The establishment of a high-resolution chronostratigraphic framework for the whole Phanerozoic Eon would be extremely important in stratigraphy. Previous studies generally focus on biostratigraphy, radiometric isotopic geochronology and magnetostratigraphy, e.g., a date of 251.902±0.024 Ma for the Permian-Triassic boundary was derived from U-Pb dating for zircon from boundary ash beds (Burgess et al., 2014). However, lack of volcanic ashes at other boundaries, substantially limits the establishment of a high-resolution time scale for critical geological periods (e.g., Capitanian and Wordian). The proposal and confirmation of Milankovitch cyclicities (Hays et al., 1976) has provided a new approach that allows stratigraphers to place stratigraphic successions in a chronostratigraphic framework (Gong et al., 2008; Huang et al., 2014; Wu et al., 2011). Cyclical climate changes on Earth, as a consequence of cyclical changes in astronomical forcing, are recorded in sediments. Therefore, cyclostratigraphy can provide a high-resolution astronomical time scales by tuning the cyclic stratigraphic records to astronomical solutions. Recently, orbitally-tuned age models have been widely used in stratigraphic studies (Gradstein et al., 2012), however, relatively few such studies have been applied to Paleozoic or older sediments.

The Capitanian Stage is a critical time interval that represents the fierce eruptions of the Emeishan Large Igneous Province (LIP) and the initiation of a double-phased Paleozoic-Mesozoic biocrisis, drawing the attention of researchers now working on biological and ecological responses to such environmental changes (Ali et al., 2002; Bond et al., 2010a, b; He et al., 2007; Li et al., 2015; Sun et al., 2010; Wignall, 2001; Wignall et al., 2009a, b, 2012; Zhou et al., 2002; Nestell et al., 2015). However, the duration of Capitanian is not well constrained due to scarce radiometric ages from the basal and top boundaries of Capitanian. The only available zircon age known is from an ash bed lying 37.2 m below the base of Capitanian. This provides a maximum estimate for the age of the base of the Capitanian (Bowring et al., 1998). The duration for the Capitanian has been given as 5.3~5.4 Ma (Gradstein et al., 2004; Ogg et al., 2008; Gradstein et al., 2012), with an error estimated of±0.8~1.4 Ma. The precision of this duration is far below Early or Late Permian age values, thus restricting our understanding of onset time and duration of the mass extinction during eruption of the Emeishan LIP.

Magnetic susceptibility (χ) is a proportionality constant that indicates the degree of magnetization of a material in response to an applied magnetic field (Borradaile, 1988). In marine sediments and sedimentary rocks, geochemical and experimental studies have clearly shown that χ is dominated by the ferrimagnetic and paramagnetic mineral content (Ellwood et al., 2000). These minerals come primarily from terrestrial sources, transported into the marine environment by rivers or wind. χ is controlled by cyclical climate because the climate affects the erosion and transport of terrigenous material (Crick et al., 1997 Ellwood et al., 1999). Measurement of χ is simple, fast and non-destructive. Therefore, χ is universally applied in time-series analysis for stratigraphic sequences (Ellwood et al., 2012; Guo et al., 2008; Huang et al., 2011; Jovane et al., 2006; Wu et al., 2012, 2013a, b, c; Nestell et al., 2015; and many others). In this study, we characterize the Capitanian stage in South China, and develop a floating-point time scale (FPTS) for the Capitanian using χ determinations and additional biostratigraphic data sets. We also estimate the onset time and duration of the Emeishan LIP eruptions.

The Laibin area (Fig. 1) is located at the southeastern margin of the Laibin-Heshan isolated carbonate platform, and was surrounded by deep water basins during Late-Middle Permian time (Zheng and Hu, 2010). The sections around Laibin record deposition in a narrow basin, dominated by radiolarian chert deposition during much of the Middle-to Late Permian, but with shallower-water limestone facies, more typical of the surrounding cratonic areas, developed in the latest Middle Permian (Li et al., 2015; Qiu et al., 2014; Sha et al., 1990; Yao et al., 2012). This facies change is in response to eustatic sea-level fall (Mei et al., 1998) or Dongwu uplift (Shen et al., 2007; Wang and Jin, 2000). The Tieqiao section is located along the northern bank of the Hongshui River (23°42.733'N, 109°13.533'E), 5 km southeast of the town of Laibin. This and the Penglaitan section, the global boundary stratotype section and point (GSSP) defining the GuadalupianLopingian (G-L) boundary, are about 10 km apart and situated on the flanks of the Laibin Syncline (Fig. 1; Jin et al., 1998). The Permian Chihsia, Maokou, Wuchiaping and Dalong formations are well exposed along the emerged riverbanks of the Hongshui River. The Maokou Formation (middle and upper parts of the Capitanian) is composed dominantly of grey, thin-to medium-bedded chert alternating with cherty limestone. Considerable work on conodont biostratigraphy has previously been conducted for the Tieqiao section. However, most of this biostratigraphic work has been limited to work on the G-L boundary, because the lower and middle part of the Capitanian mainly consists of chert (Henderson et al., 2002; Jin et al., 1998, 2001, 2006; Mei et al., 1998; Shen et al., 2007).

Fig. 1

Fig. 1 Paleogeographic map (revised from Zheng and Hu, 2010) and geological map with locations (revised from Jin et al., 1998) of the studied sections in Laibin, South China

A total of 45 samples (interval 3 m) were collected through a stratigraphic interval that begins 31.2 m below the Wordian-Capitanian (W-C) boundary and extends to 2.6 m above the Capitanian-Wuchiapingian (C-W) boundary in the Tieqiao section (Fig. 2); each sample weighed more than 3.5 kg. All samples were processed with acetic acid digestion. These samples were crushed into fine gravels 1~2 cm in grain size. Diluted acetic acid (10%) was used to dissolve the samples. A 2.80~2.81 g·mL⁻¹ gravity liquid made of bromoform (2.89 g·mL⁻¹) and acetone (0.79 g·mL⁻¹) was used in conodont separation for all samples (Jiang et al., 2004).

Fig. 2

Fig. 2 Lithological column with conodont zones for the Wordian-Wuchiapingian period at Tieqiao section, Laibin, South China Bed numbers of Maokou Formation are referring to subdivision in Sha et al (1990). Member numbers are from Shen et al. (2007). Sampling locations for conodont biostratigraphy are marked with short bars on the right side of the lithological column.

A total of 1456 unoriented samples (interval 0.1 m) were collected from the foregoing Wordian-Wuchiapingian interval in the Tieqiao section. All lithologies encountered were collected. Low-field bulk magnetic susceptibility (κ) measurements in this study were performed using a KLY-3S Kappa bridge magnetic-susceptibility meter in the Rock-Magnetic Laboratory at the Institute of Geophysics and Geomatics, China University of Geosciences (Wuhan). The results produced by the KLY-3S are initially given relative to volume as κ, in dimensionless SI units, and then converted to mass magnetic susceptibility (χ) after weighing (Ellwood et al., 1988). In order to evaluate the magnetic mineralogy in these samples, four representative samples were selected for thermomagnetic susceptibility measurements (TSM). All powdered samples were heated from room temperature to 700 ℃ at a heating rate of 11 ℃/min and then cooled at the same rate while simultaneously measuring the κ under an argon gas atmosphere. These measurements were performed with the KLY-4S Kappa Bridge in the Paleomagnetism and Environmental Magnetism Laboratory (PEML) at China University of Geosciences (Beijing).

Time-series analysis is a statistical method of analyzing data from repeated observations on a single stratigraphic succession where many samples are collected at regular intervals. If Milankovitch cycles are recorded in sedimentary rocks by proxies measured, then these cycles should be identifiable using time-series methods. In this paper, time-series analyses of the raw χ data (independent of smoothing) have been performed for the Tieqiao section. Because samples were collected over relatively uniform intervals, we assume that the depth of samples is linearly related with time, i.e., ∆x (change in depth) is proportional to ∆t (change in time), so that a power spectral analysis could be conducted. The less this assumption is true, due to variations in sediment accumulation rate, differential diagenesis, or other factors, the more noise that will be produced in the spectral graph, and the less well defined will be spectral peaks (less confidence, Ellwood et al., 2011). The spectral power of these χ data sets was observed via both the Multi Taper Method (MTM) and Fourier Transform (FT) analysis. The data were both detrended and subjected to a Hanning window so as to reduce spectral leakage and increase the dynamic range (Jenkins and Watts, 1968; Weedon, 2003). Incidences of high confidence level peaks (at the 90%, 95% and 99% level) in the resulting spectra are determined by employing MTM (Ghil et al., 2002) as calculated with the SSA-MTM toolkit (Dettinger et al., 1995).

For data presentation purposes, a bar-log format similar to that used in magnetic polarity stratigraphy (MPS) studies is used here (Ellwood et al., 2011; GarcÍa-Alcalde et al., 2011). As in MPS studies, a point half way between the highest and lowest point in a trend is chosen as the boundary between the magnetostratigraphic susceptibility stratigraphy (MSS) zones identified. The bar-log approach, also used in magnetic polarity stratigraphy, has always been somewhat subjective, but given that the raw data are also presented in this study, the reader can judge if the boundary point chosen is reasonable. In order to create the MSS zonation, the following bar-log plotting convention is used; if the χ cyclic trends increase or decrease and are represented by two or more data points, then this change is assumed to be significant and the highs and lows associated with these cycles are differentiated by filled (high χ values) or open (low χ values) bar-logs (Ellwood et al., 2012).

Lithology and conodont zones for the Capitanian Stage within the Tieqiao section are given in Fig. 2. The W-C boundary lies within the lower part of Bed H116, and is defined by the first appearance datum (FAD) horizon of Jinogondolella postserrata in the GSSP section. However, sampling resolution for conodont in the lower part of the Tieqiao section was relatively low due to the intense silicification, and limited the precision of definition for the boundary there. The actual W-C boundary may be 0~11.8 m down from the boundary in this study. The C-W boundary is placed between Bed 6j and 7a-b at the top of Bed H119. This is based on the lowest observed occurrence point (LOOP) of Clarkina postbitteri postbitteri (Jin et al., 2006). Seven conodont zones are identified within the Capitanian (Fig. 2), including J. postserrata, J. shannoni, J. altudaensis, J. prexuanhanensis, J. xuanhanensis, J. granti, and C. postbitteri hongshuiensis, from the base to the top of the Capitanian (Fig. 3), conforming to previous work by Jin et al.(2001, 2006) and Mei et al. (1998).

Fig. 3

Fig. 3 Key conodonts from the Tieqiao section, Laibin, South China (a) Upper view; (b) Lateral view; (c) Lower view. Scale bar 100 µm. 1 Jinogondolella shannoni (Wardlaw, 1994), SP 012; 2 Jinogondolella granti (Mei and Wardlaw, 1994), 06-70 027b; 3 Jinogondolella altudaensis (Kozur, 1992), S3 033; 4 Jinogondolella xuanhanensis (Mei and Wardlaw, 1994), S3 055; 5 Clarkina postbitteri postbitteri (Mei and Wardlaw, 1994), S6 054.

Bed H115 to H118 (Fig. 2; Member Ⅲ) are dominantly composed of limestone intercalated with thinmedium bedded chert or silicified limestone, and include many cycles representing turbidite deposits. The proportion of limestone or chert in this limestone-chert alternation was used for dividing these beds (Figs. 2 and 4). The limestone beds in Member Ⅲ are grey, mainly mudstone-wackstone, with foraminifers, crinoid ossicles and ostracodes. Bedded chert or siliceous limestone are mostly kermesinus, and contain numerous radiolarians and sponge spicules. The turbidite deposits in Beds H115 to H118 are obvious within limestone beds, and identified by a normal graded bedding representing incomplete Bouma sequences. These are generally 3~6 cm in thickness. The rock structure and biological composition of turbidite deposits is distinct from the background. These beds are usually composed of bioclastic packstone with allochthonous bioclasts of fusulinaceans, crinoid ossicles, bryozoans and tubiphytes that were transported from the relatively shallow shelf margin or platform nearby. Bioclasts are strongly eroded. Beds H115 to H118 were considered to have been deposited mainly in a basinal/slope environment.

Fig. 4

Fig. 4 Raw MS profile from the Wordian-Wuchiapingian stages for the Tieqiao section, Laibin, South China Time duration of the Emeishan LIP emplacement event is marked with a grey shaded box (Ali et al., 2001; Sun et al., 2010). See lithology legend in Fig. 2.

Bed H119 (the Laibin Limestone Member; Member Ⅳ in Fig. 2) is mainly composed of grey medium-thick massive limestone that includes abundant fossils, and is quite different from the underlying interbedded chert and limestone beds (H115 to H118), which have been divided into seven units (Unit 2-8 in Jin et al., 2001, 2006). The facies change from Member Ⅲ to Member Ⅳ is considered to be a record of regression during the Late Guadalupian (Chen et al., 2009; Jin et al., 2001, 2006; Qiu et al., 2014; Wignall et al., 2009b). This sea level fall ultimately led to an unconformity that caps the Maokou Formation in most places in South China.

The base of the Laibin Limestone Member (Unit 2) is composed of massive limestone with chert lenses in the lower part, containing normal graded bedding. Radiolarians are abundant in the chert, and the massive limestone is rich in sponges and foraminifers. This unit is interpreted to have been deposited as a basinal facies. Unit 3, overlying Unit 2, consists of thin-bedded limestone. Slump folding was present in this unit, indicating a slope facies. Units 4 and 5 are dominated by a massive, pale, pink limestone, rich with bryozoans, corals, crinoids and fusulinaceans. In addition, a layer of brachiopod shells is parallel to bedding surfaces in Unit 5. These two units (4 and 5) are interpreted to have been deposited as a shallow water carbonate platform facies. The limestone in Unit 6 is grey, and contains abundant crinoids of various sizes that indicate burial in situ. Unit 6 can be explained as a shoaling sedimentary facies. Unit 7 to Unit 8 becomes a thinly dark grey limestone, indicating the beginning phases of transgression (Li et al., 2015).

The Wuchiaping Formation, immediately above the Maokou Formation, consists of dark, thinly bedded chert and cherty limestone, containing radiolarians, and calcareous spicules, but lacking in benthic organisms. This is interpreted to indicate a low-energy slope or basinal facies during Early Wuchiapingian transgression.

In summary, Member Ⅲ and the lower part of the Wuchiaping Formation were deposited as a relatively deep, slope-basinal facies. In contrast, the Laibin Limestone Member mainly represents a shallow water facies. Thus, it can be characterized as an “intersequence” limestone between bedded chert, deposited during a maximum sea-level lowstand (Jin et al., 2001, 2006).

χ data are reported in Figs. 4 and 5. The χ values for samples from Tieqiao section are generally low, ranging from –2.22×10⁻⁹ m³·kg⁻¹ to 5.92×10⁻⁸ m³·kg⁻¹, with a mean value of 1.48×10⁻¹⁰ m³·kg⁻¹, which is significantly lower than values expected for most marine rocks (most observed χ values are 1×10⁻⁹ m³·kg⁻¹ to 2×10⁻⁷ m³·kg⁻¹; Ellwood et al., 2000), suggesting less detrital input into the Laibin area at that time. The χ profile can be divided into two distinct parts. Below 630 m, most χ values are diamagnetic (negative), while above 630 m there is a major shift in χ to positive values (Figs. 4, 5). Examination of Fig. 4 shows that there are a number of samples within the diamagnetic field that are paramagnetic. We believe these are the results of single ash falls and therefore these χ values are not controlled by climatic fluctuations, are considered abnormal samples, and therefore were removed for time-series analysis.

Fig. 5

Fig. 5 κ-T curves for selected samples from the Tieqiao section, Laibin, South China

The results of TSM are presented in Fig. 5. At low temperatures (20~200 ℃), all samples exhibit low κ values with generally flat or slightly upward trends. When temperatures exceed 350 ℃, κ values rapidly increase with a sharp peak at 500 ℃ (Fig. 5a). This is followed by rapid decay ending near or slightly above 580 ℃, the Curie temperature of magnetite (Dunlop and Özdemir, 1997). Sharply increasing κ values at 580 ℃ in the cooling curves also proves the existence of magnetite. Cooling to room temperature yields κ values that are significantly higher than initial values (Fig. 5b), indicating the formation of substantial magnetite during the heating process. Previous research suggested that this magnetite is generated by the transformation of iron silicate, clay minerals (such as illite) and iron sulfide (Dunlop and Özdemir, 1997; Ellwood et al., 2007).

In this study, the raw (unsmoothed) χ data set for the Tieqiao section were used in the MTM and FT analysis. Fig. 6 shows the spectral power for the χ data set, as indicated in both FT and MTM results. To check for confidence level, the MTM result is compared with deviation from red noise (colored bands). With such a large data set many wavelengths will have high confidence levels through chance. These results are supportive of the following interpretation, but we wish it clear that the likely variation in deposition rates over time in the sample makes this analysis non-definitive.

Fig. 6

Fig. 6 Analysis result of spectral power with Multi-Taper (MTM) and Fourier Transform (FT) methods for the raw (unsmoothed) data sets from Tieqiao section, Laibin, South China

If we assign the peak at the 0.78 cycles·m⁻¹ point to the Milankovitch Obliquity 2 (O2) cycle (44.1 ka; here all obliquity and precessional peaks are calculated from Berger et al. (1992) for an age of 262 Ma for the middle of the Capitanian Stage), we derive an age range of 5.07 Ma and mean sediment accumulation rate (SAR) of 2.91 cm·ka⁻¹ for the entire section (including 31.2 m of the Wordian and 2.6 m of the Wuchiapingian). With the O2 assignment and 5.07 Ma duration, all other Milankovitch peaks can be calculated and are plotted in Fig. 6. Reported in Fig. 6 are six Milankovitch bands with relatively high confidence levels, four exhibiting confidence levels >99% (Fig. 6). These peaks correspond to long-term eccentricity, E2 (405 ka, at 0.08 cycles·m⁻¹), shortterm eccentricity, E1 (100 ka, at 0.35 cycles·m⁻¹), obliquity, O2 (44.1 ka, at 0.78 cycles·m⁻¹), and O1 (35.00 ka with low confidence); and precessional peaks, P2 (20.95 ka, at 1.69 cycles·m⁻¹), and P1 (17.7 ka, at 1.94 cycles·m⁻¹).

Long-term Milankovitch bands (eccentricity) are well developed in older sediments than short-term bands (obliquity and precession). It is therefore expected that longer-term climate cycles are more readily identifiable in older stratigraphic sequences. Given that a 405 ka eccentricity climate cyclicity can be expected in long data sets, we smoothed the χ by splines (λ=5), to conform to a 405 ka cyclicity observed in Fig. 6. Bar-logs were then drawn down from the smoothed data and reported in Fig. 7. We argue that the cyclicity we see in this data set represents the control of climate cyclicity on the flux into the marine environment of paramagnetic and ferrimagnetic particles, either through erosion by increased rainfall, due to base-level changes due to eustacy, or by eolian processes. The deposition of these particles has resulted in the χ highs and lows that are reflected in the MSS bar-log reported in Fig. 7.

Fig. 7

Fig. 7 χ data and MSS zonation for the Wordian-Wuchiapingian stages represented by the Tieqiao section, Laibin, South China Dashed lines represent raw χ data with anomalous data points removed and substituted by adjacent data points via linear interpolation. Solid lines are smoothed using splines. The close-up of the middle-lower interval (520~630 m) of the profile is on the right side

The generally low values of that dominate in the Tieqiao section indicate these samples are composed mainly of diamagnetic calcite and quartz, with very little content of paramagnetic and ferrimagnetic minerals. This indicates that the Laibin area was isolated from terrigenous materials during the Capitanian. This isolation is supported by the fact that the Laibin area was separated from the Cathaysia Land or Kangdian Land in the South China Block (Fig. 1), and that the South China Block was isolated from Pangea in the Paleo-Tethys Ocean during the Guadalupian Series (Kasuya et al., 2012; Ziegler et al., 1997).

The variation of χ near the G-L boundary in the Tieqiao section is essentially the same as the results for χ in the Penglaitan Section (G-L) GSSP (Fig. 8). For both sections, as the boundary level is approached from 2 m below the GSSP, χ increases and reaches a high, 1~2 m below the boundary coinciding with the regression in the Late Capitanian. At or near the boundary, χ decreases through the boundary interval owing to the beginning of transgression observed through the Unit 7-Unit 8 interval. χ values from both sections are low, indicating that the input of paramagnetic and ferrimagnetic minerals into the Gan-Xiang-Gui Basin from terrestrial sources are enhanced just below and through the G-L boundary, but then falls immediately above the boundary.

Fig. 8

Fig. 8 The χ profile across the Capitanian-Wuchiapingian (C-W) boundary at the Tieqiao section in comparison with the Guadalupian-Lopingian (G-L) GSSP at Penglaitan section, Laibin, China Sampling position and conodont zones of the G-L GSSP are modified from Clark (2012). Raw χ data are represented as dashed curves with solid circles. Solid curves are smoothed using splines (splines are calculated using the JMP statistical software package by SAS Institute Inc.).

The large excursion of χ values above 630 m in the upper Capitanian (Fig. 4) indicates a sudden increase in terrigenous paramagnetic and ferrimagnetic minerals into the marine environment in the Late Capitanian stage, although Late Capitanian regression may have contributed to this increase in χ. However, the onset of regression is recorded within the Laibin Limestone (656.5 m) (Jin et al., 2001, 2006), which is later than the increase of χ in the middle of Bed H117 (Figs. 2, 4). Therefore, the regression couldn’t be totally responsible for this increase. For the following reasons, we propose that this excursion in χ is correlated with the eruptions of the Emeishan LIP during the Late Guadalupian. First, extensive eruptions of the Emeishan LIP (0.3×10⁶~0.5 × 10⁶ km³ from Ali et al., 2005) and subsequent weathering would have increased the input of terrestrial detrital minerals into the marine system, and the χ of coeval or subsequent marine sediments recorded this increase. Second, biostratigraphic controls suggest that the onset of Emeishan LIP eruptions began in the J. altudaensis zone of the Middle Capitanian Stage and greatly increased in extent and volume in the J. xuanhanensis Zone (Ali et al., 2001; Sun et al., 2010). This corresponds with the sharp increased χ at the top of the J. altudaensis Zone (Fig. 4). Third, in the Upper Capitanian of the nearby Penglaitan GSSP section, thin, pale-cream ash beds and sandy pyroclastic debris are interbedded with cherts, which are considered to be of aeolian origin. These ash fall beds are reported within both Tieqiao and Penglaitan sections (Wignall et al., 2009b). All of these observations indicate that during the Upper Capitanian, Emeishan volcanism resulted in an enhanced χ signal recorded in the Tieqiao section. In turn, it appears to have had an impact on the biota, as evidenced by the extinction and changeover of large fusulinacean foraminifers and calcareous algae in the upper J. altudaensis zone (Bond et al., 2010a; Wignall et al., 2009a). If this assumption is correct, the sharp increase in χ should represent the onset of Emeishan LIP eruptions. The onset time of Emeishan LIP can be accurately constrained based on our floating-point time scale (FPTS) model. In addition, a regional maximal regression, recorded in the Laibin Limestone, may further contribute to the second pulse of χ increase in the latest Capitanian (J. granti-C. postbitteri hongshuiensis zone; Li et al., 2015).

A standard reference zonation (SRZ) climate model (Fig. 9) was established for the W-C sequence examined here, by fitting 405 ka (E2) eccentricity cyclicity to the 5.07 Ma duration determined from the time-series data set (Fig. 6). From these E2 data, twenty five 200 ka eccentricity half cycles were produced. The MSS zones developed in Fig. 7, were designed to conform to a frequency of 0.08 cycles·m⁻¹, representing the E2 Milankovitch cyclicity in the eccentricity band shown in Fig. 6. Therefore, using these MSS zones, it is possible to make a direct visual (graphic) comparison to the SRZ climate model that is designed to reflect uniform E2 cyclicity and is independent of the MSS zonation, and thus to establish a FPTS for the interval sampled. Here we use the MSS zones (y axis in Fig. 9) to graphically compare the χ data to the uniform E2 cycle model (x axis in Fig. 9). We have projected the bottoms and tops of corresponding MSS zones for the Tieqiao section (y axis) and the SRZ (x axis) in Fig. 9 to their intersection point (dots in Fig. 9), and drawn a series of best-fit lines of correlation (LOC) through those data points that define straight-line segments. This is similar to the Graphic Correlation method used for biostratigraphic data sets (Shaw, 1964).

Fig. 9

Fig. 9 Graphic comparison of the MSS zonation from Tieqiao section to a standardized reference zonation (SRZ) model for the Wordian-Wuchiapingian period

Using this approach, it is possible to develop high-resolution ages for the Wordian-Wuchiapingian sedimentary successions in the Tieqiao section. This FPTS is independent of radiometric ages, and can be used to calculate durations for biostratigraphic zones as well as to estimate the onset time for Emeishan LIP eruptions (discussed in section 5.4, below). Absolute timing for the FPTS is based on assigning a value to the base of the Capitanian Stage of 265.1 Ma (Gradstein et al., 2012). If this value changes due to additional work by others, then it is a simple matter to adjust the FPTS by replacing the value for the base of the Capitanian with a new value, and recalculating durations.

We recognize nineteen 200 ka half E2 cycles in the Capitanian data set, yielding an age range of 3.85 Ma for the Capitanian Stage (Fig. 9). Because of the imprecise W-C boundary, there is an error of +0~0.28 Ma possibly existing in the duration of Capitanian Stage. This duration is 1.5 Ma shorter than estimated values by other workers (e.g., 5.4 Ma, Gradstein et al., 2004; Ogg et al., 2008; 5.3 Ma, Gradstein et al., 2012), but given the 0.8~1.4 Ma error in those estimates, the FPTS approach is reasonable.

Overall, graphic comparison using the MSS zonation from the Tieqiao section in comparison with a standardized SRZ model for the latest Wordian through earliest Wuchiapingian results in five LOC segments with different trends (A-E in Fig. 9), which represent slightly different sediment accumulation rates (SARs) at different times. These SAR differences reduce the power in the MTM and FT data sets (Fig. 6). While LOC A (SAR, 1.77 cm·ka⁻¹) and C (SAR, 1.27 cm·ka⁻¹) have relatively low SARs, LOC D (SAR, 3.10 cm·ka⁻¹) and E (SAR, 2.66·ka⁻¹) have medium SARs, and LOC B (SAR, 4.76 cm·ka⁻¹) has a relatively high SAR. These differences are clear in the MSS zonation, with MSS zones Wor U-Wor Y, Cap5-Cap7 being shorter and Wor Z-Cap4 being longer than the equivalent SRZ elements (Fig. 9).

Based on the SRZ climate model presented here, the FPTS (Fig. 9) and using Milankovitch E2 eccentricity (405 ka) as the climate-based cyclicity (Fig. 6), we have estimated durations for the conodont zones given in Figs. 2 and 4. These data are given in Table 1. The shortest C. postbitteri hongshuiensis zone lasted for only 26.6 ka, while the longest J. altudaensis zone lasted for over 2.3 Ma.

Table 1

Table 1 Duration of Capitanian conodont zones in the Tieqiao section

Table 1 Duration of Capitanian conodont zones in the Tieqiao section

The absolute age of the first eruption of basalts and the duration for all eruptive events are still inconclusive, because it is hard to obtain precise zircon U-Pb ages for the beginning or termination of these eruptions (Liu and Zhu, 2009). Biostratigraphic correlation offers an alternative solution. For the Emeishan LIP, the onset of eruptions in southwest China has been well dated using fusulinids (He et al., 2003, 2006) and conodonts (Sun et al., 2010). Here, the onset age of Emeishan LIP eruptions was estimated using the conodont results, combined with the FPTS established in Fig. 9. The eruptions are identified to fall at the top of J. altudaensis zone, where χ exhibits a sudden increase (see Section 5.2). Using the FPTS, the onset age of eruptions is assigned an age of~262.7 Ma, and falls~1.4 Ma prior to the G-L boundary age. Furthermore, at~261.9 Ma, Sun et al. (2010) argued that within the J. xuanhanensis zone the eruptions greatly increased in extent and volume. Previous studies indicated that the eruption of Emeishan LIP may have terminated near in time to the G-L boundary (Ali et al., 2002; Sun et al., 2010). Therefore, the calculated duration of the Emeishan LIP eruption, based on the FTPS data set, is~1.4 Ma. This is significant because eruption of the Emeishan LIP was considered to be the lead factor in the late Middle Permian mass extinction. Therefore, onset of the Emeishan LIP eruption and the late Middle Permian mass extinction should be synchronous.

Seven conodont zones are identified within the Capitanian stage, including Jinogondolella postserrata, J. shannoni, J. altudaensis, J. prexuanhanensis, J. xuanhanensis, J. granti, and Clarkina postbitteri hongshuiensis.

Initial mass-specific magnetic susceptibility (χ), known to be a useful proxy for paleoclimate variability, has been shown to record Milankovitch orbital-scale cycles in sedimentary successions collected in South China from the Middle Permian Wordian Stage, through the entire Capitanian Stage and extending into the Upper Permian Wuchiapingian Stage. Many of the χ values are negative (diamagnetic), suggesting that the Laibin area was isolated from terrigenous detrital sediment supply during deposition of this succession. A sudden increase in χ during the Upper Capitanian appears to be attributed to the eruption of the Emeishan LIP and eustatic regression, increasing paramagnetic and ferrimagnetic components into the marine sediments. The results of thermomagnetic susceptibility measurements suggest that the dominant magnetic minerals in this study are both paramagnetic and ferrimagnetic.

Time-series analysis supports the observations of a generally well-developed cyclical pattern. Five Milankovitch bands are reported with eccentricity peaks at E2 (405 ka) and a frequency of 0.08 cycles·m⁻¹; E1 (100 ka) at 0.35 cycles·m⁻¹; obliquity peaks, O2 (44.1 ka) at 0.78 cycles·m⁻¹, and precessional peaks P2 (20.95 ka) at 1.69 cycles·m⁻¹; and P1 (17.7 ka) at 1.94 cycles·m⁻¹. Four peaks exceed 99% confidence levels. Based on the magnetostratigraphic susceptibility (MSS) zones developed from smoothing the χ data to conform to the 405 ka (E2) eccentricity band identified in the time-series data set, a high-resolution floating point time scale (FPTS) was developed for the entire Tieqiao succession reported here. This FPTS allows a temporal resolution of~200 ka within the studied section because each MSS zone represents half of the E2 eccentricity cycle. The results indicate that the duration of the Capitanian Stage is estimated at 3.85 Ma (with +0~0.28 Ma error). The mean sediment accumulation rate (SAR) for the succession was 2.91 cm·ka⁻¹. From these data, conodont ranges within the Capitanian were estimated, yielding the shortest temporal range of 26.6 ka and the longest temporal range of 2.3 Ma. In addition, the onset age of Emeishan LIP eruptions was estimated at~262.7 Ma, 1.4 Ma below the Guadalupian-Lopingian boundary. This age could also be used in calculating timing of the onset of the late Paleozoic mass extinctions.

Conodont identification was done by Dr. Sun Yadong. Yan Yajuan, Li Aozhu, Yao Yao and Hu Zongjie helped with field sampling and laboratory processing. This research was supported by grants from the Key Laboratory of Marine Mineral Resources, Ministry of Land and Resources (KLMMR-2014-A-12), the National Natural Science Foundation of China (41472087, 41072078), the National 973 Project of China (2011CB808804), and National Science Foundation of the United States (EAR-0745393003). We thank M. Clark for collecting the Penglaitan GSSP samples used in this study.