Palaeogeography, Palaeoclimatology, Palaeoecology 353–355 (2012) 73–86
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Palaeogeography, Palaeoclimatology, Palaeoecology
journal homepage: www.elsevier.com/locate/palaeo
Organic accumulation in the lower Chihsia Formation (Middle Permian) of South
China: Constraints from pyrite morphology and multiple geochemical proxies
Hengye Wei a, b, Daizhao Chen a,⁎, Jianguo Wang a, Hao Yu a, Maurice E. Tucker c
a
b
c
Key Laboratory of Petroleum Resources Research, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China
College of Earth Science, East China Institute of Technology, Nanchang 330013, China
School of Earth Sciences, University of Bristol, BS8 1RJ, England, United Kingdom
a r t i c l e
i n f o
Article history:
Received 26 January 2011
Received in revised form 30 June 2012
Accepted 3 July 2012
Available online 20 July 2012
Keywords:
Chihsia Formation
Middle Permian
Primary productivity
Redox state
Organic matter
Yangtze platform
a b s t r a c t
The organic-rich Chihsia Formation (Middle Permian) on the Yangtze platform, western Hubei, South China,
is considered to be the major hydrocarbon source rock in South China. The lower part of the Chihsia
Formation, reported here, was deposited in an intrashelf basin and is characterised by black laminated
marlstones/shales intercalated with dark-grey limestones/dolomites in which metre-scale shallowing (or
cleaning)-upward cycles are ubiquitous. Each cycle is composed of basal marlstone/shale and upper limestone. The organic matter is concentrated exclusively in the shaley bases of these cycles, and shows systematic, cyclic variations in abundance up through the succession. An integrated approach of pyrite morphology
and multiple geochemical proxies was used to determine the mechanism of organic accumulation. In the
organic-rich sediments, many redox proxies, including pyrite framboid size distribution, C–S–Fe relationships, iron speciation, and δ34S values of pyrite suggest deposition beneath the dysoxic bottom water,
which evolved into a more oxic water body when the organic-poor limestones were deposited. However,
the coincidence of increased abundance of total organic carbon (TOC) with increased Ba and Mo contents indicates that organic matter accumulation was controlled mainly by primary productivity. These data indicate
that increased organic accumulation resulted mainly from higher organic carbon export with increased nutrient flux during superimposed sea-level rises of different orders. This in turn could have increased the oxygen
consumption by respiration of microorganisms and organic matter decay, leading to the depletion of oxygen
in the water columns. Furthermore, the rapid sedimentation rate further limited bioturbation and organic
matter decomposition, facilitating the preservation of buried organic matter.
Crown Copyright © 2012 Published by Elsevier B.V. All rights reserved.
1. Introduction
Much attention has been paid to organic-rich sediments, not only
because of their potential as hydrocarbon source rocks, but also because of their time-specific occurrences at critical intervals in geological history. However, the mechanisms that control the enrichment of
organic matter in sediments remain debatable (e.g. Demaison and
Moore, 1980; Tyson and Pearson, 1991; Arthur and Sageman, 1994;
Wignall, 1994; Murphy et al., 2000; Hetényi et al., 2004). The debates
focus on whether the enrichment of organic matter is mainly
controlled by primary productivity (organic carbon flux) (Pedersen
and Calvert, 1990; Caplan and Bustin, 1998; Sageman et al., 2003;
Gallego-Torres et al., 2007), or by particular conditions allowing
preservation of organic matter (i.e. depletion of dissolved oxygen in
the water column and the sedimentation rate) (Demaison and
Moore, 1980; Canfield, 1989; Ingall et al., 1993; Arthur and
Sageman, 1994; Ingall and Jahnke, 1994; Arthur et al., 1998; Mort et
⁎ Corresponding author. Tel.: +86 10 82998112.
E-mail address: [email protected] (D. Chen).
al., 2007). The preservation model emphasises the importance of
dysoxia/anoxia in the water column as the cause of enhanced organic
accumulation. In contrast, the productivity model favours a higher
settling flux of organic carbon, caused by a high primary productivity
rate in the surface water, as the main control on organic accumulation. However, several lines of evidence show that the processes of
anoxia and high productivity are not necessarily mutually exclusive.
Although deposition of organic-rich sediment in shelf basins has
been extensively documented, little attention has been paid to the
organic-rich sediments deposited in intrashelf basins (Abdallah and
Meister, 1997; van Buchem et al., 2001; Hetényi et al., 2004; Lüning
et al., 2004) due to their limited occurrence in the stratigraphic
record. Thus, in general, the mechanisms of organic accumulation in
such a setting have not been well documented. The organic-rich sediments in the lower part of the Chihsia Formation in western Hubei
Province, South China, provide a useful example for understanding
the mechanisms of organic accumulation in an intrashelf basin
where the water column may have not been highly oxygen-depleted.
In organic-rich sediments, pyrite occurs in a wide spectrum of
morphologies and size distributions, which are commonly regarded
0031-0182/$ – see front matter. Crown Copyright © 2012 Published by Elsevier B.V. All rights reserved.
doi:10.1016/j.palaeo.2012.07.005
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H. Wei et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 353–355 (2012) 73–86
as good indicators of the redox state of the fluids from which it
formed (Wilkin et al., 1996; Wilkin et al., 1997; Wignall and
Newton, 1998; Wilkin and Arthur, 2001). In addition, geochemical
parameters such as carbon–sulphur–iron (C–S–Fe) relationships, Fe
speciation, δ 34S values of pyrite (δ 34Spy), and the abundance of trace
metals, are commonly used to assess the redox condition of the
water column and early diagenetic environment (Berner, 1978;
Jørgensen, 1982; Berner and Raiswell, 1983; Jørgensen, 1983;
Berner and Raiswell, 1984; Raiswell and Berner, 1985; Raiswell et
al., 1988; Canfield, 1989; Wignall, 1994; Raiswell and Canfield,
1998; Raiswell et al., 2001). On the other hand, the contents of TOC,
Ba and Mo are widely used as proxies for palaeoproductivity
(Brumsack, 1986; Schmitz, 1987; Stroobants et al., 1991; Dymond et
al., 1992; Shimmield, 1992; Francois et al., 1995; Paytan and
Kastner, 1996; Werne et al., 2002; Wilde et al., 2004; Challands et
al., 2009).
This study aims to assess which factors controlled the enrichment
of organic matter in the intrashelf basin of western Hubei during
Early Chihsian time. Pyrite morphology and multiple geochemical parameters (i.e. C–S–Fe relationships, Fe speciation and δ 34Spy values)
are used to determine the redox state of bottom waters. In addition,
the abundance of TOC and trace elements Ba and Mo are used as
proxies of primary productivity.
2. Geological and sedimentological setting
During the Early Middle Permian, a widespread tropical to
subtropical carbonate shelf developed on the Yangtze block, South
China. A deeper-water, intrashelf basin formed in an antecedent depression in western Hubei (Fig. 1) (Zhao et al., 1992; Feng et al.,
1997; Wang and Jin, 2000). The transition from the shallow-marine
carbonate shelf into the intrashelf basin (or sag) was generally gradual, through a carbonate ramp system influenced dominantly by
storm currents, rather than by gravity flows. Previous studies have
suggested a water depth of approximately 100 m in this intrashelf
basin (Zhao et al., 1992). Palaeogeographically, the studied section
(Figs. 2 and S1) within the intrashelf basin was located near Enshi,
in western Hubei Province (Fig. 1).
The lower Middle Permian, unconformably overlying Carboniferous strata after a long period of karstification and subaerial exposure,
is composed of a basal siliciclastic unit (Liangshan Formation) and an
overlying carbonate succession (Chihsia Formation) (Figs. 2 and 3).
The Liangshan Formation overlies the regional unconformity and is
dominated by pedogenic claystone with abundant rootlets, coal or
carbonaceous shale, and quartzitic sandstone and siltstone with
abundant vertical burrows, reflecting a littoral swamp/marsh environment. The overlying carbonate succession of the Chihsia Formation (150–200 m thick) is characterised by an organic-rich lower
section (~ 30 m thick) and a thicker organic-poor upper part. This
study focuses on the organic-rich lower part of the Chihsia Formation.
The relatively thin organic-rich lower part (~30 m thick) is
characterised by black calcareous shales/marlstones intercalated
with bioclastic limestones (mudstones/wackestones) which are locally dolomitized (Figs. 2 and 3). Calcareous shales are laminated and
contain rare bioclastic grains that originated from thin-shelled articulated brachiopods, foraminifers and ostracods. However, bioclastic
Fig. 1. Map shows lithofacies palaeogeography during the Chihsian time in South China and the location of the studied section (star), south of Enshi City in the intrashelf basin developed
in western Hubei Province. Inset box shows the global palaeogeography of Middle Permian. Note the location of South China block (SC; arrowed) and the studied section (star).
Modified from Feng et al., 1997; Wang and Jin, 2000.
H. Wei et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 353–355 (2012) 73–86
75
Fig. 2. Panoramic view for the studied section at Tanjiaba, south of Enshi City in western Hubei (30°14′12.70″ N, 109°28′30″ E). Note the vertical stratigraphic succession and stratal
cyclicity of organic-rich sediments in the lower Chihsia Formation. Car at the lower right for scale. Refer to the index map for the location of studied section (star, arrowed) during
the Middle Permian.
grains, mostly disarticulated, are common in the non-laminated carbonate intercalations, indicative of more intense bioturbation and/or
current agitation during deposition. This scenario indicates that the
black calcareous shales/marlstones were deposited in a relatively
quiet dysoxic water condition in contrast to the carbonate intercalations that were deposited in a more agitated and oxygenated water
setting.
Vertical facies variations within the 30 m-thick section show a
well-developed stratal cyclicity (Fig. 3). The basic stratal unit (cycle)
is composed of two components: the lower carbonaceous calcareous
shale/marlstone and the upper carbonate, which constitute a
metre-scale upward-cleaning (or shallowing) cycle. Such stratal cyclicity is commonly regarded as having been driven by short-term
orbitally-driven sea-level fluctuations (Goodwin and Anderson,
1985; Read et al., 1986; Goldhammer et al., 1987, 1990; Masetti et
al., 1991; Montañez and Read, 1992; Chen and Tucker, 2003). Two
such cycles vertically stack into a cycle set (CS) and two cycle sets further group into a larger composite cycle set (CCS) (Fig. 3). Thus, the
ratio of cycles to cycle sets, to a composite cycle set, is 4:2:1. Such a
hierarchy of stratal cyclicity suggests an orbital forcing control related
to the short (100 kyr), moderate (200 kyr) and long eccentricity
(400 kyr) rhythms (Borer and Harris, 1991; Osleger and Read, 1991;
Chen and Tucker, 2003). In the studied interval, four composite
cycle sets (CCS) are recognised, which constitute two third-order depositional sequences (Fig. 2). The lower sequence is composed of
CCS1 and CCS2, the upper sequence comprises CCS3 and CCS4 (Fig. 3).
3. Methods
Thirty samples were taken from a newly-cut roadside outcrop
south of Enshi City (Fig. 2). Approximately sixteen polished
thin-sections were prepared to examine the pyrite morphology and
measure the framboid diameter and area of pyrite using a scanning
electron microscope (SEM; JEOL JSM-5610LV). In terms of size, only
the pyrite framboids with a primary texture (i.e. uniform microcrystals and abundant intercrystal nanopores) were measured (Fig. 4A,
B; see Section 4) and more than 50 measurements were taken from
each sample (Table 1) to ensure reliability and comparability. Pyrite
framboids were easily distinguished by their shape and fabric, and
their size was measured directly from the SEM screen. In area measurement, in addition to the normal pyrite framboids, partially
recrystallized framboids (with their outlines or ghosts; Fig. 4C2–E1)
were included since they originated from the former. The proportion
of framboidal pyrite (or framboid ghost) to other types of pyrite, i.e.
overgrowth crystals upon framboids (Fig. 4D) and anhedral to
euhedral pyrite grains (Fig. 4E2, F), was determined by area measurement in SEM photomicrographs. The pictures were scanned at 600 ×
magnification. Typically, 30 pictures (~ 200 pyrite grains) were
taken for each sample. These pictures were then imported into
Adobe Photoshop to determine the proportion of pixels interpreted
to be pyrite either as primary framboids or diagenetic-altered crystals. The ratio of framboid pixels to pyrite pixels approximates the
proportion of pyrite as framboids (Zhang et al., 2009).
After washing and drying, the selected rock chips were crushed
and ground into powder in an agate mortar. Aliquots (200 mg) for
total organic carbon (TOC) were first treated with 10% (volume)
hydrochloric acid (HCl) at 60 °C to remove carbonate, and then
washed with distilled water to remove the HCl. Afterwards, the samples were dried overnight (50 °C) and then analysed using a LECO
CS-400 analyzer. Major elements, including total iron (FeT) in the sediment, were determined by XRF spectrometer on fused glass discs.
The precision and accuracy of the major element analysis are ≤3%
and 5% (2σ), respectively. Samples (powder) for trace element (including Ba and Mo) analysis were first digested in mixed acid
(HF + HNO3 + HClO4) at ~ 200 °C within pressure-tight Teflon cups,
then the dissolved solution was measured on a Finnigan MAT Element II mass spectrometer. Precision for trace elements is better
than 5%.
Pyrite was extracted using the chromium reduction method described by Canfield et al. (1986). The sample split was treated with
1 M CrCl2 solution and 6 N HCl + 10 ml alcohol. The hydrogen sulphide was immediately purged by the N2 stream, and was trapped
and collected in zinc acetate (ZnAc) solution, to which 2%
AgNO3 + 6 N NH4OH solutions were added to precipitate Ag2S
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H. Wei et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 353–355 (2012) 73–86
Fig. 3. Stratigraphic succession of the studied section (see Fig. 2 for a panoramic view). Note the depositional cycles and their stacking patterns and inferred sea-level changes.
Shaded areas correspond to the organic-rich intervals. S, shale; M, mudstone; W, wackestone; P, packstone. C, depositional cycles; CS, cycle sets; CCS, composite cycle sets.
which was then filtered, rinsed, dried and weighed. The reproducibility of replicate analyses is generally better than 1%. Pyrite sulphur (SP)
and iron (FeP) contents were calculated by stoichiometry from the
extracted pyrite. Afterwards, the dried Ag2S was mixed with
V2O5 + pure quartz sand and then converted to SO2 by combustion
with copper turnings at 1050 °C. The purified SO2 was used to analyse
the sulphur isotopic ratio in a Finnigan Delta-S mass spectrometer.
Sulphur isotopic ratios are expressed as standard δ-notation relative
to Cannon Diablo Troilite (CDT) standard. The analytical precision is
better than ±0.2‰.
Iron species were extracted using the dithionite method of
Canfield (1989) and the boiling HCl method of Berner (1970). The
amount of iron released from the sediments was determined by
atomic absorption spectroscopy (AAS) after dilution. Canfield and
Thamdrup (1994) concluded that dithionite quantitatively extracts
the iron oxide/oxyhydroxide phases (lepidocrocite, ferrihydrite, goethite, and hematite) with only relatively minor amounts extracted
from any iron silicates. The boiling HCl method also quantitatively extracts Fe from the same oxides but, in addition, iron is partially
extracted from silicate phases. Highly-reactive iron (FeHR) comprises
iron extracted by dithionite (FeD) and pyrite iron (FeP). Reactive iron
(FeR) comprises iron extracted by boiling HCl (FeH) and the FeP. The
difference between FeH and FeD represents an iron fraction which
reacts slowly with dissolved sulphide, and is referred to as
poorly-reactive iron (FePR). The difference between total iron content
(FeT) and FeP + FeH represents the remaining fraction of iron, mainly
in silicates, which is essentially unreactive and is expressed as FeU
(Raiswell and Canfield, 1998). The iron liberated in acid is assumed
to be potentially reactive toward sulphide, and the degree of
pyritization (DOP) of the sediment (Berner, 1970; Raiswell and
Berner, 1985; Raiswell et al., 1988) is given as:
DOP ¼ FeP =ðFeP þ FeH Þ:
4. Results
Pyrite morphologies in this study mainly include: (1) the normal
spherical framboid that consists of pyrite microcrystal aggregates
commonly with intercrystal nanopores (Fig. 4A, D); (2) clustered
H. Wei et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 353–355 (2012) 73–86
77
Fig. 4. SEM photomicrographs of pyrite morphology. (A) Normal framboids with uniform microcrystal and abundant intercrystal nanopores, ES31. (B) Clustered framboids, ES22.
(C) Irregular aggregation of pyrite microcrystals (arrow 1) and weakly recrystallized framboidal pyrite with a slight increase in crystal size and reduction in intercrystal nanopores
(arrow 2), ES25. (D) Annulated framboids with pyrite overgrowth annulets, and framboid clusters, ES31. (E) Partially recrystallized framboids with relict nanopores within
framboidal ghosts (arrow 1) to recrystallized pyrite spherules with framboidal ghost but an absence of nanopores (arrow 2), to a more intensely recrystallized euhedral pyrite crystal (arrow 3), ES39. (F) Irregular anhedral (arrow) to subhedral pyrite crystals, ES39.
Table 1
Pyrite content and statistics of framboid texture and framboid size distribution in the
sediments from Enshi section, western Hubei, South China.
Sample
no.
Pyrite Framboidal
(wt.%) pyrite
(%)
ES18
ES20
ES21
ES22
ES23
ES24
ES25
ES27
ES28
ES29
ES30
ES31
ES34
ES36
ES39
ES40
0.19
0.32
1.10
0.15
0.35
0.27
0.21
0.26
0.39
0.47
1.15
0.51
0.29
0.29
0.15
0.01
Mean MFD Standard Skewness Framboids
(μm) (μm) deviation
measured
10
8.0
34.5
5.9
2.8
74
36
44
47
70
72
64
69
34
41
27
60
59
37
18
6.5
10.7
7.6
7.2
7.8
6.6
9.0
7.5
6.2
9.4
7.8
8.4
8.2
17.7
30.2
17.9
20.6
21.4
24.7
38.6
18.3
13.4
40.3
17.6
28.6
30.8
2.6
7.7
2.6
2.7
3.4
2.8
6.3
3.1
1.7
6.3
3.1
5.0
3.3
1.9
1.2
1.3
2.2
1.8
3.7
3.3
1.6
2.0
2.4
1.1
1.9
4.4
88
78
94
85
95
133
88
89
83
95
89
88
76
MFD: maximum framboid diameter.
framboids (Fig. 4B, D); (3) irregular aggregation of pyrite microcrystals (Fig. 4C1); (4) annulated framboid with overgrowth rings
(Fig. 4D); (5) partially recrystallized framboid with minor internal
nanopores (Fig. 4C2, E1); (6) pyrite spheroid, a framboidal ghost
with no intercrystal nanopores (Fig. 4E2); (7) irregular anhedral to
subhedral mass (Fig. 4F); and (8) euhedral pyrite grain (Fig. 4E3).
Of these, the normal framboids or their aggregates (e.g. Fig. 4A, B)
are the most common. However, they were commonly subject to
variable diagenetic alterations in morphology, as indicated by the
increased microcrystal size and decreased intercrystal nanopores
(Fig. 4C2, E1), and/or the occurrence of overgrowth annulets
(Fig. 4D). In some cases, the primary framboidal forms were totally
obliterated (Fig. 4E3, F) in the course of prolonged and/or enhanced
diagenesis.
Size statistics of framboids indicate that their diameter is generally
smaller in black marlstones and larger in grey limestones (Figs. 5 and
6). The proportions of pyrite as framboid in the basal and top limestone
are less than 20%, whereas in the dark grey marlstones/shales values
are greater than 40%, except for those within the interval from 15 m
to 19.5 m, where the proportions are ~40% (Fig. 7; Table 1).
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H. Wei et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 353–355 (2012) 73–86
Total organic carbon (TOC) contents range from 0.2 to 1.33%
(mean 0.73%) (Fig. 7; Table 2). There are systematic variations in
TOC which follow the stratal cyclicity, with higher values in the
basal shaley portion and lower values in the upper limestone of
each cycle. Carbonate contents vary between 55 and 95% (mean
71%), and show a decrease from the base to upper part of CCS1,
followed by a long-term increase further upward. Pyrite sulphur
(Sp) contents are generally low, ranging between b0.01 and 0.59%
(mean 0.15%) (Table 2; Fig. 7). Accordingly, S/C ratios are also low,
ranging between b 0.01 and 1.76 (mean 0.33), particularly in the
black marlstones/shales (Table 2).
Fe speciation data are illustrated in Fig. 7 and Table 2. Total iron
(FeT) contents generally range between 0.10 and 1.0% (mean
0.41%), and relatively are higher in the lower parts of CCS1 and
CCS2. FeP contents range between b 0.01 and 0.54% (mean 0.14%).
FeR contents range between 0.01 and 0.75% (mean 0.27%). FeHR contents range between 0.02 and 0.55% (mean 0.2%), and are relatively
higher in the middle part of CCS2 (at depth of ~ 18 m). FeU contents
range between 0.01 and 0.47% (mean 0.16%), and are relatively high
at the base of CCS1 (Fig. 3). FeT/Al ratios range between 0.37 and
1.23 (mean 0.64) (Table 2), with most between 0.5 and 0.68; this is
slightly higher than that of the average shale (Taylor and McLennan,
1985). The FeHR/FeT ratios range between 0.17 and 0.84 (mean
0.56), and are relatively lower in the basal part of CCS1, that is opposite to the FeU variations, which indicates a decoupling between the
FeHR/FeT ratios and FeU contents. DOP values vary between 0.05 and
0.9, with most within the range of 0.4 to 0.9.
The δ34Spy values range between −17.0 and −6.1‰ (Fig. 7;
Table 2), with most being lower than −10.0‰ except for those in the
lowermost and uppermost limestones. Cyclic variations in δ34Spy are
prominent, and are generally lower in organic-rich intervals (i.e. in
the lower parts of the depositional cycles).
The concentrations of Ba range between 10 and 60 ppm (mean
27 ppm), and are generally higher in the lower part of the section
(CCS1 and CCS2), and decrease upward. Mo concentrations range between 0.3 and 4.5 ppm, with most greater than 1.5 ppm. Fluctuations
of Mo generally coincide with TOC variations, being higher in the black
marlstones/shales and lower in the dark-grey limestones/dolomites.
Moreover, the Mo abundance is higher in the lower parts of CCS1 and
CCS3 (Figs. 3 and 8).
Fig. 5. Box-and-whisker plots showing framboid-size variations in sediments at the studied section. The boxes extend from quartile Q = 0.25 to quartile Q = 0.75. The median value
is marked in each box. The lines extending to the left and right of each box delineate the minimum and maximum values, respectively.
H. Wei et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 353–355 (2012) 73–86
79
Fig. 6. (A) Plot of the mean vs. the standard deviation of the framboid size; (B) Plot of the mean vs. the skewness of the framboid size.
5. Discussion
5.1. Redox conditions
5.1.1. Pyrite morphology
Most sedimentary pyrite minerals form immediately below the
O2/H2S redox boundary. The location of this redox interface relative
to the sediment–water interface controls the texture and size of pyrite framboids. Raiswell and Berner (1985) distinguished between
the syngenetic pyrite that formed in the anoxic water column above
the sediment–water interface, and the diagenetic pyrite that formed
in sediments. In this study, the presence of clustered framboids and
recrystallized framboids suggests precipitation in anoxic pore waters
within the sediments (Fig. 4B–E1, 2); thus they are diagenetic in
origin.
Many studies (Wilkin et al., 1996, 1997; Wilkin and Arthur, 2001)
have shown that the size distribution of pyrite framboids in modern
sediments reflects the redox condition of the bottom waters.
Syngenetic framboids are relatively small in size, because of their
short residence and growth time in the water column prior to settling
onto the sediments beneath. In contrast, diagenetic framboids tend to
be coarser as a result of the prolonged time available for growth within the sediments. Thus, pyrite framboids that form in anoxic waters,
on average, are smaller and less variable in size than those that
form in sediments under dysoxic or oxic bottom waters.
The mean size of pyrite framboids which formed in anoxic water
columns is generally less than 4.7 ± 0.5 μm (Wilkin and Barnes,
1997; Wilkin et al., 1997; Wignall and Newton, 1998). In this study,
the mean framboid diameter is 7.3 μm in laminated marlstones/
shales, and 9.3 μm in limestones/dolomites (Fig. 5), both of which
are much greater than 4.7 μm, indicating a diagenetic origin. This scenario implies that the pyrite framboids formed in the reduced pore
waters within sediments below the oxic–anoxic interface. However,
those from the laminated marlstones/shales generally precipitated
from more oxygen-depleted fluids in view of their smaller size
(Fig. 5). Moreover, all of the maximum framboid diameter (MFD)
values are greater than 18 μm (except for sample ES30, 13.4 μm;
Fig. 5), indicating a much longer residence time for their growth,
which could not therefore have taken place in an anoxic or euxinic
water column (Wignall and Newton, 1998).
Cross-plots of the mean size of framboids vs. the standard deviation from the mean (Fig. 6A) and skewness (Fig. 6B) suggest that
most of the pyrite framboids formed within the sediments below
the dysoxic–oxic bottom water, except for one sample (ES30). This
generally agrees with the scenario derived from the data of MFD
and box-and-whisker plots (Fig. 5), as documented above. Therefore,
it can be concluded that the organic-rich sediments in the lower part
of Chihsia Formation were deposited mostly in a dysoxic, or oxic
water column.
Although the size parameters of one sample (ES30) fall in the
range of anoxic or euxinic waters, they fall very close to the threshold
for dysoxic conditions (Fig. 6B). On the other hand, the lower proportion of framboidal pyrites (~40%; Fig. 7) argues against anoxic or
euxinic conditions. Moreover, the mean diameter of framboids of
this sample is 6.2 μm (Table 1), much coarser than that from modern
anoxic or euxinic environments (4.7 ± 0.5 μm) (Wilkin and Barnes,
1997; Wilkin et al., 1997). Accordingly, it is likely that this sample
was not deposited under anoxic or euxinic conditions.
The proportion of pyrite present as framboids is more or less related to the redox conditions of the water column, higher in oxygendeficient waters and vice versa (Wignall and Newton, 1998). Higher
proportions of framboidal pyrite in CCS1 and CCS3 indicate
more oxygen-depleted waters during intervals of transgression of
sequences 1 and 2; this scenario points to an ultimate control of
sea-level fluctuations on the seawater redox state (Fig. 7).
5.1.2. C–S–Fe relationship
The pyrite abundance in sediments depends on organic flux, availability of reactive iron and dissolved sulphate in the water (Berner
and Raiswell, 1984). Normal-marine sediments deposited in oxic
waters generally show a positive correlation between TOC and SP contents, which defines a regression line passing through the origin
(Berner and Raiswell, 1984). By contrast, sediments deposited beneath
an anoxic water column are characterised by high ratios of sulphur to
carbon (S/C ratio> 0.36), so that the data are plotted above the
normal-marine regression line (Leventhal, 1983) with positive intercepts on the S-axis in S–C plots if the iron is persistently available, or
by fairly constant sulphur contents independent of carbon variations
if the iron is deficient (euxinic or sulphidic; Hofmann et al., 2000). In
this study, some samples yield S/C ratios along the normal-marine
regression line (S/C = 0.36), but most of them have S/C ratios lower
than 0.36, i.e. beneath the normal-marine line (Fig. 9A), suggesting
deposition in a spectrum from oxic to possibly dysoxic marine water
(Berner, 1984; Berner and Raiswell, 1984).
It is worthy to note that organic carbon and iron contents are often
correlated positively (Berner, 1970, 1984), partly due to the adsorption of fine clay particles to both colloidal organic matter and
iron-oxide particles. Deposition of more clay, with its consequently
higher surface area, entails more deposition of both organic carbon
and iron. However, in this study, the abundance of TOC is not related
to the FeR availability (Fig. 9C). Therefore, pyrite formation is probably limited by reactive iron, to some extent, rather than by organic
80
H. Wei et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 353–355 (2012) 73–86
Fig. 7. Vertical variations of TOC contents, cabonate contents, total Fe contents (FeT), FeT/Al ratios, pyrite Fe (FeP), high reactive iron (FeHR), unreactive iron contents (FeU), FeHR/FeT ratios, DOP values, pyrite sulphur contents (Sp), δ34Spy
values, and ratios of pyrite framboids in sediments from the Enshi section. The dashed line in the panel of FeT/Al ratio indicates the value of the average shale (Taylor and McLennan, 1985). The two dashed lines in DOP panel marks at
0.46 and 0.75, respectively (Raiswell et al., 1988).
H. Wei et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 353–355 (2012) 73–86
81
Table 2
Variations of multiple geochemical parameters at studied section in Enshi, western Hubei, South China.
Sample no.
Depth
(m)
TOC
(%)
ES18−
ES18
ES19
ES20
ES21
ES22
ES23
ES24
ES25
ES26+
ES27
ES28
ES29
ES30
ES30+
ES31+
ES31
ES32
ES33
ES34−
ES34
ES35
ES36
ES37
ES37+
ES38
ES39
ES40
ES41
ES42
ES43
ES44
10.56
11.31
11.96
12.11
12.84
13.22
13.92
14.32
14.52
15.03
15.12
15.73
16.5
17.55
18.2
19.12
19.49
20.15
20.69
21.19
21.64
22.37
23.02
23.77
24.52
25.21
25.83
26.95
27.15
28.02
28.58
28.98
0.05
0.11
1.15
0.73
0.54
0.33
0.82
1.14
1.12
1.33
0.56
0.14
1.18
0.72
0.82
0.66
0.64
0.88
0.15
0.78
1.00
0.20
0.48
1.15
0.61
0.62
1.29
0.68
0.17
0.22
0.15
0.08
Sp
(%)
FeP
(%)
0.09
0.06
0.22
0.33
0.29
0.08
0.24
0.26
0.26
0.08
0.06
0.20
0.30
0.26
0.07
0.22
0.24
0.23
0.27
0.18
0.15
0.59
0.18
0.16
0.26
0.13
0.01
0.24
0.16
0.14
0.54
0.16
0.15
0.24
0.11
0.01
0.15
FeH
(%)
FeD
(%)
FeHR
(%)
FeR
(%)
δ34S
(‰)
TiO2
(%)
Al2O3
(%)
−6.08
−10.53
0.03
0.03
0.66
0.56
0.14
0.16
0.05
0.12
0.10
0.09
0.22
0.10
−15.30
−11.69
−9.32
−10.76
−10.64
−16.96
−16.38
−15.21
−10.54
−14.44
−10.18
FeU
(%)
MgO
(%)
CaO
(%)
CO2
(%)
FeT
(%)
0.62
1.51
51.36
51.39
41.04
42.04
0.31
0.36
2.69
3.19
0.89
2.24
1.76
1.61
6.38
8.16
2.35
9.14
18.08
18.13
31.55
30.27
50.56
23.18
14.03
12.71
31.81
32.76
42.31
28.27
30.91
29.93
0.85
0.94
0.38
0.64
0.69
0.57
0.08
1.50
12.09
20.85
29.68
0.42
0.07
0.16
1.26
3.15
16.26
20.85
19.43
6.81
33.15
28.29
0.43
0.83
DOP
FeT/Al
FeHR/FeT
0.45
0.25
0.88
1.23
0.79
0.43
0.56
0.55
0.44
0.54
0.47
0.51
0.60
0.56
0.80
0.54
0.74
0.68
0.40
0.30
0.17
0.37
0.42
0.47
0.48
1.27
0.13
0.83
0.22
0.25
0.41
0.14
0.05
0.72
0.59
0.66
0.73
0.53
0.62
0.64
0.50
0.36
0.66
0.57
0.73
0.24
0.50
0.37
0.66
S/C
0.10
0.17
0.16
0.10
0.24
0.16
0.18
0.23
0.13
0.24
0.21
0.09
0.19
0.27
0.22
0.04
0.02
b0.01
0.02
0.05
0.04
0.34
0.28
0.07
0.24
0.29
0.27
0.54
0.47
0.16
0.41
0.51
0.45
0.32
0.47
0.22
0.23
0.18
0.12
0.09
0.11
0.07
0.20
0.02
0.03
0.02
0.01
0.26
0.19
0.15
0.55
0.33
0.27
0.20
0.73
0.09
0.18
0.04
0.02
0.13
0.08
0.05
0.37
0.20
0.05
0.42
0.15
0.03
0.14
−9.84
0.11
0.05
2.08
1.05
10.21
20.11
28.78
13.53
33.84
32.75
0.55
0.20
0.14
0.08
0.01
0.15
0.21
b0.01
−14.19
0.04
0.62
16.44
21.00
34.58
0.19
0.15
0.64
0.64
0.79
0.07
0.09
0.06
0.08
0.03
0.03
b0.01
b0.01
0.05
0.08
0.09
0.11
0.09
b0.01
−11.29
−15.77
0.03
0.02
0.54
0.39
12.72
15.60
28.77
25.50
36.60
37.20
0.18
0.10
0.14
0.05
0.68
0.72
0.64
0.48
0.30
0.84
0.04
b0.01
b0.01
b0.01
b0.01
0.01
0.03
b0.01
b0.01
b0.01
b0.01
0.01
0.01
0.06
0.06
0.03
0.01
0.02
0.05
0.08
0.06
0.07
0.01
0.03
0.05
0.06
0.06
0.03
0.01
0.02
0.06
0.07
0.06
0.08
−12.57
−10.25
−10.36
−11.36
−8.64
−9.62
0.02
0.03
0.02
0.02
0.37
0.47
0.33
0.26
7.92
14.62
2.42
13.29
38.74
27.13
47.76
30.76
39.15
37.40
40.19
38.79
0.11
0.13
0.12
0.11
0.03
b0.01
0.02
0.01
0.02
0.08
0.71
0.04
0.05
0.08
0.23
0.27
0.54
0.54
0.68
0.76
0.42
0.61
0.53
0.70
0.011
0.08
0.06
0.07
0.01
0.03
1.76
0.57
0.19
0.45
0.53
0.24
0.30
0.23
0.23
Fig. 8. Vertical variations of Ba and Mo concentrations in sediments from the Enshi section, compared with TOC contents.
82
H. Wei et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 353–355 (2012) 73–86
Fig. 9. Relationships between pyrite sulphur (Sp) and TOC contents (A), DOP values and TOC contents (B), reactive iron (FeR) and TOC contents (C), and pyrite sulphur (SP) and
reactive iron (FeR) contents (D). In plot A, S/C ratio = 0.36 indicates the normal marine values (Berner, 1970). In plot D, the solid line is the regression line, and the dashed line
represents the stoichiometric ratio of Fe and S in pyrite.
carbon. This is also supported by the low FeT content (Fig. 7; Table 2).
Moreover, the content of SP is linearly related to the concentration of
FeR (R 2 = 0.92) (Fig. 9D); the regression line has an intercept that is
very close to the origin and a slope that is slightly lower than that
for ideal pyrite, suggesting that most FeR was converted into pyrite.
Iron limitation usually occurs in euxinic (or sulphidic) environments
in which excess H2S may not have been sequestrated to form pyrite,
escaping from the organic-rich sediments to waters as a result of bacterial sulphate reduction (BSR) (Raiswell and Berner, 1985; Arthur
and Sageman, 1994). However, the dysoxic condition, as documented
above, implies that the BSR-mediated H2S may have been oxidised
rapidly by ventilated bottom waters so that less pyrite precipitated
as well. In this case, the depletion of oxygen in bottom waters may
have been caused mainly by organic matter decay.
5.1.3. Fe speciation
Iron speciation is widely used to assess the redox conditions of the
water column. Fe/Al ratio, DOP and related FeHR/FeT ratio are useful inorganic geochemical proxies of palaeo-redox conditions (Lyons and
Severmann, 2006). Numerous studies have shown that sediments
deposited from oxygenated bottom waters generally have DOP
valuesb 0.45; those deposited beneath a dysoxic water column have
DOP of 0.45–0.75, and those beneath anoxic/euxinic waters have DOP
values> 0.75 (Raiswell et al., 1988; Wignall, 1994; Raiswell et al., 2001).
At the studied section, with the exception of those from the basal
and top carbonate horizons, most samples show DOP values between
0.45 and 0.75, pointing towards overall dysoxic depositional conditions. This is consistent with the scenario inferred from framboid morphology, size distribution and C–S relationships as documented above.
Black marlstones/shales in the bases of variable-order cycles generally
have higher DOP values with respect to the upper carbonate-rich horizons. Thus DOP values vary episodically, more or less, following the
patterns of the stratal cyclicity from organic-rich marlstones/shales to
limestones. Such stratal cyclicity is considered to have been driven by
glacioeustatic sea-level fluctuations in response to the Earth's orbital
rhythms (procession, obliquity, eccentricity) (e.g. Fischer, 1964; Read
and Goldhammer, 1988; Goldhammer et al., 1990; Osleger and Read,
1991; Strasser, 1991; Preto et al., 2001). Under this circumstance,
dysoxic to oxic fluctuations in the water column, as indicated by the
DOP variations, were linked to fluctuations in sea level (Figs. 3 and
7), being dysoxic during times of sea-level rise, with this condition enhanced during superimposed intervals of sea-level rise, and vice versa.
Only slightly higher FeT/Al ratios (0.5 to 0.68) compared to the
average-shale value (0.4 to 0.5) (Lyons and Severmann, 2006)
would be expected in the absence of syngenetic pyrite (i.e. lack of
FeR). However, the FeHR/FeT ratios reported here are anomalously
high (0.17 to 0.84, average 0.56) (Fig. 7), implying the relatively
high potential for pyrite formation. A FeHR/FeT ratio of 0.38 represents
the maximum proportion of iron which can be extracted from the terrigenous fraction for pyrite formation in normal-marine waters. Thus,
ratios of FeHR/FeT exceeding 0.38 indicate deposition within an anoxic
water column (Raiswell and Canfield, 1998). The high FeHR/FeT ratios
(mostly > 0.5) in this study (Fig. 7) may suggest anoxic conditions,
but this scenario is inconsistent with the dysoxic conditions inferred
from other redox parameters. Enhanced deposition of biogenic matter is a possible way to bring about this scenario since this process
would dilute the poorly-reactive iron (FePR) and unreactive iron
(FeU) contribution of terrigenous particles, and enhance the deposition of FeHR which, according to the model of Canfield et al. (1996),
is itself derived indirectly from the decomposition of the organic carbon associated with inorganic particles (Raiswell and Canfield, 1998).
The decoupling between FeU (unreactive iron) and FeHR/FeT (Fig. 7)
also supports the interpretation of dilution. This example shows the
ratio of FeHR/FeT is unlikely an appropriate proxy to assess the redox
states of carbonate-rich sediments.
5.1.4. Sulphur isotopes
Bacterial sulphate reduction (BSR) is the major pathway of organic
remineralization in marine and terrestrial deposits (Jørgensen, 1982),
and is the fundamental process linking the biogeochemical cycles of
carbon, sulphur and oxygen. The BSR activity induces a fractionation
of stable isotopes 34S and 32S, resulting in a depletion of 32S in the
H. Wei et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 353–355 (2012) 73–86
sulphate and enrichment of 32S in the produced sulphide. A wide
spectrum of sulphur isotope fractionation occurs during BSR which
is ultimately controlled by the sulphate availability. High sulphur isotope fractionation (> 20–22‰) can only be achieved with unlimited
sulphate supply within open systems (Zaback et al., 1993; Habicht
and Canfield, 1997). However, some studies (Schwarcz and Burnie,
1973; Chambers, 1982; Habicht and Canfield, 1996, 1997) concluded
that, when sulphur isotope fractionation is less than 25‰, BSR activity
occurs in a low sulphate environment under a closed system.
In this study, δ34Spy values range from −17.0 to −6.1‰ (Fig. 7;
Table 2). Previous studies have shown that the δ34S value of seawater
sulphate in the Early Middle Permian (Early Chihsia) was about 12‰
(Holser and Kaplan, 1966; Claypool et al., 1980; Strauss, 1997;
Kirkland et al., 2000). Assuming that bottom waters and/or pore
waters in sediments were derived from coeval open-marine waters
(cf. Hudson, 1982; Ebli et al., 1998), then sulphur isotope fractionations
here range from 18 to 29‰, suggesting a partial, open pore-water system in which pyrite precipitated. Furthermore, δ34Spy values fluctuate,
being lighter in the black marlstones/shales and heavier in the limestones, especially in the basal and uppermost thick limestones. These
fluctuations generally follow the patterns of the depositional cyclicity,
and are interpreted here to have been driven by glacioeustatic
sea-level fluctuations. In other words, the BSR occurred in relatively
open pore-waters, with the necessary sulphate supply, during deposition of organic-rich marlstones/shales while sea-level rose rapidly. By
way of contrast, the pore-water system was relatively closed with
much less sulphate available during deposition of the limestones
associated with sea-level stillstand and fall as a result of a downward
movement of the oxic/anoxic interface into the sediments.
5.2. Primary productivity: constraints from TOC, Ba and Mo
The TOC abundance, although being reduced during remineralization
and burial, is generally considered to be correlated with the
83
photosynthetic primary productivity (Nara et al., 2005) and accumulation rate of organic matter (Tyson, 1995; Kuypers et al., 2002; Twichell
et al., 2002; Kuypers et al., 2004; Meyers and Arnaboldi, 2005). Thus
TOC abundance can serve as a proxy for productivity when supported
by other indicators (e.g. Ba and Mo). The cyclic variations of TOC abundance are likely a reflection of primary productivity changes. Further
examination of the data show that TOC fluctuations closely follow the
stratal cyclicity, being higher in the basal and lower parts of the depositional cycles, and being particularly prominent in the lower- order composite cycle sets (CCS) (Fig. 7). It has been hypothesised widely that
depositional cyclicity on different hierarchies was driven by orbital forcing rhythms of different orders; TOC tends to be enriched when sea-level
rises rapidly and to be depleted (or diluted) when sea-level falls. This
scenario was notably prominent during the longer-term superimposed
sea-level fluctuations (Fig. 7).
Biogenic barium (Babio) is a widely-used tracer to estimate variations of palaeoproductivity in marine sediments (Schmitz, 1987;
Stroobants et al., 1991; Dymond et al., 1992; Shimmield, 1992;
Francois et al., 1995; Paytan and Kastner, 1996); it is associated
with the precipitation of barite in reducing microenvironments within settling organic particles (Chow and Goldberg, 1960; Dehairs et al.,
1980; Bishop, 1988). Most particulate barium is composed of barite
micro-crystals, which act as the main carrier for barium into sediments (Bishop, 1988; Dehairs et al., 1991). However, other sources
may also contribute to the Ba content of marine sediments: 1) detrital
aluminosilicates; 2) hydrothermal precipitates, and 3) secretion by
some benthic organisms (Dymond et al., 1992).
In this study, there is no evidence for an input of barium from hydrothermal sources or secretion by benthic organisms. A normalised
approach is commonly used to distinguish between detrital barium
and ‘bio’-barium (Dymond et al., 1992; Gingele and Dahmke, 1994;
Nürnberg et al., 1997; Schroeder et al., 1997; Bonn et al., 1998):
Babio ¼ Batot −ðAl Ba=Alalu Þ ¼ Batot −ðAl 0:0075Þ:
Fig. 10. (A) Cross-plots of Ba vs. SP contents; (B) Cross-plots of Ba vs. TOC contents; (C) Cross-plots of Mo vs. TOC contents; (D) Cross-plots of Mo vs. SP contents.
ð1Þ
84
H. Wei et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 353–355 (2012) 73–86
This equation assumes that all the aluminium in sediments is of
aluminosilicate origin. Various compilations of elemental abundances
in crustal rocks (Taylor, 1964; Rösler and Lange, 1972) suggest that
the Ba/Alalu ratios of aluminosilicate detritus in crustal rocks range
between 0.005 and 0.01. A Ba/Alalu ratio of 0.0075 is used to calculate
a normalised Babio content (see Dymond et al., 1992). As shown
in Fig. 8, the similarity between unnormalized Ba (ppm) and
bio-barium abundances demonstrates that total Ba contents are
mainly composed of bio-barium with very low detrital contributions.
The good correlation between Ba and Sp contents (Fig. 10A) suggests
that most Ba was related to BSR upon organic decomposition, implying that barite microcrystals are the main carrier of barium in sediments. The abundance of Ba fluctuates in the CCS1 and CCS2
(Fig. 8), being higher in the lower part of composite cycle sets (CCS)
deposited during the stage of sea-level rise, and lower in the upper
part of a CCS deposited during the stage of sea-level fall. However,
there is a decrease upwards in Ba abundance in the overlying composite cycle sets CCS3 and CCS4, indicating that the preservation of
Ba was also influenced by environmental factors other than the biological productivity. The much lower FeT contents in these two horizons
(Fig. 7) indicate iron depletion in pore-waters, which may have resulted
from more H2S presence in pore-waters that resulted in thicker BSR
zones than those in CCS1 and CCS2. In this case, prolonged BSR activity
would have consumed more sulphate in pore-waters, resulting in
sulphate depletion, thereby dissolving barite and reducing the potential
for Ba preservation in the sediment (Brumsack and Gieskes, 1983;
Torres et al., 1996). This scenario is consistent with a relatively weak
correlation between Ba and TOC abundances (R2 =0.29; Fig. 10B).
The positive correlation between organic matter and Mo contents
in many black shales and sediments of oxygen-deficient marine basins is well documented (Brumsack, 1986; Werne et al., 2002;
Wilde et al., 2004). Mo may have been directly incorporated into
the kerogen or precipitated with sulphides (Bertine, 1972; Ripley et
al., 1990). Accordingly, the Mo content in organic-rich sediments
can be used as an indicator of primary productivity. In this study,
the Mo abundance is positively correlated with TOC content (R 2 =
0.62; Fig. 10C) but has no relationship with Sp content (R 2 = 0.08;
Fig. 10D), suggesting a more direct coupling between organic matter
and Mo burial through reactions between Mo-bearing dissolved species and organic matter (Helz et al., 1996). The cyclic variations of Mo
concentrations correlate with TOC variation patterns (Fig. 8); both
follow the patterns of the stratal cyclicity, particularly those of CCS
as shown by the higher contents in the lower part of the CCS (CCS1
and CCS3), and lower contents in the upper part of the CCS (CCS2
and CCS4). Thus, Mo enrichment was temporally coincident with
major (third-order) sea-level rises (i.e. transgressions) (Figs. 3 and
8), during which the increased nutrient fluxes could have enhanced
primary productivity.
5.3. Controls on the accumulation of organic carbon
The accumulation of organic matter in marine sediments requires
particularly favourable conditions which are commonly considered to
be related to high bioproductivity (Pedersen and Calvert, 1990;
Caplan and Bustin, 1998; Sageman et al., 2003; Gallego-Torres et al.,
2007), and/or oxygen-deficiency in bottom waters (Demaison and
Moore, 1980; Canfield, 1989; Ingall et al., 1993; Arthur and
Sageman, 1994; Mort et al., 2007), or a combination of both (Calvert
and Fontugne, 2001; Rinna et al., 2002). Within the studied section,
several lines of evidence indicate that all of the organic-rich sediments were deposited beneath an oxic to dysoxic water column
which is usually considered unfavourable for organic matter preservation. Under these circumstances, the enrichment of organic matter
in sediments requires high primary productivity. The weak relationship between DOP values and TOC contents (R 2 = 0.38; Fig. 9B) and
the good correlation between the productivity-related Mo and TOC
contents (R 2 = 0.62; Fig. 10C) suggest that accumulation of organic
carbon mainly resulted from increased primary productivity, which
in turn could have led to the dysoxic conditions. This means that increased primary productivity by blooms of marine photoautotrophs
was the principal agent that led to deposition of the organic-rich
sediments in the lower Chihsia Formation.
The coincidence of high TOC contents with stratal cyclicity as indicated by geochemical and sedimentological evidence (Fig. 7), particularly the high TOC contents at the bases of depositional cycles that are
superimposed on lower-frequency cycles (i.e. third-order sequences),
indicates that substantial organic matter accumulation took place
during the sea-level rise (transgression), notably when it was
superimposed on a third-order sea-level rise. Under this condition,
enhanced oceanic upwelling currents could have carried the
nutrient-rich deep waters into surface waters, leading to the
flourishing of photosynthetic phytoplankton and so increased organic
carbon generation. The increased oxygen demands for respiration of
microorganisms and the decomposition of organic matter could
have resulted in oxygen deficiency in the water column, especially
in bottom waters. Furthermore, the configuration of an intrashelf
basin would greatly slow down the intrabasinal ventilation of the
water mass, particularly below storm wave base, enhancing the oxygen depletion in the deeper and quieter water mass, thus favouring
organic matter preservation. In contrast, during sea-level stillstand
and subsequent fall, the seafloor upon the carbonate shelf was greatly
ventilated and agitated due to enhanced wave and storm currents,
which would have reduced the potential of organic preservation
due to both direct decomposition of organic matter in the water column and within the sediments, further leading to increased erosion
and bioturbation in the sediments. Furthermore, increased carbonate
production would have further diluted the organic matter content in
the sediments (Fig. 7). All of these factors could account for the
lower abundance of organic matter in the carbonate horizons as
observed.
6. Conclusions
This study has provided sedimentological, mineralogical and geochemical data to constrain the processes of accumulation of organic
matter in the lower Chihsia Formation in western Hubei Province,
South China. The lower Chihsia Formation is mainly composed of
dark grey marlstones/shales intercalated with thin limestones deposited in an intrashelf basin. Organic matter mainly accumulated in the
marlstones/shales, constituting the lowermost parts of metre-scale
shallowing-upward cycles. The organic-rich sediments were deposited mainly under dysoxic water conditions. Organic matter enrichment was controlled mainly by primary productivity, which
increased as a result of increased nutrient flux during a time of
sea-level rise, particularly when this was superimposed on a
longer-term (third-order) sea-level rise. This, along with the deeper
and quieter water mass within the intrashelf basin, surrounded
by shelf carbonates, would have further enhanced the oxygen
deficiency within the basin, favouring organic matter accumulation
and preservation.
Acknowledgements
This research was supported by the National Natural Science Foundation of China (NSFC) through grant 40839907 and the National Key
Basic Research Project (973 Project) through grant 2005CB422101.
Yong Fu is thanked for his assistance with field work. The authors
thank Fangyuan Chen for the technical assistance with the SEM work.
Thanks also go to the chief editor Finn Surlyk for his insightful comments and the time spent on this paper.
H. Wei et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 353–355 (2012) 73–86
Appendix A. Supplementary data
Supplementary data associated with this article can be found in
the online version, at http://dx.doi.org/10.1016/j.palaeo.2012.07.005.
These data include Google maps of the most important areas
described in this article.
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