Payne et al. (2007) suggest that there is widespread evidence for a carbonate dissolution surface at the level of the end-Permian mass extinction, and that this crisis was at least partly due to ocean acidification that preferentially affected heavily-calcified marine invertebrates. These ideas were previously mooted as the Deev Jahi Model of Heydari and Hassanzadeh (2003), although the majority of Permo-Triassic boundary workers have emphasized the evidence for a dramatic increase in carbonate saturation of oceans in the immediate aftermath of the extinction (Kershaw et al. 1999; Ezaki et al., 2003; Groves et al., 2005; Kershaw et al., 2007; Pruss et al. 2006; Baud et al., 2007; Mu et al., in press). We would like to argue here firstly, that the evidence for undersaturation is more parsimoniously (and less sensationally) interpreted as a karstic surface, and secondly, that this surface does not coincide with the main mass extinction level but predates it.
Karst or submarine dissolution surface?
Payne et al. (2007) present evidence for a limestone truncation surface of marine origin from three regions, the eastern Tethyan Great Bank of Guizhou (GBG) in SW China, a a western Tethyan section in southern Turkey and a Panthalassan atoll section now located in Japan. In the GBG sections a sharp, erosional surface is seen to truncate grains and marine phreatic cements in bioclastic limestones of latest Changhsingian age. This is overlain by alternating microbial limestones and molluscan packstone lenses. A submarine phase of carbonate dissolution is favored over a karstic origin for the surface because there is “no evidence of subaerial exposure” (Payne et al., 2007, p.776). However, dissolution and an irregular erosion surface are typical features of subaerial exposure. Payne et al. report that freshwater vadose and phreatic cements and pedogenic features are missing. This is offered as further evidence for a non-subaerial origin; but these features are often absent from karstic surfaces, particularly where emergence is brief and/or marine cement has occluded porosity before emergence. Both are likely to be the case in this example. Payne et al. (2007, p. 776) tacitly acknowledge this possibility: “If subaerial exposure had occurred…it would necessarily have occurred abruptly and briefly….and left no diagnostic lithological [i.e. diagenetic] signature”. In fact, Reinhardt (1988) described and illustrated meteoric phreatic and vadose fabrics from the uppermost Changhsing Formation in Sichuan, South China. This shows that, in some sections at least, there was subaerial exposure before the deposition of the microbialite.
The best evidence for the GBG surface being of karstic origin comes from its location in a very shallow-water setting and by comparison with Permo-Triassic boundary sections from throughout the world. These reveal that a late Changhsingian type 2 sequence boundary (and its correlative conformity) is widespread in both carbonate and clastic sections (Hallam and Wignall, 1999; Yin et al., 2007). In the shallow-water sections of the GBG such a eustatic event should be well developed – so where is it? In the Turkish example (the Taskent section), this type 2 sequence boundary clearly occurs at the sharp base of a 20 cm-thick grainstone that caps the bioclastic packstones of the Cekic Dagi Formation (Unal et al., 2003). However, at this location, Payne et al. (2007) identify three potential dissolution surfaces developed within the upper 10 cm of the grainstone, although the illustration – their Fig. 4E – is not sufficiently clear to permit unequivocal identification of the nature of these surfaces. They could, for example, be simply marine erosion surfaces unrelated to dissolution, or perhaps subaerial exposure surfaces formed briefly within the oolite unit. Moreover, this level is an oosparite and is the local manifestation of a widespread “oolite event” in western Tethys that occurred in the latest Permian immediately after the mass extinction. Oolites are developed in virtually all western Tethyan boundary sections at this time (Groves et al., 2005). They are also widespread in eastern Tethys (e.g. Algeo et al., 2007). This phenomenon can be explained as a phase of prolific abiotic carbonate precipitation following the extinction of most skeletal carbonate producers (e.g. Groves et al., 2005). This raises the question, why is evidence for multiple submarine dissolution events found within a bed that provides evidence for carbonate supersaturation?
What is the age of the surface?
A key facet in the argument for the unique nature of the truncation surface is the observation that “biostratigraphic data from South China demonstrate that the surface postdates the uppermost Permian sequence boundary” (Payne et al., 2007, p.771) and therefore cannot be confused with it. This claim is supported by their biostratigraphic data. They report the composition of a conodont assemblage from immediately below the truncation surface at one of the GBG sections. Within this assemblage, the only age-diagnostic conodont is Hindeodus eurypyge which first appears at the base of the topmost Permian H. changxingensis Zone (Jiang et al., 2007). Payne et al. therefore suggest that the dissolution event occurred very close to the end of the Permian, within the H. changxingensis Zone, probably at the boundary between beds 27a and 27b of the Meishan stratotype (Fig. 1). For comparison, the sequence boundary, noted above, occurs in the later part of the underlying H. praeparvus Zone, between beds 24d and 24e at Meishan (Yin et al., 2007).
As Payne et al. recognize, this H. changxingensis age causes a problem when trying to link the carbonate dissolution event to the mass extinction because all recent studies show that the main extinction losses occurred during the earlier H. praeparvus Zone (e.g. Jin et al., 2000; Fang, 2004; Xie et al., 2005; Yin et al., 2007; Metcalfe et al., 2007). At Meishan this levels corresponds to Bed 24e. Thus, they note (p. 781) that their “observations could indicate a diachronous extinction pattern, they could also reflect local lithofacies controls on fossil occurrences that result in offsets between last occurrences among sections and/or minor sampling and facies-related offsets in conodont ranges.” Implicit in this suggestion of a “diachronous extinction” is that the diachroneity only applies to their GBG study sections. Therefore the carbonate dissolution event can only have affected taxa in the GBG region. Payne et al.’s apparently preferred H. changxingensis age for the GBG extinction overlooks the fact that a distinctive aftermath fauna, characterized by a “mixed assemblage” of Permian holdover taxa and “Triassic” progenitors, had appeared by this time everywhere else in South China (Yin et al., 1996; Fang, 2004; Metcalfe et al., 2007). No sections record an H. changxingensis interval characterized by a low diversity, under-sampled, pre-extinction assemblage.
A further problem with Payne et al’s age assignment for the truncation surface/dissolution event is that it postdates the onset of deposition of abiotic precipitates, including crystal fans, varied microbialites and ooids. All these provide evidence of carbonate saturation in ocean surface waters at a time when Payne et al. want under saturation. The pre-H. changxingensis age of the microbialites is provided by both chemostratigraphic and biostratigraphic evidence. Studies in South China have shown that a major negative excursion of C isotope values begins at the base of the post-extinction microbialites and reaches a low point within these strata (Krull et al., 2004; Mu et al., in press). In the Meishan type section this excursion begins within Bed 24 and reaches its lowpoint in Bed 25 (e.g. Xie et al., 2007) indicating that the onset of microbialite deposition coincides with the main phase of the mass extinction. In Iran the start of microbialite deposition may have even been a little earlier because the onset of the C isotope excursion occurs within this facies (Wang et al., 2007). Alternatively, there may be a brief hiatus at the base of the Chinese microbialites at the level of the late H. praeparvus sequence boundary (Fig. 1). This chemostratigraphic evidence is reinforced by conodont discoveries from South China. Thus, Qi and Liao (2007) report Clarkina yini from the base of a microbialite section at the well-known Tudiya section near Chongqing. This species went extinct at the end of the H. praeparvus Zone (Jiang et al. 2007), once again indicating that the onset of microbialite deposition occurred within this zone.
A simpler explanation of Payne et al.’s age data is that the report of H. eurypyge either represents a misidentification or an earlier occurrence of this species than hitherto recorded. This is supported by the fact that the most common hindeodid of the H. changxingensis Zone, the eponymous zonal species, is absent from their assemblage. If the pre-truncation limestones of GBG belong to the H. praeparvus Zone this creates the problem (for Payne et al.) of an older date for their truncation surface. This then becomes potentially synchronous with the H. praeparvus sequence boundary and a more prosaic origin therefore becomes available.
A pH-sensitive extinction scenario?
In keeping with many recent attempts to understand the origin of Permo-Triassic d13C fluctuations, Payne et al. (2007, p. 782) use the better-understood perturbations of the Paleocene-Eocene Thermal Maximum (PETM) event as a “similar scenario”. This rather vitiates their argument. Although there were interesting faunal migrations and some brief, temporary changes in oceanic plankton composition, the PETM coincides with an interval of low extinction rates by Phanerozoic standards. To compare it to the worst crisis in the history of life requires that PETM-like conditions must have been exacerbated to an extraordinary degree. Alternatively, this interval provides a poor scenario for understanding the end-Permian extinction.
If lowered seawater carbonate saturation was a causal factor in the end-Permian extinction, then it would “have acted synergistically in preferentially eliminating heavily calcified marine taxa with limited ability to compensate physiologically for rapid changes in the carbonate chemistry of surrounding seawater” (Payne et al. 2007, p. 782). They then claim that “Heavily calcified genera were the most strongly affected”. In fact, with such wholesale losses at the end of the Permian, it is possible to argue for almost any selectivity and any cause and different stresses could have the same outcome. For example, the widespread development of dysaerobic facies (Wignall & Twitchett, 2002) would have made the secretion of thick, calcareous shells particularly difficult (Rhoads and Morse, 1971). Soft-bodied taxa also suffered major losses as witnessed by the loss of many types of burrowing activity (Twitchett and Barras, 2004). All siliceous taxa (radiolarians, sponges) also suffered devastating losses with the result that there is an Early Triassic “chert gap”. Neither of these last two observations are readily attributable to lowered pH levels.
There are at least three different possible interpretations of the observations of Payne et al. (2007):
Return to list.