Tianlin (Guangxi)
In section 1 and 2, the δ13C are relatively stable with values varying between 2.7 and 3.8 ‰ and between 2.5 and 3.3 ‰, respectively (Fig. 5).
In section 3, the δ13C values of samples originating from the Early Tournaisian exhibit mean values around 1.2‰. At the end of the Tournaisian, a positive δ13C shift up to +1.3‰ is recorded (from 1.2 to 2.5‰). During the Early Viséan, the δ13C values remain low (around 1.5‰) but increase gradually to reach a relatively stable level comprises between 3 and 4‰ of δ13C. Overall a global positive δ13C trend from the Tournaisian to Early Viséan is displayed. In the Mid-Late Viséan, the δ13C values decrease from 3.5 to 0.8‰. Subsequently, during the latest Viséan-Early Serpukhovian, δ13C values increase to remain constant between 2.5 and 3‰. During the Early Serpukhovian, the
Figure 4: Photomicrographs of brachiopod shells. Most of shells display partly preserved layers (A), which exhibit numerous patches of red-yellow-orange luminescence under cathodoluminescence (B). Conversely, the well-preserved shells exhibit well-define microstructures under natural light (C), and appear as non-luminescent under cathodolumi-nescence (D).
δ13C values record short-term variations (- 1‰; +1‰), followed by a drastic decrease at the late Early Serpukhovian (from 2.5‰ to 0‰). During the Late Serpukhovian, the δ13C values increase gradually, from 0‰ to 1.5‰, followed by a decline, from 1.5 to 0‰. During the latest Serpukhovian-Early Bashkirian, the δ13C values increase by +1,5‰ (from 0 to 1.5‰). During the Early Baskhirian, δ13C values range from 2 to 2.5‰. At the Mid Bashkirian, δ13C values decrease from 2.5 to 0.5‰, followed at the late Bashkirian by a positive δ13C shift of +1.6 ‰. During Moscovian, δ13C values decrease drastically by -3‰ (from 2.1 to -0.9‰). During the Kasimovian-Gzelian, the trend in δ13C values is reverse with a major shift of +4.5‰ (from -0.9 to 3.3‰).
Ziyun (Guizhou)
In section 1, during the Early and Mid Kasimovian, the δ13C is globally stable and range between 3.5 and 4.9 ‰. At the Late Kasimovian, the δ13C values decrease by -1.5‰ (from 3.6 to 2.1‰).
From the end of the Kasimovian to earliest Gzhelian, the δ13C is virtually constant with a mean value around 2.5‰, with a maximum of 3.0 ‰. At the end of the Gzhelian, δ13C values decrease from 2.9 to 0.9‰ (-2‰; Fig. 6).
In section 2, the part of the section dated as Gzhelian records δ13C values of about 2.5‰, followed at the late Gzhelian by a decline of -1.3‰ (from 2.5 to 1.2 ‰). During the Early Permian (Asselian), δ13C values increase progressively to reach 3.5 ‰. During the Sakmarian, δ13C values record a decline of -1.1‰ (from 3.9 to 2.8 ‰; Fig. 6).
INTERPRETATIONS AND DISCUSSION
The δ13C of carbonates vary in response to changes in the global carbon cycle, which has led to use δ13C variations for stratigraphic purposes (Swart and Eberli, 2005). Numerous studies have doc-umented and used δ13C oscillations to correlate sediments (e.g. Glumac and Walker, 1998; Immen-hauser et al., 2002; Krull et al., 2004; Saltzman and Thomas, 2012). However, δ13C can be subject to limitations showing a considerable spread of values at any given time, due to the spatial variability and biological factors (e.g. Kroopnick, 1985; McConnaughey and Whelan, 1997; Swart and Eberli, 2005).
In this part, carbon isotope values of southern China are compared to contemporaneous data, re-corded from China, western Europe, North America and Russia (Bruckschen and Veizer, 1997; Mii et al., 1999, 2001; Grossman et al., 2008; Buggisch et al., 2011; Liu et al., 2015). Bruckschen and Veizer (1997), Mii et al. (1999, 2001) and Grossman et al. (2008) realized carbon isotopes analysis on well-preserved brachiopod shells. Buggisch et al. (2011) performed carbon isotopes analysis on brachiopod shells as well as on limestones. Limestones deposited in deeper water slope environments were interpreted to have mainly retained their carbon isotope ratios. Liu et al. (2015) performed carbon isotopes analysis on micrite samples.
Figure 5: Carbon and Oxygen isotope data, records from Guangxi.
Figure 6: Carbon and oxygen isotope data, records from Guizhou.
Mississippian (Tournaisian – Early Serpukhovian)
In Tianlin (section 3, Guangxi), the first major shift in δ13C is recorded at the Late Tournaisian, (+1.3‰; interval 2; Fig. 7) but appears to be low compared to other values reported worldwide (e.g.
+3.4‰ in US, from 2‰ to 5.4‰; Grossman et al., 2008). In North America and western Europe, causes of the Tournaisian positive shift is still debated but this variation may have global implications.
However, better isotopic and biostratigraphic resolution are needed to evaluate the cause of the shift (Mii et al., 1999).
From the Late Tournaisian to Early Viséan, the δ13C values of the section 3 record a positive shift (+2‰; interval 4; Fig. 7), widely reported in other locations (e.g. North American Mid-Continent, western Europe, Russian platform; Bruckschen and Veizer, 1997; Mii et al., 1999; Grossman et al., 2008). The global δ13C positive shift fits with the proliferation of terrestrial plants in Early Carbon-iferous time (Mii et al., 1999; Peters-Kottig et al., 2006). Peters-Kottig et al. (2006) have recently reported an increase in the fractional burial of light (terrestrial) organic matter (TOM) in the Late Paleozoic sediments. More precisely, δ13C exhibit high values during the Early Mississippian (near -22‰) followed by a phase of relatively 13C depleted TOM in the Late Mississippian and Pennsyl-vanian (near -24‰). The increase of δ13CTOM values is explained as a consequence of an evolution of higher land plants.
In section 3, the δ13C shift is punctuated at the Mid-Late Viséan by a negative shift (-2.7‰; in-terval 5; Fig. 7). A contemporaneous decline is reported in the Youjiang basin (South China) and in the subequatorial western Euramerica, whereas coeval sections in subequatorial eastern Euramerica shows consistently elevated δ13C values across the entire Viséan (Liu et al., 2015). The δ13C decline in South China and western America is currently attributed to oceanic current circulation changes, related to water-depth variability (Mii et al., 1999; Saltzman, 2003; Liu et al., 2015). Indeed, phyto-plankton preferentially utilizes 12CO2 for photosynthesis, leading to an increase of the 13C:12C ratio in surface waters. Conversely, at depth, respiration exceeds photosynthesis and there is a net release of 13C-depleted carbon during the oxidation of organic matter, inducing the decrease of the 13C:12C ratio. Consequently, the oceanic δ13C is higher in the ocean surface than the deep ocean (Anderson and Arthur, 1983). Thus, the decline of δ13C water surface can be explained by intensified upwelling which would have brought 13C-depleted water to the surface (Mii et al., 1999). The causes of the Mid-Late Viséan reorganization of oceanic circulation can be attributed to the closure of the Rheic Ocean (e.g. Jastrzębski et al., 2013; Korn et al., 2012; Qiao and Shen, 2014, 2015; Liu et al., 2015).
In Tianlin, from the Late Viséan to Early Serpukhovian, the δ13C values increase and stay constant around 2.5-3‰ (interval 6, section 3, Fig. 7; sections 1 and 2), before recording a decline (-1‰;
interval 7). The δ13C positive shift (interval 6, section 3) coincides with short-lived glacial events (Late Viséan: South America, Caputo et al., 2008; latest Viséan-earliest Serpukhovian: Argentina, eastern Australia, Henry et al., 2008; Fielding et al., 2008b). Conversely, the following δ13C negative shift (-1‰; interval 7; section 3) in the Early Serpukhovian could be explained by a pulse of climate warming. Periods of cooling and ice sheet formation are largely reported as related to a positive shift
of δ13C (Wenzel and Joachimski, 1996; Kump and Arthur, 1999; Saltzman, 2002; Lehnert et al., 2007;
Buggisch et al., 2008, 2011; Frank et al., 2008; Grossman et al., 2008). The causes of these shifts during cooling are still debated but are commonly explained by the deposition and storage of organic carbon (Grossman et al., 2002). However, it seems difficult to explain how the terrestrial ecosystem might have stored more carbon during a glacial time period, since there was a massive expansion of ice sheets over land. Hence, one hypothesis to explain the storage of carbon during cooling, is based on the ocean-atmosphere exchange: ocean have a higher CO2 uptake during a cold glacial than during a warmer interglacial, affecting the carbon storage in the deep ocean (Williams and Follows, 2011).