mj section corresponds to the δ13Ccarb increase across the Spathian–Anisian boundary of the Palaeotethyan sections (P4; Fig. 1). Furthermore, small fluctuations (within Beds 29 to 49) and a subsequent decrease (in Beds 49 to 53) in the δ13Corg in the upper part of the Mj section also resemble the δ13Ccarb profile just above the base of the Anisian. These carbon isotopic correlations between the deep-sea Panthalassa and shallow-water Palaeotethys suggest that marine photosynthesis controls the isotopic variation in dissolved marine inorganic carbon, even for the central pelagic Panthalassa, in the case of stable carbon isotopic fractionation of primary producers in the pelagic Panthalassa during this time interval. According to the lithologic column of Payne et al. (2004) and the time scale by Sun et al. (2012), Goudemand et al. (2013) and Song et al. (2014), the above-described carbon isotopic swings across the upper Spathian and lowest Anisian (from N4 to P4) accord with ca. 100-m thickness and ca. 1–1.5 Ma in the shallow-water carbonate section of the Great Bank of Guizhou from South China. In contrast, in the pelagic deep-sea record of the Mj section, these corresponding carbon isotope variations occur within a ca. 3-m-thick interval (Fig. 1). Using these numeric data, the roughly calculated linear sedimentation rate throughout the Mj section is 2–3 mm/1000 year (= 3000-mm thickness / 1–1.5 × 106 years). This result is conformable with the general pelagic sedimentation rate (reviewed in Hüneke and Mulder, 2011) and that of the Late Triassic pelagic bedded cherts (ca. 1 mm/1000 year; Onoue et al., 2012). This agreement supports our carbon isotopic correlation between the pelagic Panthalassic deep sea and shallow Palaeotethyan carbonate sections.

4.4. Sulphide sulphur isotope ratio (δ34Ssulphide)

5.2. Redox conditions inferred from elemental composition

The δ34Ssulphide values from this interval vary between − 40.7 and − 19.8‰. The δ34Ssulphide values vary in Beds 1 to 11 but are stable in the Spathian part of the section from chert Bed 12 to black chert Bed

Sedimentary redox-sensitive elements such as Mn, V, U and Mo are a useful tool for evaluating depositional history and early diagenesis (as the compilation by Tribovillard et al., 2006). We discuss the

4.3. Pyrite observation

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Table 1 Geochemical data of the Momotaro–Jinja section. Samples of Bed 4 (Mj 85), Bed 20 (Mj 102black), and Bed 53 (Mj134) are considered to be influenced by hydrothermal input based on the discrimination diagram of Adachi et al. (1986). See Fig. 5. Sample ID

Bed group number of

Bed number

Lithology

Rock colour

Takahashi et al., 2009a Mj 134 Mj 133 Mj 132

15 53 Chert 15 51 Siliceous claystone 15 50 Muddy chert 15 Black claystone Mj 131 15 49 Muddy chert 15 Black claystone Mj 130 15 48 Chert Mj 129 15 47 Muddy chert Mj 128 15 46 Muddy chert Mj 127 15 45 Chert Mj 126 15 44 Chert Mj 125 15 43 Chert Mj 124 15 42 Chert Mj 123 15 41 Muddy chert Mj 122 15 40 Chert Mj 121 15 39 Siliceous claystone with black laminae Mj 120 15 38 Siliceous claystone Mj 119 15 37 Siliseous claystone 15 Black claystone Mj 118 15 36 Siliceous claystone Mj 117 15 35 Siliceous claystone Mj 116 15 34 Siliceous claystone Mj 115 15 33 Chert Mj 114 15 32 Chert 15 31 Chert Mj 113 Mj 112 15 30 Chert Mj 111 15 29 Muddy chert Mj 110 15 28 Chert Mj 109 15 27 Muddy chert Mj 108 15 26 Muddy chert Mj 107 15 25 Muddy chert Mj 106 14 24 Siliceous claystone Mj 105 13 23 Siliceous claystone Mj 104 13 22 Siliceous claystone Mj 103 13 21 Siliceous claystone Mj 102black 12 20 Black claystone Mj 100 11 19 Black chert Mj 99 11 18 Black claystone Mj 98 10 17 Siliceous claystone with black clay fragments Mj 97 10 16 Siliceous claystone Mj 96 10 15 Siliceous claystone Mj 95 10 14 Siliceous claystone Mj 94 10 13 Siliceous claystone with black laminae Mj 93 9 12 Chert with black laminae Mj 92 8 11 Siliceous claystone Mj 91 8 10 Siliceous claystone Mj 90 8 9 Siliceous claystone Mj 89 8 8 Siliceous claystone Mj 88 8 7 Siliceous claystone Mj 87 8 6 Siliceous claystone Mj 86 8 5 Siliceous claystone Mj 85 8 4 Siliceous claystone Mj 84 8 3 Siliceous claystone Mj 83 8 2 Siliceous claystone Mj 82 7 1 Chert Average value of upper continental crust (McLennan, 2001)

sedimentary redox conditions through the late Spathian to the early Anisian using the following geochemical proxies. Mn forms insoluble trivalent (Mn[III]) or tetravalent (Mn[IV]) hydroxides or oxides (e.g., MnO2) that are deposited rapidly in particulate form in oxic environments (Calvert and Pedersen, 1993; Sholkovitz et al., 1994). However, under anoxic conditions, Mn is reduced to divalency (Mn[II]) and forms soluble cations (e.g., Mn2+, MnCl+). Consequently, low Mn concentrations in marine sediment suggest reducing depositional conditions. MnEF has very low values (0.36–1.17) in most of the samples throughout the study section, but high MnEF values of more than 10 in Bed 1 and more than 2 in Beds 7 and 26 were detected

Dark grey Dark greenish grey Dark greenish grey Black Dark greenish grey Black Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Black Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Black Black Black Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey Dark greenish grey

δ13Corg

Height (cm) L

a

b

Base

Top

(VPDV ‰)

20 37 35

−2 −6 −3

2 6 4

−5

4

34

−7

2

33

−5

7

29 36 35 30 32 31 37

−2 −3 −6 −2 −4 −4 −5

2 3 11 7 2 2 3

35 31 34

−5 −5 −5

3 1 5

31 29 28 30 24 28 33 31 29 30 32

−5 −6 −6 −4 −4 −4 −5 −4 −4 −5 −5

3 6 5 0 6 1 0 0 2 6 5

15 14 12

−3 −2 0

0 0 0

30 29 28 27 27 33 30 33

−2 −2 −2 −4 −3 −6 −5 −4

0 0 0 2 3 5 5 4

30 26

−4 −3

0 2

33

−3

2

27 28

−6 −5

3 3

310.4 307.4 304.3 299.1 298.3 295.2 294.9 286.8 281.8 277.8 273.8 272.8 267.8 265.8 264.3 260.3 255.3 249.3 244.3 243.8 237.8 226.8 224.8 222.8 219.3 211.3 201.3 198.3 188.3 186.5 184.5 182.5 178.5 176.5 173.5 162.5 140.5 124.5 108 105.5 103.4 100.8 93.7 88.1 82 71 67.6 62.1 60 49 36 29 23 17 10

−28.2 −29.7 −27.2

36

307.4 304.3 299.1 298.3 295.2 294.9 286.8 281.8 277.8 273.8 272.8 267.8 265.8 264.3 260.3 255.3 249.3 244.3 243.8 237.8 226.8 224.8 222.8 219.3 211.3 201.3 198.3 188.3 186.5 184.5 182.5 178.5 176.5 173.5 162.5 140.5 124.5 108 105.5 103.4 100.8 93.7 88.1 82 71 67.6 62.1 60 49 36 29 23 17 10 0

−28.6 −26.8 −27.1 −29.1 −26.0 −27.9 −28.8 −25.5 −27.7 −27.4 −27.7 −28.3 −29.9 −30.1 −26.3 −26.3 −27.5 −29.7 −28.3 −27.5 −27.1 −25.3 −27.8 −28.8 −27.3 −28.3 −27.8 −27.8 −31.0 −30.9 −30.7 −30.9 −31.5 −30.9 −31.0 −31.2 −31.1 −29.3 −29.4 −29.0 −29.6 −30.7 −29.0 −31.4 −28.8 −28.9 −28.1 −30.0

(Fig. 3). Based on a comparison with Palaeozoic red-coloured bedded chert with higher MnEF of 4.4 to 49.9 interpreted as an oxic deposit (Kato et al., 2002), the lower MnEF values in the Mj section are exclusively lower than the oxic trends, and the higher values are near the oxic trend. Therefore, the low MnEF values in the Mj section rule out the possibility of an oxic depositional environment, but temporal relatively oxic conditions did occur recurrently in the late Spathian to early Anisian. As in the following paragraphs, higher redox proxies (such as VEF, UEF, MoEF) indicate a suboxic depositional environment associated with late Spathian horizons of Beds 14, 16, 18 and 19. Therefore, low MnEF values with no other trace elemental reducing signals

S. Takahashi et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 432 (2015) 15–28

Q1

21

Table 1 Geochemical data of the Momotaro–Jinja section. Underlined samples are considered to be influenced by hydrothermal input based on the discrimination diagram of Adachi et al. (1986). See Fig. 3. Al

Si

S

(%)

(%)

(%)

0.20 6.90 6.24

35.64 38.44 39.40

0.25 0.24 0.12

7.58

35.71

0.19

6.43

38.74

6.64 3.15 4.92 5.35 2.57 4.75 4.78 4.41

Mn

MoEF

δ34Ssulphide

EF

EF

10.69 0.03 0.34

48.76 1.26 0.87

0.87

0.86

−29.2

0.78

0.01

0.74

−30.9

0.77

0.83

0.88

0.75

0.85 0.88 0.90 0.97 0.99 0.85 0.77

1.58 0.95 1.11 1.94 0.94 0.96 1.09

0.27 0.22

0.87 0.77 0.76 0.96 0.84 0.75 0.76

−22.1 −20.5 −22.5 −23.6 −21.4 −23.0 −23.1 −24.0 −23.8 −23.6

Fe

V

U

Mo

MnEF

FeEF

VEF

(%)

(%)

(ppm)

(ppm)

ppm

EF

EF

EF

0.02 0.02 0.02

0.83 1.99 1.85

31.86 102.07 69.17

0.76 0.08 0.73

1.86 1.63 1.01

13.71 0.36 0.42

9.29 0.66 0.68

11.77 1.11 0.83

0.02

2.36

116.54

2.30

1.22

0.41

0.71

1.16

0.04

0.02

1.97

66.92

0.02

0.88

0.39

0.70

37.64

0.07

0.02

2.23

73.23

2.03

0.93

0.47

40.55 39.26 39.97 41.03 38.45 40.52 40.34

0.14 0.06 0.04 0.30 0.31 0.11 0.13

0.02 0.02 0.03 0.02 0.04 0.03 0.02

1.17 1.89 2.09 1.09 2.06 1.76 1.48

66.29 62.53 79.00 66.29 59.54 61.11 63.96

0.29 0.38 undetectable 0.31 1.64 0.91 0.58

0.51 0.71 0.76 0.46 0.74 0.67 0.62

0.76 0.65 0.76 1.17 1.03 0.74 0.52

UEF

0.34 0.99 0.55 0.38

DBT (dibenzothiophene)

DBT/TOC

(VCD ‰)

(ng/g sediment)

(ng/g sediment TOC)

−21.7 −20.0 −26.8

Takahashi et al., 2009a

5.76 5.15 3.64

38.08 40.02 39.70

0.00 0.05 0.05

0.02 0.02 0.02

2.33 1.89 1.26

72.77 58.97 80.29

1.18 1.43 1.81

0.76 0.74 0.56

0.58 0.54 0.60

0.93 0.85 0.80

0.95 0.86 1.66

0.59 0.80 1.43

0.71 0.78 0.83

−23.6 −23.1 −22.1

4.68 5.43 6.42 2.06 5.83 6.31 1.69 5.25 5.81 6.32 5.92

39.35 38.34 36.92 41.03 37.54 38.10 39.88 39.22 38.22 37.97 38.18

0.00 0.06 −0.02 0.06 −0.01 −0.02 0.12 0.03 −0.01 0.05 0.14

0.02 0.02 0.03 0.01 0.03 0.02 0.03 0.04 0.02 0.02 0.02

1.91 2.55 2.81 0.98 3.39 2.75 0.91 1.97 2.22 2.60 2.28

50.78 55.50 94.03 55.00 54.11 79.59 48.58 46.53 61.47 71.00 74.23

1.62 0.92 0.57 0.20 0.10 1.13 0.77 1.58 0.24 0.29 1.70

0.68 0.83 0.93 0.47 0.88 0.89 0.45 0.70 0.88 0.90 0.83

0.62 0.57 0.55 0.76 0.68 0.51 2.63 1.11 0.46 0.51 0.49

0.94 1.08 1.00 1.09 1.34 1.00 1.23 0.86 0.88 0.95 0.89

0.81 0.77 1.10 2.01 0.70 0.95 2.15 0.67 0.79 0.84 0.94

0.99 0.48 0.25 0.27 0.05 0.51 1.31 0.86 0.12 0.13 0.83

0.78 0.82 0.78 1.22 0.81 0.75 1.43 0.71 0.81 0.76 0.75

−28.5 −35.8 −22.0 −19.8 −37.4 −28.4 −24.0 −26.0 −34.2 −39.4 −40.7

0.20 1.75 6.46 5.14 6.95 3.06 5.13 6.16 6.19 6.13 6.55 6.58

34.88 39.99 34.58 40.03 34.65 42.06 38.07 38.34 35.16 38.59 37.51 37.82

0.37 0.48 2.06 1.30 4.36 0.60 0.70 0.87 9.60 −0.01 −0.01 −0.02

0.09 0.01 0.02 0.02 0.02 0.01 0.03 0.02 0.02 0.02 0.02 0.02

0.82 0.95 3.95 1.98 4.56 1.03 1.89 2.32 6.50 2.16 2.33 2.45

74.18 85.03 462.12 71.20 484.08 75.21 113.49 66.06 74.93 69.54 102.11 101.23

1.84 1.08 3.56 undetectable 4.98 undetectable 3.51 0.11 undetectable 0.24 0.43 0.41

1.45 3.06 17.59 2.62 25.60 1.12 1.03 2.13 16.42 0.89 0.96 0.99

56.37 1.13 0.45 0.53 0.48 0.47 0.69 0.42 0.52 0.46 0.51 0.50

9.25 1.24 1.40 0.89 1.51 0.78 0.85 0.86 2.41 0.81 0.82 0.86

27.27 3.65 5.38 1.04 5.23 1.85 1.66 0.81 0.91 0.85 1.17 1.16

25.88 1.77 1.58

−30.5 −27.7 −26.2 −27.6 −28.8 −28.9 −27.5 −29.4 −31.0

0.11 0.19 0.18

38.11 9.39 14.59 2.73 19.74 1.95 1.07 1.85 14.21 0.78 0.79 0.80

3.59 6.06 5.87 0.20

40.67 37.08 37.27 36.97

0.03 −0.01 −0.01 0.06

0.06 0.04 0.03 0.04

1.20 3.32 3.08 0.59

49.50 79.74 82.52 40.10

0.98 0.10 1.07 0.10

0.55 0.91 0.94

2.08 0.91 0.74 28.95

0.77 1.26 1.21 6.60

1.04 0.99 1.06 14.75

0.78 0.05 0.52 1.42

0.82 0.80 0.85 0.00

−28.0 −38.2 −31.7

5.54 2.34 8.04

36.11 38.57

−0.02 −0.01

0.06 0.18 0.06

3.73 1.20

71.60 37.30 107

1.34 1.83 2.8

0.93 0.50 1.5

1.37 10.42

1.55 1.18

0.97 1.20

0.70 2.24

0.90 1.13

−23.9 −30.0

in the Mj section, except for these suboxic horizons, would reflect dysoxic depositional conditions (0.2–2.0 ml O2 / 1.0 L H2O; Tyson and Pearson [1991]), which is a lower than suboxic reducing rank. V enrichment in sediments is a useful proxy for reducing depositional conditions, ranging from mildly to strongly reducing. V exists as and H2VO−4) under oxic soluble vanadate ionic species (e.g., HVO−4 2 water conditions. When conditions change from oxic to mildly reducing, V converts to tetravalent V(IV) state and forms the vanadyl ion (VO2+), related hydroxyl species (e.g., VO(OH)−3) and insoluble hydroxides. Under more strongly reducing conditions, further reduced trivalent V(III) precipitates as a solid oxide (V2O3) or hydroxide (V(OH)3)

2.06 1.96 0.05

1.6

4900

2.2

3900

0.24

700

0.053

160

18, 14

6800, 2700

1.8

2000

and/or geoporphyrin (Breit and Wanty, 1991; Wanty and Goldhaber, 1992). In the study section, V and VEF show increases in the siliceous claystone bed with black laminae (Bed 16) and the black claystone and overlying black chert and grey siliceous claystone (Beds 18–19). These increases in V coincide with increases in U, Mo and S related to redox changes (Figs. 3, 6), suggesting that reducing conditions occurred at that time. However, enrichments of U and Mo are not very high compared with typical cases of the modern sulphidic ocean (described in the following paragraphs). The redox conditions of V deposition in the study section would be within the mildly reduced condition.

22

S. Takahashi et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 432 (2015) 15–28

Fig. 4. Photographs of thin sections of example sedimentary rock samples with carbonaceous veinlets. In the thin section from Bed 20, carbonaceous materials (relatively dark in open nicols, bright colours in crossed nicols) are filled out into the siliceous sediments. On the other hand, in the thin section from Bed 19, such carbonaceous materials are less abundant.

U concentration can be used as a redox proxy because U is removed from seawater into sediments under reducing depositional conditions (Algeo and Maynard, 2004; Tribovillard et al., 2006). Under reducing conditions, soluble hexavalent U(VI) is reduced to insoluble tetravalent U(IV). In this state, U removal to the sediment may be accelerated by the formation of organometallic ligands and by the enhanced influence of organic substrates on U uptake. U reduction occurs under conditions close to those required for the conversion of Fe(III) to Fe(II). In the study section, UEF is constantly low or similar to the AUCC value (Fig. 3). Higher UEF overwhelming the AUCC value was detected from Bed 1 and Beds 14, 16, 18 and 19. As the high UEF in Bed 1 coincides with high MnEF, this increase should be caused by input of U-bearing material, such as Mn particles. High UEF in other horizons coincides with high V, Mo and S (described in the following paragraphs), suggesting them to reflect reducing water conditions. But its strength would not be so strong. The UEF in the modern anoxic–sulphidic basins overwhelms value of 5 (Algeo and Tribovillard, 2009) (Fig. 8). Considering this criteria, UEF increases up to 2.0 in the study section suggests the development of suboxic conditions (0.2–2.0 ml O2 / 1.0 L H2O; Tyson and Pearson, 1991) rather than anoxic conditions. Mo is present in seawater as molybdate (MoO2− 4 ), although, under anoxic conditions, MoO24 − may be released from organic matter through its decay by sulphate-reducing bacteria and reduced to Mo(V) (e.g., MoO2+) or Mo(IV) (e.g., thiomolybdate and MoOxS2− 4− x) species (Calvert and Pedersen, 1993; Zheng et al., 2000). Under hydrogen sulphide (H2S)-rich euxinic conditions, more rapid uptake of Mo by authigenic sulphides forming from free H2S in the water column becomes possible (Huerta-Diaz and Morse, 1992; Morse and Luther, 1999; Adelson et al., 2001). For this reason, Mo enrichment in sediments requires no O2 levels and high H2S levels in the depositional environment. In the study section, Mo and MoEF show increases in the late Spathian horizons of Beds 12, 16, 18 and 19, their magnitude being clearly beyond the range of analysing error (38% of absolute value of several ppm; Fig. 3). Similar increasing trends are recognised in the total S profile and occurrences of pyrite framboids (Fig. 6), also implying reducing conditions in which sulphate reduction would at least occur in the sediments. These coincidences imply Mo hosted in sulphide and/or

thiomolybdate. The first increase in MoEF in Bed 12 coincides with high S but not other trace elements such as U and V. A possible explanation of the independent Mo increase has been provided by Tribovillard et al. (2008). They argue that pyrite in sediment above the redox chemocline can capture only Mo, not U and V, resulting in selectively enriched Mo in the geologic record. In this case, the pyrite came from above the depositional depth, and the pyrite sink selectively absorbed Mo from pore water. Other Mo enrichment horizons around the late Spathian black chert coincide with increases in VEF and UEF. This combination suggests reducing seawater conditions at least at the sediment– water interface. However, the strength of the conditions would not have reached anoxic–sulphidic because the magnitude of the MoEF increase is not sufficiently high to interpret anoxic–sulphidic evidence. MoEF from typical strongly anoxic and sulphidic oceans such as the Black Sea and Cariaco Basin exceed 100–1000 because of high H2S (~20 μM; reviewed in Algeo and Tribovillard [2009]). MoEF values reaching hundreds to thousands in the geologic past have been interpreted as evidence for sulphidic conditions (Algeo and Maynard, 2004). Comparing these facts, the maximum MoEF reaching a value of 16 in the Mj section cannot be interpreted as evidence for anoxic–sulphidic conditions. A cross-plot of UEF and MoEF is shown in Fig. 8. Accepting the criteria based on modern seawater conditions (Algeo and Tribovillard, 2009), the three highest UEF and MoEF values in the Mj section (filled circles in Fig. 8) do not place within the area of the “open marine redox conditions” including modern suboxic–sulphidic cases. The maximum UEF values remain level with the suboxic conditions as described in the previous paragraph, and the MoEF values place towards the “active Mo particulate shuttle”, whereas Mo-hosted material originates from the water column. Considering these trends, the development of reducing conditions in the end-Spathian black-coloured horizons remained suboxic in the depositional area, did not reach anoxic–sulphidic, and additional Mo was derived from the water column. Although Algeo and Tribovillard, 2009 gave the active Mn oxyhydroxide cycle as an example of Mo transportation, the most likely material of the additional Mo in the study section is pyrite framboids, because these two materials show clear coincidences at Beds 12, 16 and 18. Synthesis of pyrite framboids in the water column (in fact, these pyrites are of small

23

5.4 4.0 5.7 – 4.0 – 4.1 8.4 5.1 4.8 5.5 4.9 – 5.7

6.0

10 4 20 – 4 – 8 30 12 20 40 40 – 40

6

4 4 2 – 4 – 2 2 2 2 2 2 – 2

6



μm μm μm

1.6 – – – – – – – 1.6 2.1 3.1 2.1 – 3.8

Al

Non-hydrothermal Bed12

Bed53

Total number of framboidal pyrite

Bed20

99 5 21 0 2 0 31 12 112 259 295 734 0 143

Bed4

Hydrothermal

1

Minimum Maximum Average Standard diameter diameter diameter deviation

S. Takahashi et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 432 (2015) 15–28

Mn

5.6 0.0002 0.0002 0.0029 16.4

0.0056 0.0005 0.0011 0.0000 0.0001 0.0000 0.0061 0.0011 0.0520 0.0518 0.4683 0.4647 0.0000 0.0859 0.0090 0.0005 0.0021 0.0000 0.0001 0.0003 0.0070 0.0081 0.0637 0.0707 0.8853 0.6231 0.0000 0.1788 0.0092 0.2140 0.0105 0.1159 0.0123 0.0651 0.1689 0.0286 0.3293 0.0341 0.3666 0.1626 0.1606 1.4244 39.6 26.8 45.6 27.0 55.4 29.6 11.4 17.8 5.2 11.4 1.4 3.6 17.4 4.0

% mm2

8 Mj 87

46 40 34 31 29 28 26 23 22 19 18 16 14 12 15 15 15 15 15 15 15 13 13 11 11 10 10 9 Mj 128 Mj 122 Mj 116 Mj 113 Mj 111 Mj 110 Mj 108 Mj 105 Mj 104 Mj 100 Mj 99 Mj 97 Mj 95 Mj 93

Takahashi et al., 2009a

6

Lithology

Muddy chert Chert Siliceous claystone Chert Muddy chert Chert Muddy chert Siliceous claystone Siliceous claystone Black chert Black claystone Siliceous claystone Siliceous claystone Chert with black laminae Siliceous claystone

Areal % of euhedral pyrite on thin sections Observed area of thin section

%

%

30.6 0.2 8.6 0.0 0.8 0.0 3.5 2.9 13.2 49.5 37.4 59.1 0.0 5.4

Fig. 5. Triangle diagram of Al–Fe–Mn composition for estimating the formation environment of the study samples. Areas are after Adachi et al. (1986).

Bed number Bed group number of Sample ID

Table 2 Areal % of Pyrite and size distributions of pyrite framboids of the Momotaro−Jinja section.

Areal % of framboidal pyrite on thin sections

Areal % of small framboidal pyrite (less than 6 μm) on thin sections

Ratio of small framboidal pyrite against other pyrite

Fe

diameters, and possibly derived from the water column; this is described in the following section) and/or Mo capture from porewater by framboids during early diagenesis proposed by Tribovillard (2008) can function in the Mo increase, in addition to their accumulation under suboxic depositional conditions. 5.3. Occurrences of pyrite framboids Pyrite is one of the minerals that indicate somewhat reducing conditions at the time of deposition and diagenesis. Wignall et al. (2010) and Bond and Wignall (2010) associate pyrite occurrences with redox conditions. Pyrites are absent in fully oxic conditions (8.0–2.0 ml O2 / 1.0 L H2O); pyrites including euhedral pyrite and pyrite framboids appear from around the upper dysoxic conditions (0.2 ml O2 / 1.0 L H2O) to the anoxic–sulphidic conditions (no O2 with H2S). The size distribution of pyrite framboids provides information on water mass redox conditions (Wilkin et al., 1996). When the redox interface between the oxic and sulphidic water masses locates within the water column, pyrite framboids with a diameter of a few micrometres sink rapidly to the sediment because of their high density. In contrast, when pyrite framboids form at a redox interface within the sediment, they can grow to a larger size due to the support of the sediment substrate. Thus, small pyrite framboids with a narrow size range (low standard deviation) indicate euxinic water presence in the water column (Wilkin et al., 1996; Wilkin and Barnes, 1997; Wignall and Newton, 1998; Bond and Wignall, 2010; Wignall et al., 2010). Bond and Wignall (2010) provide the concrete criteria for redox reproduction by pyrite framboid occurrences, average diameters less than 6.5 μm and standard deviation less than 2.5 account for sulphidic water-column conditions. In contrast, a 6–10 μm average diameter and standard deviation of 2.5–8.0 are features of suboxic–dysoxic boundary conditions (around 0.2 ml O2 / 1 L H2O, based on Tyson and Pearson [1991]). In the study section (Fig. 6), euhedral pyrite is common in all observed horizons denying oxic conditions, as the Mn-proxy suggested previously. Abundant pyrite framboid occurrences are detected from the uppermost Spathian horizons of Beds 12, 16, 18, 19 and 22, and an Anisian horizon of Bed 46. All of these samples show small average diameters less than 6.5 μm (4.8–5.7 μm). Beds 12 and 18 show relatively higher standard deviations because of 40-μm pyrite framboid grains which are close to suboxic–dysoxic trends. Beds 16, 19 and 22 with low standard deviations (1.6–2.1) correspond to sulphidic trends. Abundant pyrite framboids with greater standard deviations coincide with MoEF increases in Beds 12, 16, 18 and 19 (Fig. 6). As explained in

24

S. Takahashi et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 432 (2015) 15–28

Momotaro Jinja section (Japan, the Panthalassa, deep sea) m

total S

MoEF 1

10

(%) 100

δ34Ssulfide (VCDT‰)

0.1 1 10 -40

-30

50 40

Pyrite

30

(area % on thin sections) 20

-20 0

0.5

1.0

1.5

n = 99 average = 5.4 µm std=1.6

Size distributions of pyrite framboids

anoxic-sulphidic trend

(%)

3

2 4 6 8 10 12 14 16 18 20 (µm)

Anisian H. gifuensis

Middle Triassic

48

50

45

few framboids

40

15

few framboids few framboids

30

few framboids few framboids few framboids few framboids

2 28

14

30

n = 112 average = 5.1 µm std=1.6 anoxic-sulphidic trend

(%) 0

2 4 6 8 10 12 14 16 18 20 (µm) 50 40 30 20

21

euhedral 13 20

P. nakatsugawaensis

Lower Triassic

Bed 22

10

35

Upper Olenekian (Spathian)

40

20 36

31

1

10 0

Bed 46

framboidal

12 11 19 18

no framboids

(%)

0

30 20

9 11

n=259 average = 4.8 µm std = 2.0 anoxic-sulphidic trend

10

50 40

10

Bed 19

2 4 6 8 10 12 14 16 18 20 (µm)

Bed 18 n = 295 average = 5.5 µm std = 3.1 disoxic-suboxic trend

10 (%) 0

7

8

few framboids

6 5 4 3 2

0

7

30 20

1

1

10

Mo 6

-0.5

50 40

continue

Bed group number of Takahashi et al., 2009

Bed16 n = 734 average = 4.9 µm std = 2.1 disoxic-suboxic trend

10

100 Legend

(ppm)

2 4 6 8 10 12 14 16 18 20 40(µm)

chert (dark greenish grey) black chert muddy chert (dark greenish grey) siliceous claystone (dark greenish grey) black claystone black laminae

(%) 0

50 40 30 20

2 4 6 8 10 12 14 16 18 20 (µm)

Bed 9 n = 143 average = 5.7 µm std = 3.7 disoxic-suboxic trend

10 (%) 0

2 4 6 8 10 12 14 16 18 20 40 (µm)

Fig. 6. Lithologic column and vertical plots of Mo concentration and its enrichment factors (MoEF), total sulphur (total S), sulphur isotope ratio of sulphide (δ34Ssulphide), and areal ratio of euhedral pyrite and pyrite framboids observed in thin sections, and size distributions of pyrite framboids. In the diagrams of size distributions of pyrite framboids, “n”, “average”, and “std” are the observed number of pyrite framboids, average diameter of pyrite framboids and standard deviations of the diameters, respectively.

the previous section, increased MoEF values are lower than in the case of the typical sulphidic conditions. These facts are consistent with the wide size distribution of pyrite framboids in Beds 12 and 18, suggesting a suboxic–dysoxic trend. However, the narrow size distribution in Beds 16, 19, 22 and 46 suggesting sulphidic conditions conflicts with trace elemental features of UEF, VEF and MoEF, which are clearly lower than sulphidic levels in Beds 16, 19, 22 and 46 (Figs. 3 and 6). A possible explanation of these trends is that sulphidic water develops high in the water column far from the bottom depositional area. A similar interpretation was provided by Algeo et al. (2011) on redox conditions in the Griesbachian pelagic Panthalassa inferred from low redox-sensitive elemental compositions with small pyrite framboid occurrences. As mentioned, the oxygen-minimum zone would have developed high in the water column over the suboxic bottom-water conditions, where small pyrite framboids would be synthesised during the late Spathian to early Anisian pelagic Panthalassa. Considering this possibility, horizons of abundant small framboids with a relatively wide size distribution in Beds 12 and 18 might also be mixtures of small pyrite framboids from a sulphidic water column and larger pyrite framboids from suboxic sediments. If so, a MoEF higher than UEF in these horizons, which implies an additional Mo supply by pyrite framboids to the suboxic depositional conditions (Fig. 8; described in former section on MoEF), could be consistently explained.

5.4. Sulphur isotope ratio of sulphide Sulphur isotope ratios in sedimentary rocks provide information related to the anaerobic sulphate-reducing process (e.g., Canfield, 2001). In anoxic water conditions with enough nutritious organic matter, bacterial sulphate reducers thrive and accelerate the production of 32 S-enriched sulphide (H2S), causing isotopic fractionation (− 20 to − 40‰; Canfield et al., 1998; Canfield, 2001). In the case of sulphateunlimited situations, such as a sulphidic water column in the open ocean, a greater proportion of isotopic light sulphur from the oceanic sulphate reservoir is included in the sulphides. In a sulphate-limited condition, such as diagenesis in closed sediment, isotopic disproportionation could occur between initially reduced sulphide with a light isotope ratio and residual heavy sulphide. Song et al. (2014) suggested that carbonate-associated sulphate (CAS) reflects seawater sulphate composition across the late Spathian and early Anisian, showing isotopic values of +30‰ at the late Spathian and decreased values of + 20‰ at the early Anisian. Adopting their results, all samples from the Mj section, ranging from − 40.7 to −19.8‰, are sufficiently fractionated, by −60 to 45‰, compared with ambient seawater sulphate recorded in CAS values (Fig. 6). These facts suggest that the main sulphide material accounts for sulphate reduction in an open system such as around the sediment–water interface and/or

S. Takahashi et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 432 (2015) 15–28

25

Fig. 7. Photographs of pyrite in the thin sections observed using the backscattered electron mode of a scanning electron microscope. (A) An example of highly frequent pyrite framboids (Bed). (B) An example of less frequent pyrite framboids (Bed 26). (C) Examples of euhedral pyrites (left) and small pyrite framboids (centre). (D) Closed small pyrite framboids (Bed 18). (E) A typical example of euhedral pyrite overgrowth around pyrite framboids (Bed 16). (F) fossil-shaped pyrite (Bed 18).

the contemporaneous water column. The analysed δ34Ssulphide values represent a mixture of sulphides from several different formation processes; variations in δ34Ssulphide were possibly controlled by the contribution ratios of syngenetic sulphide and diagenesis sulphide and the occurrence of repeated sulphate reduction around the chemocline (such as H2S-rich water and oxic water). Increases in pyrite framboids (more than 50% against other pyrites; Table 2) are detected in Beds 16, 18, 19. However, their δ34Ssulphide did not differ from neighbouring euhedral pyrite-dominated horizons (e.g., Bed 14). These trends suggest that sulphur isotopic fractionation during the formation of these pyrite framboids would not be more significant than the process of euhedral pyrite formation in early diagenesis. It suggests that relatively reduced water conditions at the end Spathian did not reach the welldeveloped sulphidic condition, whereas sulphate reduction is repeated. Drastic decreases in δ34Ssulphide by ca. 20‰ occurred at the uppermost Spathian and earliest Anisian of Beds 22, 23, 28 and 31 (Fig. 6). Sulphides in these horizons are dominated by euhedral pyrite with no

or little pyrite framboids (Fig. 6; Table 2). Because of the lack of pyrite framboids, which possibly have light δ34Ssulphide, largely isotopic disproportions of sulphide in diagenetic processes are implied as causes of these low δ34Ssulphide values. Further understanding of the sulphur cycle across the Spathian–Anisian transition requires isotopic and chemical features isolated from sulphide minerals from each δ34Ssulphide features. 5.5. Anoxic water development at the Spathian–Anisian transition in the central Panthalassa and other oceanic regions Summarising the previous sections, the geochemical redox proxies indicated that the depositional redox conditions of the late Spathian to early Anisian pelagic Panthalassa were mostly dysoxic (0.2–2.0 ml O2 / 1 L H2O; Tyson and Pearson, 1991), as inferred from low Mn, V, U, Mo and pyrite presence. At the end of the Spathian, reducing conditions became relatively stronger suboxic conditions (0–0.2 ml O2 / 1 L H2O;

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MoEF 100

euxnic particulate

low Mo

anoxic

10

suboxic redox

1

0.1

1

10

UEF

Plot area of modern oceanic sediment (Algeo and Tribovillard, 2009; Tribovillard et al., 2012) redox

Open marine redox conditions

particulate

Active Mo particulate shuttle

low Mo

Low Mo concentration

Fig. 8. X–Y plot of MoEF and UEF. SW is the Mo/U weight ratio in modern seawater, which ranges from 3.0 to 3.1 (modern Pacific and Atlantic, respectively; Anderson, 1987; Bruland, 1983; Chen et al., 1986; Emerson and Huested, 1991).

Tyson and Pearson, 1991), as inferred from low Mn, higher V, U, Mo and large-sized pyrite framboids. Takahashi et al. (2009a) detected a sulphur-bearing organic compound (dibenzothiophene [DBT]; Fig. 3) and a high S/C ratio from late Spathian black chert from the same section (Bed 19). They postulated that DBT compounds were created in the presence of sulphur during oxygen-poor depositional conditions and/or reduced diagenesis. According to the suboxic conditions inferred by trace-element proxy information, it is more likely that the increased DBT in the black chert was formed by free sulphur produced during reducing diagenesis. Wignall et al. (2010) reported pyrite framboid occurrences based on sporadically sampled Spathian–Anisian deep-sea sediment from the Inuyama area of central Japan, including the study section. According to their report, most of the pyrite framboid occurrences show average diameters smaller than 6.5 μm and low standard deviation less than 2, suggesting a sulphidic water column from late Spathian chert beds and shale beds. These trends agree with results of this study in terms of pyrite framboid distribution. They also report large pyrite framboids (average of 8.1-μm diameter) and high standard deviation (3.1), from a Spathian black shale bed. This pyrite occurrence implies another development of suboxic bottom water during the Spathian. According to our carbon isotopic correlation, the late Spathian reducing bottom-water conditions coincided with carbon isotope minimum values, followed by an increase towards the early Anisian (Fig. 1). At a similar time interval, anoxic seawater evidence exists from several oceanic areas. High Mo/Al ratios have been reported from

the high-latitude boreal sea (Sverdrup Basin; Grasby et al., 2012) and the Palaeotethyan shallow platform (Saito et al., 2014). Their maximum values of Mo/Al accord with MoEF values of 48.6 and 118.93. These values are substantially higher than MoEF values from the Mj section of the deep-sea Panthalassa, despite its smaller water mass. These differences suggest that the oxygen-poor depositional conditions were weaker in the pelagic ocean and more severe in the shallow and continental margin oceanic regions. Size distribution patterns of pyrite framboids from the Paleotethyan isolated platform (Great Bank Guizhou) change from an oxic to a dysoxic trend at the late Spathian (Tian et al., 2014). These data also support simultaneous development of a reducing ocean. If the shallow-water and deep-seawater low-oxygen conditions were linked, oxygen consumption might have proceeded from the shallow and continental-sided oceanic regions. Sustainable mechanisms of large-scale oceanic anoxia require high organic matter production supported by continuous nutrient flux (Ozaki et al., 2011; Winguth and Winguth, 2011). If the late Spathian reducing ocean event was controlled by high surface productivity, the gradient of suboxic pelagic bottom-water conditions and relatively stronger continental-area anoxia is consistent with distance from the continental nutrient supply. In fact, the 87Sr/86Sr ratio as a continental weathering proxy peaks at the late Spathian (Korte et al., 2003; Sedlacek et al., 2014). However, it is not so clear if weathering intensity was high during the late Spathian. Because the most rapid increase in the 87Sr/86Sr ratio occurred during the Griesbachian to Smithian; i.e., earlier than the Spathian, the continuously increasing 87Sr/86Sr during the early Triassic (a constant increase from Griesbachian to Spathian) is considered to account for intense weathering until the late Smithian, earlier than the late Spathian (Sedlacek et al., 2014). According to δ18O of conodont apatite, decreased sea-surface temperatures from Smithian–Spathian maximum values increase again at the late Spathian, with timing similar to that of the 87 Sr/86Sr maxima (Romano et al., 2012; Sun et al., 2012). Warming climate could increase oceanic surface productivity. For instance, it might cause a change in oceanic circulation and patterns of nutrient supply by deep-water upwellings. The ultimate cause of this warming trend during the late Spathian is still uncertain. Although, the radiometric ages of the Siberian Trap volcanic activity range within Spathian and Anisian (Reichow et al., 2009), direct stratigraphic coincidence has not been confirmed to date. Therefore, more detailed and widespread information of oceanic redox conditions and associated environmental records would aid understanding of the mechanism of early Triassic oxygen-poor ocean development, an understanding essential for understanding the recovery of life after the mass extinction event. 6. Conclusion This study presented continuous δ13Corg, elemental composition and sulphide mineral occurrences and their isotopes for a pelagic deep-sea Spathian (upper Olenekian)–Anisian section. δ13Corg showed a 2–6‰ increase within the Spathian–Anisian transition. Because a 6‰ carbonatecarbon isotopic increase has been reported from this interval in the shallow Paleotethys, the carbon isotopic variation in the study section and in the Palaeotethyan sections are in accordance. Most horizons during the late Spathian–early Anisian show features of not fully oxic but dysoxic redox conditions, inferred from low Mn, U, V, Mo and euhedral pyrite-dominated occurrences. Conversely, in the end-Spathian blackcoloured beds and underlying siliceous claystone beds, relatively higher concentrations of redox-sensitive elements such as U, V, Mo and abundant pyrite framboids are detected. As enrichment factors of redoxsensitive elements are not much higher than the typical anoxic– sulphidic trend and large pyrite framboid occurrences are recognised, these trends suggest suboxic conditions rather than strong anoxic conditions. At similar end-Spathian timing, reducing seawater conditions have been also reported in continental margin sections. These coincidences imply global environmental perturbations that may have been

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related to the delayed recovery of life after the end-Permian mass extinction. To understand these coincidences and the global environmental perturbations related to the delayed recovery of life after the endPermian mass extinction, further studies are required. Acknowledgement This research was supported by the Japan Society for the Promotion of Science (JSPS#24740340). We greatly appreciate Akira Yao for helpful field information. Hideto Yoshida Aya Ohkubo help observation of thin sections by the electron microscopy. Thomas Algeo of University of Cincinnati and an anonymous reviewer provided helpful comments on the manuscript. Appendix A. Supplementary data Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.palaeo.2015.04.018. References Adachi, M., Yamamoto, K., Sugisaki, R., 1986. Hydrothermal chert and associated siliceous rocks from the northern Pacific their geological significance as indication od ocean ridge activity. Sediment. Geol. 47, 125–148. Adelson, J.M., Helz, G.R., Miller, C.V., 2001. Reconstructing the rise of recent coastal anoxia; molybdenum in Chesapeake Bay sediments1. Geochim. Cosmochim. Acta 65, 237–252. Algeo, T.J., Maynard, J., 2004. Trace-element behavior and redox facies in core shales of Upper Pennsylvanian Kansas-type cyclothems. Chem. Geol. 206, 289–318. Algeo, T.J., Tribovillard, N., 2009. Environmental analysis of paleoceanographic systems based on molybdenum–uranium covariation. Chem. Geol. 268, 211–225. Algeo, T.J., Chen, Z.-Q., Fraiser, M.L., Twitchett, R.J., 2011. Terrestrial-marine teleconnections in the collapse and rebuilding of Early Triassic marine ecosystems. Palaeogeogr. Palaeoclimatol. Palaeoecol. 308, 1–11. Algeo, T.J., Henderson, C.M., Tong, J., Feng, Q., Yin, H., Tyson, R.V., 2013. Plankton and productivity during the Permian–Triassic boundary crisis: an analysis of organic carbon fluxes. Glob. Planet. Chang. 105, 52–67. Algeo, T.J., Hinnov, L., Moser, J., Maynard, J.B., Elswick, E., Kuwahara, K., Sano, H., 2010. Changes in productivity and redox conditions in the Panthalassic Ocean during the latest Permian. Geology 38, 187–190. Anderson, R.F., 1987. Redox behavior of uranium in an anoxic marine basin. Uranium 3, 145–164. Ando, A., Kodama, K., Kojima, S., 2001. Low-latitude and Southern Hemisphere origin of Anisian (Triassic) bedded chert in the Inuyama area, Mino terrane, central Japan. J. Geophys. Res. Solid Earth 106, 1973–1986. Bond, D.P.G., Wignall, P.B., 2010. Pyrite framboid study of marine Permian–Triassic boundary sections: a complex anoxic event and its relationship to contemporaneous mass extinction. Geol. Soc. Am. Bull. 122, 1265–1279. Breit, G., Wanty, R.B., 1991. Vanadium accumulation in carbonaceous rocks: a review of geochemical controls during deposition and diagenesis. Chem. Geol. 91, 83–97. Bruland, K.W., 1983. Trace elements in sea-water. In: Riley, J.P., Chester, R. (Eds.), Chemical Oceanography vol. 8. Academic Press, New York, pp. 157–220. Calvert, S., Pedersen, T., 1993. Geochemistry of recent oxic and anoxic marine sediments: implications for the geological record. Mar. Geol. 113, 67–88. Canfield, D.E., 2001. Biogeochemistry of sulfur isotopes. Rev. Mineral. Geochem. 43, 607–636. Canfield, D.E., Raisewell, R., Westrich, J., Reaves, C., Berner, R.A., 1986. The use of chromium reduction in the analysis of reduced inorganic sulfur in sediments and shales. Chem. Geol. 54, 149–155. Canfield, D.E., Thamdrup, B., Fleischer, S., 1998. Isotope fractionation and sulfur metabolism by pure and enrichment cultures of elemental sulfur-disproportionating bacteria. Limnol. Oceanogr. 43, 253–264. Chen, Z.-Q., Benton, M.J., 2012. The timing and pattern of biotic recovery following the end-Permian mass extinction. Nat. Geosci. 5, 375–383. Chen, J.H., Edwards, R.L., Wasserburg, G.J., 1986. 238U, 234U and 232Th in seawater. Earth Planet. Sci. Lett. 80, 241–251. Emerson, S.R., Huested, S.S., 1991. Ocean anoxia and the concentration of molybdenum and vanadium in seawater. Mar. Chem. 34, 177–196. Galfetti, T., Bucher, H., Brayard, A., Hochuli, P., Weissert, H., Guodun, K., Atudorei, V., Guex, J., 2007. Late Early Triassic climate change: insights from carbonate carbon isotopes, sedimentary evolution and ammonoid paleobiogeography. Palaeogeogr. Palaeoclimatol. Palaeoecol. 243, 394–411. Goudemand, N., Romano, C., Brayard, A., Hochuli, P.A., Bucher, H., 2013. Comment on “Lethally Hot Temperatures During the Early Triassic Greenhouse”. Science 339, 1033-1033. Grasby, S.E., Beauchamp, B., Embry, A., Sanei, H., 2012. Recurrent Early Triassic ocean anoxia. Geology 41, 175–178. Huerta-Diaz, M.A., Morse, J.W., 1992. Pyritization of trace metals in anoxic marine sediments. Geochim. Cosmochim. Acta 56, 2681–2702.

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