Ice Volume and Paleoclimate History of the Late Paleozoic Ice Age from Conodont Apatite Oxygen Isotopes from Naqing (Guizhou, China)

PALAEO-07632; No of Pages 11 Palaeogeography, Palaeoclimatology, Palaeoecology xxx (2016) xxx–xxx

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Palaeogeography, Palaeoclimatology, Palaeoecology

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Ice volume and paleoclimate history of the Late Paleozoic Ice Age from conodont apatite oxygen isotopes from Naqing (Guizhou, China)

Bo Chen a,⁎, Michael M. Joachimski b, Xiang-dong Wang c, Shu-zhong Shen a,Yu-pingQic, Wen-kun Qie c a State Key Laboratory of Palaeobiology and Stratigraphy, Nanjing Institute of Geology and Palaeontology, Chinese Academy of Sciences, 39 East Beijing Road, Nanjing 210008, PR China b GeoZentrum Nordbayern, Universität Erlangen-Nürnberg, Schlossgarten 5, 91054 Erlangen, Germany c Key Laboratory of Economic Stratigraphy and Paleogeography, Nanjing Institute of Geology and Palaeontology, Chinese Academy of Sciences, 39 East Beijing Road, Nanjing 210008, PR China article info abstract

Article history: A high-resolution and continuous conodont apatite oxygen isotope record spanning the late Viséan to Middle Received 5 August 2015 Permian is reported from South China, which is interpreted with respect to the ice volume and/or tropical seawa- Received in revised form 21 December 2015 ter temperature history of the Late Paleozoic Ice Age (LPIA). The presented δ18O record shows significant fluctu- Accepted 1 January 2016 ations in δ18O from the late Viséan to Middle Permian with highest values observed in the Bashkirian (Early Available online xxxx Pennsylvanian). The δ18O maximum coincides with a major eustatic sea level fall recorded in low-latitude successions and postdates the significant increases in 87Sr/86Sr and δ13C measured on well-preserved brachiopod Keywords: carb fl fi Oxygen isotopes calcite, which can be interpreted as re ecting intensi ed weathering as consequence of the closure of Rheic Conodont apatite Ocean as well as enhanced carbon burial. Both processes may have contributed to lower greenhouse gas levels Late Palaeozoic Ice Age and cooled down the Earth's surface, triggering the maximum glaciation. The high Bashkirian δ18O values are Bashkirian interpreted to represent the glacial maximum of the LPIA. A coeval change in faunal composition and a decreasing Maximum glaciation diversity in climate-sensitive marine invertebrates can be ascribed to icehouse cooling and/or loss of habitat. De- Naqing section spite inconsistencies with earlier interpretations that the Early Permian represented the glacial maximum of the LPIA as inferred from Gondwanan glacial sediments records, the suggested Bashkirian glacial maximum agrees well with ice extent estimates based on the regional tectonic history in Gondwana, which suggests that the Bashkirian glaciation occurred during Gondwana interior uplift promoting maximum ice cover of the entire LPIA. However, maximum glaciation is only poorly represented in the depositional record because large parts of the glacial deposits were possibly removed by erosion as outlined by a major regional unconformity. © 2016 Elsevier B.V. All rights reserved.

1. Introduction of Gondwana (Crowell and Frakes 1970; Caputo and Crowell 1985; Dickins 1997; López-Gamundí 1997; Isbell et al. 2003a; Fielding et al. The Late Palaeozoic Ice Age (LPIA) has been widely recognized as the 2008a, 2008b, 2008c; Gulbranson et al. 2010). Ice volume changes in longest ice age during the Phanerozoic, which occurred at a time while this dynamic LPIA model are critical for assessing the LPIA's extent, the Earth's physical, chemical, and biological systems experienced pace as well as its climatic influences. major changes (Heckel 1994, 2008; Falcon-Lang 2004; Joachimski Based on the palaeogeographic distribution of glacial sediments, the et al. 2006; Clapham and James 2008; Grossman et al. 2008; Isbell ice volume of the LPIA was generally thought to reach its first peak dur- et al. 2008a; Falcon-Lang and DiMichele 2010). Early studies considered ing the Bashkirian, followed by a contraction in the Late Pennsylvanian. the LPIA glaciation as a single massive ice sheet that waxed and waned Maximum ice volume was suggested to occur in the Early Permian, be- continuously across Gondwana for more than 100 million years fore the demise in the middle Sakmarian (Isbell et al. 2003a, 2003b, (Veevers and Powell 1987; Frakes and Francis 1988; Frakes et al. 2012) or the end of the Capitanian and early Wuchiapingian (Fielding 1992; Ziegler et al. 1997; Hyde et al. 1999; Blakey 2008; Buggisch et al. 2008a, 2008b, 2008c; Mory et al. 2008; Waterhouse and Shi et al. 2011). However, this concept was challenged recently due to in- 2010; Frank et al. 2015). Besides the sedimentary record, indirect prox- creasing evidence that indicated a more dynamic LPIA characterized ies, such as eustatic sea level changes inferred from low-latitude sedi- by multiple, short-lived glaciations of 1–8 million years duration alter- mentary sequences, are considered as alternative and potentially more nating with non-glacial or warm periods having approximately identi- reliable proxy to reconstruct the ice volume history during the LPIA. Sig- cal duration (Fielding et al. 2008a, 2008b). Further, onset and demise nificant eustatic sea level changes in low-latitude successions are gener- of these glaciations were found to be asynchronous in various parts ally coincident with variations in continental ice volume (Rygel et al. 2008), but many uncertainties remain in estimating both magnitude ⁎ Corresponding author. of sea level and ice volume change (Fielding et al. 2008b; Rygel et al. E-mail address: [email protected] (B. Chen). 2008). As a consequence, disparities are usually found while comparing

A Late Carboniferous (320Ma)

Euramerica South China Paleo Tethys Panthalassic Ocean Gondwana

Ice Cover

B Hefei 200km Nanjing Shanghai

Chengdu Wuhan Hangzhou Yangtze Land Chongqing Nanchang Changsha m onater b o Dian Xi tf ar a ysian Land C Pl Yangtze Guiyang ha t Carbonate Fuzhou Luodian Ca Guilin Kunming Platform Taipei Land Dian-Qian-Gui Platform

Guangzhou Deep Basin Nanning

Carbonate Platform

Fig. 1. Carboniferous palaeogeographic reconstruction showing the position of the South China Block (A) and the location of the Naqing Section in South China (B) (reconstruction from Saltzman 2003; Feng et al. 1998).

Henderson and Mei (2003). Recently, detailed studies of the conodont as well as fusulinid and non-fusulinid foraminifer biostratigraphy 750 were conducted on the Carboniferous part of this section in order to establish GSSPs for several stage boundaries (Qi et al. 2014, 2015). δ18O values of conodonts from Naqing 700 Locﬁt regression with 3. Materials and methods 95% conﬁdence interval Roadian Cherty limestone Guadalupian 650 Conodonts were extracted by dissolving carbonate rocks in 10% acetic acid. Sodium polytungstate was used for heavy liquid separation Chert of conodont elements that subsequentially were picked from the 600 Limestone heavy fraction under a binocular microscope. Mono-generic conodont Brecciated assemblages were preferentially chosen for analysis from samples Limestone with abundant conodont elements; in case of a lower abundance of co- 550 nodont elements, multi-generic assemblages were used (see Appendix for taxonomic composition of each sample analyzed). 500 The conodont color alteration index (CAI) is considered as a proxy for the maximum thermal overprint of the sediments (Rejebian et al., Kungurian 1987), with a high thermal overprint (high CAI) suggested to result in 450

aggrading crystallization of conodont apatite (Nöth 1998), and thus po- Cisuralian tentially overprinting the primary oxygen isotope ratios. The CAI value 400 for most conodont samples from Naqing section is up to 5, corresponding to 200 °C to 300 °C thermal overprint. However, several studies have Artinskian 350 shown that heating of conodont apatite does not necessarily result in an Sakmarian exchange of oxygen in the phosphate group of conodont apatite even at Asselian high CAI value up to 5–7(Buggischetal.2008;Joachimskietal.2009). 300 Approximately 0.5 to 1.0 mg of conodont apatite was weighed into Gzhelian small polyethylene beakers and dissolved by adding 5 ml 2 M HNO3. 250 n Kasimovian The phosphate group was reprecipitated as Ag3PO4 as described by Joachimski et al. (2009).TheobtainedAg3PO4 crystals were washed in- ﬁ tensively and ground to ne powder using an agate mortar. Ag3PO4 Moscovian 200

(0.2–0.3 mg) was weighed into silver foil and transferred to the sample nnsylvania carousel of a TC-EA (temperature conversion elemental analyser) Pe coupled online with a ThermoFisher Delta V Plus isotope ratio mass Bashkirian 150 spectrometer. Samples and internal standards were generally analyzed in triplicate (helium flow rate was 80 ml/s, reactor temperature was set Serpuk- 100 to 1450 °C, column temperature was 90 °C). Conodont apatite oxygen hovian isotope values are reported in ‰ relative to VSMOW. The reproducibility ssipian of triplicate sample measurements was generally better than ±0.25‰ i 50 Viséan (1 SD). The average oxygen isotope composition of standard NBS 120c Miss was measured as 21.7‰ ±0.25‰ (1 SD). Palaeotemperatures are calcu- 0m lated from δ18O values using the temperature equation published by Pucéat et al. (2010). The nonparametric locally weighted regression 19 20 21 22 23 24 25 “ fi ” 18 method Loc t is used to calculate isotope trend lines (Loader, 1999), δ O(‰VSWOM) which produces a smoothed curve retaining local minima and maxima and yields good results, even with unevenly spaced data points. All cal- Fig. 2. Oxygen isotopes of conodont apatite from the Naqing section (red solid squares). culations were performed with the open source statistic software “R” Solid line and dashed line—Locfit regression with 95% confidence interval. (version 3.0.2, Ihaka and Gentleman 1996).

4. Results 5. Discussion The conodont apatite oxygen isotope record from the Naqing section is based on a total of 312 analyses, among them, 58 Permian data points 5.1. Oxygen isotope composition of Carboniferous–Permian seawater were reported previously in Chen et al. (2013). The description of the main trends in conodont δ18O is based on the calculated Locﬁt re- The δ18O values of conodont apatite are dependent on temperature gression. In general, the oxygen isotope ratios are between 20‰ and the oxygen isotope composition of seawater in which the conodont and 24‰ (Fig. 2). For the late Viséan, average δ18Ovaluesare animals thrived. The latter is a function of variations in salinity and con- slightly above 22.0‰ (Fig. 2). In the early Serpukhovian, the values tinental ice volume, or may be affected by a secular change in the oxy- show a minor decrease to around 21.2‰ followed by a pronounced gen isotope composition of seawater (Veizer et al. 1999; Veizer and positive shift from 21.2‰ (early Serpukhovian) to the maximum Prokoph 2015). An increase in δ18O values of conodont apatite can be value of 23.3‰ in the Middle Bashkirian. The δ18O values decrease to interpreted as cooling, higher salinities, and/or continental ice buildup. 22.0‰ at the Bashkirian–Moscovian boundary. During the Moscovian A decrease in conodont apatite δ18Ovaluesmayreﬂect warming, and Kasimovian, average δ18O values are around 22.0‰ and decrease lower salinities, and/or deglaciation. from 22.0‰ to 21.0‰ from the early Gzhelian to Asselian. No major The relationship between the oxygen isotope composition of apatite 18 18 18 change in conodont apatite δ O is observed from the Sakmarian phosphate (δ Op), δ O of seawater, and temperature is expressed in throughout Artinskian, until a further decrease to 20.5‰ occurs in the the phosphate–water oxygen isotope fractionation equation originally Roadian. given by Longinelli (1966) and Longinelli and Nuti (1973a, 1973b),

Please cite this article as: Chen, B., et al., Ice volume and paleoclimate history of the Late Paleozoic Ice Age from conodont apatite oxygen isotopes from Naqing (Guizhou, Chin..., Palaeogeogr. Palaeoclimatol. Palaeoecol. (2016), http://dx.doi.org/10.1016/j.palaeo.2016.01.002 4 B. Chen et al. / Palaeogeography, Palaeoclimatology, Palaeoecology xxx (2016) xxx–xxx refined later by Kolodny et al. (1983), and more recently by Pucéat et al. The positive δ18O shift during the Serpukhovian could be interpreted (2010) for modern analytical techniques: as combined effect of the waxing of ice sheets and climate cooling; however, the exact contribution of climate cooling and ice volume increase on the 1.9‰ increase in δ18O is difficult to constrain. A potential ana- ðÞ¼∘ : – : δ18 þ : −δ18 −δ18 ð Þ T C 118 7 4 22 Op 22 6 ONBS120c Oseawater 1 logue is the Pleistocene Last Glacial Maximum (LGM) with surface- dwelling foraminifers from the equatorial Pacific and Atlantic recording approximately 1.8‰ higher δ18O values during the LGM in comparison However, the oxygen isotope ratio of ancient seawater is usually un- to the Holocene (Ignacio Martinet et al. 1997). The contribution from known and has to be assumed when calculating seawater temperatures. ice buildup was suggested as 1‰ (Schrag et al. 1996), corresponding An oxygen isotope ratio of −1‰ VSMOW is generally assumed for nor- to 120–130 m sea level lowering (Fairbanks 1989; Peltier and mal marine waters in an ice-free world (Savin 1977). For comparison, Fairbanks 2006; Yokoyama et al. 2000; Mitrovica 2003), which is com- the Modern Ocean has an average oxygen isotope ratio of 0‰ parable with a sea level fall of ~100 m estimated for the Serpukhovian 18 VSMOW, whereas average δ Oseawater during the Pleistocene Last Gla- to Bashkirian time interval (Rygel et al. 2008). The remaining 0.8‰ in- cial Maximum is suggested to have been around +1‰ VSMOW crease in δ18O could be interpreted as reflecting cooling of ~3 °C–4°C, (Schrag et al. 1996). Since the Naqing section represents a slope setting which generally agree with 1 °C–6 °C cooler conditions for low- suggesting fully marine conditions, significant evaporation or fresh latitude surface seawaters during the LGM as constrained by Mg/Ca water inputs by rivers are implausible. Thus, the observed changes in ratios of foraminifera and further independent evidence (Bard et al. the oxygen isotope ratios of apatite phosphate are interpreted to reflect 1997; Pelejero et al. 1999; Lea et al. 2002; Rosell-Melé et al. 2004; Barker primarily changes in ice volume and seawater temperature. et al. 2005; Hippler et al. 2006; Beck et al. 1992; Guilderson et al. 1994; 18 Assuming a δ Osea for Carboniferous and Permian seawater of Stute et al. 1995; Thompson et al. 1995; McCulloch et al. 1999). −1‰,0‰,or1‰ VSMOW, calculated palaeotemperatures are between We suggest that the highest δ18O values in the Middle Bashkirian 22 °C and 33 °C, between 17 °C and 28 °C, and between 13 °C and 24 °C, reflect maximum glaciation of the entire LPIA. A coincident increase in 13 respectively. Temperatures ranging from 17 °C to 28 °C or from 22 °C to δ Ccarb (Grossman et al. 2008, Fig. 3) supports this assumption as an 13 33 °C are comparable to modern tropical sea surface seawater tempera- increase in δ Ccarb is normally interpreted to reflect an increase in C tures which are generally between 20 °C and 28.5 °C (Locarnini et al. sequestration and a consequent decrease in CO2 levels culminating in 18 2006). However, δ Osea values between 0‰ and 1‰ VSMOW are global cooling. In addition, nitrogen isotopes of marine sediments higher than values inferred from carbonate clumped isotopes. The latter show a pronounced increase in marine δ15N starting from the are a proxy for temperature independent of the oxygen composition of Tournaisian and culminating in the Serpukhovian–middle Bashkirian the solution. Combining carbonate δ18O and clumped isotope values (ca. 358.9–323.2 Ma) representing δ15N peak values of the entire allows to calculate the oxygen isotope composition of the solution Phanerozoic (Algeo et al. 2014; Yao et al. 2015; Fig. 3). Strontium from which the carbonate precipitated. Clumped isotopes measured isotopes exhibit a similar but later increase starting in the middle on brachiopod calcite translate into Middle Pennsylvanian mean tropi- Viséan (ca. 341 Ma) and reaching a maximum in the middle Bashkirian cal seawater temperatures of 24.9 °C ± 1.7 °C corresponding to a seawa- (ca. 320 Ma) (Bruckschen et al. 1999; Fig. 3). ter δ18Ovaluesof−1.6‰ ±0.1‰ (VSMOW; Came et al. 2007)and Algeo et al. (2014) explained the high δ15Nvaluesinthe 18 Carboniferous–Permian δ Oseawater values ranging from −0.7‰ to Serpukhovian to middle Bashkirian as consequence of enhancing −2.2‰ using the equation proposed by Veizer and Prokoph (2015). water-column denitrification resulting from the glacio-eustatic sea According to Veizer and Prokoph (2015), the Early Palaeozoic ocean level fall or change in the oceanic circulation as a consequence of the 18 should have had lower δ Osea than the Late Palaeozoic ocean. Interest- closure of the Rheic Ocean. This explanation is consistent with a low ingly, for the Hirnantian glacial maximum, clumped isotopes indicate sea level as suggested by highest oxygen isotope ratios (maximum ice 18 87 86 that tropical seawater temperatures were around 28 °C and δ Osea volume effect). The shift in Sr/ Sr is seen in context with the final was between 2‰ and 3‰ (VSMOW), which would suggest that closure of the Rheic Ocean and the Ouachita–Alleghanian–Variscan Hirnantian ice volume was comparable or higher as during the last orogeny (ca. 370–330 Ma, Nance et al. 2010) but precedes the shifts in (Pleistocene) glacial maximum (Finnegan et al. 2011). carbon and oxygen isotopes and thus might suggest that the orogeny lead to intensified weathering of continental rocks delivering radiogenic 5.2. The Early Pennsylvanian glacial maximum 87Sr to the ocean. Enhanced silicate weathering may have consumed carbon dioxide and/or released nutrients to the oceans, providing an The oxygen isotope record from South China shows a significant additional mechanism for lowering greenhouse gas levels and cooling δ18Oincrease(~1.9‰) during the Mississippian–Pennsylvanian transi- the Earth's surface (Raymo 1994; Raymo and Ruddiman 1992). tion with the δ18O maximum occurring in the middle Bashkirian It is interesting to note that seawater 87Sr/86Sr remains relatively (Figs. 2 and 3). This increase in δ18O agrees well with previously pub- high during the Pennsylvanian (Fig. 3) with a substantial decline only lished oxygen isotope records. For example, an increase in δ18Oof occurring in the latest Pennsylvanian. Similar high seawater 87Sr/86Sr well-preserved brachiopod calcite was observed from the Serpukhovian ratios were also observed in the Cenozoic. For example, the past to Middle Bashkirian, both on the U.S. Midcontinent as well as Russian 2.5 Myr experienced a significant rise in seawater 87Sr/86Sr of platform (Fig. 3; Grossman et al. 2008; Mii and Grossman 1999; Mii ~1.4 × 10−5 (Hodell et al. 1991; Capo and DePaolo 1990). This rise is et al. 2001). Grossman et al. (2008) noted that the Bashkirian brachio- time equivalent to the onset of the Northern hemisphere glaciation pod δ18O values are comparable or higher than those of the Asselian. (e.g., Shackleton et al. 1984; Tiedemann et al. 1994), with an increased Aridification as a potential cause of the high δ18O values of Bashkirian flux of radiogenic strontium delivered by the weathering of fine- brachiopods is ruled out since the Mississippian/Pennsylvanian transi- grained glacial tills being responsible for this long-term rise in tion in the USA normally coincides with a shift toward more humid 87Sr/86Sr (Blum and Erel 1995). However, recent high-precision conditions (Cecil 1990). Conodont apatite oxygen isotope records mea- 87Sr/86Sr analyses on planktonic foraminifers show no resolvable varia- sured from low-latitude Euramerican sections (Buggisch et al. 2008) tion through the last glacial cycle (Mokadem et al. 2015), thus show a significant δ18O increase in the Serpukhovian (Fig. 3). The questioning this interpretation. Nevertheless, the resemblance of high coincidence of the Early Pennsylvanian δ18O maximum in oxygen 87Sr/86Sr ratios of seawater during both the LPIA and the Pleistocene gla- isotope records of both conodont apatite and brachiopod calcite from ciation suggest that glaciation may have played a role for the Sr cycle. various low-latitude sections suggests that this maximum is of global The coincidence of highest δ18O values and a significant sea level significance. lowering documented in various low-latitude sedimentary sequences

87Sr/ 86 Sr

0 0. 0 0 0. 0. 0

0.70800 Magnitude of

. . .

706

7 70 70 7 7

18 18 13 7 Sea-Level Coastal onlap

0 0760 07 δ δ δ 0 high-frequency

O (‰V-SWOM) O(carb ‰V-PDB) C(carb ‰V-PDB) Eustasy

7 70 7

8 Changes Ukraine

8 4 8 2

2 U. S. eustatic sea-level

0 0 0 0 0 18 20 22 24 -6 -4 -2 0 0 2 4 6 0 (m above PD) (km) change (m) Changhsin. 255 Wuchiapin.

260 Loping. Relief an Capitanian (carbonates) pi

265 u Relief al Wordian (clastics) Facies 270 Roadian juxtaposition Guad Cycle thickness 275 xxx (2016) xxx Palaeoecology Palaeoclimatology, Palaeogeography, / al. et Chen B. Kungurian Geochemical proxies 280

Artinskian 285

290 Cisuralian Sakmarian Moscow Basin Long-term 295 sea-level Asselian onlap curve lowstands + 0 300 Gzhelian

305 Kasimovian ian n a

310 lv Moscovian

315 nnsy e P http://dx.doi.org/10.1016/j.palaeo.2016.01.002 Bashkirian 320

325 Serpuk- hovian 330

335 an

Viséan 340 – xxx 345 Mississippi 350 35 30 25 Toutnaisian 30 25 20 355 25 20 15 0 4 8 12 (°C) 100 0m 200 100 0 20015010050 20 60 100140 δ15N(‰)

Fig. 3. Comparison of oxygen isotopes of conodont apatite from the Naqing section and Euramerica with published oxygen and carbon isotope data of brachiopod shells, 87Sr/86 Sr, δ15 N, as well as eustatic sea level. Blue line—Naqing section (this study), red line—Buggisch et al. (2008). Yellow-shaded zones indicate interval characterized by a significant coeval oxygen isotope increase during the Mississippian–Pennsylvanian transition. Trend lines are calculated using a Locfit regression with 95% confidence interval absolute ages after the Geological Time Scale 2012. Temperatures calculated using equation of Pucéat et al. (2010) assuming a δ18 O for Carboniferous and Permian seawater of −1‰(blue), 0‰(red), or 1‰ (black) VSMOW. Oxygen and carbon isotope data of brachiopod shells from Grossman et al. (2008): green line—data from USA, dark blue line—data from Russia. 87Sr/86Sr from McArthur et al. (2012) and Bruckschen et al. (1999), δ15 NafterAlgeo et al. (2014),global eustatic curve after Haq and Schutter (2008), U.S. Midcontinent eustatic curve from Ross and Ross (1987), coastal onlap curve for the Moscow Basin after Alekseev et al. (1996), coastal onlap curve for the Donets Basin after Eros et al. (2012).Mag- nitudes of eustatic sea level changes modified from Rygel et al. (2008) after removing the model-based estimate. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.) 5 6 B. Chen et al. / Palaeogeography, Palaeoclimatology, Palaeoecology xxx (2016) xxx–xxx further supports maximum glaciation in the Bashkirian. In general, the Warm-water benthic foraminifers are generally considered as a magnitude of eustatic sea level changes inferred from low-latitude se- good indicator for palaeoclimate changes during the Carboniferous quences is thought to represent a more sensitive proxy for ice volume and Permian. The lowest fusulinid diversity on the North American con- changes than the geographic distribution of high-latitude glacial de- tinental shelves was observed in the Bashkirian. Davydov (2014) argued posits (Montañez and Poulsen 2013). For example, ice volume estimat- that the Bashkirian represented the coldest climate conditions during ed from the extent of glacial deposits may be biased due to uncertainties the Permo-Carboniferous, at least on the North American continental in chronostratigraphy, difficulties in determining the type of the ancient shelves. In addition, major ammonoid lineages disappeared at the ice cover and the potential override of former glacial deposits by subse- Mississippian–Pennsylvanian transition (Saunders and Ramsbottom quently advancing ice sheets. The Mississippian–Pennsylvanian bound- 1986). Conodonts underwent a crisis and showed low diversity at ary is recognized as one of the major eustatic low-stands of the Late the Mississippian–Pennsylvanian boundary (Ziegler and Lane 1987; Paleozoic (Haq and Schutter 2008, Fig. 3), with lowest sea level being Nemirovskaya and Nigmadganov 1994). The vegetation patterns from observed in eustatic sea level or coastal onlap curves reconstructed northern mid-latitude Angara indicate climate cooling during the from the U.S. Midcontinent (Ross and Ross 1987, Fig. 3), the eastern Serpukhovian–Bashkirian and Early Permian (Fig. 4), with the former and southern Great Basin in North America (Bishop et al. 2009; event known as the Ostrogsky Cooling Episode (Meyen 1982, 1987; Martin et al. 2012), the Moscow Basin in Russia (Alekseev et al. 1996; Durante 2000). The expansion of Paleozoic tropical rainforests is Fig. 3), and the Donets Basin in Ukraine (Eros et al. 2012; Fig. 3). In ad- found to have started in the Serpukhovian and was interpreted as the dition, the Mississippian–Pennsylvanian unconformity is well known in consequence of the exposure of continental shelves due to eustatic sea cratonic sequences of North America and regarded as a global eustatic level lowering. However, the maximum expansion occurred only in event (Saunders and Ramsbottom 1986), resulting in the formation of the late Bashkirian (Cleal and Thomas 2005) and thus postdated the 35–150 m deep paleovalleys in eastern North America (Siever 1951; proposed glacial maximum. Bristol and Howard 1971; Rice 1984; Howard and Whitaker 1988; Our interpretation of the LPIA glacial maximum occurring in the Droste and Keller 1989; Beuthin 1994; Archer and Greb 1995; Blake Bashkirian is in contrast with the generally accepted view that ice vol- and Beuthin 2008) and 100 m of palaeokarst in North Africa ume climax of the LPIA occurred in the Late Pennsylvanian (Gzhelian) (Lemosquet and Pareyn 1983). The erosion of latest Serpukhovian stra- to Early Permian (Sakmarian) (Laskar and Mitra 1976; Veevers and ta and the occurrence of exposure surfaces in carbonate sequences in Tewari 1995; López-Gamundí 1997; Visser 1997a; Isbell et al. 2003a, South China are interpreted to correspond to this eustatic event 2003b, 2008b, 2008c; Fielding et al. 2008a-c; Martin et al. 2008; Mory (Wang et al. 2006, 2013). This sea level low-stand is generally et al. 2008; Rocha-Campos et al. 2008; Stollhofen et al. 2008; Melvin interpreted as reflecting ice mass buildup in high-latitude areas, with et al. 2010; Fig. 4). Evidence supporting an Early Permian ice volume cli- an estimated lowering of sea level of up to 140 m according to Rygel max is mainly based on the geographic distribution and extent of glacial et al. (2008). It is worth to note that except for model-based estimates, deposits. Pennsylvanian to Early Permian glacial sediments were de- Rygel et al. (2008) report the highest amplitudes in sea level change in scribed to have been deposited over highlands throughout middle and excess of 100 m to have occurred in the earliest Bashkirian and high palaeolatitudes of the southern hemisphere including Antarctica Kasimovian, the latter predating the proposed climax of the LPIA during (Isbell et al. 2003a), South America (Rocha-Campos et al. 2008; Holz the Gzhelian to Sakmarian (Fielding et al. 2008b; Isbell et al. 2003b). et al. 2008), South Africa (Visser 1997a, 1997b; Stollhofen et al. 2008), Several significant bio-events occurred at the Mississippian– India (Wopfner and Casshyap 1997), eastern and Western Australia Pennsylvanian transition, with climate cooling and/or loss of habitat (Fielding et al. 2008a, 2008c; Mory et al. 2008), the Arabian Peninsula resulting from the large scale glaciation and glacio-eustatic sea level (Martin et al. 2008), and the South Asian crustal blocks (Taboada fall seen as the main triggers of these events. Rates of origination and 2010). By contrast, Serpukhovian–Bashkirian glacial deposits were only extinction of marine genera dropped sharply (Stanley and Powell reported from western and southern South America (López-Gamundí 2003) during the so-called “Serpukhovian bio-crisis.” This crisis is 1997; Isbell et al. 2003b; Limarino and Spalletti, 2006; Limarino et al. ranked as the fifth of the Phanerozoic biodiversity crises in terms of 2006; Henry et al. 2008; Holz et al. 2008; Rocha-Campos et al. 2008), ecological impact, suggestive of having a larger impact to the ecosys- eastern and Western Australia (Fielding et al. 2008a, 2008c; Mory et al. tems than e.g. the end-Ordovician mass extinction (e.g. McGhee et al. 2008), as well as South Africa (Isbell et al. 2008a). As a consequence, 2012, 2013). A significant turnover in taxa composition or fossil di- the Early Pennsylvanian glaciation was interpreted to represent an al- versity was identified for several climate-sensitive marine inverte- pine glaciation with considerable smaller ice volume in comparison brate groups, however, the severity and duration of these events with the latest Carboniferous to Early Permian glaciation (Isbell et al. may vary for different faunal groups. For example, major ecological 2003a, 2003b, 2012). changes are observed for corals in South China, with the dominating The obvious contradiction between the glacial deposit and oxygen large, solitary corals being replaced by colonial rugose corals. This re- isotope records raises the question whether the most widely distributed placement was suggested to be time equivalent with a major regres- glacial deposits represent largest ice volume. This question was previ- sion considered to have been triggered by the glaciation in ously addressed by González-Bonorino and Eyles (1995),whosug- Gondwana (Wang et al. 2006). A coeval decline in coral diversity gested an inverse relationship between ice extent and the Late was also recognized on the northern Island of Novaya Zemlya Paleozoic Gondwanan glacial record. The authors argued that the (Kossovaya 1996). Palaeoreef studies show a 97% loss of reefs at Bashkirian ice cover represented maximum ice volume throughout the Mississippian–Pennsylvanian transition (Flügel and Kiessling the entire Late Paleozoic, despite the distribution of glacigenic sedi- 2002). Further, the initiation and following development of the ments of this age is limited compared to the latest Carboniferous to LPIA is thought to have played a significant role for brachiopod diver- Early Permian. This inverse relationship was explained in the context sity and faunal composition (Qiao and Shen 2014, 2015). From the of regional tectonism. The Bashkirian glaciation was interpreted to Viséan to Serpukhovian, global brachiopod diversity centers migrat- have occurred immediately after wide regional uplift of the interior of ed northward followed by an enhanced faunal provincialism in the Gondwana as consequence of intraplate compressional stresses origi- Serpukhovian (Qiao and Shen 2014). In addition, a distinct low nating in protracted collisions with Laurussia (Variscan orogeny, brachiopod diversity is recognized at the Serpukhovian–Bashkirian Ziegler 1988), and cratonward subduction and microcontinent colli- transition based on brachiopod data from South China and the Cen- sions along the Pacific margin. Tectonic compression resulted in an up- tral Appalachian Basin in the USA (Shen et al. 2006; Powell 2008). lift of many regions within Gondwana, which became glaciated and The well-known Gigantoproductus faunas with a low thermal toler- covered by ice sheets, potentially reflecting the most extensive ice ance became extinct during this time interval (Qiao and Shen 2015). cover of the Late Paleozoic Ice Age. However, in many areas, this

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Fig. 4. Comparison of conodont apatite δ18 O records from the Naqing section (South China, this study) and Euramerica (Buggisch et al. 2008) with the glacial record from Gondwana (Montañez and Poulsen 2013), the palaeoclimate reconstruction from South America (Limarino et al. 2014) and Angara (Meyen 1982, 1987 and Durante 2000), regional tectonism, as well as ice extent estimation for Gondwana (González-Bonorino and Eyles 1995). Geochemical data and yellow zone is the same as in Fig. 3. Glacial phases are color coded: early Carboniferous—pink, late Carboniferous—green, latest Carboniferous to Early Permian—blue, late Early–Middle Permian—orange. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.) 7 8 B. Chen et al. / Palaeogeography, Palaeoclimatology, Palaeoecology xxx (2016) xxx–xxx extensive ice cover may have been not preserved as evidenced by a re- continental rocks and increasing input of radiogenic 87Sr (higher seawa- gional unconformity separating Early Permian glacial (Glacial III) de- ter 87Sr/86Sr ratios) to the oceans. Enhanced silicate weathering may posits from underlying Devonian and older strata as well as crystalline have consumed carbon dioxide and/or released nutrients to the oceans basement rocks (Collinson et al. 1994; López-Gamundí et al. 1994; promoting higher carbon burial rates, providing a mechanism for lower- Veevers et al. 1994a, 1994b, 1994c; Veevers and Tewari 1995; ing greenhouse gas levels and cooling the Earth's surface. The following López-Gamundí 1997; Wopfner and Casshyap 1997). In contrast, during ice buildup is indicated by sea level lowering recorded in various low- late Bashkirian to Gzhelian times, the change to an extensional stress re- latitude sedimentary successions. Climate cooling and loss of habitat gime resulted in the development of intracratonic basins (Ziegler 1988). may have influenced the biosphere. Several bio-events characterized The extension lead to enhanced subsidence, and as a result, a number of by faunal replacement or diversity decrease occurring around this basins were flooded, allowing the deposition and preservation of thick time interval may be related to the proposed glacial maximum. Howev- glacially influenced marine deposits accumulating at the margins of er, the exact cause–effect relationship between glaciation development ice-covered areas (Isbell et al. 2003a). As a consequence, glacio- and bio-events need further work for clarification. marine strata were preserved across a large area of the Gondwana su- percontinent. According to this interpretation, increasing areal distribu- Acknowledgments tion and stratigraphic thickness of the Early Permian glacial deposits can be seen as a consequence of glaciers decay rather than evidence for peak This study was financially supported by the Natural Science Founda- glacial activity. tion of China (grant nos. 41302019, 41290260, J1210006). We thank The view that the pre-Permian unconformity resulted from wide- two anonymous reviewers for valuable comments that greatly im- spread sub-glacial erosion and, hence, reflects occurrence of large ice proved the manuscript. sheets (Veevers and Powell 1987; González-Bonorino and Eyles 1995) is challenged by weathering profiles and soft-sediment deformation observed immediately below the unconformity across the Transantarctic Appendix A. Supplementary data Mountains in Antarctica. 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