SUPPLEMENTARY INFORMATIONDOI: 10.1038/NGEO1667

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Supplementary Information

Climate and biotic upheavals following the end-Permian mass extinction

Carlo Romano, Nicolas Goudemand, Torsten W. Vennemann, David Ware, Elke Schneebeli-Hermann, Peter A. Hochuli, Thomas Brühwiler, Winand Brinkmann and

Hugo Bucher

To whom correspondence should be addressed: carlo.romano@pim.uzh.ch (C.R.), goudemand@pim.uzh.ch (N.G.), hugo.fr.bucher@pim.uzh.ch (H.B.)

This file includes: Page 2: Geological setting and material Page 3: Preparation of samples and measuring technique Page 4: Biochronology Page 4: Extended results Page 6: Calculation of temperature changes Page 6: Additional notes on data used in Figure 2 Page 10: Supplementary figure S1: Distribution of ammonoids in the Griesba-

chian and Dienerian of Nammal Nala Page 11: Supplementary figure S2: Oxygen isotope compositions of conodont

pectiniform elements (δ18Ocp) from the Early Triassic of Nammal Nala plotted against lithology and time

Page 13: Supplementary figure S3: Oxygen isotope compositions of whole fish teeth of the ‘Saurichthys type’ (δ18Oft) from the Early Triassic of Nammal Nala plotted against lithology and time

Page 15: Supplementary figure S4: Bivariate plot between δ13Corg and δ18Ocp values from the late Griesbachian to the late Smithian of Nammal Nala

Page 16: Supplementary table S1: Oxygen isotope compositions of conodont pectiniform elements (δ18Ocp) and whole fish teeth of the ‘Saurich-thys type’ (δ18Oft) from Nammal Nala

Page 18: Supplementary table S2: Average temperatures calculated from δ18Ocp values using different temperature equations and based on dif-ferent estimates of seawater δ18O values

Page 19: Supplementary table S3: Data matrix for the bivariate plot in Sup-plementary figure S4

Page 21: Supplementary table S4: Smithian ammonoid taxa of the Northern Indian Margin and their distribution from the Dienerian-Smithian boundary to the Smithian-Spathian boundary

Page 24: Supplementary references

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Geological setting and material The conodont elements and fish teeth used for the present oxygen isotope study

were sampled from the Lower Triassic Mianwali Formation of Nammal Nala (Nammal Gorge, Salt Range, Pakistan, Fig. 1). The Salt Range is located about 200 km southwest of Islamabad. During the Early Triassic the land masses were assembled into the super-continent Pangaea and the Salt Range was located on the northern Gondwanan shelf (Northern Indian margin, Fig. 1), at a palaeolatitude of ca. 30° South1.

Carbonate rock samples from the Mianwali Formation were collected between 2007 and 2010 during four seasons of extensive fieldwork by our research group in col-laboration with the Pakistan Museum of Natural History, Islamabad. At Nammal Gorge, the Mianwali Formation is about 118 m thick and it consists of siltstones, sandstones, limestones, and dolomites2. The formation can be subdivided into seven lithological units: the Kathwai Member, the Lower Ceratite Limestone, the Ceratite Marls, the Ceratite Sandstone, the Upper Ceratite Limestone (including the Bivalve Beds at its top), the ‘Niveaux intermédiaires’, and the Topmost Limestone (see e.g. Hermann et al.2 for detailed information on the study area). The material used for the current analyses comes from all lithological units of the Mianwali Formation except the Topmost Lime-stone (Fig. 2).

Several studies have already focused on variations in the oxygen isotope record of marine Early Triassic whole rock carbonates3-7. However, biogenic carbonates (for ex-ample, articulate brachiopod shells or foraminifers) and phosphates (e.g. vertebrate apa-tite) have a much better potential to retain their primary isotopic composition than bulk rock carbonates8. Apatite (for instance, conodont elements or gnathostomian tooth en-amel) has proven to be especially useful for palaeotemperature studies as this mineral is generally resistant to post-depositional alteration8-13. Korte et al.14 published nine oxy-gen isotope values of Early Triassic conodont elements (tooth-like structures of basal jawless vertebrates that lived from the Cambrian to the end of the Triassic) from sec-tions in Iran and Italy. However, their data are too spotty and are thus inadequate given the high resolution against which Early Triassic events must be scaled. Meanwhile, Joachimski et al.10 presented a high-resolution conodont-based 18O record for the latest Permian to early Griesbachian interval of south China. Their data indicate that tempera-tures were rising at the end of the Permian leading to warm climate at the onset of the Triassic. The main extinction phase in the classic sections of Meishan and Shangsi near the close of the Permian coincided with a rise in temperature, indicating that global warming played an important role10.

Conodont elements are ubiquitous in the Early Triassic of the Salt Range and are exceptionally well preserved (Conodont Alteration Index15 of 1, which is very rare for the Early Triassic). Therefore, these translucent conodont elements represent an excep-tional archive for determining the Early Triassic oxygen isotope evolution. Microscopic ray-finned fish teeth are not as abundant as conodont elements in the studied interval of the Mianwali Formation. The fish teeth could not be identified with maximal taxonomic accuracy. However, we selected only teeth of a distinct type for our analysis, that is, high-conical teeth with a smooth enamel cap and a striated basal portion. Such teeth are usually assigned to the basal ray-finned fish Saurichthys Agassiz, 1834. Because of their small size, whole fish tooth samples including the dentine base were used for the analy-sis. Because dentine is porous and has an originally higher organic matter content, it is more prone to diagenetic alteration compared to enamel12; hence, the oxygen isotope

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record of fish teeth is herein treated separately and is neither used for temperature calcu-lations nor for palaeoclimatic interpretations. As the amount of dentine relative to that of enamel was variable in the fish teeth samples, different degrees of potential alteration can be assumed (post-depositional modification normally leads to 18O depletion, i.e. lower δ18O values12). Preparation of samples and measuring technique

The conodont elements and actinopterygian fish teeth were obtained by dissolution of bedrock-controlled carbonate samples. This work has been done at the Palaeontologi-cal Institute and Museum, University of Zurich, Switzerland. Rock samples were dis-solved with 10% acetic acid for about one week. The residues were density-fractionated using a sodium-polytungstate solution to separate higher-density minerals (for instance, conodont elements and fish teeth) from those with lower density (for example, quartz) using the technique of Jeppsson & Anehus16. About 600 to 4600 µg of conodont pectini-form elements (one to three taxa per sample, see Supplementary Table S1) and between 900 to 3100 µg of microscopic fish teeth of the ‘Saurichthys type’ were then picked from the heavy fraction of each sample. The conodont elements and fish teeth were treated as two independent datasets during further analysis.

The conodont elements and fish teeth were chemically processed and measured at the Institute of Mineralogy and Geochemistry, University of Lausanne, Switzerland, according to a method adopted from Dettman et al.17. First, the samples of conodont elements and fish teeth were cleaned in 2.5% NaOCl (for about 24 h) to remove pos-sible organic compounds and then rinsed twice with distilled/deionised water. About 800 µl of 2 n hydrofluoric acid (HF) was added and the samples were dissolved over-night, leaving behind a precipitate of CaF2. During dissolution, the samples were re-peatedly treated with ultrasound. The phosphate solution was then decanted and mixed with 143 µl of ammonium hydroxide (NH4OH) to obtain a neutral pH. Subsequently, 800 µl of 2 n silver nitrate (AgNO3) was added and a yellowish precipitate of tri-silver phosphate (Ag3PO4) was instantly observed. This rapid precipitation method ensures an unbiased isotopic composition of the Ag3PO4. After the crystals had settled, the over-standing solution was discarded and the precipitates were washed twice with distilled water. The samples were then dried overnight in the oven at 70°C. Finally, small por-tions (about 300 to 1000 µg per subsample) of homogenized Ag3PO4 powder were weighed into silver capsules for isotopic analysis. One to several phosphate subsamples of the same bed were analysed for oxygen isotope composition depending on the amount of material available.

Oxygen isotope analyses of the phosphate subsamples follows the procedures summarised by Vennemann et al.18 using the technique of on-line, high-temperature re-duction (TC-EA) of the previously prepared Ag3PO4 in the presence of He carrier gas. The CO produced by reduction at temperatures of about 1450°C is passed into a Finni-gan MAT Delta Plus XL mass spectrometer. This technique minimizes the amount of sample required, so that a few hundred micrograms of Ag3PO4 per analysis will suffice. Analysed sequences were normalised (corrected for drift and different reference gas values for samples run several months apart) by linear regression to the values reported for the TU-1 and TU-2 standards run with the same sequence as the samples18. When normalized to TU-1 and TU-2 of 21.1‰ and 5.5‰, respectively, NBS120c prepared in the same way as the samples gives values of 21.6±0.3‰.

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Stable oxygen isotope compositions of phosphates are expressed as the difference in ratio between 18O and 16O of the sample and the Vienna Standard Mean Ocean Water (VSMOW), relative to VSMOW and reported in per mill as the δ18O value19. Stable carbon isotopes are expressed as the δ13C value (in ‰), relative to the Vienna Pee Dee Belemnite (VPDB) international standard.

Biochronology The Early Triassic is divided into four substages: Griesbachian, Dienerian,

Smithian, and Spathian20. U-Pb ages indicate a total duration of ca. 5 million years for the Early Triassic and the Spathian substage represents slightly more than half of the duration of this epoch21-23.

For the lower part of the Mianwali Formation (Griesbachian-Dienerian interval), new and robust correlations of the ammonoid sequence with other coeval faunas from the northern Indian margin are granted thanks to profound similarities at the species level. However, how the local Salt Range sequences24,25 compare to high paleolatitude successions known from Canada20, north-east Siberia26, and South Primorye27 (south-east Russia) still requires improving the taxonomic consistency of the entire data set.. During field work, more than 1700 ammonoid specimens have been collected from 22 different beds of the lower part of the Mianwali Formation. Based on a revised tax-onomy a new biostratigraphic scheme was constructed (Supplementary Fig. S1) based on the unitary association method of Guex25. This led us to differentiate fourteen resid-ual maximal horizons (MH) within the lower Mianwali Formation at Nammal Nala, three in the Griesbachian (MH-G1 to 3) and eleven in the Dienerian (MH-D1 to 11). The Griesbachian-Dienerian boundary is here placed using the first occurrence of the conodont Sweetospathodus kummeli (Sweet, 1970) (Supplementary Fig. S1, Supple-mentary Table S1).

A refined ammonoid zonation based on the unitary association method of Guex25 for the Smithian of the Northern Indian Margin, including the Salt Range, was recently published by Brühwiler et al.28,29. These authors distinguished thirteen ammonoid zones for the Smithian of the Salt Range and they divided this substage into early, middle, and late Smithian (Fig. 2; Supplementary Figs S2−S3). However, in Nammal Gorge, two of these thirteen zones, the Radioceras evolvens Zone and the Shamaraites rursiradiatus Zone, could not be identified, but there is rock bracketed by the zones above and be-low28. The base of the Smithian was defined by the first appearance of typical Smithian ammonoids (Flemingitidae and Kashmiritidae)30. The Smithian-Spathian boundary is here placed at the base of the Bivalve beds. For each Smithian ammonoid zone of the Northern Indian Margin an average maximal duration of 67 kiloyears was estimated28.

The Spathian zonation of the Salt Range and Surghar Range (Fig. 1) is in construc-tion. The ranges of some of the ammonoids from the Bivalve Beds and the ‘Niveaux intermédiaires’ are indicated in Supplementary Figs S2−S3. Basal Spathian faunas char-acterized by the Tirolites-Columbites association are locally missing in Nammal. Their absence is probably due to a local hiatus near the base of the Bivalve Beds2.

Extended results Oxygen isotope values of conodont pectiniform elements (δ18Ocp) and fish teeth of

the ‘Saurichthys type’ (δ18Oft) from the Lower Triassic Mianwali Formation of Nammal

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Gorge (Salt Range, Pakistan, Fig. 1) are plotted against lithology and time (means with standard deviations) in Supplementary Fig. S2 (for δ18Ocp) and Supplementary Fig. S3 (for δ18Oft). The data for both records are summarised in Supplementary Table S1.

In the upper Kathwai Member and the basal Lower Ceratite Limestone (LCL, upper Griesbachian-lowest Dienerian, oxygen isotope phase Ia) of the Mianwali Forma-tion, mean δ18Ocp values start around 17‰ (all δ18O values are reported relative to VSMOW) and mean δ18Oft values near 14‰ (Supplementary Figs S2−S3). At the base of the LCL, a remarkable positive shift occurs, with peak mean values of 20.5‰ (δ18Ocp) and 18.65‰ (δ18Oft) within MH-G3 (Supplementary Figs S2−S3). Because it is documented by both conodont elements and fish teeth (both from bed LCL-C1, see Supplementary Table S1), this positive δ18O shift could be a valid signal but its signifi-cance needs further evaluation.

Near the boundary between the LCL and the Ceratite Marls mean δ18Ocp values de-crease below 17‰ and average δ18Oft values below 15‰. Mean values of both records remain low for most of the basal third of the Ceratite Marls (Ib, Supplementary Figs S2−S3). This interval of the Mianwali Formation corresponds to the middle and upper part of the Dienerian and the lowest Smithian28.

A distinct positive shift in the δ18Ocp record is observed in the earliest Smithian, i.e. from the Flemingites bhargavai Zone to around the Flemingites nanus Zone (just above Ib-Ic boundary, Supplementary Fig. S2). Despite the gap in the δ18O record between the upper Ceratite Marls and the lower Ceratite Sandstone due to the scarcity of carbonate beds within this part of the Mianwali Formation, relatively high δ18Ocp values (>17‰) seem to prevail up to almost the end of the early Smithian (oxygen isotope phase Ic, Supplementary Fig. S2).

A marked negative δ 18Ocp shift of approximately 1.3‰ is observed within the upper Ceratite Sandstone (Ic-Id transition, near the early Smithian-middle Smithian boundary, Supplementary Fig. S2, Supplementary Table S2). This prominent decrease leads to a plateau of relatively low δ18Ocp values in the basal part of the Upper Ceratite Limestone (Id, middle and late Smithian, Supplementary Fig. S2). Mean δ18Ocp values for the middle Smithian are 16.2‰; mean values for the late Smithian Glyptophiceras sinuatum Zone are 15.5‰ for δ 18Ocp, and 16.7‰ for δ 18Oft (Supplementary Tables S1−S2). Note that currently no δ 18Oft data are available between the F. nanus Zone (early Smithian) and the Wasatchites distractus Zone (late Smithian) because fish teeth are rare in the upper Ceratite Marls, the Ceratite Sandstone, and the Upper Ceratite Limestone at Nammal Gorge (Supplementary Fig. S3).

A large increase in δ18Ocp and δ 18Oft values is observed across the Smithian-Spathian boundary, which is also used to separate oxygen isotope phases I and II. A shift of 2.7‰ is registered between beds Nam42 (Glyptophiceras sinuatum Zone) and BV1 for the δ18Ocp record (~1.7‰ difference between means of phases Id and II, Sup-plementary Fig. S2, Supplementary Table S2), and a rise of ~1.5‰ is recorded between beds Nam32 (Glyptophiceras sinuatum Zone) and BV3 for the δ18Oft data (Supplemen-tary Table S1, Supplementary Fig. S3). Relatively high δ18Ocp and δ18Oft values (mostly >18‰) are recorded in the Bivalve Beds and the base of the ‘Niveaux intermédiaires’ (II, early Spathian). The δ18O trend within the Bivalve Beds is characterised by abrupt oscillations (Supplementary Figs S2−S3). Depositional hiatuses within the Bivalve Beds were first documented by sedimentological analyses2 and the δ18O excursions observed within the Bivalve Beds thus represent a degraded signal. The overlying ‘Niveaux in-

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intermédiaires’ contain only a limited number of carbonate beds. Hence, only a few δ18O values are available for this part of the section.

Calculation of temperature changes The oxygen isotope composition of apatite precipitated by marine animals is con-

trolled by temperature, but also by the δ18O value of the ambient seawater8. Therefore, the oxygen isotope composition of Early Triassic seawater has to be known in order to convert the measured δ18O values from biogenic phosphate into absolute temperatures. Although knowledge of δ18Oseawater may no longer be needed with the new ‘carbonate clumped isotope’ method of Ghosh et al.31, this approach to calculate palaeotempera-tures requires pristine biogenic carbonates (for instance, shells of articulate brachiopods or bivalves), which are not readily available for the Early Triassic.

There is still an ongoing dispute whether ancient seawater δ18O was buffered close to present day values (0±1‰32) through seawater-oceanic crust isotope exchange or whether it has evolved during Earth’s past33-36. Because the δ18O value of Early Triassic seawater is not known, we only provide relative temperature differences in the main text. Supplementary Table S2 gives absolute temperatures calculated from mean δ18Ocp values of the various oxygen isotope phases, with different estimates of δ 18Oseawater values and using either the calibration of Kolodny et al.8 or a new equation that was re-cently established by Pucéat et al.37. Note that with the calibration of Pucéat et al.37 a secular offset in the seawater δ 18O value through time has to be assumed to obtain plausible temperatures for the various Early Triassic oxygen isotope phases (see also discussion in Pucéat et al.37).

Apart from temperature, the oxygen isotope composition of phosphates may also be controlled by the salinity of the ambient seawater and the amount of ice stored on the continents. The presence of ammonoids and conodonts throughout the studied section indicates normal marine conditions and there is no reported evidence for glacial events during the Early Triassic.

Additional notes on data used in Figure 2 Due to the mixed siliciclastic-carbonate lithology of the Mianwali Formation, only

a few δ 13C values from carbonates are currently available2,3. However, a much higher resolution exists for the organic carbon isotope record from the Salt Range2. The δ13Corg curve from the Salt Range shows the same trends as numerous other globally distributed Early Triassic δ13C records3-5,7,14,21,38-44. Early Triassic carbon isotope records are char-acterised by large oscillations, which began with two45 distinct negative shifts closely spaced in time near the Permian-Triassic boundary2,3,5,10,39,40,42. The carbon isotope data (after Hermann et al.2) shown in Figure 2 were obtained from samples from Nammal Gorge (Griesbachian-Smithian), Salt Range, and from Chitta-Landu (Spathian) in the neighbouring Surghar Range (cf. Fig. 1).

The trend lines for the carbon and oxygen isotope compositions from the Lower Triassic Mianwali Formation were calculated using the nonparametric locally weighted regression method locfit of Loader46, with a 2 σ confidence interval. Because of the sedimentary gap at the very base of the Spathian (see above), separate locfit curves for both the carbon and oxygen isotope data were determined, one for the Griesbachian-Smithian interval and one for the Spathian substage. The two curves were then com-

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bined in Fig. 2. The high mean δ 18O value of LCL-C1 (MH-G3, latest Griesbachian) was omitted for the calculation of the oxygen isotope locfit curve of the Griesbachian-Smithian phase. A smoothing parameter α of 0.28 and a global confidence band were used for both datasets. Calculations were performed with the open source statistic soft-ware ‘R’47 (version 2.13.0, ‘R’ package version 1.5-6).

The calculated locfit curves for the δ13C record from organic carbon and the δ18O data from biogenic apatite of Nammal Gorge (Fig. 2) show a good correlation within the Griesbachian-Smithian interval (Spearman’s ρ = 0.7928, p = 0.001; Kendall’s τ = 0.5985, p = 0.001; Supplementary Fig. S4, Supplementary Table S3). However, near the end of the Smithian the oxygen isotope and carbon isotope curves indicate divergent trends: while δ18Ocp values remain high in the late Smithian and do not increase before the onset of the Spathian, δ13C values already begin to rise. Note that the duration of this lag is currently difficult to constrain. It might have been much shorter than suggested in Figure 2 because unitary associations zones have been shown to be of much shorter dur-ation on average during extinction phases48. These different trends in the oxygen and carbon isotope compositions coincide with the late Smithian anoxic event (see below). Similar lags between δ18O and δ13C values are known from other episodes of increased organic carbon burial, for example, the early Aptian oceanic anoxic event49. However, note that for the upper Smithian only one measurement is currently available for the δ18Ocp record due to only limited amounts of sample material within this part of the sec-tion.

Based on the positive correlation between the δ13Corg and the δ18Ocp records as well as parallel trends in the spore-pollen ratio of palynomorph assemblages at high and low palaeolatitudes40,50,51 and the diversity patterns of conodonts52 and ammonoids28,29,40,53, the temperature trends derived from our oxygen isotope data are inferred to reflect glo-bal climatic changes. However, different mean values and amplitudes of excursions in δ18O values are to be expected for other palaeolatitudes. According to climate model-ing54, for example, an increase in atmospheric pCO2 during Permian-Triassic times would have lead to a greater warming of the polar regions compared with lower palaeo-latitudes, thereby reducing the latitudinal sea-surface temperature gradient. The recur-rent emergence of latitudinally flat ammonoid diversity gradients during the Early Tri-assic (e.g. during the late Smithian) supports episodic modification of the pole-to-equator sea-surface temperature gradient53. Therefore, larger temperature changes may be expected for higher palaeolatitudes.

The original conodont samples used for measuring the δ 18Ocp values consisted of several dozen of pectiniform elements corresponding to one, two or three taxa (see Sup-plementary Table S1). Each data point in Figure 2 and Supplementary Fig. S2 repre-sents the mean value measured from Ag3PO4 samples obtained from the chemical pre-paration of several conodont pectiniform elements of the same layer (see above). With a few exceptions (samples Nam326, Nam330, Nam526), only one round of measurement was conducted per layer (each based on replicate analyses).

No conodont genus ranges across the entire studied interval, thus, it is impossible to use the same genus for all measurements. Although conodont taxa may have dwelled at different depths and, hence, may have recorded differing seawater temperatures, we have tried to minimise this potential bias by selecting specimens that are monogeneric or can be considered as such (for example, Neospathodus and Novispathodus). Al-though the positive δ18Ocp shift at the end of the Smithian (Id-II boundary) concurs with a taxonomic change of the analysed conodonts (Supplementary Table S1), it also over-

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laps in time with a regional regression (see below). Contrary to the documented positive δ18Ocp trend, any genus-specific depth bias should yield more negative oxygen isotope values because of the regression straddling the Smithian-Smithian boundary. This obvi-ous decoupling between water depth and temperature speaks against any important con-trol from the habitat depth of the conodonts. Even though habitat differences and spe-cies-specific effects cannot be theoretically ruled out, the correlation (see Supplemen-tary Fig. S4) of the oxygen isotope curve with the global carbon isotope record and the parallel trends with other proxies suggest that the main trends within the δ18Ocp values represent genuine environmental signals.

The seawater temperatures inferred from the δ18O values of Nammal Gorge do not consistently track sea level changes. Although the negative δ18O shift (warming) be-tween the early and middle Smithian (Ic-Id boundary) coincides with a local regression2, sea-level change cannot account for the increases in δ18O values (cooling) at the begin-ning of the Smithian (base of Ic) and the SSB (I-II boundary), both of which are also concomitant with local regressive events2. We therefore reasonably assume that the measured temperature variations are largely independent from fluctuations of water depth.

The oxygen isotope curve from bulk rock carbonates from part of the same section published by Baud et al.3 shows conspicuous noise and comprises partially very low values. Baud et al.3 therefore assumed that their samples have been strongly affected by diagenesis. Nevertheless, there are some resemblances between the oxygen isotope re-cord from bulk rock carbonates and the one from biogenic phosphates, even though a bed-by-bed comparison is not possible. For instance, δ18O values of both bulk rock car-bonates and conodont pectiniform elements from the middle Ceratite Marls (early Smithian) are generally higher than in the lower part of the Upper Ceratite Limestone (middle Smithian). Additionally, large oscillations within the Bivalve Beds are charac-teristic for both the bulk rock carbonate and the biogenic apatite δ18O record. These similarities suggest that at least some trends seen in the bulk rock carbonate δ18O data from Nammal Gorge may have withstood diagenesis, although the values and ampli-tudes are definitely altered. However, without control from diagenetically unaltered biogenic apatite, the reliability of bulk rock carbonate oxygen isotope records for palaeotemperature reconstructions remains questionable. The nine δ18O values from conodont elements of the western Tethys (Italy and Iran) obtained by Korte et al.14 do not allow reconstruction of any clear trend but do show a comparable range of mean values when corrected for the difference in NBS120c values (22.4‰ in Korte et al.14, note that the record of Korte et al.14 covers a much smaller interval of the Early Tri-assic). δ18O values published by Joachimski et al.10 for the early Griesbachian of south China are approximately 0.5‰ higher than our values from the late Griesbachian when corrected for the difference in NBS120c values (22.6±0.16‰ in Joachimski et al.10).

Although some authors proposed widespread anoxic conditions for vast parts of the Early Triassic55-58, evidence for oxygen depletion (high abundance of amorphous or-ganic matter and total organic carbon) at Nammal Gorge is only found within the lower third of the Ceratite Marls2 between MH-D3 and MH-D11. This interval of the Mian-wali Formation falls within the middle and upper parts of the Dienerian. The middle-late Dienerian anoxic event was a widespread phenomenon that affected the Tethys (south China, northern India, Pakistan)5,59 and the eastern equatorial Panthalassa (Nevada, USA)60. In the Salt Range sections, well-oxygenated conditions prevail during the Gri-esbachian and the early part of the Dienerian as well as during the Smithian and

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Spathian2. However, low oxygenation conditions in shallow waters re-occur during the late Smithian of south China and the Indian Himalaya5 (also included in Fig. 2). In the Salt Range, the overriding clastic flux due to synsedimentary tectonics is likely to have prevented the development of an anoxic water-sediment interface2,5. Modifications in oceanic circulations at the SSB inferred from ammonoid palaeogeography61 are prob-ably linked with the coeval climatic changes.

The diversity curves for the Smithian ammonoids of the Northern Indian Margin (generic richness and turnovers) were produced by combining the data of Brühwiler et al.28 and Brühwiler et al.29 (see Supplementary Table S4). The data indicate an increase in originations at the genus level during the cooler early Smithian (phase Ic) as well as an increase in turnover at the genus level (originations and extinctions) during the sub-sequent warming (Ic-Id boundary, early-middle Smithian boundary). Generic turnover remains elevated and genus richness begins to decrease during the following warm phase (Id), leading to the well-documented global diversity collapse of ammonoids (and also conodonts52) near the end of the Smithian28,40. Seawater temperatures near the up-per tolerance limit of many marine metazoans might have been reached during phase Id (compare Supplementary Table S2). On the family level, emergence of seven new ammonoid families is recognized for the early Smithian opposed to only four new fami-lies being added during the warmer middle Smithian28,29. Eleven families are lost during phase Id (middle-late Smithian) but only one family during Ic

28,29. Increase in turnover rates with warming and low diversity at the genus and family levels during episodes of greenhouse climate conditions is consistent with patterns observed throughout the Phan-erozoic62. The increase in taxonomic diversity at the base of the Smithian (Ic) and the beginning of the Spathian (II) coinciding with cooling from probably high temperature levels resembles patterns documented, for example, for the Great Ordovician Biodi-versification Event36. This indicates that during these time intervals cooling of seawater temperatures may have enabled major diversification pulses.

The presently documented sequence includes a diversification during cool intervals followed by increased turnover rates during the main warm phase and ending with a di-versity crash. Such a three steps cycle in response to global warming is documented for the first time. The first reaction to warming is a massive increase of turnover rates. Mass extinction may then follow only if the biological pump is over-reacting. This Smithian sequence is here referred to as the Smithian model against which other extinctions can be compared to.

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Supplementary figure S1. Distribution of ammonoids in the Griesbachian and Diener-ian of Nammal Nala (Nammal Gorge, Salt Range, Pakistan). Ammonoid taxonomy and distribution are based on more than 1700 specimens (see also text). MH = residual Maximal Horizon; G = Griesbachian; D = Dienerian.

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Supplementary figure S2. Oxygen isotope compositions of conodont pectiniform ele-ments (δ18Ocp, mean values with standard deviation, data given in Supplementary Table S1) from the Lower Triassic Mianwali Formation at Nammal Gorge (Salt Range, Paki-stan), plotted against lithology (left column) and against time (right column). Ammon-oid zonation for the Smithian after Brühwiler et al.28 and Brühwiler et al.29. The early Smithian Radioceras evolvens Zone and Shamaraites rursiradiatus Zone (stipple-lined boxes) are present at other sections within the Salt Range but not documented in Nam-mal Nala. For the Griesbachian and the Dienerian a preliminary biostratigraphic scheme was established (compare Supplementary Fig. S1). For the Bivalve Beds and the ‘Niveaux intermédiaires’ (Spathian) the ranges of identifiable ammonoid taxa are indi-cated (see also text). U-Pb ages for the end-Permian extinction horizon23, the early Smithian Kashmirites kapila beds21 and the early Spathian Tirolites/Columbites beds22 were used to calibrate the δ18Ocp values against time. Timescale: Chang., Changxingian; Perm., Permian. Lithological Units: K, Kathwai Member; L, Lower Ceratite Limestone; CM, Ceratite Marls; CS, Ceratite Sandstone; U, Upper Ceratite Limestone; B, Bivalve Beds; N, ‘Niveaux intermédiaires’. Drawing: Conodont. For abbreviations of the zona-tions for the Griesbachian and Dienerian see caption of Supplementary Fig. S1.

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Supplementary figure S3. Oxygen isotope compositions of whole fish teeth of the ‘Saurichthys type’ (δ18Oft, mean values with standard deviation, data given in Supple-mentary Table S1) from the Lower Triassic Mianwali Formation at Nammal Nala (Nammal Gorge, Salt Range, Pakistan), plotted against lithology (left column) and against time (right column). Oxygen isotope phases of the conodont pectiniform ele-ment δ18O record (see Supplementary Fig. S2) are indicated on the right. Drawing: The Triassic lower actinopterygian Saurichthys Agassiz, 1834. For abbreviations and further information see captions of Supplementary Figs S1−S2.

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Supplementary figure S4. Bivariate plot between δ13Corg and δ18Ocp values from the late Griesbachian to the late Smithian of Nammal Nala (Nammal Gorge, Salt Range, Paki-stan). Raw data is given in Supplementary Table S3. The data indicate a positive corre-lation between the carbon isotope compositions of organic carbon2 and the oxygen iso-tope compositions of biogenic phosphate (conodont pectiniform elements) for the con-sidered interval. The statistical tests (Spearman’s ρ , Kendall’s τ) and plot (“robust method”) were calculated using the PAST data analysis software63.

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Supplementary table S1. Oxygen isotope compositions of conodont pectiniform ele-ments (δ18Ocp, plotted in Supplementary Fig. S2) and whole fish teeth of the ‘Saurich-thys type’ (δ18Oft, plotted in Supplementary Fig. S3) from the Lower Triassic Mianwali Formation of Nammal Nala (Nammal Gorge, Salt Range, Pakistan). The ammonoid zonation for the Smithian follows Brühwiler et al.28 and Brühwiler et al.29. Note that the early Smithian Radioceras evolvens Zone and Shamaraites rursiradiatus Zone, which are present in several other sections of the Salt Range but not at Nammal Gorge, are omitted. For the Griesbachian and Dienerian a preliminary biostratigraphic scheme was established (see Supplementary Fig. S1). Conodont taxa that were used for the oxygen isotope analysis are indicated.

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Supplementary table S2. Mean temperatures of the Early Triassic oxygen isotope phases (Ia–Id, II, see text) calculated from their mean δ18Ocp values using either the phosphate temperature equation of Kolodny et al 8 or Pucéat et al.37 (with correction for difference in NBS120c), and based on different estimates of seawater δ18O values. Note that when using the new calibration of Pucéat et al.37 a lower δ18Oseawater has to be as-sumed to obtain realistic mean temperatures for the oxygen isotope phases. Oxygen iso-tope phase Ia corresponds approximately to the late Griesbachian-earliest Dienerian in-terval, Ib to the middle and late Dienerian and the earliest Smithian, Ic largely to the early Smithian, Id to the middle-late Smithian, and II to the early part of the Spathian (compare Fig. 2; Supplementary Figs S2−S3). δ18O value of LCL-C1 (MH-G3) was omitted for calculation of the mean δ18O value of phase Ia.

Mean δ18Ocp

VSMOW (δp)

Kolodny et al.8 T[°C]=113.3−4.38·(δp−δw)

Pucéat et al.37 T[°C]=118.7−4.22·(δp+1−δw)

δ18Oseawater VSMOW (δw)

0‰ -1‰ -2‰ -3‰ -1‰ -2‰ -3‰ -4‰

Phase Ia 17.1‰ 38.4 34.0 29.6 25.3 38.1 33.9 29.7 25.4 Phase Ib 16.6‰ 40.6 36.2 31.8 27.5 40.2 36.0 31.8 27.5 Phase Ic 17.5‰ 36.7 32.3 27.9 23.5 36.4 32.2 28.0 23.8 Phase Id 16.2‰ 42.3 38.0 33.6 29.2 41.9 37.7 33.5 29.2 Phase II 17.9‰ 34.9 30.5 26.1 21.8 34.7 30.5 26.3 22.1

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Supplementary table S3. Data matrix of the locfit carbon and oxygen isotope curves (see Fig. 2) used for the bivariate plot in Supplementary Fig. S4. Data was obtained using the function <predict.locfit> in ‘R’, which is based on a cubic hermite polynomial interpolation. Ages are estimated on the basis of published U-Pb ages21-23 and under the assumption of equal durations of ammonoid zones between two U-Pb dating points (see Fig. 2). Gries. = Griesbachian.

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Age [Ma] δ18Ocp VSMOW [‰] σ δ13Corg VPDB [‰] σ 250.60 15.5 0.4 -28.1 0.5 250.65 15.8 0.3 -29.8 0.4 la

te

250.70 16.0 0.3 -31.2 0.5 250.75 16.2 0.3 -32.2 0.6 250.80 16.3 0.3 -32.8 0.7 250.85 16.3 0.3 -33.2 0.7 250.90 16.3 0.3 -33.2 0.8 250.95 16.1 0.3 -32.7 0.8 251.00 16.0 0.3 -31.4 0.8

mid

dle

251.05 16.1 0.2 -29.8 0.8 251.10 16.5 0.2 -28.3 0.7 251.15 17.1 0.2 -26.5 0.6 251.20 17.6 0.3 -25.3 0.5 251.25 17.8 0.3 -24.2 0.4 251.30 17.8 0.3 -25.0 0.4 251.35 17.7 0.3 -25.2 0.4 251.40 17.5 0.2 -25.1 0.4 251.45 17.2 0.2 -27.2 0.4 251.50 17.0 0.2 -28.0 0.5

Smith

ian

early

251.55 16.8 0.2 -27.6 0.6 251.60 16.8 0.3 -27.5 0.6 251.65 17.0 0.2 -27.7 0.5 251.70 17.0 0.2 -28.2 0.5 251.75 17.0 0.3 -28.6 0.5 251.80 16.8 0.3 -28.8 0.5 251.85 16.5 0.4 -28.8 0.5 251.90 16.2 0.5 -28.8 0.5 251.95 16.2 0.4 -28.7 0.5 252.00 16.4 0.2 -28.5 0.5 252.05 16.5 0.2 -28.2 0.5 252.10 16.5 0.3 -28.0 0.5 252.15 16.4 0.3 -27.7 0.5 252.20 16.3 0.3 -27.4 0.6 252.25 16.3 0.3 -27.2 0.6 252.30 16.7 0.3 -26.9 0.6

Die

neria

n

252.35 17.2 0.2 -26.6 0.7 252.40 17.4 0.2 -26.3 0.8 252.45 17.2 0.2 -25.9 0.9 la

te

Grie

s.

252.50 16.8 0.3 -25.6 1.1

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Supplementary table S4. Smithian ammonoid taxa (left column) of the Northern Indian Margin (NIM) and their distribution from the Di-enerian-Smithian boundary to the Smithian-Spathian boundary (data used for the turnover and generic richness curves in Fig. 2). Data from Brühwiler et al.28, Brühwiler et al.29, and Supplementary Fig. S1. Ammonoid zonation: F. bha., Flemingites bhargavai; S. rur., Sha-maraites rursiradiatus; X. per., Xenodiscoides perplicatus; F. nan., Flemingites nanus; R. evo., Radioceras evolvens; Rohil., Rohillites ro-hilla zone;C. sup., Clypeoceras superbum; E. cir., Euflemingites cirratus; B. com., Brayardites compressus; Escar., Escarguelites horizon; Na. pil., Nammalites pilatoides; P. mul., Pseudoceltites multiplicatus; Ny. ang., Nyalamites angustecostatus; W. dis., Wasatchites distrac-tus; Subvish., Subvishnuites beds; G. sin., Glyptophiceras sinuatum28,29; SM-1 to SM-14, Unitary Association zones for the Smithian of the NIM after Brühwiler et al.28. Ammonoid zones that are absent in the Salt Range (Pakistan) are shaded.

Legend: Smithian

early middle late

1 known occurrence

1 virtual occurrence

SM

-1

SM

-2

SM

-3

SM

-4

SM

-5

SM

-6

SM

-7

SM

-8

SM

-9

SM

-10

SM

-11

SM

-12

SM

-13

SM

.14

1

questionable occur-rence

Die

neria

n

F. bha. S. rur. X. per. F. nan. R. evo. Rohil. C. sup. E. cir. B. com Escar. Na. pil. P. mul. Ny. ang. W. dis. Subvish. G. sin. S

path

ian

Anasibirites 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0

Anaxenaspis 0 0 0 0 0 0 0 0 0 0 0 0 1 1 0 0 0 0

Arctoceras 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0

Aspenites 0 0 0 0 0 0 0 1 1 1 0 0 0 0 0 0 0 0

Brayardites 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0

Clypeoceras 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0

Dieneroceras 0 0 0 0 0 0 0 0 1 1 1 1 1 0 0 0 0 0

Eoptychitoides 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0

Escarguelites 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0

Euflemingites 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0

Flemingites 0 1 1 1 1 1 1 1 1 0 0 0 0 0 0 0 0 0 "Flemingites" griesbachi group 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0

Galfettites 0 0 0 0 0 0 0 0 0 1 1 1 0 0 0 0 0 0

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Gen. et sp. indet A 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0

Gen. et sp. indet B 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0

Glyptophiceras 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0

Hanielites 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0

Hedenstroemia 0 0 0 0 0 0 0 1 1 0 0 0 0 0 0 0 0 0

Hemiprionites 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0

Hermannites 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0

Hochuliites 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0

Jinyaceras 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0

Juvenites 0 0 0 0 0 0 0 0 1 1 1 1 0 0 0 0 0 0

Kashmirites 0 1 1 1 1 1 1 1 1 0 0 0 0 0 0 0 0 0

Kashmiritidae gen. et sp. nov. 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0

Kingites 1 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0

Koiloceras 1 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0

Kraffticeras 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0

Leyeceras 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0

Meekoceras 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0

Mianwaliites 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 1 1 0

Monneticeras 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0

Mudiceras 0 1 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0

Nammalites 0 0 0 0 0 0 0 0 0 0 1 1 0 0 0 0 0 0

Nuetzelia 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0

Nyalamites 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0

Owenites 0 0 0 0 0 0 0 0 0 0 0 1 1 1 0 0 0 0

Parakymatites 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0

Paranannites 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0

Paranorites 0 0 0 0 1 1 0 0 0 0 0 0 0 0 0 0 0 0

Paraspidites 0 0 0 1 1 0 0 0 0 0 0 0 0 0 0 0 0 0

Preflorianites 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0

Prionites 0 0 0 0 0 0 0 0 0 0 0 0 1 1 0 0 0 0 Proptychitidae gen. et sp. indet. 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0

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Pseudoaspenites 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0

Pseudoaspidites 0 0 0 0 0 0 1 1 1 1 1 0 0 0 0 0 0 0

Pseudoceltites 0 0 0 0 0 0 0 0 0 0 0 1 1 0 0 0 0 0

Pseudosageceras 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 Pseudoflemingites 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0

Punjabites 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0

Radioceras 1 1 1 1 1 1 0 0 0 0 0 0 0 0 0 0 0 0

Rohillites 0 0 0 1 1 1 1 1 0 0 0 0 0 0 0 0 0 0

Shamaraites 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0

Shigetaceras 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0

Steckites 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0

Stephanites 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0

Subflemingites 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0

Subinyoites 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0

Subvishnuites 0 0 0 0 0 0 0 0 0 0 0 0 0 1 1 1 0 0

Truempyceras 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0

Tulongites 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0

Urdyceras 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0

Ussuriidae gen. indet. 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0

Vercherites 0 0 0 0 1 1 1 1 1 0 0 0 0 0 0 0 0 0

Wasatchites 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0

Xenoceltites 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 1 0

Xenodiscoides 0 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0

Xiaoqiaoceras 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0

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Supplementary references 1 Golonka, J. & Ford, D. Pangean (Late Carboniferous-Middle Jurassic) paleoenvi-

ronment and lithofacies. Palaeogeography, Palaeoclimatology, Palaeoecology 161, 1–34 (2000).

2 Hermann, E. et al. Organic matter and palaeoenvironmental signals during the Early Triassic biotic recovery: The Salt Range and Surghar Range records. Sedimentary Geology 234, 19–41 (2011).

3 Baud, A., Atudorei, V. & Sharp, Z. Late Permian and Early Triassic evolution of the Northern Indian margin: Carbon isotope and sequence stratigraphy. Geodinamica. Acta 9, 57–77 (1996).

4 Brühwiler, T. et al. The Lower Triassic sedimentary and carbon isotope records from Tulong (South Tibet) and their significance for Tethyan palaeoceanography. Sedimentary Geology 222, 314-332 (2009).

5 Galfetti, T. et al. Late Early Triassic climate change: Insights from carbonate carbon isotopes, sedimentary evolution and ammonoid paleobiogeography. Palaeogeogra-phy, Palaeoclimatology, Palaeoecology 243, 394–411 (2007).

6 Kearsey, T., Twitchett, R. J., Price, G. D. & Grimes, S. T. Isotope excursions and palaeotemperature estimates from the Permian/Triassic boundary in the Southern Alps (Italy). Palaeogeography, Palaeoclimatology, Palaeoecology 279, 29–40 (2009).

7 Korte, C., Kozur, H. W. & Veizer, J. δ13C and δ18O values of Triassic brachiopods and carbonate rocks as proxies for coeval seawater and palaeotemperature. Palaeo-geography, Palaeoclimatology, Palaeoecology 226, 287-306 (2005).

8 Kolodny, Y., Luz, B. & Navon, O. Oxygen isotope variations in phosphate of bio-genic apatites, I. Fish bone apatite ― rechecking the rules of the game. Earth and Planetary Science Letters 64, 398–404 (1983).

9 Joachimski, M. M. & Buggisch, W. Conodont apatite δ18O signatures indicate cli-matic cooling as a trigger of the Late Devonian mass extinction. Geology 30, 711–714 (2002).

10 Joachimski, M. M. et al. Climate warming in the latest Permian and the Per-mian−Triassic mass extinction. Geology 40, 195–198 (2012).

11 Joachimski, M. M. et al. Devonian climate and reef evolution: Insights from oxygen isotopes in apatite. Earth and Planetary Science Letters 284, 599–609 (2009).

12 Sharp, Z. D., Atudorei, V. & Furrer, H. The effect of diagenesis on oxygen isotope ratios of biogenic phosphates. American Journal of Science 300, 222–237 (2000).

13 Luz, B., Kolodny, Y. & Kovach, J. Oxygen isotope variations in phosphate of bio-genic apatites, III. Conodonts. Earth and Planetary Science Letters 69, 255–262 (1984).

14 Korte, C. et al. Massive volcanism at the Permian-Triassic boundary and its impact on the isotopic composition of the ocean and atmosphere. Journal of Asian Earth Sciences 37, 293–311 (2010).

© 2013 Macmillan Publishers Limited. All rights reserved.

25

15 Epstein, A. G., Epstein, J. B. & Harris, L. D. Conodont Color Alteration – an index to organic metamorphism. Geological Survey Professional Paper 995, 1-27 (1977).

16 Jeppsson, L. & Anehus, R. A new technique to separate conodont elements from heavier minerals. Alcheringa: An Australasian Journal of Palaeontology 23, 57–62 (1999).

17 Dettman, D. L. et al. Seasonal stable isotope evidence for a strong Asian monsoon throughout the past 10.7 m.y. Geology 29, 31-34 (2001).

18 Vennemann, T. W., Fricke, H. C., Blake, R. E., O’Neil, J. R. & Colman, A. Oxygen isotope analyses of phosphates: A comparison of techniques for analysis of Ag3PO4. Chemical Geology 185, 321–336 (2002).

19 Coplen, T. B. Reporting of stable hydrogen, carbon, and oxygen isotopic abundan-ces. Geothermics 24, 707-712 (1994).

20 Tozer, E. T. Lower Triassic stages and ammonoid zones of Arctic Canada. Paper of the Geological Survey of Canada 65, 1–14 (1965).

21 Galfetti, T. et al. Timing of the Early Triassic carbon cycle perturbations inferred from new U-Pb ages and ammonoid biochronozones. Earth and Planetary Science Letters 258, 593–604 (2007).

22 Ovtcharova, M. et al. New Early to Middle Triassic U-Pb ages from South China: Calibration with ammonoid biochronozones and implications for the timing of the Triassic biotic recovery. Earth and Planetary Science Letters 243, 463–475 (2006).

23 Mundil, R., Ludwig, K. R., Metcalfe, I. & Renne, P. R. Age and timing of the Per-mian mass extinctions: U/Pb dating of closed-system zircons. Science 305, 1760–1763 (2004).

24 Waagen, W. Salt Ranges Fossils 2: Fossils from the Ceratites formation — Part I — Pisces, Ammonoidea. Palaeontologia Indica 13, 1–323 (1895).

25 Guex, J. Le Trias inférieur des Salt Ranges (Pakistan): problèmes biochronologi-ques. Eclogae Geologicae Helvetiae 71, 105-141 (1978).

26 Dagys, A. S. in Recent developments on Triassic Stratigraphy (eds Jean Guex & Aymon Baud) 15–23 (Imprivite S.A., 1994).

27 Shigeta, Y., Zakharov, Y. D., Maeda, H. & Popov, A. M. The Lower Triassic sys-tem in the Abrek Bay area, South Primorye, Russia. 38 (National Museum of Nature and Science Monographs, 2009).

28 Brühwiler, T., Bucher, H., Brayard, A. & Goudemand, N. High-resolution bio-chronology and diversity dynamics of the Early Triassic ammonoid recovery: The Smithian faunas of the Northern Indian margin. Palaeogeography, Palaeoclimatol-ogy, Palaeoecology 297, 491–501 (2010).

29 Brühwiler, T., Bucher, H., Roohi, G., Yaseen, A. & Rehman, K. A new early Smithian ammonoid fauna from the Salt Range (Pakistan). Swiss Journal of Palae-ontology 130, 187-201 (2011).

30 Brühwiler, T., Bucher, H. & Goudemand, N. Smithian (Early Triassic) ammonoids from Tulong, South Tibet. Geobios 43, 403–431 (2010).

31 Ghosh, P. et al. 13C-18O bonds in carbonate minerals: A new kind of paleothermom-eter. Geochimica et Cosmochimica Acta 70, 1439–1456 (2006).

32 Gregory, R. T. & Taylor, H. P., Jr. An oxygen isotope profile in a section of Creta-ceous oceanic crust, Samail ophiolite, Oman: Evidence for δ18O buffering of the

© 2013 Macmillan Publishers Limited. All rights reserved.

26

oceans by deep (>5 km) seawater-hydrothermal circulation at mid-ocean ridges. Journal of Geophysical Research 86, 2737-2755 (1981).

33 Lécuyer, C. & Allemand, P. Modelling of the oxygen isotope evolution of seawater: implications for the climate interpretation of the δ18O of marine sediments. Geo-chimica et Cosmochimica Acta 63, 351–361 (1999).

34 Veizer, J. et al. 87Sr/86Sr, δ13C and δ18O evolution of Phanerozoic seawater. Chemi-cal Geology 161, 59–88 (1999).

35 Jaffrés, J. B. D., Shields, G. A. & Wallmann, K. The oxygen isotope evolution of seawater: A critical review of a long-standing controversy and an improved geo-logical water cycle model for the past 3.4 billion years. Earth-Science Reviews 83, 83–122 (2007).

36 Trotter, J. A., Williams, I. S., Barnes, C. R., Lécuyer, C. & Nicoll, R. S. Did cooling oceans trigger Ordovician biodiversification? Evidence from conodont thermom-etry. Science 321, 550–554 (2008).

37 Pucéat, E. et al. Revised phosphate-water fractionation equation reassessing paleotemperatures derived from biogenic apatite. Earth and Planetary Science Let-ters 298, 135–142 (2010).

38 Baud, A., Magaritz, M. & Holser, W. Permian-Triassic of the Tethys: Carbon iso-tope studies. Geologische Rundschau 78, 649-677 (1989).

39 Corsetti, F. A., Baud, A., Marenco, P. J. & Richoz, S. Summary of Early Triassic carbon isotope records. Comptes Rendus Palevol 4, 473-486 (2005).

40 Galfetti, T. et al. Smithian-Spathian boundary event: Evidence for global climatic change in the wake of the end-Permian biotic crisis. Geology 35, 291–294 (2007).

41 Korte, C. et al. Carbon, sulfur, oxygen and strontium isotope records, organic geo-chemistry and biostratigraphy across the Permian/Triassic boundary in Abadeh, Iran. International Journal of Earth Sciences 93, 565-581 (2004).

42 Payne, J. L. et al. Large perturbations of the carbon cycle during recovery from the end-Permian extinction. Science 305, 506–509 (2004).

43 Retallack, G. J. et al. Multiple Early Triassic greenhouse crises impeded recovery from Late Permian mass extinction. Palaeogeography, Palaeoclimatology, Palaeo-ecology 308, 233–251 (2011).

44 Sakuma, H. et al. High-resolution lithostratigraphy and organic carbon isotope stra-tigraphy of the Lower Triassic pelagic sequence in central Japan. Island Arc 21, 79–100, doi:10.1111/j.1440-1738.2012.00809.x (2012).

45 Hermann, E. et al. A close-up view of the Permian–Triassic boundary based on ex-panded organic carbon isotope records from Norway (Trøndelag and Finnmark Plat-form). Global and Planetary Change 74, 156–167 (2010).

46 Loader, C. R. Local Regression and Likelihood. (Springer, 1999). 47 R: A language and environment for statistical computing v. 2.13.0 (Vienna, Austria,

2011). 48 Escaguel, G. & Bucher, H. Counting taxonomic richness from discrete biochrono-

zones of unknown duration: a simulation. Palaeogeography, Palaeoclimatology, Palaeoecology 202, 181–208 (2004).

49 Bellanca, A. et al. Palaeoceanographic significance of the Tethyan 'Livello Selli' (Early Aptian) from the Hybla Formation, northwestern Sicily: biostratigraphy and

© 2013 Macmillan Publishers Limited. All rights reserved.

27

high-resolution chemostratigraphic records. Palaeogeography, Palaeoclimatology, Palaeoecology 185, 175–196 (2002).

50 Hermann, E. et al. Climatic oscillations at the onset of the Mesozoic inferred from palynological records from the North Indian margin. Journal of the Geological So-ciety of London 169, 227–237, doi:10.1144/0016-76492010-130 (2012).

51 Hochuli, P. A. & Vigran, J. O. Climate variations in the Boreal Triassic ― Inferred from palynological records from the Barents Sea. Palaeogeography, Palaeocli-matology, Palaeoecology 290, 20–42 (2010).

52 Orchard, M. J. Conodont diversity and evolution through the latest Permian and Early Triassic upheavals. Palaeogeography, Palaeoclimatology, Palaeoecology 252, 93–117 (2007).

53 Brayard, A. et al. The Early Triassic ammonoid recovery: Paleoclimatic signifi-cance of diversity gradients. Palaeogeography, Palaeoclimatology, Palaeoecology 239, 374–395 (2006).

54 Kiehl, J. T. & Shields, C. A. Climate simulation of the latest Permian: Implications for mass extinction. Geology 33, 757–760 (2005).

55 Wignall, P. B. & Hallam, A. Griesbachian (earliest Triassic) palaeoenvironmental changes in the Salt Range, Pakistan and southeast China and their bearing on the Permo-Triassic mass extinction. Palaeogeography, Palaeoclimatology, Palaeoecol-ogy 102, 215–237 (1993).

56 Hallam, A. Why was there a delayed radiation after the end-Palaeozoic extinctions? Historical Biology: An International Journal of Paleobiology 5, 257–262 (1991).

57 Wignall, P. B. & Twitchett, R. J. Extent, duration, and nature of the Permian-Triassic superanoxic event. Geological Society of America Special Papers 356, 395–413 (2002).

58 Isozaki, Y. Permo-Triassic boundary superanoxia and stratified superocean: Re-cords from lost deep sea. Science 276, 235–238 (1997).

59 Galfetti, T. et al. Evolution of Early Triassic outer platform paleoenvironments in the Nanpanjiang Basin (South China) and their significance for the biotic recovery. Sedimentary Geology 204, 36-60 (2008).

60 Ware, D., Jenks, J., Hautmann, M. & Bucher, H. Dienerian (Early Triassic) am-monoids from the Candelaria Hills (Nevada, USA) and their significance for pa-laeobiogeography and palaeoceanography. Swiss Journal of Geosciences 104, 161–181 (2011).

61 Brayard, A., Escarguel, G., Bucher, H. & Brühwiler, T. Smithian and Spathian (Early Triassic) ammonoid assemblages from terranes: Paleoceanographic and pa-leogeographic implications. Journal of Asian Earth Sciences 36, 420–433 (2009).

62 Mayhew, P. J., Jenkins, G. B. & Benton, T. G. A long-term association between global temperature and biodiversity, origination and extinction in the fossil record. Proceedings of the Royal Society B: Biological Sciences 275, 47–53 (2008).

63 Hammer, Ø., Harper, D. A. T. & Ryan, P. D. Paleontological statistics software package for education and data analysis. Palaeontologia Electronica 4, 1–9 (2001).

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