Microbially induced sedimentary structures from the Mesoproterozoic Huangqikou Formation, Helan Mountain region, northern China

Precambrian Research 233 (2013) 73–92

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Microbially induced sedimentary structures from the Mesoproterozoic Huangqikou Formation, Helan Mountain region, northern China Zhong-Wu Lan a,b,∗ , Zhong-Qiang Chen c , Xian-Hua Li a , Kuino Kaiho d a

State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing, China Key Laboratory of Petroleum Resources Research, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing, China State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences (Wuhan), Wuhan, China d Institute of Geology and Paleontology, Tohoku University, Sendai, Japan b

c

a r t i c l e

i n f o

Article history: Received 6 November 2012 Received in revised form 3 April 2013 Accepted 5 April 2013 Available online 22 April 2013 Keywords: Microbially induced sedimentary structures Biogenicity Huangqikou Formation Mesoproterozoic North China

a b s t r a c t Seven types of microbially induced sedimentary structures (MISS): wrinkle structures, levelled ripple marks, organic carbonaceous laminae, microsequences, polygonal sand cracks, polygonal sand crack fills, and gas domes are documented from the Mesoproterozoic Huangqikou Formation of the Helan Mountain region, western part of North China. Facies analysis suggests that the MISS-hosting rocks were deposited in a supertidal to intertidal setting. Both wrinkle structures and gas domes are morphologically different from those previously documented from other ancient siliciclastic rocks, but probably formed by the same microbially mediated process as those published elsewhere. Other MISS resemble their ancient and modern counterparts in both morphology and size. The presence of carbonaceous laminae, coupled with laser Raman and carbon isotopic data (−25 ± 0.5‰), indicates a biogenic origin of the wrinkle structures. The absence of intercalating mudstone layers as well as different lithologic composition between polygonal sand cracks/crack fills and host rock indicate that both the polygonal sand cracks and polygonal sand crack fills likely resulted from dehydration and desiccation of organic rich microbial mats. The biogenicity of gas domes is supported by the different lithologic composition between infillings and wall. © 2013 Elsevier B.V. All rights reserved.

1. Introduction The Proterozoic witnessed the flourishing of global microbial communities (Walter, 1976; Knoll, 2003). The fossilized microbes, however, are seldom preserved due to their typically tiny size (Noffke and Paterson, 2008) and tendency to degrade (Jones, 2000). This is particularly common in siliciclastic rocks. As such, the role played by microbes in the formation of siliciclastic sedimentary structures has long been underestimated until the 1980s (Schieber, 1986; Gehling, 1986; Gerdes and Krumbein, 1987). Since then, siliciclastics predominantly occurring in peritidal depositional environments have been extensively and intensively investigated to provide better understanding of microbial roles in the formation of resultant microbial mats and preserved proxy structures in the clastic record (Hagadorn et al., 1999; Noffke et al., 2001; Prave, 2002; Pruss et al., 2004; Noffke, 2005; Schieber et al., 2007; Sarkar et al., 2006, 2008; Noffke, 2010; Lan and Chen, 2012, 2013).

∗ Corresponding author at: State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing, China. Tel.: +86 82998445. E-mail address: [email protected] (Z.-W. Lan). 0301-9268/$ – see front matter © 2013 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.precamres.2013.04.006

Microbial mats are believed to be formed by biostabilization, baffling, trapping, and binding with their ambient siliciclastic sediments (Noffke et al., 2003a). Various types of microbial mat features thus formed are termed “microbially induced sedimentary structures” (MISS sensus Noffke et al., 2001). The bio-physical processes generate a variety of within-bed and bed-surface structures such as bed-surface wrinkle structures, bed-surface levelled ripple marks, and within-bed organic carbonaceous laminae (Noffke, 2010). Microbial mats commonly grow in the settings where the deposition of clastic sediments was intermittent and the episodic nature of the sedimentation regime permitted adequate time for mats to grow and become established (Noffke et al., 2003a). While MISS are regarded as an ideal proxy for detecting ancient microbial life traces, caution must be exercised in interpreting the genesis of MISS-like structures as some purely physical structures without any microbiological interference are morphologically indistinguishable from MISS. As a result, several criteria have been developed to test the biogenicity of MISS and distinguish them from similar physical and/or chemical structures (Noffke, 2010). These criteria are generally categorized into external macroscopic morphological features and internal microscopic features. Wrinkled surface, loaded ripple marks, polygonal sand cracks/crack fills, and gas domes characterize macroscopic morphology of MISS that resulted from microbial mat growth, whereas carbonate or pyrite

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laminae, organic carbonaceous laminae, and negative ı13 C isotope ratios are typical of microscopic characteristics of MISS that indirectly or directly indicate the existence of microbial biomass (Schieber et al., 2007; Noffke et al., 2003b; Noffke, 2010). Although the Proterozoic MISS have been widely reported worldwide, most examples have been documented from the Neoproterozoic (Gehling, 1999; Noffke et al., 2002; Schieber et al., 2007; Noffke, 2010; Lan and Chen, 2012, 2013). Our knowledge of Mesoproterozoic and Paleoproterozoic MISS is limited solely to several reports from Africa and India (Eriksson et al., 2000; Sarkar et al., 2006; Parizot et al., 2005). Mesoproterozoic successions are well developed in North China and comprise siliciclastics and carbonates. In China some MISS have also been sporadically documented from the Mesoproterozoic successions in Henan, Jixian, Beijing, and Ningxia areas, North China (Shi et al., 2008a,b,c; Xing et al., 2011). Previously, these MISS have incorrectly been treated as trace fossils based solely on macroscopic morphologic analyses (Gao et al., 1993; Hua et al., 1993; Yang and Zhou, 1995; Qi, 2005; Zheng et al., 2009). When combined with microscopic internal microstructure analyses, most of, if not all, these supposed trace fossil structures are, however, actually seen to be MISS (Shi et al., 2008a,b,c; Xing et al., 2011). Although scanning electron microscope (SEM), laser Raman and other high resolution instruments have become routine methods for studying MISS, little research have been conducted on microstructures and genesis of Mesoproterozoic MISS. Accordingly, this paper aims to present a catalogue of MISS as represented by wrinkle structures, levelled ripple marks, organic carbonaceous laminae, microsequences, polygonal sand cracks, polygonal sand crack fills, and gas domes from the Mesoproterozoic strata of the Helan Mountain region, western part of the North China Craton, and to test their biogenicity mainly by microscopic microstructure analysis using SEM and other high resolution instruments.

2. Geological and stratigraphic settings On the basis of structural, petrological and geochronological data, the Archean to Paleoproterozoic basement of the North China Craton has been divided into three major tectonic blocks, i.e. the Western Block, the Eastern Block, and the Trans-North China Orogen (Zhao et al., 2005, 2010; Fig. 1a and b). The Western Block was further divided into the Yinshan Block in the north and the Ordos Block in the south (Fig. 1c). The Helanshan region is situated in the western Ordos Block. Exposed rock components in the Helanshan region include the Paleoproterozoic khondalites which underlies unconformably the Mesoproterozoic unmetamorphosed sedimentary cover of the Huangqikou Formation (Fig. 1d). MISS documented below were collected from the Mesoproterozoic successions of the Huangqikou Formation exposed at the Baisikou section (38◦ 44 58.68 N, 105◦ 55 18.34 E), which is located about 50 km northwest of Yinchuan, capital city of Ningxia Huizu Autonomous Region (Fig. 1e). In Baisikou, the MISS-bearing strata are assignable to the lower Huangqikou Formation (Fig. 2; Fig. 3a). Geographically, the studied section belongs to the Helan Mountain region (Fig. 1), which once was an aulacogen during the Proterozoic period (Lu et al., 2008). Following the formation of the North China Craton during the Lüliangian Orogeny at 2.0–1.8 Ga, the North China Craton underwent a series of extensional and rifting events at 1.8–1.6 Ga which induced the formation of aulacogens and marginal rift basins (Lu et al., 2008). These aulacogens accommodated onshore shallow marine deposits during a regional transgression (Qiao and Gao, 2007). The sedimentary rocks in the Helan Mountain region are dominated by shallow marine siliciclastics and carbonates. They are assigned to, in ascending order, the Mesoproterozoic Huangqikou,

Minjiagou, and Wangquankou Formations and the Neoproterozoic Zhengmuguan Formation (Wang et al., 2002; Figs. 1 and 2). Underlying disconformably the Huangqikou Formation, the Zhaochigou Group of Paleoproterozoic age is composed of quartz schist and leptynite. A SIMS zircon U-Pb age of ca. 1.95 Ga has been obtained from this unit (Dan et al., 2011). The Huangqikou Formation is dominated by peritidal siliciclastic rocks (Hua et al., 1993; Shi et al., 2008a). Its lower portion comprises purple, greyish quartzite, quartzitic sandstone intercalated with subsidiary variegated silty slate (Fig. 2); the middle portion is composed of coarse quartz sandstone, silty slate, and glauconite quartz sandstone; the upper part consists of cherty dolomite and grey siliciclastic rocks. The quartz sandstone yields some problematic structures on bedding surface interpreted by Hua et al. (1993) as trace fossils. The silty slate contains microfossil plants such as Taeniatum crassum, Leipsophosphaera sp., and Trematosphaeridium sp. Glauconite quartz sandstone gave a K–Ar age of 1291 Ma (Chen et al., 1999). In Baisikou, the Huangqikou Formation is about 180 m thick (Fig. 2), but it varies greatly in terms of lithofacies and thickness from place to place around the Helan Mountain region. The Wangquankou Formation, whose name originated from the previous “Wangquankou Group” (NBGMR, 1990), is dominated by dolomite and cherty limestone with minor quartzite, siltstone, and calcareous slate at its lower part. Glauconite from this unit gave a K–Ar age of 1289 Ma (Lu et al., 2002; Fig. 2). Stromatolite biostratigraphy suggests that the Wangquankou Formation (cf. “Group”) is subdivided into the Minjiagou and Wangquankou Formations. The former unit, about 200 m thick, yields stromatolites Conophyton garganicum, Colonnela sp., and Gaoyuzhuangia sp., while the latter formation contains Pseudogymnosolen sp., Lochmecolumnella sp., and Pareacolonnella sp. (Hua and Qiu, 2001). These two units correlate with the Gaoyuzhuang and Wumishan Formations of the Jixian area, Yanliao aulacogen, North China, respectively (Hua and Qiu, 2001). The above stromatolite biostratigraphic correlation is also strengthened by the similar lithologies and abundant MISS in both the Huangqikou Formation of the Helan Mountain region and the Dahongyu Formation of the Jixian area, Yanliao aulacogen (Shi et al., 2008a). A zircon TIMS U-Pb age of 1625 ± 6 Ma (Lu and Li, 1991), zircon SHRIMP U-Pb ages of 1626 ± 9 Ma (Gao et al., 2008), and 1622 ± 32 Ma (Lu et al., 2008) have been obtained from the volcanic interbeds within the Dahongyu Formation (Fig. 2). Thus, if the above Huangqikou and Dahongyu correlation is accepted, the Huangqikou Formation is of ca. 1.6 Ga age. Although a younger age (ca. 1.3 Ga) was reported from glauconite of the Huangqikou Formation (Chen et al., 1999), glauconite is vulnerable to diagenetic alteration (Obradovich, 1988), and thus can only produce accurate ages for Mesozoic and younger samples (Derkowski et al., 2009). The sandstone of the Huangqikou Formation has subjected to diagenesis and recrystallization. The K–Ar system in the glauconite therefore is highly likely to have been reset with the loss of certain radiogenic argon so as to give a younger ca. 1.3 Ga age. As a result, we consider that the depositional age of the Huangqikou Formation is probably 1.3–1.6 Ga.

3. Methodology The materials under consideration were studied in thin section and polished slab using optical microscope as well as SEM. Selected slabs with fresh surface were finely ground to expose the flat surface of longitudinal and/or cross sections, and cleaned ultrasonically with distilled water to get rid of surface pollutants. Gold coating was applied before SEM photography and EDS analysis. A Zeiss optical microscope equipped with a camera was used to

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Fig. 1. (a) Map showing the position of North China Craton and the study area. (b) Inset showing the three subdivisions of the North China Craton and the approximate position of (c). (c) Simplified geological map showing distributions of the Archean to Paleoproterozoic metamorphic rocks and granites in the Western Block of North China Craton. (d) Simplified geological map showing outcrops of the Paleoproterozoic granites and Mesoproterozoic–Phanerozoic sedimentary rocks around the study location. (e) Traffic map showing the location of the studied section in the Helan Mountain region, western North China. (a) Modified from Dan et al. (2011).

observe and photograph mineral composition of MISS, problematic structures, and their host rocks. A Leo1450VP SEM equipped with Inca Energy 300 energy dispersive X-ray analysis attachment was used to characterize the chemical compositional difference between the problematic structures and surrounding matrix. Both the optical microscope and SEM are housed at the State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences (SKLLE-IGGCAS). Samples for ı13 Corg analysis were crushed into powder and sieved with an 80–100 mesh. The sieved sample powder was dissolved with 5 N HCl overnight to remove carbonate. The decalcified samples (1–1.3 g) were then put into a centrifuge beaker, repeatedly washed with deionized water to make the solution approach neutrality. The neutralized samples were ground after drying them overnight in the heating oven and then put into the decalcified, hydrochloric acid; the removed samples (100–200 mg) + CuO wire (4 g) were then put into a quartz tube, and combusted at 750 ◦ C for 3 h. Three standards of GBW4405, GBW4407, and GBW4416

were used for correcting isotopic ratios. The organic carbon isotopic ratios were determined using cryogenically purified CO2 in a Finnigan MAT-252 mass spectrometer at SKLLE-IGGCAS, and reported in standard ı-notation as per mil (‰) deviation relative to Vienna Peedee Belemnite (VPDB). Analytical precision for ı13 Corg of the unknowns is better than ±0.06%. An invia-Reflex Laser Raman Confocal Microscope equipped with a digital pulse scanning near-field optical microscope based at the Technical Institute of Physics and Chemistry, Chinese Academy of Sciences (TIPCCAS) was used to identify organic carbon in the dark laminae of thin sections. A laser beam with a diameter of 1um and a wavelength of 633 nm was used as to excite the samples to collect the backscattered radiation. The laser beam was focused through a 100× ocular lens. The irradiation power used was is 8 mW. Surface laser powers of 1.5–2.0 mW were applied to minimize laser induced heating of the organic carbon. An accumulation time of 10 s was utilized to produce adequate signal-to-noise ratio of the spectra. A

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Fig. 2. (a) Composite columnar section showing the Proterozoic succession in the Helan Mountain region. (b) Logged section of part of the Mesoproterozoic Huangqikou Formation from the Baisikou area of Yinchuan, Ningxia Huizu Autonomous Region, western North China.

silicon wafer (1 1 1) was utilized to calibrate the instrument against the Raman signal of Si at 520 cm−1 . The spectral resolution and scanning range were 1 cm−1 and 1000–2500 cm−1 , respectively. Reproducibility of spectrum was better than ±0.2 cm−1 .

4. Facies analysis and paleoenvironmental interpretation Integration of physical sedimentary structures features such as sediment texture, composition, bedding surface morphology, and

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Fig. 3. (a) Field photo showing the unconformable contact between the Huangqikou Formation and Paleoproterozoic biotite granite. (b) Tabular cross-bedding in orange coarse sandstone. (c) Herringbone cross-bedding in reddish coarse sandstone. (d and e) Muddy drapes developed in purple coarse sandstone. Coin is 2.5 cm in diameter in (b) and (c), and 2 cm in diameter in (d) and (e).

cross-bedding structures (Walker and Plint, 1992) permits recognition of a total of seven facies from the lower part of the Huangqikou Formation. They are listed in Table 1 and categorized into two facies associations (Fig. 2).

4.1. Facies association 1 (FA 1) FA 1 is characterized by coarse sandstone at its base followed by a fining-up sedimentary package of coarse sandstone (Sc),

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Fig. 4. (a) Muddy flaser bedding in purple coarse sandstone. (b) Muddy rip-up clasts within purple coarse sandstone. (c) Symmetrical, straight-crested ripples recorded in orange medium sandstone. (d) Symmetrical, branching-crested ripple marks in reddish medium sandstone. (e) Asymmetrical ripples in purple fine sandstone. (f) Mud cracks in grey mudstone. Coin is 2.5 cm in diameter in (a), (b), (c) and (e) and 2 cm in diameter in (d) and (f).

quartzite (Q), tabular cross-bedded coarse sandstone (Sxc) (Fig. 3b), medium sandstone (Sm), herringbone cross-bedded medium sandstone (Sxm) (Fig. 3c), and fine sandstone (Sf) (Fig. 2). The strata of FA 1 are arranged into vertically stacked fining-upward packets with facies combination of Sc–Q–Sxc, Sxm–Sm–Sf and Q–Sf. The packets are up to 5.5 m thick, typically 2.5–3.5 m thick. Muddy rip-up clasts, muddy drapes, and muddy flaser bedding are present (Fig. 3d

and e; Fig. 4a and b). A sharp erosional base occurs at the bottom of the packet of Sc–Q–Sxc. Ripples are common throughout FA 1, typical of wavy ripple type (Fig. 4c and d). Symmetrical, straight, and round-crested ripples characterize the upper part, whereas the symmetrical, branching, and round-crested ripples typify the lower part. Sand cracks are common on the bedding surface of coarse and medium sandstone.

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Medium to thickly bedded (0.3–1 m)

Thinly to medium bedded (0.1–0.5 m)

Medium to thickly bedded (0.2–0.25 m)

0.5–2.5 m thick

Herringbone cross-bedded reddish medium sandstone, gradual base, gradual and sharp top. Sets are 3–6 cm high. Individual cross bed is 0.05–0.2 cm, and dips at an angle of 4–20◦ .

Massive orange or reddish fine sandstone. Gradual base and sharp top contact. Asymmetrical ripple marks (RI = 6; RSI = 7) are present. Some sandstone beds have recrystallized to form quartzite.

Herringbone cross-bedded orange and reddish fine sandstone, sharp or gradual base and top. Sets are 5–8 cm high. Individual cross bed is 0.3–0.9 cm, and dips at an angle of 8–12◦ .

Grey or purple tabular units. Mud cracks are present.

Herringbone cross-bedded medium sandstone (Sxm)

Fine sandstone (Sf)

Tabular cross-bedded fine sandstone (Sxf)

Mudstone (M)

Suspension settling of mud (George, 1994)

Migration of straight-crested bed forms (George, 1994)

Thinly to thickly bedded (0.09–1.2 m) Massive orange or reddish medium sandstone, sharp bases and tops. Muddy drapes and flaser bedding are present, 0.1–0.3 cm thick. Symmetrical branching ripple marks (RI = 4.5; RSI = 1.1) and symmetrical straight crested ripple marks (RI = 5.2; RSI = 1.2) are present. Some sandstone beds have recrystallized to form quartzite. Medium sandstone (Sm)

Wrinkle structures, gas domes, loaded ripples

Lower flow regime conditions generated by shoaling wave (George, 1994)

Bipolar migration of straight-crested bed forms (George, 1994)

Lower flow regime conditions generated by shoaling wave (George, 1994) Wrinkle structures, polygonal sand cracks/crack fills, gas domes

Migration of straight-crested bed forms (George, 1994) Medium bedded (0.13–0.25 m) Tabular cross-bedded reddish coarse sandstone, sharp, eroded base and top. Set is 15–35 cm high. Individual cross bed is 0.3–1.4 cm thick, and dips at an angle of 4–20◦ Tabular cross-bedded coarse sandstone (Sxc)

Upper flow regime conditions generated by shoaling wave (George, 1994) Polygonal sand crack/crack fills Medium to thickly bedded (0.1–0.7 m) Purple, grey or reddish coarse sandstone, gradual or sharp bases and tops. Muddy rip-up clasts are present up to 0.4–2.6 cm. Some sandstone beds have recrystallized to form quartzite. Coarse sandstone (Sc)

Interpretation of sedimentary process MISS Bed thickness Description Lithofaces

Table 1 Sedimentary facies recognized from the Huangqikou Formation. Lithofacies codes after Miall (1978).

4.2. Facies association 2 (FA 2) FA 2 is characterized by a fining-up package composed of fine sandstone (Sf), tabular cross-bedded fine sandstone (Sxf), and mudstone (M) (Fig. 2). When compared with FA 1, FA 2 is characterized by thinly bedded, finer sediments. Facies combination of FA 2 is composed of Sf, Sxf, and M. The packets are 2.5–5 m thick. Tabular cross-beds in Sxf are in sets of 0.2–0.25 m thickness. Asymmetrical ripples with flat crests are present on the bedding surfaces of the fine sandstone (Fig. 4e). Mud cracks are present on the top of the mudstone (Fig. 4f). 4.3. Paleoenvironmental interpretation The combination of fining-up package together with tabular cross-bedding, herringbone cross-bedding, muddy drapes, muddy flaser bedding, and muddy rip-up clasts in FA 1 suggests an upper intertidal to lower supertidal zone of a sandy beach setting. In particular, the association of muddy drapes, muddy flaser bedding and herringbone cross-bedding indicates bipolar migration of dunes or sand waves typically present in intertidal sandy deposits (Reineck, 1975; Ehlers and Chan, 1999; Tucker, 2003). Similar lithological associations with identical cross-bedding styles have also been recorded from other ancient intertidal to supertidal successions (Druschke et al., 2009; Lan and Chen, 2012). Symmetrical, straight or branching, round-crested ripples indicate migration of paleowaves at different sites where differential migration and erosion resulted in branching or coalescence of ripples, which commonly occur in a coastal rather than a fluvial system (Walker and Plint, 1992; Krassay, 1994). The presence of muddy rip-up clasts implies that the tidal currents were strong enough to have peeled off the muddy rip-up clasts from muddy sediments (e.g. Roux et al., 2004), and then mixed them into siliciclastic sediments. The combination of Sf, Sxf and M of FA 2 is typical of an intertidal setting where sandy flats gradually transitioned into muddy flats (Dalrymple, 1992). This interpretation is further reinforced by the wide occurrence of fine-grained siliciclastic sediments typical of the lower intertidal zone (Dalrymple, 1992; Poiré, 1993). The mud cracks in FA 2 indicate episodic exposure of muddy sediments in the intertidal to supertidal zone. As such, the Huangqikou Formation succession overall was most likely deposited in an intertidal to supertidal setting. The fining-up pattern in FAs 1 and 2 indicates episodic rise in sea level that possibly resulted from tide-influenced marine transgression. The initial transgression may have accounted for the formation of the erosional base at the basal Huangqikou Formation. 5. Major types of MISS and interpretation 5.1. Wrinkle structures Wrinkle structures are characterized by abundant pits and bulges preserved as positive epirelief on sandstone bedding surfaces. The pits and bulges are 0.3–1 cm wide and 0.05–0.4 cm deep (Fig. 5a–d). Wrinkle surfaces are somewhat weathered as indicated by dark materials in the low relief areas. Fresh rock surface appears to be either orange or reddish in colour. Thin sections cut perpendicular to bedding surface show that the wrinkled area is characterized by alternating elevations and depressions stretching like a sinusoidal curve (Fig. 5e). Further beneath the wrinkled surface are dark brown or grey laminae mainly composed of carbon (Fig. 6a). Carbonaceous laminae are smooth, 0.15–0.25 mm thick, nearly parallel to the bedding surface, and extend continuously laterally. No mutual cross-cuttings occur among individual laminae. One thick, dark lamina is surrounded by two thin laminae,

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Fig. 5. (a and c) Field photos showing abundant pits and bulges on reddish sandstone bedding surfaces. (b) and (d) are close-ups of (a) and (a), respectively. (e) Thin section view showing the wrinkled surface is composed of alternation of elevations and depressions. Coin is 2 cm in diameter.

0.05–0.1 mm in thickness. The latter are, however, less smooth, and show an unconformity with the bedding surface and zigzagging along its strike (Fig. 6b and c). Quartz grains are 0.1–0.7 mm in size, but concentrated in the size range of 0.2–0.45 mm. Quartz grains have a subangular to angular morphology, but are dominated by the latter. Large quartz grains loosely float in fine grains, exhibiting clearly a floating grain texture. Carbonaceous laminae are typically sandwiched between coarse and fine grain layers. Three microsequences (sensus Noffke et al., 1997) are also recognizable based on mineral compositions (Fig. 6c). Sequence 1 is composed mainly of medium quartz grains that increasingly graded upward into fine quartz grains and subsidiary clay minerals. The former grains occupy ca. 60% of whole rock, whereas the latter occupy ca. 35%. No carbonaceous laminae occur in the transition. Sequence 2 possesses similar texture to Sequence 1, but has a thin dark carbonaceous lamina separating medium quartz grains from fine quartz grains and clay minerals. Moreover, Sequence 2 saw a decrease in content of medium quartz and an increase in content of fine quartz and clay minerals, occupying ca. 30% and 65%, respectively. A thin dark carbonaceous lamina is pronounced (Lamina 1 in Fig. 6c). Sequence 3 is comprised of medium-fine quartz grains and clay minerals. Medium quartz grains occupy ca. 40% in the whole

rock, while fine quartz grains and clay minerals occupy ca. 55%. A thin dark carbonaceous lamina (Lamina 3 in Fig. 6c) is also distinct in Sequence 3, separating the medium and fine quartz grains from clay minerals. It should be noted that a thick, dark carbonaceous lamina (Lamina 2) form a distinct boundary between Sequences 2 and 3. Laminae 1 and 2 cross one another at an acute angle. Lamina 2 is nearly parallel to Lamina 3. Of these, Lamina 2 is composed of filamentous structures oriented differently (Fig. 6d–f). About 56% of filamentous structures are aligned with their long axis nearly parallel to the bedding surface with a dip angle of less than 10◦ . Filamentous structures having dip angles of 10–20◦ , 20–30◦ , and 30–50◦ make up 18.6%, 13.4%, and 12%, respectively of all observed filaments. Minor iron and titanium oxides are also present between the filamentous structures. The gold coated filamentous structures display intense carbon peak in the EDX spectrum (see inset in Fig. 6a). The Raman spectra from filamentous structures show vibrational bands at ≈1112 cm−1 , ≈1185 cm−1 , ≈1452 cm−1 , and ≈1606 cm−1 , respectively (Fig. 7). The approximately 1606 cm−1 vibrational band is distinct and indistinguishable from the standard Raman spectrum peak of organic carbonaceous materials at 1620 cm−1 . The band at approximately 1185 cm−1 is indistinguishable from organic

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Fig. 6. (a) Polished slab showing three prominent carbonaceous laminae beneath the wrinkled surface. Inset in the lower right corner is the EDX spectrum indicating a dominant carbon composition of the laminae. (b) Thin section view showing one distinct carbon lamina sandwiched between two thin carbon laminae. (c) Line drawings of B showing the lithologic composition of the carbonaceous lamina-hosting rock. (d) Dark lamina in Fig. 6B is composed chiefly of fossil elongate structures. (e) Line drawings of D showing elongated structures. (f) Rose diagram showing the dip angles of the elongate structures relative to bed surfaces.

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Fig. 7. (a) Photomicrograph showing the laser Raman spot position in the studied carbonaceous lamina. (b) Laser Raman spectrum showing Raman bands at 1112 cm−1 , 1185 cm−1 , 1452 cm−1 , and 1606 cm−1 .

polyene structures that make up the aromatic organic compounds (Nemanich and Solin, 1979; Allwood et al., 2006). The band at approximately 1452 cm−1 increases slightly relative to the band at approximately 1340 cm−1 of the standard organic carbon materials, whereas the band at approximately 1112 cm−1 declines slightly relative to the band at approximately 1185 cm−1 . The fresh carbonaceous laminae attain carbon isotopic values of −25.29 ± 0.49‰ to −25.29 ± 0.55‰ in two separate analyses. Thus, we consider −25 ± 0.5‰ as the approximate carbon isotopic value of the carbonaceous lamina. Filamentous structures in the carbonaceous lamina resemble, in morphology and size, filamentous benthic bacteria documented from other ancient and modern MISS (Noffke et al., 2003a, 2006a,b). We interpret them as fossilized filamentous bacteria that could have been partly replaced with inorganic minerals such as iron oxide. Although subjected to diagenesis, most of the organic carbon remains unaffected, as indicated by the laser Raman spectrum. Those filamentous bacteria by binding siliciclastic grains, formed a mat fabric that was similar to the organic mesh-like fabrics of modern microbial mats. The dominant, nearly parallel to bedding surface alignment suggests filamentous bacteria weaving microbial mats that experienced compaction in the presence of syndepositional load pressure (cf. Noffke et al., 2008). Some filamentous bacteria remained unaffected and remained dipping at a large angle to the bedding surface. Microsequences (Sequences I, II and III in Fig. 6c) resemble those recorded from sediment cores of modern mats covering tidal deposits (Noffke et al., 1997). Medium to fine quartz grain transition in the microsequences could indicate a hydrodynamic change from low energy/calm to high energy environment. The number of energy-changing events therefore can be indicated by the number of microsequences. Thus, three microsequences of the Huangqikou Formation reveal three small-scale, low to high energy changing events probably resulting from tide-induced regression–transgression. Organic carbon laminae in Sequences II and III occur at the low to high energy transitional boundary where MISS commonly occur (Noffke et al., 2006a). The subangular to angular morphology of quartz grains suggests moderate erosional

and oscillating depositional conditions which also favor the development of wrinkle structures (Noffke et al., 2002). The decrease of hydraulic energy permits microbial mat to form on the sedimentary surface. By baffling and trapping, the microbial mats are enriched in fine quartz grains and clay minerals, and thus form the microsequences. Carbonaceous Laminae 1 and 3 in Sequences II and III could not have been stylolites which are usually generated by pressure solution. First, Laminae 1 and 3 are composed mainly of carbonaceous materials with minor iron and titanium oxides. These mineral assemblages have been frequently documented from fossilized microbial mat laminae in peritital siliciclastic rocks (see Noffke et al., 2008 and references therein). Secondly, all carbonaceous laminae are laterally continuous and lack clay minerals, whereas stylolites are laterally discontinuous and result from tectonic overprint and are commonly enriched in clay minerals (Middleton, 2003). Thirdly, all carbonaceous laminae occur between medium and fine quartz grain layers, whereas stylolites typically occur in rocks with a homogeneous lithological composition (Golding and Conolly, 1962). Accordingly, most, if not all, carbonaceous laminae in Sequences II and III can be interpreted as degraded microbial mats. Carbonaceous Laminae 1 and 3 are much thinner than Lamina 2 (Fig. 6c), indicating lesser involvement of bacterial biomass in the mat fabric. Sequence 1 lacks organic carbon laminae, indicating no development of microbial mats. This is probably because the duration of seafloor exposure was inadequate to allow the development of microbial mat fabrics. As a result, only biofilms were produced enveloping single quartz grains. Large porosity between the quartz grains easily introduced oxygen for biofilms to be decayed prior to diagenesis. 5.2. Levelled ripple marks Inferred fossilized microbial mats cover the troughs of ripple marks (Fig. 8). Despite the development of microbial mats, some ripple troughs are still easily visible (Fig. 8a and b) while others are largely smoothed out and thus difficult to distinguish from ripple crests (Fig. 8c and d). The ripple troughs can be retrieved

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Fig. 8. Levelled ripple marks. (a) Microbial mat sealed ripples on orange sandstone surface. Note that these microbial mat features have evolved into erosional remnants and pockets due to later multiple erosion events. (b) Close-up of erosional pockets shown in (a). (c) Fragile mat chips covering on both the ripple crests and troughs. (d) Close-up of fragile mat chips. Coin is 2.5 cm in (c).

only after breaking and removing the covering inferred fossilized microbial mat-bound clastic sedimentary rock. Fresh rock surface is either orange (Fig. 8a and b) or reddish colour (Fig. 8e and f). On orange sandstone the levelled ripple marks are preserved as erosional remnants, 2–4 mm in thickness, and as pockets on the bedding surface (Fig. 8a and b) with ripple heights of 0.8–1.5 cm and widths of 4.5–5 cm. On reddish sandstone, the levelled ripple marks occur on both the ripple crests and troughs (Fig. 6e and f). The levelled ripple marks broke up into fragile chips of 1–3 mm thick when lightly hit by a hammer. The exposed ripples are 0.1–0.3 cm in height and 2–2.5 cm in width. The ripple crests are straight or slightly curved. Thin sections cut perpendicular to the bedding surface show the levelled ripple mark surface is composed of a series of alternating elevations and depressions (Fig. 9a and c). A lamina of 0.06–0.1 mm thickness occurs about 0.5 cm beneath the preserved bedding surface (Fig. 9b and d). EDX spectrum shows the lamina is dominated by carbon. Minor iron and titanium oxides are also associated with the carbonaceous lamina, which is laterally consistent and has a zigzagging appearance along its strike. The lithology beneath the carbonaceous lamina is dominated by medium quartz grains with subsidiary clay and accessory minerals. Quartz grains range from 0.2 to 0.5 mm in diameter with grains having size range of 0.25–0.35 mm being most common. Above the carbonaceous lamina, the medium quartz gradually grades into fine quartz and clay minerals. These quartz grains range from 0.1 to 0.3 mm in

diameter but those with grain size in the range of 0.13–0.25 mm occur most frequently. The transition from medium to fine quartz grains in microsequence (Fig. 9d) indicates a hydrodynamic change from high energy to low energy/calm conditions, probably resulting from localised tide-induced sea level changes (Noffke et al., 2006a). A similar microsequence has been recognized from sediment cores of modern mat covered tidal deposits (Noffke et al., 1997), and from the Mesoarchean Mozaan Group, Pongola Supergroup, South Africa (Noffke et al., 2003b). The presence of a carbonaceous lamina, coupled with the iron and titanium oxide mineral assemblage between the medium and fine quartz grain layers, is suggestive of a mat fabric resulting from degradation of filamentous bacterial microbial mat. Some ripple troughs and crests (Fig. 6a and b) are visible, possibly indicating a relatively thin mat overgrowth and concomitant trapped clastic sediment. Other ripple troughs and crests (Fig. 8e and f) are barely indistinguishable from one another, indicating a thick mat overgrowth. The preservation of a rippled surface as erosional remnants and pockets (Fig. 8a and b) indicates a differential erosional process of microbial mat and trapped sediment following their initial formation (Noffke et al., 2002). Similar microbial mats contouring ripple troughs as ornament-like wrinkle structures have been documented from the Late Neoproterozoic peritidal sandstones in Western Australia (Lan and Chen, 2012, 2013), and the

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Fig. 9. (a) Slab showing thin section cutting position relative to the inferred microbial mat covered surface. (b) Cross section view of the polished slab showing dark grey lamia beneath the bedding surface. (c) Wrinkled surface is composed of alternating elevations and depressions. (d) Carbonaceous lamina is sandwiched between fine and medium quartz grain layers. Inset on the left of (d) is the EDX spectrum showing a dominant carbon composition for the lamina.

Lower Arenigian (Ordovician) shallow marine sandstones of the Montagne Noire, France (Noffke, 2000).

5.3. Polygonal sand cracks They are preserved as negative epirelief on fine sandstone bed surfaces (Fig. 10a–c; Fig. 11a and b). The polygons include tetragonal, pentagonal, and hexagonal geometries which are each composed of curved to slightly curved margins (cracks). Polygon width varies from 0.9 cm to 2.4 cm. Crack width varies from 0.1 to 0.3 cm and remain constant along their elongation direction. They penetrate 0.45–1.1 mm below the bedding plane. The crack-bearing sandstone beds are 2–2.5 cm thick (Fig. 10a). No mudstone layers are present between the crack-bearing beds. In contrast, the polygonal cracks occur repeatedly on the successive thin bedded sandstone bed surfaces (Fig. 10a). Thin sections cut perpendicular to the bedding surface show that the crack displays a near flattened ‘U’ (Fig. 10d and e) or ‘U’ morphology (Fig. 11c and d). The size of quartz grains in the crack-bearing sandstone ranges from 0.08 to 0.5 mm, but is dominated by 0.12–0.26 mm sizes (Fig. 10e; Fig. 11d). Grain contacts are point contacts showing a floating grain texture (Fig. 10e) or point contacts and long contacts (Fig. 11d). Minor clay mineral grains and iron oxides are also visible in thin section. The ‘U’-shaped cross section of the cracks differs clearly from the v-shaped cross section of mud cracks (Allen, 1982). Further, the absence of mudstone layers and mudstone remnants in the polygonal crack-bearing sandstones imply that these polygonal cracks are less likely to have been desiccation mud cracks (Eriksson et al., 2007). The interconnected polygonal morphology also differs clearly from syneresis cracks of discontinuous or spindle-shaped morphology (Allen, 1982). Rather, the polygonal cracks were probably formed because of desiccation and subaerial shrinkage of microbial biomass (Lan and Chen, 2012). Microbes

interacted with siliciclastic sediment via their sticky mucilages and thus the organic-enriched microbial biomass contributed cohesion to detrital quartz grains. Upon subaerial exposure, a loss of cohesion between quartz grains resulting from microbial biomass decomposition allowed the sand cracks to form. Decay of organic material increased spaces between quartz grains thus making them loosely distributed (Fig. 8e). The cracks would probably have initially formed as ‘triple junctions’ at the early stage of their growth (Eriksson et al., 2007). Subsequent propagation, combination and alignment of ‘triple junctions’ were also crucial in the formation of polygonal cracks (Eriksson et al., 2007). The repeated occurrence of polygonal cracks in thin bedded fine sandstones (Fig. 10c) may suggest multiple and successive periods of colonization of microbial biomass (Noffke et al., 2006a).

5.4. Polygonal sand crack fills These pronounced polygonal ridges are preserved as positive epirelief on the upper surface of fine sandstone and they usually project 0.8–1.2 mm up from the bedding plane (Fig. 12a–d). Most ridges have been flattened. Each polygon is composed of 3–5 ridges. Polygon width varies from 1 cm to 4 cm. The host rock is massive fine sandstone, and is dominated by fine quartz grains with subsidiary clay minerals and iron oxides (Fig. 12e and f). The size of quartz grains ranges from 0.05 mm to 0.48 mm, but is dominated by the 0.13–0.25 mm sizes. The size of quartz grains in the ridge areas ranges from 0.07 mm to 0.27 mm, but is dominated by 0.15–0.26 mm. Grain contacts in either host rock or ridge areas are floating grain to long contact in texture. Thin sections cut perpendicular to the bedding surface show that the ridges have different lithological composition from the host rock (Fig. 12e and f). These inferred polygonal sand crack fills are morphologically similar to the polygonal sand cracks (Figs. 10 and 11). The minor

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Fig. 10. (a–c) Polygonal sand cracks preserved as negative epirelief on sandstone bed surface. (d) Slab cut through the cracks showing the cutting position of thin sections. (e) A nearly flattened ‘U’-shape morphology of the crack area. Coin is 2 cm in diameter in (a) and (b), and 2.5 cm in diameter in (c).

difference that we can observe is that the former are preserved in the form of positive ridges instead of being negative features. Absence of the laterally extending mudstone layers overlying the crack-bearing bed tops (Section 5.3) indicates that these irregular sand crack fills should not be related to sand-filled mudstone desiccation cracks (Eriksson et al., 2007). The different sandstone composition between the sand crack fills and host rock and absence of deformation of sandstone laminae distinguish these polygonal sand crack fills from petees/petee ridges (Noffke et al., 2002, 2008). The formation process of polygonal sand crack fills is analogous to

the polygonal cracks presented in Figs. 10 and 11. Desiccation of microbial mats initially caused a series of ‘triple junction’ cracks. Subsequent propagation, combination and alignment of ‘triple junctions’ facilitated the formation of polygonal cracks (Eriksson et al., 2007). Following deposition of another bed of sandy deposits by aqueous currents, the sand cracks were filled with siliciclastics which have the same composition with the overlying sandstone bed. Later, the sand crack fillings together with the original microbial mats bounding the cracks experienced flattening in the presence of

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Fig. 11. (a and b) Polygonal sand cracks preserved as negative epirelief on sandstone bed surface. (c) Slab cut through the cracks showing the cutting position of thin sections. (d) A nearly ‘U’-shape morphology of the crack area.

loading from overlying beds (Eriksson et al., 2007). Later erosion of the overlying sandy deposits and microbial mats left only the polygonal sand crack fills, as also inferred from the Mesoproterozoic Chorhat Sandstone in central India (Sarkar et al., 2006). 5.5. Gas domes The inferred gas domes are subcircular to circular in cross section (Fig. 13) and occur at different horizons. The gas dome component parts include relatively light coloured (mostly grey) infillings enclosed by a dark wall (Fig. 14). In cross section, the wall and infillings of inferred preserved gas domes, and host rock have the same relief, i.e. they are present at the same flat horizon. In

the longitudinal section, they exhibit near ‘U’-shaped morphology (Fig. 14). Representative measurements reveal that individual gas domes are 0.8–1.5 cm long, 0.4–1.6 cm wide, appearing to have a spheroidal to ellipsoidal morphology. Wall thickness of gas dome features ranges from 0.06 to 0.18 cm. The host rocks are composed exclusively of subangular to subrounded quartz grains. The size of quartz grains ranges from 150 to 430 ␮m, but is dominated by 300–400 ␮m. Some fine quartz grains have coalesced to form composite grains. Grain contacts are mainly concave–convex, suggesting the rock has been subjected to diagenetic compaction. The walls of inferred gas domes comprise finer angular to subangular quartz grains and clay minerals, which are composed geochemically of K, Fe, Si, Al and O. The size

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Fig. 12. (a–d) Polygonal sand crack fills preserved as positive epirelief on sandstone bed surface. (e) Slab cut through the crack-fills showing the cutting position of thin sections. (f) Different lithologic compositions/grain sizes between the sand crack fills/ridge area and host rock. Coin is 2 cm in (a) and (b), and 2.5 cm in (c).

of quartz grains ranges from 60 to 400 ␮m, but is dominated by 150–200 ␮m. Quartz grain contacts are floating or tangent, but are dominated by the former. In contrast, the infillings of inferred gas domes are composed of angular to subrounded quartz grains in the size range of 50–420 ␮m, but most of them fall in the range of 100–180 ␮m. Quartz grain contacts are long or tangential. Also present in infillings, in variably quantities and sizes, are clay mineral plus subsidiary amorphous silicon. The inferred gas domes are regular and circular in cross section, differing from weathering fronts which could only produce

irregular, uncertain structures (Seilacher, 2007). In addition, the fresh quartz-dominated mineral composition in petrographic images is inconsistent with weathering on burrow-bearing quartz sandstone. Accordingly, abiogenic structures produced from purely physical processes can not explain the gas domes. The gas domes could not have been animal burrows due to the age of the rocks in question, either. In any case, burrows are commonly characterized by vertical extending decorations and concave upward or convex downward laminae (Chakrabarti, 1990; Yuan, 1993; Liu, 2003; Qi, 2005), both of which are absent in our gas domes. The only

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Fig. 13. (a–d) Inferred gas domes are circular to subcircular in cross section. Note that each gas dome possesses a dark, clay rich wall entombing light coloured infillings. Scale bar is 0.5 cm and applicable to each figure.

similarity between the inferred gas domes and burrows is the presence of a wall dominated by clay minerals. The clay mineral-dominated wall of burrows formed by means of animals intentionally segregating and preferentially digesting and concentrating fine sand grains and clay minerals from coarse ones when making intrastratal trails (Clifton and Thompson, 1978). However, the infillings of the studied gas domes has the same quartz grain dominated composition as the host rock, only the wall being dominated by clay minerals. Therefore, we agree with Shi et al. (2008a) that these gas domes resulted from microbial mat influences. The inferred degradation of microbial mats from the Helan Mountain region was not accompanied by the formation of pyrite or ferroan dolomite, as widely recorded by others (Dornbos et al., 2007; Lan and Chen, 2012), but by clay minerals. This is probably because pore waters in sandy sediments were devoid of S anions but were rich in K, Fe, Al ions which determined the formation of dominant clay minerals during the degradation of the microbial mats. Following diagenesis, the mucus organics were decayed to precipitate clay minerals under localized anaerobic microenvironment (Petrovich, 2001). Worthy of note is that gas domes occurring on bed surfaces are characterized either by a radial surface morphology or by an elevated circular

margin (Pickerill and Harris, 1979; Dornbos et al., 2007), both of which are absent in our materials. However, the intrastratal occurrence of these gas domes hindered us from observing bed surface morphologies. 6. Discussions In the study area, the investigated MISS bearing siliciclastic successions of the Huangqikou Formation have experienced metamorphism of less than prehnite–pumpellyite facies (Shi et al., 2008a). MISS from the Huangqikou Formation mainly occur in fine sandstone beds, the siliciclastic grain of which is neither too coarse nor too fine and is favorable for the development and preservation of MISS. The MISS bearing successions were all deposited at turning points of regression–transgression where peritidal hydraulics predominates. Except for wrinkle structures and gas domes, all the other microbially induced sedimentary structures morphologically resemble their ancient and modern equivalents whose biogenicity has been well demonstrated (e.g. Noffke, 2010). Accordingly, all the prerequisites are in accordance with the biogenic criteria catalogued by Noffke (2009), and thus all MISS presented in this paper should be biogenic in origin.

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Fig. 14. (a) Hand specimen showing the occurrence of gas domes in different azimuths. (b) A gas dome was cut at various positions. (c) A slab photo showing the longitudinal section of the gas dome. (d) Cross section of a gas dome. (e) Different lithologic compositions between gas dome wall and infilling. Inset is the EDX spectrum showing the dominant K, Fe, Si, Al and O composition for clay minerals.

The biogenicity of gas domes is evidenced by its differential lithological composition between infillings and wall, as shown in similar studies (Gerdes et al., 2000; Dornbos et al., 2007). Many ancient and modern examples of wrinkle structures have been

suggested to be biogenic origin (e.g. Gerdes et al., 2000; Schieber et al., 2007; Noffke, 2010). A biogenic origin for similar biosedimentary structures preserved in the Huangqikou Formation is strengthened by the presence of organic carbonaceous laminae

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combined with laser Raman and negative carbon isotope shifts (−25 ± 0.5‰). The carbon peak from the EDX spectrum suggests the existence of carbonaceous materials in carbonaceous laminae. Typical organic carbonaceous materials commonly have two major standard Raman spectrum bands at 1620 cm−1 and 1340 cm−1 (Allwood et al., 2006; Potgieter-Vermaak et al., 2011). We interpret the distinct vibrational band at approximately 1606 to be associated with organic carbonaceous materials (cf. Kudryavtsev et al., 2001). The 1185 cm−1 band appears only in highly disordered materials, possibly resulting from C C bond stretching vibration in polyene or olefin-like organic compounds (Quirico et al., 2005), and has been attributed to a fossil-specific spectrum (Kudryavtsev et al., 2001). The slight increase of the band 1452 cm−1 from the band 1340 cm−1 could indicate the change of polyaromatic structures accompanied by loss of peripheral hydrogen during the conversion of organic carbon to inorganic carbon (Negri et al., 2002), a postulate which also applies to the band 1112 cm−1 . That the band 1452 cm−1 shows a larger width and a low intensity suggests the presence of abundant disordered carbon in organic materials. The original organic carbon experienced partial transformation into inorganic carbon during diagenesis and thus displays transitional band characteristics such as in bands 1452 cm−1 and 1112 cm−1 . It should be pointed out that inorganic carbonaceous materials such as graphite have only one 1580 cm−1 Raman band at first-order Raman spectrum and one approximate 2700 cm−1 Raman band at second-order Raman spectrum (Zinner et al., 1995). As such, the carbonaceous materials in the laminae were unlikely inorganic in origin, although the possibility that some organic carbon has transformed into inorganic carbon cannot be totally discounted. Instead, most carbon was likely organic in origin. The organic origin of the carbonaceous materials is also strengthened by carbon isotope evidence. The fresh unaltered carbonaceous laminae have 0.02–0.04 wt% carbon with a ı13 C value of −25 ± 0.5‰. Repeated measurements suggest that the total amount of measured carbon is positively correlated with the amount of fresh rock powders, which means this negative carbon isotope value is most probably not pollutants from the sample surface. Rather, the negative carbon isotope excursion is suggestive of biogenicity of Precambrian organic matter, which could indicate a photoautotrophic process in the inferred microbial mats (Schidlowski et al., 1983; Strauss et al., 1992). Worthy of note is that the ı13 C value of −25‰ just falls in the range of Mesoproterozoic siliciclastic sediments, the ı13 C values from organic matters of which range from −20‰ to −30‰ (Strauss et al., 1992; Des Marais et al., 1993). Comparable carbon isotope ratios and carbon abundances have also been documented from other wrinkle structures developed in peritidal siliciclastic successions, and were unanimously interpreted as organic carbon resulted from degradation of organic microbial mats woven by dominant filamentous bacteria (Noffke et al., 2003b, 2006a,b, 2008). As a result, we also interpret the ı13 C value of −25 ± 0.5‰ to represent a biological origin for these laminated microsequences in wrinkle structures from the Huangqikou Formation.

7. Conclusions An association of wrinkle structures, levelled ripple marks, organic carbonaceous laminae, microsequences, polygonal sand cracks, polygonal sand crack fills, and gas domes is documented from the Mesoproterozoic Huangqikou Formation, northern China. The typical fine grain composition and much lower metamorphism of the MISS bearing successions together with the presence of organic carbon and microtextures are consistent with the established biogenic criteria and thus all MISS presented in this paper should be biogenic in origin. Most of these microbially induced

sedimentary structures morphologically resemble their ancient and modern equivalents. The absence of intercalating mudstone layers and the different lithological composition between the polygonal sand crack/or sand crack fills and host rock indicate that both the polygonal sand crack and polygonal sand crack fills resulted from hydration and desiccation of organic rich microbial mats. The biogenicity of gas domes is supported by its differential lithological composition between infillings and wall. A biogenic origin for wrinkle structures is supported by the presence of organic carbonaceous laminae combined with laser Raman and negative carbon isotope shifts (−25 ± 0.5‰). Acknowledgments Constructive comments from Nora Noffke, an anonymous reviewer, and the handling editor Guochun Zhao have greatly improved the quality of this paper. We thank Saihong Yang and Xin Yan for assistance during SEM imaging. Junfang Zhao is appreciated for assisting in acquiring the Raman spectrum. We are also indebted to Jianhua Wang for making thin sections, and Hongwei Li for analyzing carbon isotopes. This work was supported by NSFC Grants (41173010 to XHL, and 41273069 to ZWL), China Postdoctoral Science Foundation Grant (2012M510537 to ZWL), and a grant in aid from the State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences (Wuhan) (grant number: GPMR200901 to ZQC). References Allen, J.R.L., 1982. Sedimentary structures, their character and physical basis. Part II. In: Developments in Sedimentology Series. Elsevier, Amsterdam, pp. 1–676. Allwood, A.C., Walter, M.R., Marshall, C.P., 2006. Raman spectroscopy reveals thermal palaeoenvironments of ca. 3.5 billion-year-old organic matter. Vibrational Spectroscopy 41, 190–197. Chakrabarti, A., 1990. Traces and dubiotraces: examples from the so-called Late Proterozoic siliciclastic rocks of the Vindhyan Supergroup around Maihar, India. Precambrian Research 47, 141–153. Chen, J.B., Zhang, P.Y., Gao, Z.J., Sun, S.F., 1999. Chinese StratigraphyMesoproterozoic. Geological Publishing House, Beijing, pp. 89 (in Chinese with English Abstract). Clifton, H.E., Thompson, J.K., 1978. Macaronichnus segregatis: a feeding structure of shallow marine polychaetes. Journal of Sedimentary Research 48, 1293–1302. Dalrymple, R.W., 1992. Tidal depositional systems. In: Walker, R.G., James, N.P. (Eds.), Facies Models: Response to Sea-Level Change. Geological Association of Canada, pp. 195–218. Dan, W., Li, X.H., Guo, J.H., Liu, Y., Wang, X.C., 2011. Integrated in situ zircon U–Pb age and Hf–O isotopes for the Helanshan khondalites in North China Craton: juvenile crustal materials deposited in active or passive continental margin? Precambrian Research, http://dx.doi.org/10.1016/j.precamres.2011.07.016, in press. ´ ˇ Derkowski, A., Srodon, J., Franus, W., Uhlík, P., Bana´s, M., Zielinski, G., Caploviˇ cová, M., Franus, M., 2009. Partial dissolution of glauconitic samples: implications for the methodology of K–Ar and Rb–Sr dating. Clays and Clay Minerals 57, 531–554. Des Marais, D.J., Strauss, H., Summons, R.E., Hayes, J.M., 1993. Proterozoic carbon cycle. Nature 362, 117–118. Dornbos, S.Q., Noffke, N., Hagadorn, J.W., 2007. Mat-decay features. In: chieber, J., Bose, P.K., Eriksson, P.G., Banerjee, S., Sarkar, S., Altermann, W., Catuneau, O. (Eds.), Atlas of Microbial Mat Features Preserved Within the Siliciclastic Rock Record, Atlases in Geosciences. Elsevier, Amsterdam, pp. 106–110. Druschke, P.A., Jiang, G.Q., Anderson, T.B., Hanson, A.D., 2009. Stromatolites in the Late Ordovician Eureka Quartzite: implications for microbial growth and preservation in siliciclastic settings. Sedimentology 56, 1275–1291. Ehlers, T.A., Chan, M.A., 1999. Tidal sedimentology and estuarine deposition of the Proterozoic Big Cottonwood Formation, Utah. Journal of Sedimentary Research 69, 1169–1180. Eriksson, P.G., Simpson, E.L., Eriksson, K.A., Bumby, A.G., George, L., Steyn, G.L., Sarkar, S., 2000. Mat roll-up structures in siliciclastic interdune beds of the ca. 1.8 Ga Waterberg Group, South Africa. Palaios 15, 177–183. Eriksson, P.G., Porada, H., Banerjee, S., Bouougri, E., Sarkar, S., Bumby, A.J., 2007. Mat-destruction features. In: Schieber, J., Bose, P.K., Eriksson, P.G., Banerjee, S., Sarkar, S., Altermann, W., Catuneau, O. (Eds.), Atlas of Microbial Mat Features Preserved Within the Siliciclastic Rock Record, Atlases in Geosciences. Elsevier, Amsterdam, pp. 76–105. Gao, J.H., Cai, K.Q., Yang, S.F., Li, Y.W., Lei, H.P., 1993. Discovery of the oldest trace fossils from the Changcheng Group, Jixian, North China. Chinese Science Bulletin 38, 1891–1895 (in Chinese with English Abstract). Gao, L.Z., Zhang, C.H., Yin, C.Y., Shi, X.Y., Wang, Z.Q., 2008. SHRIMP zircon ages: basis for refining the chronostratigraphic classification of the Meso- and

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