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Journal of Asian Earth Sciences 36 (2009) 459–480

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Journal of Asian Earth Sciences

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Integrated ‘‘plume winter” scenario for the double-phased extinctionduring the Paleozoic–Mesozoic transition:The G-LB and P-TB events from a Panthalassan perspective

Yukio Isozaki *

Department of Earth Science and Astronomy, The University of Tokyo, 3-8-1 Komaba, Meguro, Tokyo 153-8902, Japan

a r t i c l e i n f o a b s t r a c t

Article history:Received 10 October 2007Received in revised form 9 May 2009Accepted 23 May 2009

Keywords:Permian–Triassic boundaryGuadalupian–Lopingian boundaryMass extinctionPangeaSuperplumeIllawarra Reversal

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* Tel.: +81 3 5454 6608; fax: +81 3 5465 8244.E-mail address: [email protected]

The event across the Paleozoic–Mesozoic transition involved the greatest mass extinction in historytogether with other unique geologic phenomena of global context, such as the onset of Pangean riftingand the development of superanoxia. The detailed stratigraphic analyses on the Permo-Triassic sedimen-tary rocks documented a two-stepped nature both of the extinction and relevant global environmentalchanges at the Guadalupian–Lopingian (Middle and Upper Permian) boundary (G-LB, ca. 260 Ma) andat the Permo-Triassic boundary (P-TB, ca. 252 Ma), suggesting two independent triggers for the globalcatastrophe. Despite the entire loss of the Permian–Triassic ocean floors by successive subduction, somefragments of mid-oceanic rocks were accreted to and preserved along active continental margins. Theseprovide particularly important dataset for deciphering the Permo-Triassic paleo-environments of theextensive superocean Panthalassa that occupied nearly two thirds of the Earth’s surface. The accreteddeep-sea pelagic cherts recorded the double-phased remarkable faunal reorganization in radiolarians(major marine plankton in the Paleozoic) both across the G-LB and the P-TB, and the prolonged deep-sea anoxia (superanoxia) from the Late Permian to early Middle Triassic with a peak around the P-TB.In contrast, the accreted mid-oceanic paleo-atoll carbonates deposited on seamounts recorded clear dou-ble-phased changes of fusuline (representative Late Paleozoic shallow marine benthos) diversity and ofnegative shift of stable carbon isotope ratio at the G-LB and the P-TB, in addition to the Paleozoic mini-mum in 87Sr/86Sr isotope ratio in the Capitanian (Late Guadalupian) and the paleomagnetic IllawarraReversal in the late Guadalupian. These bio-, chemo-, and magneto-stratigraphical signatures are concor-dant with those reported from the coeval shallow marine shelf sequences around Pangea. The mid-oce-anic, deep- and shallow-water Permian records indicate that significant changes have appeared twice inthe second half of the Permian in a global extent. It is emphasized here that everything geologically unu-sual started in the Late Guadalupian; i.e., (1) the first mass extinction, (2) onset of the superanoxia, (3)sea-level drop down to the Phanerozoic minimum, (4) onset of volatile fluctuation in carbon isotope ratio,5) 87Sr/86Sr ratio of the Paleozoic minimum, (6) extensive felsic alkaline volcanism, and (7) IllawarraReversal.

The felsic alkaline volcanism and the concurrent formation of several large igneous provinces (LIPs) inthe eastern Pangea suggest that the Permian biosphere was involved in severe volcanic hazards twice atthe G-LB and the P-TB. This episodic magmatism was likely related to the activity of a mantle superplumethat initially rifted Pangea. The supercontinent-dividing superplume branched into several secondaryplumes in the mantle transition zone (410–660 km deep) beneath Pangea. These secondary plumesinduced the decompressional melting of mantle peridotite and pre-existing Pangean crust to form severalLIPs that likely caused a ‘‘plume winter” with global cooling by dust/aerosol screens in the stratosphere,gas poisoning, acid rain damage to surface vegetation etc. After the main eruption of plume-derived floodbasalt, global warming (plume summer) took over cooling, delayed the recovery of biodiversity, andintensified the ocean stratification. It was repeated twice at the G-LB and P-TB.

A unique geomagnetic episode called the Illawarra Reversal around the Wordian–Capitanian boundary(ca. 265 Ma) recorded the appearance of a large instability in the geomagnetic dipole in the Earth’s outercore. This rapid change was triggered likely by the episodic fall-down of a cold megalith (subducted oce-anic slabs) from the upper mantle to the D00 layer above the 2900 km-deep core-mantle boundary, in tightassociation with the launching of a mantle superplume. The initial changes in the surface environment in

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the Capitanian, i.e., the Kamura cooling event and the first biodiversity decline, were probably led by theweakened geomagnetic intensity due to unstable dipole of geodynamo. Under the low geomagneticintensity, the flux of galactic cosmic radiation increased to cause extensive cloud coverage over the pla-net. The resultant high albedo likely drove the Kamura cooling event that also triggered the unusuallyhigh productivity in the superocean and also the expansion of O2 minimum zone to start the superanoxia.

The ‘‘plume winter” scenario is integrated here to explain the ‘‘triple-double” during the Paleozoic–Mesozoic transition interval, i.e., double-phased cause, process, and consequence of the greatest globalcatastrophe in the Phanerozoic, in terms of mantle superplume activity that involved the whole Earthfrom the core to the surface biosphere.

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1. Introduction

The greatest mass extinction in the Phanerozoic marks the ma-jor era boundary between the Paleozoic and Mesozoic across whichthe Earth’s biosphere drastically changed from the older Paleozoicregime to the Mesozoic-Modern one (e.g. Sepkoski, 1984; Erwin,1993, 2006; Fig. 1A). A huge proportion of Paleozoic marine andterrestrial invertebrates was terminated and the Mesozoic pioneertaxa filled the empty ecospace in the aftermath. The Paleozoic–Mesozoic boundary or Permian–Triassic boundary (P-TB) is out-standing among the Big 5 extinction-related boundaries in thePhanerozoic, not only in the greater magnitude of mass extinctionper se (over 80% diversity loss on generic level) but also in record-ing quite unique geologic phenomena that are rarely recognized inthe rest; i.e., chert gap, reef gap, coal gap, sea-level minimum, ze-nith of Pangea, prolonged anoxia, and volatile changes in variousisotope ratios. Nonetheless, the ultimate cause is still the centerof discussion (e.g. Erwin, 2006; Bottjer et al., 2008).

The double-phased nature of the end-Paleozoic mass extinctionhas been much emphasized during the last decade (e.g. Isozakiet al., 2004; Racki and Wignall, 2005; Isozaki, 2007b; Retallacket al., 2006); however, before 1994 not much attention was paidto the Middle-Late Permian boundary (or Guadalupian–Lopingianboundary; G-LB) event by the great shadow of the widely-recog-nized P-TB event. This was partly due to the lesser awarenessamong geologists outside China where the G-LB event is best doc-umented. Since the first claim made by Jin et al. (1994) and Stanleyand Yang (1994), the largest massacre in the Phanerozoic has beenunderstood as a combined effect of two independent mass extinc-tion events that took place back to back within a short-time period;i.e., first at the G-LB (ca. 260 Ma; Gradstein et al., 2004) and secondat the P-TB sensu stricto (ca. 251–252 Ma; Bowring et al., 1998;Mundil et al., 2004) (Fig. 1B). The interval of nearly 8 million yearsbetween the two independent extinction events was probably tooshort for various Late Permian biota to retrieve their original highdiversity and population size (Stanley and Yang, 1994). This dou-ble-phased extinction pattern suggests that unusually harsh condi-tions appeared at least twice in the Permian shallow seas aroundPangea.

The G-LB has profound implications that are quite distinct fromthose of the P-TB. First the G-LB marks the initial major declineafter the Late Paleozoic long-term biodiversity stability in a highplateau since the end-Devonian extinction, as clearly shown inFig. 1A. Second, the G-LB extinction apparently coincided in timingnot only with the onset of the long-term oxygen-depletion in ocean(superanoxia; Isozaki, 1997; Fig. 1B), but also with the lowest sealevel in the Phanerozoic (Hallam and Wignall, 1999; Miller et al.,2005; Haq and Schutter, 2008) and the initial breakup of Pangea(Isozaki, 2007b) (Fig. 2). Third, the isotopic systems of carbonand strontium had started to change drastically in the Late Guad-alupian (Veizer et al., 1999; McArthur and Howarth, 2004; Isozakiet al., 2007a,b; Kani et al., 2008), and kept fluctuating during theLate Permian and Early Triassic (e.g. Holser et al., 1989; Baudet al., 1989; Payne et al., 2004; Horacek et al., 2009) (Fig. 2). These

varied geologic phenomena suggest that the main story of thePaleozoic–Mesozoic transition was not simple, like the end-Creta-ceous (K-T boundary) scenario, such as with one large bolideimpact and the sudden great mass-death, but was a long-termcomplicated event composed of at least two sub-stories andspanned for over 10 million years. A major reorganization of theglobal oceanographic regime likely started in the late Middle Perm-ian, and the perturbation in biosphere culminated at the P-TB todrive the biggest biodiversity loss in the Phanerozoic (Fig. 1B).

In addition to the conventional studies on continental shelf sed-iments in S. China and peri-Tethyan regions, the double-steppedfaunal turnover of fusulines and radiolarians was recently docu-mented at the G-LB and the P-TB also in the mid-superocean(Ota and Isozaki, 2006; Isozaki, 2007a), and these results suggestthe global context of the two distinct events. This article at first re-views the current knowledge on the mid-oceanic records spanningacross both G-LB and P-TB derived mostly from Japan, with respectto the double-phased extinction and relevant environmentalchanges of global context. Then, on the basis of all data for the dou-ble-phased changes in mid-Panthalassa and peri-Pangean do-mains, their geological implications are discussed with regard tothe Permian mantle superplume activity. At the end, an integratedversion of the ‘‘plume winter” scenario is introduced for explainingthe cause, processes, and effects of the greatest biosphere catastro-phe of the Phanerozoic that took place in the Paleozoic–Mesozoictransition interval (ca. 265–240 Ma).

2. Mid-oceanic records

The oldest extant ocean floor has an age of about 200 Ma (EarlyJurassic), which is exposed in the deep-sea floor just off the MarianaTrench in the western Pacific. In other words, there is no pre-200 Ma ocean floor remaining in modern oceanic domains, simplybecause non-stop oceanic subduction processes have kept consum-ing older ocean floors along active continental margins as regu-lated by continuous plate tectonics. Hence, we cannot retrieveany deep-sea record prior to the Early Jurassic from moderndeep-sea floors, and one of the popular approaches in late Meso-zoic–Cenozoic paleoceanography like deep-sea drilling is not appli-cable to studies of pre-Jurassic events including the P-TB event. Onthe contrary to such unfriendly conditions for us geologists, how-ever, crucial oceanic subduction processes somehow did pay backbenefits in the form of an accretionary complex (AC).

Ancient ACs include numerous remnants of pre-existing oceanfloor material, such as deep-sea sediments, paleo-seamount/oce-anic plateau basalts with capping reef/atoll carbonates, and some-times ophiolites from mid-oceanic ridge (Isozaki et al., 1990; seeupper right diagram of Fig. 3) as typically observed in Japan. Dueto their subduction–accretion-related tectonic history, these rocksoccur as allochthonous (exotic) blocks/lenses within youngermatrices of trench-fill turbidites, and they are usually deformed,dismembered, and metamorphosed. Nonetheless, reconstructionof primary stratigraphy of mid-oceanic sediments is possibleby utilizing high-resolution microfossil dating (e.g. conodonts,

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Fig. 1. The Phanerozoic biodiversity change punctuated by major mass extinction events (A, redrawn from Sepkoski, 1984) and the Permo-Triassic details (B, redrawn fromKnoll et al., 1996; Isozaki, 1997). The greatest mass extinction in the Phanerozoic comprises two distinct mass extinction events; i.e., one at the Guadalupian–Lopingianboundary (G-LB) and the other at the Permian–Triassic boundary (P-TB). Particularly noteworthy is (1) the sharp decline of marine animals with low metabolic rates (mostlysessile benthos or the Paleozoic fauna by Sepkoski, 1984 as shown in (A) with respect to those with high metabolic rates (animals with gills and strong internal circulationsystems; i.e. Modern fauna by Sepkoski) at the G-LB, and (2) the greater magnitude of the selectivity at the G-LB than at the P-TB. Also note the low biodiversity interval fromthe G-LB to the mid-Anisian (Middle Triassic) for nearly 20 million years by and large overlaps the superanoxic period in the deep-sea (Isozaki, 1997).

Y. Isozaki / Journal of Asian Earth Sciences 36 (2009) 459–480 461

radiolarians, fusulines etc.) as first demonstrated in Japan (e.g.Kanmera and Nishi, 1983; Matsuda and Isozaki, 1991). Suchancient accreted mid-oceanic sediments provide an importantsource of information for reconstructing paleo-oceanography andglobal environments not only of the extinction-related G-LB andP-TB cases (e.g. Isozaki, 1994, 1997) but also of the pre-Jurassic(=pre-200 Ma) world, as recently demonstrated in the Upper

Neoproterozoic to Cambrian rocks by Uchio et al. (2004, 2008),Ota et al. (2007), and Kawai et al. (2008).

Under the unique paleogeographic configuration in the Permiancharacterized by a pair of single supercontinent Pangea and super-ocean Panthalassa, the ambient oceanographic systems at thattime may have been considerably different from those of today.In particular, Panthalassa occupied nearly 70% of the Earth’s

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Fig. 2. Long-term changes in C and Sr isotope ratios and in sea level during the Phanerozoic (compiled from Veizer et al., 1999; Saltzman, 2005; Katz et al., 2005; McArthurand Howarth, 2004; Miller et al., 2005; Haq and Schutter, 2008). Note the G-LB as a unique timing with unusual geological signatures; i.e., the onset of volatile fluctuation in Cisotope, and the Phanerozoic minimum in Sr/Sr ratio and in sea level. The overall patterns of these three aspects do not appear in harmony with each other throughout thePhanerozoic; however, note that the G-LB marks a unique timing, at which these three independent phenomena synchnonized to record the onset of unusual events of globalscale.


surface then (Fig. 3 upper left map). For reconstructing the globalenvironmental changes related to the end-Permian double-phasedmass extinction, therefore, information from mid-Panthalassa ob-tained through analysis of accreted oceanic sediments is inevitable.

The Mesozoic ACs in Japan and elsewhere in the circum-Pacific(Miyashiro-type) orogenic belts contain allochthonous blocks andlenses of Permo-Triassic mid-oceanic deep-sea cherts and paleo-atoll carbonates. The Late Permian AC in the Akiyoshi belt andthe Jurassic AC in the Mino-Tanba belt (plus its lateral equivalentsin the Chichibu belt and North Kitakami–Oshima belt) in Japancontain numerous allochthonous blocks of oceanic rocks (Isozakiet al., 1990). Among these, several large blocks of ancient mid-oce-anic sediments retain primary stratigraphic sequences includingboth the P-TB and the G-LB intervals. Critical datasets for Permo-Triassic studies are two-fold; one from mid-oceanic deep-seacherts and the other from mid-oceanic shallow-sea carbonates(Fig. 3), as described below.

In general, the primary deposition of these accreted oceanicrocks is difficult to locate but we can estimate their approximatepositions in the superocean on the basis of some plate tectonicconstraints. For example, in the case of those contained in theJurassic AC in Japan, a ca. 100 million year-long age gap exists be-tween the Late Permian sedimentation in the mid-ocean and theMiddle Jurassic accretion at the trench. Given a conservative aver-age plate convergence rate of about 3 cm per year, a ca. 3000 kmseparation from the eastern margin of the South China is assumedfor the primary sedimentary site, prior to the long-term horizontaltransport and subduction–accretion. In addition, paleomagneticdata established that the lower Middle Triassic deep-sea chertswere deposited in low-latitudes within 10 degrees of the paleo-equator (Shibuya and Sasajima, 1986; Oda and Suzuki, 2000; redstar in the upper left map of Fig. 3). Preliminary paleomagneticmeasurements for the Middle–Upper Permian paleo-atolllimestone also suggest its origin in a low-latitude (12 degree)domain in the southern hemisphere (Kirschvink and Isozaki,2007). During the Middle Permian to Early Triassic, therefore,

the relevant oceanfloor was likely located somewhere in the equa-torial Panthalassa.

In the following two sections, the essentials of bio- and che-mostrtaigraphical aspects of deep-sea chert and atoll limestoneare summarized separately.

3. Deep-sea chert

The stratigraphy of Late Permian to Early Triassic deep-sea chertin Japan, primarily derived from a mid-oceanic, deep-water domainof low-latitude Panthalassa has been analyzed in many sectionsduring the last two decades (e.g. Yamakita, 1987; Isozaki, 1994;Kuwahara et al., 1998; Ezaki and Yao, 2000). Fig. 4 depicts the com-posite column of the deep-sea chert facies in Japan. Additional data,in more or less the same context with local variability, were alsoobtained from British Columbia and New Zealand (Isozaki, 1997;Spörli et al., 2007). These chert beds mostly consist of radiolariantests with a minor amount of clay minerals, as reflected in theirhigh silica (ca. 95 wt%) and low aluminum (less than 3 wt%) con-tents, whereas coarse-grained terrigenous clastics are absent. Thevery low average sedimentation rate of less than 4 mm/ky was cal-culated based on the high-resolution conodont zoning and chrono-stratigraphic calibration (Matsuda and Isozaki, 1991), and each bedcomposed of a couple of vitric chert layer (>10 cm) and siliceousmudstone film (>5 mm) represents a 23 and 40 ky Milankovitch cy-cle (Hori et al., 1993). As to the P-TB interval of the uppermostChanghsingian to Olenekian, such well-bedded Milankovitch-tuned chert sequences are replaced by less clearly bedded siliceousclaystone (gray shale) and carbonaceous claystone (black shale)(e.g. Isozaki, 1994; Yamakita et al., 1999; Kakuwa, 1996; Takahashiet al., 2009). The ubiquitous occurrence of framboidal pyrite andhigh-TOC (total organic carbon) indicates that the deep-sea chertsequence records a unique paleo-redox episode of a prolonged re-duced condition in the deep superocean Panthalassa during theLate Permian to Early Triassic interval for nearly 20 million yearswith a climax across the P-TB (superanoxia; Isozaki, 1994, 1997).

Fig. 3. Composite stratigraphy and inter-facies correlation of the Middle Permian to Middle Triassic mid-oceanic deep-sea chert and shallow marine atoll carbonate in Japanthat originated from the mid-superocean Panthalassa (modified from Isozaki, 2007b). The red star in the upper left map indicates the primary depositional site in low-latitudemid-Panthalassa of the mid-oceanic rocks. Upper right: a simplified diagram showing a ridge-trench transect showing the primary plate tectonic setting of deep-sea chert andpaleo-atoll carbonates in mid-ocean (Panthalassa) before the subduction–accretion along the Pangean margins (modified from Isozaki et al., 1990). Columns are not to scale.Note the double-phased biotic turnover at the G-LB and the P-TB was recorded both in deep and shallow mid-Panthalassan sediments but in different manners.


3.1. Double-phased turnover of radiolarians

The Paleozoic radiolarians became almost completely extinct atthe P-TB. Two major orders of Paleozoic radiolarians, Albaillellariaand Latentifistularia, disappeared at the end of the Changhsingian(Late Lopingian), whereas the Mesozoic-type orders, Nassellariaand Spumellaria, started to radiate in the Smithian, Early Triassic

(De Weber et al., 2003). This major break in radiolarian lineagessuggests development of an environmental stress of global scaleat the P-TB, although Takemura et al. (2007) recently demon-strated a delayed radiolarian extinction in the high-latitude south-ern hemisphere. Because radiolarians, in particular the Paleozoicones before the diatom dominance in the Mesozoic–Cenozoicworld, represent the major marine plankton that regulated the

Fig. 4. Stratigraphic summary of mid-oceanic deep-sea sedimentation, diversity of Family Albaillellaria (late Paleozoic radiolarians), and secular change in stable carbonisotope ratio of organic carbon in deep-sea chert (modified from Isozaki, 2007b). Column is not to scale. Carbon isotope data are compiled from Ishiga et al. (1993) and Wadaet al. (unpublished data). Redox of seawater started to change in the Late Guadalupian (onset of superanoxia), reached an ultimate state (zenith of superanoxia) at the P-TB,and returned to normal (cancellation of superanoxia) in the mid-Anisian. Note that radiolarians recorded a double-phased turnover around the G-LB and the P-TB. The firstchange involved the major decline of the genus Pseudoalbaillella that survived for a long time since the mid-Carboniferous, suggesting the appearance of a major unfavorablecondition in mid-Panthalassa. Also noteworthy is the secular change in the stable carbon isotope ratio of organic matter (d13Corg) that shows a positive shift at the P-TB, inremarkable contrast to that of carbonates (d13Ccarb) with negative shift (Fig. 5).


main silica factory in the ocean at the time, their collapse may indi-cate a considerable disturbance of the oceanic food web and a de-cline in marine productivity at various depths of the water columnin the superocean.

Not much attention has previously been paid to; however, therewas another radiolarian turnover among the order Albaillellariaimmediately before the G-LB, in particular from the MiddlePermian genus Pseudoalbaillella to the Late Permian Neoalbaillellavia a transient form Follicucullus (Isozaki, 2007a; Fig. 4). Pseudoal-baillella is a long-lived genus that diversified and constantlypredominated in the Late Carboniferous to Middle Permian andthey declined in diversity, in average body size, and also in popu-lation size around the G-LB (e.g. Ishiga, 1990; Yao and Kuwahara,1999). The genus Follicucullus appeared in the Capitanian immedi-ately before the G-LB. According to the biostratigraphy in WestTexas and South China (Ormiston and Babcock, 1979; Nestellet al., 2006; Sun and Xia, 2006), the Follicucullus scholasticus-F. ven-tricosus (radiolarian) Zone, together with the underlying F. mon-acanthus Zone and the overlying F. charveti Zone, is correlatedwith the Capitanian. The precise G-LB horizon is still difficult to lo-cate by radiolarians; however, it is tentatively placed at the top ofF. charveti Zone.

In the Capitanian F. monacanthus Zone, Pseudoalbaillella becamerelatively smaller in body size and less abundant, whereas Follicucullusbecame dominant. In turn, Follicucullus became less dominant in thelatest Capitanian (upper part of the F. charveti Zone). This quick turn-over of Albaillellaria genera in the Capitanian is remarkable in contrastto their long-term stable lineage throughout the Late Carboniferous toMiddle Permian, suggesting that the radiolarians faced a fatal change

in environment immediately before the G-LB. The monotonous dom-inance of Albaillella plus Neoalbaillella throughout the Lopingian sug-gests that the oceanic condition was relatively stable up to the finalextinction immediately before the P-TB.

In this regard, it is noteworthy that all Paleozoic-type radiolari-ans apparently became extinct during the superanoxic intervalaround the P-TB (Isozaki, 1994; Kozur, 1998).

3.2. Onset and climax of superanoxia

Another interesting observation about deep-sea chert is thedevelopment of a unique oceanographic condition called superan-oxia across the P-TB; i.e., an extraordinarily long-term oxygen-de-pleted condition in the superocean Panthalassa (Isozaki, 1994,1997). The pyritiferous dark-colored rocks across the P-TB in Japanand British Columbia represent deposition under considerably re-duced up to euxinic conditions, in remarkable contrast with theLower–Middle Permian and the Middle–Upper Triassic hematite-enriched brick-red cherts that accumulated under oxygenated con-ditions. The development of deep-sea anoxia across the P-TB hasbeen suggested by various geochemical proxies, such as sulfur iso-tope ratios (Kajiwara et al., 1994), oxidation state of iron (Kuboet al., 1996; Matsuo et al., 2003), organic geochemistry (Suzukiet al., 1998), and the rare earth element (REE) abundance pattern(Ishiga et al., 1996; Kato et al., 2002) of the boundary sequence.In addition, the unusual occurrence of deep-sea dolomite nodulesin the dark-colored boundary unit in Japan (Kondo and Adachi,1975) also suggests that diagenetic precipitation of impure carbon-ates may have taken place under reduced condition even below


CCD as observed in modern deep-sea (e.g. Pisciotto and Mahoney,1981; Kelts and McKenzie, 1982). As no Permian–Triassic seawaterremains at present, all the geochemical proxies measured from thePermo-Triassic rocks are qualitative (or semi-quantitative at best),not sufficient for determining precise paleo-redox levels. Accord-ingly, the concept of superanoxia simply indicates that the deep-sea has deviated from a normal full oxic condition to a relativelyoxygen-poor one that lasted up to late Middle Permian, and raninto a relatively reduced condition for a long time until the earlyMiddle Triassic (Isozaki, 1994).

As the zonation/correlation of the Permian radiolarians hasbeen enhanced remarkably during the last decade (e.g., Kuwaharaet al., 1998; Yao and Kuwahara, 1999; Jin et al., 2007), the onsettiming of superanoxia is better constrained than before; the super-anoxia started in mid-Panthalassa likely in the late Guadalupian(Follicucullus scholasticus–F. ventricosus Zone). On the other hand,a full oxic condition returned to the deep-sea in the Anisian, earlyMiddle Triassic (Triassocampe coronata Zone). During this Early-Middle Triassic transition from the superanoxic back to a full oxicocean, the paleo-redox fluctuated between oxic and anoxic severaltimes (Kubo et al., 1996; Matsuo et al., 2003; Fig. 4). In analogy, theonset of superanoxia may also have experienced a similar modechange in the Capitanian, from a full oxic to an oxygen-depletedcondition with multiple fluctuations (Isozaki, 1996).

In the peri-Pangean shallow shelf environments, a much shorteranoxia/euxinia appeared across the P-TB (e.g. Wignall and Hallam,1992; Wignall and Twitchett, 2002; Grice et al., 2005; Riccardiet al., 2006). The occurrence of unusual Early Triassic shelf sedi-ments, such as aragonite fans on the sea-floor (Woods et al.,1999) and anachronistic microbialites (Schubert and Bottjer,1992; Kershaw et al., 1999; Lehrmann, 1999; Pruss et al., 2005)also supports the original superanoxia story with its initial devel-opment in the deep ocean, invasion of anoxic water into a shallowshelf by chemocline upward migration around the P-TB, and re-treat to the deep ocean by the end of Anisian (Isozaki, 1994).

Kajiwara et al. (1994) and Suzuki et al. (1998) discussed theshort-term oxygenation in the deep-sea right at the P-TB even dur-ing a long-lasting anoxia; however, this interpretation is unlikelyas criticized by Wignall and Twitchett (2002). Kakuwa (2008) alsodiscussed the Late Permian oxygenated deep-sea on the basis ofichnofabrics and previously published geochemical proxies; how-ever, his conclusion is based on equivocal redox proxies such asfaint Ce-anomaly in REE abundance and ichnofabrics. As to theanalysis of REEs, we cannot confidently evaluate subtle differencesin Ce-anomaly, in particular, without measuring the critical ele-ment Pr (praseodymium) (Kato and Isozaki, 2009). The ichnofabricindices established for shallow marine, continental shelf sequenceswith photosynthesis-dependent biotic communities (Bottjer andDroser, 1991) may not be applied to anoxic mid-oceanic deep-sea chert that was probably stigmatized by distinct biota. Theclaimed variation of ichnofabrics in deep-sea chert possibly reflectthe presence of a certain geochemical gradient in deep-sea sedi-ments near the sediment-water interface, not necessarily the oxy-gen level but possibly gradient in CO2, CH4, or even H2S.

Biomarker analysis of organic matter in the P-TB black shale inJapan indicates a significant contribution from marine phyto- andzooplankton with additional input from bacteria (Suzuki et al.,1998) but the balance between them and its secular change acrossthe P-TB has not been sufficiently documented. For the G-LB, nodetailed analysis has been made to date of a C-isotope and bio-marker, and these should be undertaken.

It is noteworthy that the deep-sea anoxia likely started in thelate Guadalupian appreciably before the G-LB (Figs. 3 and 4),whereas the oxygen-depletion reached shallow seas just acrossthe P-TB. This disparity in timing between deep and shallowoceans indicates that two distinct triggers for oceanographic

changes appeared in two steps in Permian Panthalassa. In particu-lar, a potential causal link is suggested between the redox changein the deep-sea and biodiversity change (Isozaki, 1997, 2007b). Theapparent coincidence in timing between the mid-oceanic radiolar-ian turnover and the onset of the superanoxia supports the link.Along with the environmental changes in surface and moderatelydeep oceans, the Permian radiolarians were probably forced tochange their physiology thus their test morphology specificallyduring this interval of redox change.

4. Paleo-atoll limestone

Numerous allochthonous blocks of Late Carboniferous to Trias-sic carbonates occur in the Late Permian to Jurassic ACs of Japan.These carbonates are mainly composed of shallow-marine bioclas-tic limestone (packstone/wackestone) with minor amounts offramestone, and they completely lack coarse-grained terrigenousclastics. They yield abundant shallow marine fossils of a typicalTethyan assemblage that suggests tropical to subtropical warm-water environments. Their basement consists of pillowed basalticgreenstones of oceanic-island (OIB)-type geochemical composition(e.g. Nishimura et al., 1979; Tatsumi et al., 2000). These lithologiccharacteristics and pre-accretion primary assemblage suggest thatthese limestones were deposited as a carbonate buildup or atollcomplex on top of ancient mid-oceanic seamount probably of hot-spot origin (Kanmera and Nishi, 1983; Sano and Kanmera, 1988;see inset of Fig. 3). Preliminary paleomagnetic data suggest thatthe paleo-atoll limestone was deposited at 12� in southern hemi-sphere (Kirschvink and Isozaki, 2007). Some of these limestoneblocks are of kilometric size and contain both P-TB and G-LB inter-vals (Koike, 1996; Isozaki and Ota, 2001).

Lithofacies, fusuline/conodont biostratigraphy, and stable car-bon isotope signatures were analyzed for Middle Permian to LowerTriassic rocks in such accreted mid-oceanic limestone (Koike,1996; Sano and Nakashima, 1997; Musashi et al., 2001, 2007;Ota and Isozaki, 2006; Isozaki et al., 2007a,b; Horacek et al.,2009; Fig. 5). In all these three aspects, changes occur also in adouble-phased manner both at the G-LB and P-TB. Lithology ofpaleo-atoll carbonates changed gradually across the G-LB fromthe Guadalupian dark-gray limestone with TOC �0.1 wt% to theLopingian light gray limestone with lower TOC �0.01 wt% (Otaand Isozaki, 2006). The top of the Lopingian limestone wasdolomitized, and is directly covered by the Lower Triassic blackmicrobial limestone (Sano and Nakashima, 1997). The P-TB-defin-ing index conodont Hindeodus parvus was reported from ca. 50 cmabove the base of this anachronistic microbialite (Koike, 1996).

4.1. Double-phased extinction of fusulines

The extinction of the Permian biota in a mid-oceanic shallowmarine setting is best documented by fusulines (e.g. Kanmeraand Nakazawa, 1974; Sano and Nakashima, 1997; Ota and Isozaki,2006), as the occurrence of conodonts/ammonoids is extremelyrare in the Permian parts of the accreted Permo-Triassic limestoneblocks due to facies control. Among the five Permian fusuline fam-ilies, Schwagerinidae and Verbeekinidae first became extinct at theG-LB, whereas the rest three, Ozawainellidae, Schubertidae, andStaffellidae, disappeared completely at the P-TB (e.g. Kanmeraet al., 1976; Ross, 1995). This double-phased extinction patternof major fusuline families was examined not only in the mid-oce-anic paleo-atoll limestone in Japan (e.g. Kanmera and Nakazawa,1974; Ozawa and Nishiwaki, 1992; Zaw Win, 1999; Ota and Iso-zaki, 2006; Fig. 5) but also in the coeval epeiric carbonate platformsaround Pangea (Jin et al., 1994; Stanley and Yang, 1994; Wilde,2002; Yang et al., 2004; Fig. 1B).

Fig. 5. Stratigraphic summary of mid-oceanic atoll carbonate sedimentation, diversity change in fusulines, and secular change in stable carbon isotope ratios of carbonate andorganic carbon. Column is not to scale. Carbon isotope data are compiled from Musashi et al. (2001), Isozaki et al. (2007a,b), Horacek et al. (2009). Note the double-phasedextinction of fusulines and double-phased negative shift in d13Ccarb across the G-LB and the P-TB. The long-lasting fusuline families with large tests were terminatedcompletely immediately before the G-LB, and the Lopingian was dominated solely by Lilliput fusulines until their total extinction at the P-TB. Carbon isotope balance inseawater was drastically modified twice, around the G-LB and the P-TB, but the first event recorded a relatively greater negative shift, suggesting that a more profound changein global C cycle likely took place around the G-LB rather than at the P-TB.


As to the G-LB extinction, a sharp faunal turnover was recog-nized between the Capitanian (upper Guadalupian) dark-graylimestone and the overlying Wuchiapingian (lower Lopingian)light gray dolomitic limestone in the Kamura section in Kyushuand the Akasaka section in central Japan (Sakagami, 1980; Ozawaand Nishiwaki, 1992; Zaw Win, 1999; Ota and Isozaki, 2006;Fig. 5). The Capitanian limestone is enriched solely in large-shelledverbeekinid fusulines, such as Yabeina and Lepidolina (up to 1 cm indiameter), whereas the Wuchiapingian rocks lack large individualsand contain only much smaller-shelled (less than 2 mm in diame-ter) schubertellids, ozawainellids, and staffellids, such as Codono-fusiella, Reichelina, and Staffella. The Wuchiapingian interval isdominated by Reichelina and Codonofusiella, whereas the Changh-singian part by Palaeofusulina, Staffella, and Nankinella (Kanmeraand Nakazawa, 1974; Sano and Nakashima, 1997; Ota and Isozaki,2006). All these Lopingian fusulines as a whole were terminatedconsequently at the P-TB as examined at two separated sectionsin Japan, i.e., Kamura and Taho (150 km apart) (Koike, 1996).

The Kamura, Taho, and Akasaka sections in Japan, where thedouble-phased extinction of Permian fusulines was detected, arecurrently separated from each other by ca. 300–500 km. The iden-tical extinction pattern confirmed at the physiographically sepa-rated multiple sections implies that the extinction and relevantenvironmental changes in two steps were not local but regionalphenomena in low-latitude Panthalassa, and probably of globalcontext. Such a double-phased extinction pattern of fusulinessuggests that significant environmental changes affecting theshallow-water benthic protists may have appeared in the surfaceof mid-Panthalassa at least twice.

4.2. Double-phased negative shift of C-isotope

Chemostratigraphical analyses utilizing stable carbon isotopeshave also detected a double-phased change around the G-LB andthe P-TB in the paleo-atoll limestones at Kamura in Japan (Fig. 5).

First across the G-LB, the secular change in carbonate carbon iso-tope ratio (d13Ccarb) demonstrated a gradual drop of ca. 5‰ aroundthe fusuline-defined G-LB (Isozaki et al., 2007a). The underlyingCapitanian interval is characterized by extremely high positived13Ccarb values of over +5‰ up to +7‰, whereas the overlyingWuchiapingian by relatively low positive ones of +2 to +3‰. Assuch high positive values between +5 and +7‰ are unusualthroughout the Permian and unique solely to this interval, thisphenomenon was named the Kamura event that may representa high-productivity event in mid-Panthalassa, and probably ashort-term cooling in the Capitanian immediately before the G-LB (Isozaki et al., 2007a,b). It is noteworthy that the total negativeshift occurred after the main extinction and spanned the biostrati-graphical G-LB defined by the first occurrence of the Lopingiantaxa. The negative shift around the G-LB corresponds to that atthe GSSP of G-LB at Penglaitan, South China (Wang et al., 2004).Unfortunately, d13Corg values at Kamura are still under measure-ment, and the results will be reported elsewhere shortly.

The other prominent negative shift of stable carbon isotope ra-tio was recognized at the P-TB in two separate sections in Kamuraand Taho (Musashi et al, 2001, 2007). A clear negative spike oc-curred both in carbonate carbon (d13Ccarb) and organic carbon(d13Corg) (Fig. 5). Both d13Ccarb and d13Corg values decrease for ca.2‰, showing a clear parallelism between them that likely suggestsuniversal 13C-depletion in ambient seawater. This negative shifthas been detected in many sections in the world and these sectionswere mutually correlated (e.g., Holser et al., 1989; Baud et al.,1989).

Although the Lopingian interval has not yet been thoroughlydocumented, the C isotope chemostratigraphic study on theaccreted Permian limestone clarified an overall decreasing trendfrom the Late Middle Permian to the earliest Triassic punctuatedby double-phased sharp negative shifts at the G-LB and the P-TB.This trend is in general agreement with that compiled by Korteet al. (2005). As the G-LB and the P-TB are separated from each


other by ca. 8 million years, the double-phased isotopic shift re-quires two independent driving mechanisms to cause a negativeshift in C isotopic composition of seawater twice on a regional scalein the superocean.

4.3. Sr minimum

The secular change in strontium isotopes in the Guadalupian toearly Lopingian interval was recently demonstrated in mid-Panth-alassan paleo-atoll carbonates at Kamura (Kani et al., 2008).87Sr/86Sr ratios of the Wordian, Capitanian, and Wuchiapingian car-bonates fluctuate in a relatively narrow range of 0.7069–0.7074.With respect to the 87Sr/86Sr values of modern seawater and ofmany Phanerozoic carbonates in equilibrium with paleo-seawater(e.g. Veizer et al., 1999; Fig. 2), the values of the Kamura samplesare concentrated obviously in a much lower value range. Theyshow a remarkable change in a long-term trend from a gradualdecrease in the Wordian–Capitanian to an increase in the Capita-nian–Wuchiapingian. The minimum value 0.706914 ± 0.000012was recorded in the Yabeina Zone (lower–middle Capitanian) atKamura (Fig. 5). This Capitanian minimum in Sr isotope ratiodetected in mid-Panthalassa atoll carbonates readily correspondsto that detected in coeval carbonates of the peri-Pangean shelfdomains (e.g., Veizer et al., 1999; Korte et al., 2006), suggesting aglobal signature. Marking a chemostratigraphically unique turningpoint in Phanerozoic trend, this minimum indicates that a remark-able change occurred in the Capitanian global oceanography withrespect to the Sr-isotope balance of seawater between mantle-and continental fluxes. In order to explain this unique event,Isozaki (2007b) and Kani et al. (2008) suggested that the initial rif-ting of Pangea might have been responsible in providing a hugeamount of continent-derived (radiogenic Sr-enriched) terrigenousclastics into the Panthalassan Ocean on the basis of the occurrencesof LIPs in the eastern Pangea. Further 87Sr/86Sr analysis is underway with respect to the P-TB interval.

5. Discussion

Possible causes and processes of the end-Permian double-phased event are explored here on the basis of the above-intro-duced new lines of evidence from the two contrasting mid-oceanicsedimentary facies of Panthalassa, i.e., deep-sea chert and shallowmarine atoll carbonates in addition to the data from the peri-Pangean domains. The discussion starts from the geologicalimplications of the double-phased extinction mainly from thePanthalassan perspective, moves onto the interpretation of isotoperecords of carbon and strontium, and finally deals with possiblecausal mechanisms with particular emphasis on mantle plumeactivity and related environmental changes on the surface.

5.1. Double-phased biotic turnover/extinction

As to the double-phased extinction, two distinct animal groupshold the key roles in the food web of the extensive mid-superoceandomains, i.e., radiolarians and fusulines. In the Late Paleozoic, radi-olarians were the most dominant plankton in open water, whereasfusulines represent one of the major benthos in shallow marineenvironments. Their double-phased extinction pattern confirmedin the mid-Panthalassa, concordant with that in the peri-Pangeancontinental margins, suggests that major environmental stress(es)of global scale occurred twice in the Middle-Late Permian bio-sphere. Both radiolarians and fusulines likely formed relatively ba-sal parts of the Permian marine food web, therefore, the rise andfall in their diversity and in abundance probably recorded glob-

ally-averaged environmental states including nutrient availabilityin moderately deep to very shallow levels of the superocean.

5.1.1. Radiolarian turnover in the deep superoceanRadiolarians are marine plankton whose habitable zone ranges

from surface water down to more than a thousand meter depth inmodern oceans. The Late Paleozoic radiolarians probably occupiedby and large the same habitat in the ocean as modern examples. Inthe extensive Panthalassan domain, Permian radiolarians suffereddouble-phased faunal turnover; i.e., significant faunal changearound the G-LB and final extinction at the P-TB. As to the G-LBturnover, Pseudoalabaillella, the long-ranged and long-dominantgenus since the Late Carboniferous declined in abundance, indiversity, and also in average body size for the first time in thePermian (Ishiga, 1990; Kuwahara et al., 1998). The lineage wasconsequently taken over by the Lopingian genus Neoalbaillella,via a transient genus Follicucullus that first appeared in the Guad-alupian. This likely suggests that a certain kind of unfavorable sit-uation for the long-lasting radiolarian taxa, and that the relevantenvironmental stress may have emerged in mid-Panthalassa andforced the radiolarian group to change physiology in terms of testmorphology by the end of the Guadalupian. Possible triggers in-clude the oxygen-depletion in seawater, to which the radiolariantest structure may have adjusted in regard of metabolism; e.g., effi-ciency in gas exchange, in nutrient uptake, and/or in waste-dumpthrough more perforated test. It is noteworthy that the radiolarianturnover started not right at G-LB time but slightly earlier in theearly Capitanian (Fig. 4). Nonetheless, the long-term depositionof radiolarian chert per se was not much affected, if at all, becausethe Milankovitch-tuned, rhythmically-bedded chert accumulatedcontinuously all the way from the Guadalupian almost to the lateLopingian. Consequently the total radiolarian production did notchange significantly across the G-LB, despite the lineage shift.

In contrast, the P-TB event was critical for the Paleozoic radiola-rians. Except for rare survivors in the southern hemisphere(Takemura et al., 2007), all the Paleozoic forms became extinct inthe topmost Changhsingian (e.g., Yao and Kuwahara, 1999; DeWeber et al., 2003). The P-TB interval of the deep-sea facies inJapan, from the topmost Changhsingian to the lower Olenekian(Smithian, Triassic), is composed of radiolarian-free (or -poor) sili-ceous claystone/black shale without any chert (chert gap; Fig. 4).The deposition of the unique P-TB black shale in the deep-seamay have resulted from unusual bacterial blooming after theradiolarian extinction. The earliest recovery of radiolarians afterthe P-TB event was in the Dienerian in western Tethys (Kozur,1996). In other words, the radiolaria-driven Permian silica factoryin the mid-ocean collapsed almost completely in the latestChanghsingian, and became dormant during the Griesbachian (ca.1 million years immediately after the P-TB; Fig. 4). The abruptdisappearance of radiolarians at the P-TB and their retardedrecovery were confirmed in the contemporary shallow marinecontinental shelf facies in South China (e.g., He et al., 2005; Jinet al., 2007; Isozaki et al., 2007c). The direct kill mechanism hasnot yet been identified; however, their total extinction possiblyreflects a major collapse in marine productivity on a global scalebecause the Paleozoic radiolarians had a wide geographical distri-bution from low-latitudes even to higher latitudes, probably moreextensive than modern ones due to the absence of diatoms in thePaleozoic. The abrupt extinction of radiolarians also suggests thata large perturbation appeared not only in surface waters but alsoin the moderately deep-water of mid-Panthalassa. Although theslightly delayed extinction in the high-latitude southern hemi-sphere was reported (Takemura et al., 2007), this does not changethe general tend of radiolarian transition from the Paleozoic tothe Mesozoic-type.


After all, radiolarians, the representative Paleozoic marineplankton of generally cosmopolitan nature, suffered major faunalturnover twice in the second half of the Permian. The impact ofthe P-TB event was obviously more profound than that of the G-LB. This indicates that the long-lasting normal food web of the LatePaleozoic abruptly collapsed in the mid-superocean at the P-TB.Nonetheless, from a long-term viewpoint, it is noteworthy thatthe major change in the Permian plankton started not at the P-TB but around the G-LB (Fig. 4), and in particular, that the initialchange started not at the biostratigraphically-defined G-LB butsometime earlier in the Capitanian. For making comparison withthe G-LB extinction in the deep superocean, coeval extinction inshallow mid-Panthalassa is discussed next.

5.1.2. Extinction in the shallow superoceanFusulines represent a shallow-marine, warm-water, benthic

foraminifera group that probably thrived within the euphotic zone(less than 50 m deep) just like modern benthic foraminifera. Theysuccessfully flourished mainly in low-latitude areas throughoutthe Late Carboniferous–Permian not only along the peri-Pangeancontinental margins but also on isolated mid-Panthalassan paleo-atolls, and they declined in diversity in two steps by the end ofthe Permian (e.g., Jin et al., 1994; Stanley and Yang, 1994; Otaand Isozaki, 2006). The first fusuline extinction at the G-LB resultedin a sharp reduction in test size, suggesting the appearance of anenvironmental stress to screen out selectively large-tested taxa.This phenomenon may correspond to the so-called ‘‘post-extinc-tion Lilliput effect” in a broad sense (e.g., Urbaneck, 1993; Twitch-ett, 1999). According to the classic ecological observations onmodern biota, a dominance of small-sized individuals called r-strategists generally indicates the appearance of a stressful condi-tion in environments. The ‘‘Lilliput effect in fusulines” appeared atthe G-LB, and continued for another 8 million years throughout theLopingian until the total extinction of all fusulines was completedat the P-TB (Fig. 5). This suggests that the Late Permian shallowmarine environments and relevant ecological structure wereessentially different from the long-lasting stable ones during theEarly-Middle Permian.

The Early-Middle Permian fusuline faunas are generally domi-nated by large-, robust-shelled taxa, i.e., Schwagerinidae and/orVerbeekinidae (Fig. 5). In particular, the late Guadalupian fusulinefauna in low-latitude realms of Tethys, Panthalassa, and westernNorth America, consist of extraordinarily large-tested genera(e.g., Yabeina, Lepidolina, and Polydiexiodina) up to 1 cm in diame-ter, with minor amounts of small ones (e.g., Kanmera et al.,1976; Ishii, 1990; Sheng, 1990; Ross, 1995). Permian large-testedfusulines evolved to form highly complicated wall structuresincluding keriotheca, whereas the smaller ones were conservativeand retained simple and thin (primitive) wall structures. In analogywith modern foraminifers, the complicated skeletal structures withmany volutions and sophisticated wall fabrics in large fusulineslikely represent an adaptation for hosting symbiotic algae and/orcyanobacteria (Ross, 1972; Wilde, 2002; Vachard et al., 2004). Inaddition to fusulines, some Permian gigantic bivalves, corals, andbrachiopods in shallow tropical seas are also listed as fossil photo-symbiotic animals (Cowen, 1983; Seilacher, 1990; Isozaki, 2006),and their diversity change along time is also interesting as dis-cussed below.

The Permian giant fusulines became extinct abruptly at the G-LB, both in mid-oceanic paleo-atoll limestones (Ota and Isozaki,2006) and in continental shelf carbonates (e.g., Sheng, 1990; Stanleyand Yang, 1994; Ross, 1995; Wilde, 2002; Yang et al., 2004). In con-trast, the lower Wuchiapingian rocks yield only simpler-structuredand smaller fusulines less than 2 mm in diameter, e.g. Staffiella andNankinella. Wilde (2002), Yang et al. (2004), and Ota and Isozaki(2006) proposed that the sharp screening of the large fusulines at

the end of the Guadalupian might have been related to the dy-ing-out of associated symbiotic algae. It is noteworthy that the riseand fall of large fusuline diversity occurred almost at the sametime with those of the Permian gigantic bivalves and various rug-ose corals (Wang and Sugiyama, 2002; Isozaki and Aljinovic, inpress). Paying special attention to the gigantism in shallow marineinvertebrates in tropical oligotrophic environments, Isozaki andAljinovic (in press) further emphasized the synchronous extinctionof the photosymbiosis-dependent ‘‘tropical trio” (large fusulines,large bivalves, and rugose corals) in shallow Tethys and Pantha-lassa. A large-scale environmental change may have occurred inshallow marine environments at the end of the Guadalupian pref-erentially in the photic zone of low-latitude Tethys and Pantha-lassa. Possible direct kill mechanisms may include temperaturedrop, decreased insolation, change in nutrient level, and/or salinitychange; however, details are not yet clarified. Isozaki and Aljinovic(in press) proposed that the Capitanian cooling (Kamura event; Iso-zaki et al., 2007a,b) coupled with eutrophication might have termi-nated the photosymbiotic fauna that were highly adapted towarm-water, oligotrophic conditions.

On the basis of the fossil data from the peri-Pangean shelf facies,Knoll et al. (1996) pointed out the selective screening by the G-LBextinction of animals with no gills, weak internal circulation andlow metabolic rate, i.e., mostly sessile benthos with passive respi-ratory systems, such as cnidarians, echinoderms, and articulatebrachiopods. Powers and Bottjer (2007, 2009) added supportingevidence from the on-/off-shore migration pattern of the Permianbryozoans. The recent analysis on the extinction pattern of varioustaxa across the G-LB (Clapham et al., 2009) suggests a rather grad-ual turnover during the second half of the Guadalupian rather thana rapid extinction specifically at the G-LB.

The common occurrence of the barren interval in the topmostCapitanian paleo-atoll sections (Ota and Isozaki, 2006) is notewor-thy because this confirms the common sequence of stages inecological transition relevant to extinction; (1) appearance of envi-ronmental stress, (2) main extinction, (3) prolonged survival, and(4) recovery. The kill mechanism obviously started to operate al-ready in the late Capitanian at the latest, and this agrees also withthe radiolarian extinction pattern in the mid-ocean (Fig. 4), asdiscussed above. Like other major extinction cases in the Phanero-zoic, the extinction of the Guadalupian shallow marine faunastarted before the biostratigraphic G-LB that is defined by thefirst occurrence of the Lopingian fauna (Fig. 5). The environmentalstress since the late Capitanian (e.g., temperature drop) likelysuppressed the biodiversity and population size for certainduration until the environment was ameliorated in the EarlyWuchiapingian.

At the P-TB, on the other hand, various shallow marine biota,not only fusulines and rugose corals but also others (smaller fora-minifers, brachiopods, gastropods, crinoids, and calcareous algae)suffered severely both in peri-Pangean shelf environments and inmid-Panthalassan paleo-atolls. As in the case of the G-LB extinc-tion, the biotic turnover and responsible environmental changewere of global extent; however, the kill mechanism for the P-TBcase may have been slightly different from that for the G-LB. As aclassic explanation, the appearance of short-term cooling was pro-posed as a possible cause for the mass killing at the end of thePermian (Stanley, 1988; Kozur, 1998). In addition, other possiblekill mechanisms of shallow marine fauna were proposed on thesame track that essentially assumes an upward migration ofchemocline in the superocean; such as suffocation by oxygendepletion (anoxia), hypercapnia by elevated CO2, and poisoningby H2S (e.g., Wignall and Hallam, 1992; Knoll et al., 1996; Riccardiet al., 2006).

In the beginning of the Early Triassic that is claimed as aninterval of retarded recovery, the characteristic anachronistic


microbialites and unique Lilliput fauna occur both in peri-Pangeancontinental shelves and in mid-Panthalassan paleo-atolls (Sanoand Nakashima, 1997; Kershaw et al., 1999; Lehrmann, 1999;Woods et al., 1999; Twitchett, 1999; Ezaki et al., 2003; Fraiserand Bottjer, 2004; Pruss et al., 2005), indicating that the post-P-TB aftermath may have experienced a short-term return to the Pre-cambrian world characterized by shallow marine communities ofreduced paleoecological tiering under prolonged environmentalstresses. The delayed recovery after the P-TB event took muchlonger than that after the G-LB event.

5.2. Double-phased carbon isotope shift

In addition to the fossil records, the drastic environmentalchange in the Permian–Triassic mid-superocean is clearly recordedin some geochemical proxies, such as chemical state of iron, REEabundance pattern, and isotope ratios of C, S, and Sr. The secularchange in stable carbon isotope ratio monitored an unusuallyvolatile fluctuation of the global carbon cycle during the Paleo-zoic–Mesozoic transition interval (Figs. 2 and 4). In particular,the double-phased nature of the secular change is noteworthy.These data confirmed not only that the Middle Permian to LowerTriassic mid-Panthalassan limestones are bio- and chemostrati-graphically correlated with the peri-Pangean/Tethyan shelfcarbonates, but also that the Late Permian ocean seawater hasrun into an isotopically extreme condition twice across the G-LBand the P-TB in global context. These two episodic events in theglobal C cycle are discussed here one by one.

5.2.1. G-LB episodeA significant negative shift in d13Ccarb for ca. 5‰ was recently

detected across the G-LB horizon in a mid-oceanic paleo-atoll lime-stone (Isozaki et al., 2007a; Fig. 5). This shift started appreciablyafter the extinction of the Guadalupian large-tested fusulines andbefore the first appearance of the Lopingian fauna. As being corre-lated with d13Ccarb stratigraphy of peri-Pangean shelf facies at theGSSP of the G-LB at Laibin, South China (Wang et al., 2004), thisnegative shift across the G-LB likely has a global context. Thereare three interesting aspects to be noted in the C isotope stratigra-phy across the G-LB; i.e., (1) the negative excursion of d13Ccarb inmid-Panthalassa occurred in three steps, interspersed with twosteps of short-term positive shifts, (2) the total amount of negativeshift across the G-LB reached ca. 5‰, much greater than that of theP-TB (ca. 2‰), and 3) the shift occurred after the main extinctionevent. A negative shift of d13C of organic carbon across the G-LBwas detected also at the Penglaitan GSSP (Kaiho et al., 2005), andthe essentially parallel negative shift between d13Ccarb and d13Corg

values indicates that the biological pump using the same C-fixationprocesses (most likely that of oxygenic photosynthesis) was stilloperating in shallow mid-Panthalassa even after the G-LB extinc-tion. As detailed d13Corg data have not yet been available fromdeep-sea chert except for preliminary ones (Ishiga et al, 1993),the present author and his colleagues are working on this problem,and the results will be reported elsewhere.

More emphasis should be given to the Capitanian interval withunusually high positive d13Ccarb values around +5‰ to +6‰ (Fig. 5)rather than to the negative shift in d13Ccarb across the G-LB,. Thehigh d13Ccarb interval apparently lasted for ca. 3 million years, thathas never been reported from the Permian, and this unique isoto-pic episode named Kamura event likely records a > 3 m.y.-longhigh productivity in low-latitude mid-Panthalassa (Isozaki et al.,2007a,b). In general, tropical shallow sea is characterized by an oli-gotrophic condition, therefore, sufficient amounts of nutrientsshould have been fed constantly to keep the high productivity pos-sibly through an extra terrigenous flux and/or a very active oceanicmixing. Development of a cooling period in the Capitanian may

explain this. The Middle Permian has been generally regarded asa post-glaciation warming period after the Carboniferous to EarlyPermian Gondwana icehouse; however, the appearance of a coldsnap in the late Middle-Late Permian is suggested by some geolog-ical observations in high-latitudes (e.g., Ustritsky, 1973;Beauchamp and Baud, 2002; Fielding et al., 2008). The extinctionof the ‘‘tropical trio” in low-latitudes (Isozaki, 2006; Aljinovicet al., 2008; Isozaki and Aljinovic, in press) and the first migrationof the Permian mid-latitude brachiopods into the tropical domain(Shen and Shi, 2002) during the latest Guadalupian likewise sup-port this interpretation, though the global influence of the claimedKamura cooling event should be further tested elsewhere.

In contrast, the above-mentioned G-LB negative shift per se inthe aftermath of the Kamura event, likely represents the onset ofglobal warming. During the interval of a negative excursion ind13Ccarb but still with positive values, the Lopingian fauna ap-peared, and the biodiversity started a post-extinction recovery(Ota and Isozaki, 2006).

5.2.2. P-TB perturbationIn the mid-oceanic paleo-atoll carbonates spanning the P-TB,

the secular change both in d13Ccarb and d13Corg values recorded aclear negative shift (Musashi et al., 2001, 2007; Fig. 5). This nega-tive excursion in d13Ccarb is well correlated with that in many sec-tions of marine and terrestrial facies around the world (e.g., Holseret al., 1989; Baud et al., 1989; Wang et al., 1994; Twitchett et al.,2001) including the GSSP at Meishan in South China (Jin et al.,2000b; Yin et al., 2001).

The overall parallelism between d13Ccarb and d13Corg in theUpper Permian–Lower Triassic shallow marine carbonates inmid-Panthalassa and also in Pangean margins (Magaritz et al.,1992; Musashi et al., 2001, 2007; Fig. 5) proves (1) that dissolvedinorganic carbon in shallow-sea water globally became depletedin 13C across the P-TB, and on the other hand, (2) that photosynthe-sis-driven isotopic fractionation was still working generally on thesurface even after the extinction-related P-TB event. In theGriesbachian immediately after the extinction of various Permianalgae and metazoans associated with the collapse of normalecological structures, bacteria (including photosynthetic cyanobac-teria) may have been dominant in C-fixation to keep the marineproductivity, as evidenced by the anachronistic (Precambrian-like)microbialites in both Pangean shelves and mid-Panthalassa atolls(e.g., Schubert and Bottjer, 1992; Sano and Nakashima, 1997; Ker-shaw et al., 1999; Lehrmann, 1999).

In contrast, however, the P-TB deep-sea black shale recorded apositive spike in d13Corg, showing the opposite trend to the coevalsignature in shallow marine carbonates (Fig. 5). The Lopingian andDienerian–Anisian deep-sea claystone/cherts have d13Corg valuesaround �30‰ (Ishiga et al., 1993; Wada et al., unpublished data),whereas the contemporary shallow marine carbonates on paleo-atolls record d13Corg values around -25‰ (Musashi et al., 2001).On the other hand, the Griesbachian black shale uniquely recordedd13Corg values around �27‰, and the contemporary shallow mar-ine carbonates also have the same values around �27‰. This illus-trates that the Late Permian to Middle Triassic oceans generallyhad a depth-wise isotopic gradient in Corg of ca. 5‰ between theshallow and deep water (Fig. 6 upper); however, this gradient dis-appeared during the Griesbachian (Fig. 6 lower).

The development of such a remarkable isotopic gradient withdepth in the Permo-Triassic Panthalassa except the Griesbachianinterval was likely driven by selective decomposition of organicmatter in the oxygenated upper water column of the ocean. Amongthe components of organic matter, proteins (amino acids) are moreeasily decomposed by bacterial activities under oxygenated condi-tions than lipids, in particular, during their sinking as particlesthrough subphotic zone in the upper water column (e.g., Wakeham

Fig. 6. Schematic diagram showing the two distinct modes (stratified mode and superanoxic mode) of the superocean Panthallassa during the Paleozoic–Mesozoic transitioninterval with respect to carbon isotope ratios in organic carbon. Upper: the stratified ocean during the Lopingian and Dienerian to mid-Anisian, lower: the P-TB superanoxicocean during the Griesbachian. The inset in the middle schematically shows unique behavior of d13Corg in surface (green line) and deep-sea (red line) across the P-TB. Up tothe end of the Changhsingian, ca. 5‰ gradient existed in d13Corg values between the surface and deep-sea waters, probably due to the selective decomposition of protein (withheavier C isotope ratio) in organic matter. Across the P-TB, d13Corg of the surface water became lighter by 2‰ (green line) in accordance with d13Ccarb. If the samedecomposition-controlled gradient was maintained as before, d13Corg of the deep-sea should have had 2‰ lighter values (broken green line); however, it became heavier by3‰ (red line). This indicates that the clear d13Corg gradient of 5‰ persisted during the Changhsigian eventually disappeared across the P-TB, and such an unusual monotonouscondition continued at least during the Griesbachian. The absence of this gradient in the Griesbachian indicates that the preferential decomposition of protein was suppressedin the water column, likely by the upward excursion of the chemocline almost to the surface (lower). Consequently the unique oceanic condition (superanoxia) enhanced thedirect downward transport and burial of organic carbon in deep-sea.


et al., 2000). As lipids usually have relatively light C-isotope ratioswith respect to proteins, the preferential decay of proteins resultsin a relative abundance of lipids in organic matter deposited indeep-sea sediments that likely drives lower d13Corg values. Underthe ordinary oceanic conditions of the (partly or fully) ventilatedLate Paleozoic–Mesozoic deep-sea, therefore, a clear gradientdeveloped in d13Corg values between shallow and deep waters. Aslong as protein-screening in the oxic upper water column workedto maintain the same isotopic gradient, any change in the carbonsource in the surface ocean may likewise have reflected/recordedin the C-isotope of deep-sea organic matter; i.e. the negative shiftof d13Corg for 2‰ across the P-TB in shallow marine carbonates (so-lid green line in the inset of Fig. 6) should also have been recordedin deep-sea black shale (broken green line).

In remarkable contrast, however, the Griesbachian recordsshow that the shallow atoll carbonates and deep-sea black shalehave more or less the same d13Corg values (solid red and green linesin the inset of Fig. 6), in other words, the depth-wise gradient thatpersisted in the Late Permian disappeared during the Griesbachian,and then reappeared again in the Dienerian. Oceanic anoxia oreuxinia suppresses decomposition of organic matter in a middle-deep water column, instead such a condition enhances directdelivery of unaltered organic matter, in particular proteins withrelatively heavy C-isotope ratios. The P-TB interval with superan-oxia, in particular at its peak in the Griesbachian (Fig. 4), probably

corresponds to this case with more effective transportation of pro-teins to the deep-sea. In the Lopingian, deep-seated anoxic waterincreased in volume to push the chemocline upward, then aroundthe P-TB at the zenith of superanoxia, the chemocline may havereached to a very shallow level of ocean (Isozaki, 1994) as con-firmed by the biomarker for photic zone euxinia (Grice et al.,2005; Hays et al., 2007). The unique ‘‘superanoxic condition” inthe Griesbachian superocean likely drove an unusual, isotopicallymonotonous oceanic structure all the way from bottom to top(Fig. 6 lower), and it ended when the chemocline retreated backto the deep-sea; i.e., re-oxygenation of deep ocean.

5.3. Possible causes for the double-phased extinction

The above-mentioned double-phased environmental changeassociated with major extinction, in particular, the secular changesof C and Sr isotopes clearly identified the G-LB and P-TB events astwo independent events, highlighting the long-overlooked signifi-cance of the G-LB event and the uniqueness of the Late Permianinterval. The G-LB and P-TB events were clearly separated for ca.8 m.y., therefore, these two independent events should have beentriggered by two distinct causes of global influence. Given thateverything occurred in a double-phased manner, all the geologicalphenomena during the second half of the Permian can be summa-rized in the framework of two sets of cause, process, and result. As


to the possible causes of the biggest mass extinction in the Phan-erozoic, however, previous discussion naturally focused mostlyon the P-TB event.

The most popular scenario at present for the P-TB event is vol-canism-induced environmental turmoil and resultant extinction(e.g., Wignall, 2002; Erwin, 2006; Bottjer et al., 2008), becausethe apparent coincidence in timing between extinction and large-scale volcanism, in particular continental flood basalts (CFB) suchas the Siberian Traps, has often been emphasized (e.g. Campbellet al., 1992; Renne et al., 1995; Courtillot, 1999). Several varietiesof extinction scenario were derived from this assumption of theCFB-extinction link; e.g., the emission of voluminous volcanic gasinto the atmosphere and the associated melting of methane hy-drates to release more CO2 for an apocalyptic super-greenhousecondition in the biosphere (Retallack, 1999; Wignall, 2002; Hueyand Ward, 2005; Meyer and Kump, 2008). Possible direct killmechanisms include (1) suffocation by O2 depletion (Wignall andHallam, 1992), (2) hypercapnia with excess CO2 (Knoll et al.,1996), and (3) poisoning by H2S (Riccardi et al., 2006). All thesehypotheses assume volcanism-triggered global warming and asso-ciated oceanic stratification on a global scale. In accordance withthe P-TB superanoxia, the upward migration of chemocline in theworld ocean is commonly speculated. For generating such alarge-scale change in ocean structure of global extent, the temper-ature gradient between equatorial and polar regions, likely playedan important role (e.g. Hotinski et al., 2001). Nonetheless the onsetmechanism of such a unique oceanographic condition remainedenigmatic, and this will be discussed later. The hypothesis of P-TB global warming induced by CFB volcanism/methane hydratesmay explain the delayed recovery of biotic diversity during theEarly Triassic but not necessarily the cause of the extinction atthe P-TB per se. As to the timing, however, there is still an apprecia-ble time gap between the extinction and CFB volcanism. Further-more, the extinction slightly but apparently predated the CFBvolcanism. More precise dating is definitely needed for checkingthe potential cause–effect link between the CFB volcanism andextinction. An alternative scenario with volcanism-related killmechanism emphasizes the importance of felsic magmatism asso-ciated with continental rifting (‘‘plume winter” scenario by Isozaki(2007b)), and it is explained in the next section.

On the other hand, by applying the K-T boundary scenario,some have claimed an extraterrestrial bolide impact as the causeof the P-TB event (e.g., Becker et al., 2001; Kaiho et al., 2001); how-ever, failures in reproducing identical geochemical and mineralog-ical signals from the same sample and also in identifying the exactboundary horizon have hampered broad acceptance of the idea(e.g., Farley et al., 2005; Isozaki, 2001; Koeberl et al., 2002). The lastclaim for a possible contemporary impact crater off Western Aus-tralia (Becker et al., 2004) was also refuted for various reasons(e.g. Renne et al., 2004). In fact the impact crater in Australia isinconsistent with the fact that the contemporary off-shore ofNew Zealand, located not far away from Australia, served as thesole exceptional refugia for the Paleozoic radiolarians during theP-TB extinction event (Takemura et al., 2007). The apparent chro-nological similarity may suggest a possible causal link betweenextinction and extraterrestrial impact (or CFB volcanism) but can-not prove it without solid lines of material-based evidence, as weexperienced in the K-T boundary controversy until the discoveryof the culprit crater in Yucatan.

The cause of the G-LB event has been discussed less frequentlywith respect to the P-TB. The volcanism of the Emeishan Traps inwestern South China and the Panjal Traps in northern India waslikewise claimed as a possible cause of the G-LB extinction onthe basis of their chronological coincidence (e.g., Chung et al.,1998; Courtillot, 1999; Zhou et al., 2002; Ali et al., 2002; Heet al., 2007). However, the nominated CFBs in S. China and India

are apparently too young and too small to have been responsiblefor all the above-listed strange geologic phenomena that startedwithin the Capitanian. Just like the P-TB case, the significance offelsic magmatism needs more attention also for the G-LB case (Iso-zaki and Ota, 2001; Isozaki et al., 2004), as will be discussed in thenext section. There is no evidence at all for an extraterrestrialimpact around the G-LB, thus we need not to pursue the idea ofimpact of large bolides.

As mentioned above, the ultimate causes of the G-LB and P-TBevents have not been identified yet; however, the current dataindicate that no bolide impact scenario is necessary for both cases,instead that terrestrial processes, such as formation of LIPs by themantle plume activity, were somehow related. The G-LB and P-TBevents share some similarities, but on the other hand, they alsohave clear disparities. Thus it is obvious that the two events donot represent the consequences of simple repetition of the samecause and processes. In the following section, at first, common trig-ger and processes between the two events are discussed with spe-cial emphasis on non-basaltic magmatism. Then in the second nextsection, differences between the two events are reconsidered withparticular focus on the unique geomagnetic phenomenon that ap-peared in the late Guadalupian, and a possible scenario is exploredfor total explanation of all similarities and disparities between thetwo extinction-relevant events.

5.4. Double-phased ‘‘plume winter”

On the basis of all the available datasets from both shallow mar-ine and deep-sea facies in Panthalassa and Tethys, Isozaki (2007b)proposed the ‘‘plume winter” scenario, in order to explain thegreatest turmoil of the Phanerozoic biosphere that took place inthe transition interval from the Paleozoic to the Mesozoic-modernworld. This hypothesis emphasized that the greatest mass extinc-tion was led by extraordinarily large-scale magmatism probablyderived from a mantle superplume impinged beneath the super-continent Pangea, therefore, at a glance, it may appear similar tothe volcanism-related scenarios previously proposed (e.g., Camp-bell et al., 1992; Renne et al., 1995; Chung et al., 1998; Courtillot,1999). The plume winter scenario, however, focuses not on CFBas suggested by many but on plume-generated felsic alkaline vol-canism, simply because tuff beds around the G-LB and P-TB are allfelsic in composition.

In searching for the ultimate cause among various major eventsduring the Paleozoic–Mesozoic transition interval, the most impor-tant is likely the felsic volcanism that occurred uniquely aroundthe G-LB and P-TB. The concentrated occurrence of felsic volcanismat these timings suggests their origin not from steady-state magmasources but from episodic ones. The boundary felsic tuffs were rec-ognized extensively in many P-TB sections in South China (e.g., Yinet al., 1993). It is noteworthy that the onset of felsic volcanismslightly predated the timing of final extinction, thus such volcan-isms appear critical in assessing the causal mechanism of theextinction-relevant environmental changes. For example, morethan 10 tuff beds across the P-TB (within a 4 m-thick interval span-ning the uppermost Changhsingian and lowermost Griesbachian)were confirmed in northern Sichuan (Li et al., 1989; Isozaki et al.,2004, 2007c). One of these tuffs in the middle horizon is preciselydated at 251–252 Ma as the P-TB marker bed in South China (e.g.,Bowring et al., 1998; Mundil et al., 2004; U-Pb zircon age by ID-TIMS, not by SHRIMP).

The G-LB tuff in South China has long been regarded as a clay-stone under the traditional name ‘‘Wangpo bed” (Lu, 1956; Li et al.,1991); however, this unit represents a bona fide air-borne tuff thatcontains abundant volcanogenic material including euhedralphenocrysts of quartz, plagioclase, apatite, and zircon (Isozakiet al., 2001, 2004, 2008; Isozaki, 2007b). The Wangpo tuff, ca.


1–2 m thick, occurs throughout western South China from Shaanxito Guanxi, covering nearly 1000 km in an N–S direction (Isozakiet al., 2004) and this G-LB tuff in northern Sichuan has a U-Pb zir-con SHRIMP age of 260 ± 4 Ma (He et al., 2007). In addition, theoccurrence of the Capitanian fine-grained tuff beds in mid-oceanicsequences (Isozaki, 2007b) suggests much more extensive deliveryrange of ash; i.e., Western Panthalassa together with South Chinawas covered on a regional scale by air-borne tuff of rhyo-daciticcomposition in the Late Guadalupian. An unusually large-scaleand highly explosive volcanism of the Plinian to Ultraplinian-typecan solely explain such a long-distance delivery and a huge volumeof tuffaceous material. In contrast to mafic, felsic volcanism usuallyoccurs with a violent, explosive eruption that can explain the puta-tive environmental catastrophe that may have driven extinction.

The occurrence of the prominent Permian LIPs and contempo-rary volcanogenic sediments (Isozaki, 2007b) indicates that large-scale volcanism of both felsic and mafic composition occurred inthe second half of the Permian, probably peaked twice aroundthe G-LB and the P-TB. Steady-state plate boundary processesalong mid-oceanic ridges and arc-trench systems continue all thetime as long as plate tectonics is in operation on the Earth. Thusrelevant volcanisms keep active without having episodic or cata-strophic pulses that can drive unusual events like mass extinction.Instead, within-plate magmatism derived from mantle plumeslikely had a high episodicity in forming LIPs (Ernst et al., 2007).In fact, some LIPs of the Permian age accompany andesitic basaltand rhyolite units, e.g. the Emeishan and Siberian Traps; however,these felsic units occur limitedly in the middle and upper parts ofthe volcanic edifice, thus were too young to have caused theextinction-related environmental change. They are also too smallin amount to have accounted for the entire coverage of the bound-ary tuffs, as mentioned above. As to the Emeishan Traps, He et al.(2007) speculated for a voluminous rhyolites above the main ba-salt unit and their extensive erosion on the basis of geochemistryof the overlying Upper Permian rocks; however, the generation ofclaimed large amount of felsic magma from the source rocks(mostly mantle peridotite), with respect to the volume of basalticone, is unlikely because of multiple steps of partial melting re-quired for producing felsic magma. The non-basaltic compositionof the Upper Permian sediments probably reflects the erosion ofsialic (continental) crust of older basement or pre-basalt felsicvolcanics.

By considering the above issues, Isozaki (2007b) speculated thatan episodic activity of mantle superplume might have driven theG-LB and P-TB events, because the Paleozoic–Mesozoic transitioncorresponds to the time of initial rifting of Pangea prior to its mainbreakup in the Mesozoic. As to the breakup of a supercontinent, theuprising of a superplume and its impingement at the continentallithosphere are vital. At some time in the Guadalupian, a superp-lume was likely launched from the deep mantle, particularly fromthe D00 layer (ca. 2600–2900 km deep) immediately above the core-mantle boundary in order to counterbalance in mass for thedescending cold plume or megalith (subducted oceanic slabs)(Maruyama, 1994; Maruyama et al., 2007).

A mantle superplume likely branches into several secondaryplumes in the mantle transition zone (ca. 410–660 km deep) wherethe density contrast between a plume and surrounding mantle de-creases remarkably according to the mineral phase transition(Maruyama, 1994; Maruyama et al., 2007). The Middle Permiansuperplume probably branched into plural secondary plumes,which impinged at the bottom of the supercontinent in two stepsto cause a ‘‘plume winter” twice; i.e. around the G-LB and P-TB.

The impingement of a main plume head beneath Pangea prob-ably generated strong magmatism by decompressional melting ofmantle peridotite (+ recycled slabs). Prior to the main flood basalteruption, however, re-melting of pre-existing continental sialic

crust likely occurred to generate felsic magmatism. This may rea-sonably explain the order of the two distinct types of volcanismsat the G-LB and P-TB; i.e. first felsic then mafic, not vice versa. Inaddition, eruption of gaseous kimberlite may have been accompa-nied as a precursor of plume magmatism (Morgan et al., 2004;Isozaki, 2007b), in particular, when a superplume impinged at aCO2-enriched tectosphere beneath relatively older continents(Santosh et al., 2009). Ultraplinian-type eruptions of these felsic pluskimberlitic volcanoes at multiple sites in swarm may have drivensevere environmental changes in the biosphere through volcanichazards; e.g. toxic gas emissions, developing dust/aerosol screensin the stratosphere (blocking sunlight), and pouring acid rain. Theircascading effects on the environments, such as temperature drop,dim daylight, cessation of photosynthesis, and shortage in food,may have led to the decline in biodiversity both in the ocean andon land; i.e. a condition called ‘‘plume winter” (Isozaki, 2007b).Although actual kill mechanisms should be scrutinized more indetail for each animal group both in marine and terrestrial settings,all the responsible mechanisms likely became effective by the globaldevelopment of the background ‘‘plume winter” conditions.

The first arrival of the Permian superplume at the bottom ofPangea was reflected in the Phanerozoic minimum in the 87Sr/86Srratio in the Capitanian marine sediments (Kani et al., 2008) that iswell correlated with peri-Pangean shelf carbonates (Veizer et al.,1999; Korte et al., 2006; Fig. 2). Since the Cambrian, 87Sr/86Sr ratiosof marine carbonates kept decreasing for nearly 300 million yearsregardless of minor fluctuations; however, they suddenly changedthe trend in the Late Guadalupian and started to increase rapidlytoward the early Mesozoic high, suggesting that a large-scale reor-ganization had occurred in the Sr isotope balance in the Panthalas-san seawater between the mantle flux and the continental flux. Ingeneral, low sea levels induce active erosion of continental cruststhat increases 87Sr/86Sr ratio of seawater. During the Capitanianwith the lowest sea level in the Phanerozoic, however, the 87Sr/86Srratios of marine carbonates globally marked the lowest values. Iso-zaki (2007b) and Kani et al. (2008) interpreted that this contradic-tory behavior of seawater Sr isotope ratio have been related to theinitial rifting of Pangea in the late Guadalupian. By making newdrainage systems directly connected the intra-Pangean basinsand Panthalassa, the rifting may have triggered abrupt mass wast-ing of terrigenous clastics that accumulated within Pangea since itsformation in the Late Paleozoic.

As to the P-TB, no measurements on Sr isotopes have been madeof mid-oceanic rocks; however, the previous data from continentalshelf carbonates demonstrate that 87Sr/86Sr ratios uni-directionallyincreased across the P-TB (Veizer et al., 1999) merely with a short-term stand-still (Twitchett, 2007). This highlights the uniquenessof the trend change and abrupt increase of Sr-isotope around theG-LB, and proves that no major erosion/mass wasting occurred inPangea across the P-TB. In addition, it is noteworthy that the Sr-isotope fluctuation in seawater was essentially independent fromthat of the stable carbon isotope around the P-TB interval, althoughboth C and Sr isotope systems started to shift drastically at thesame time not at the P-TB but in the Capitanian.

It is emphasized here that both the G-LB and P-TB extinctionswere probably led by plume-generated felsic/kimberlitic volca-nism, because flood basalt volcanism took place slightly later thanthe main extinction in both cases. The relatively lower viscosity ofbasalt magma cannot cause explosive eruption sufficient for deliv-ering air-borne tuff over great distances. Instead, the input of ahuge amount of CO2 to the atmosphere by plume-driven volcanism(all of kimberlitic, felsic and mafic) may explain the post-extinctionapocalyptic global warming (a plume summer) coupled with a sea-level rise and the retarded post-extinction recovery in the bio-sphere in the Early Triassic (Twitchett, 1999). In short, a pair of‘‘plume winter” and ‘‘plume summer” appeared twice around the


G-LB and P-TB probably by the two-stepped activity of branchedsecondary plumes.

The mechanism of developing the 20 m.y.-long superanoxia isstill enigmatic. In general, a stratified ocean is difficult to maintainfor millions of years under the oxygenated atmosphere, as long asthe planet Earth keeps its rotation and temperature gradient existsbetween the polar and equatorial regions, as demonstrated in thenumerical modeling (e.g. Hotinski et al., 2001). The apparentconundrum of the long-term development of deep-sea anoxia,however, may be explained reasonably by assuming two steps ofoxygen depletion caused by two independent triggers (Isozaki,2007b). The earlier half of the superanoxia (during the early Lopin-gian) was probably related to the first plume summer that sup-pressed the ocean mixing after the G-LB. The anoxia may havebeen gradually relaxed with fluctuation in redox during the lateLopingian. Then, before the superocean returned to a fully oxiccondition, the second plume summer (by the Siberian Traps) ap-peared to drive the anoxia into an ultimate state across the P-TB(the zenith of superanoxia), and it took long time to be cancelleduntil the early Middle Triassic. The onset mechanism of the anoxianeeds to be discussed separately below because the anoxia startedin the late Guadalupian prior to the first plume summer, thus theinitial trigger was not related to the plume volcanism.

6. Integrated ‘‘plume winter’’ scenario: volcanic cooling andgeomagnetic cooling

As discussed above, the G-LB and P-TB events share many com-mon geological characteristics, suggesting that the two events have

Fig. 7. Chronicle of the Late Permian ‘‘triple-double”. The Paleozoic–Mesozoic transition iof volcanism (cause), environmental change (process), and extinction (effect), acrossconcentrated around G-LB and P-TB; nonetheless, note that all started not around P-TB bulikely launched from the core-mantle boundary as witnessed as the Illawarra Reversal.

similar causes probably related to mantle superplume activity.Nonetheless, clear differences exist between the two events, indi-cating that not exactly the same cause and processes repeated.By referring to the unique geologic phenomena that occurred par-ticularly in the late Guadalupian, I here present an integrated ver-sion of the ‘‘plume winter” scenario at the end of this article.

Fig. 7 illustrates the entire chronicle of the peculiar changes inthe core, mantle, crust, and surface biosphere of the planet duringthe Paleozoic–Mesozoic transition interval that was punctuatedtwice by the G-LB and P-TB events. The G-LB event, or more pre-cisely speaking, the Capitanian event should be much emphasizedsimply because everything geologically unusual appeared at first inthe Capitanian. In fact, after the long quiescence since the mid-Car-boniferous Pangean assembly and ca. 8 million years before the P-TB per se, various unique changes on the Earth’s surface startedduring the Capitanian (Isozaki, 2007b). In addition to the majorfaunal turnover (Jin et al., 1994; Stanley and Yang, 1994) and theonset of LIP-forming magmatism (Isozaki, 2007b), the followingfeatures uniquely characterize the Capitanian event; i.e., (1) onsetof superanoxia (Isozaki, 1997), (2) onset of volatile C-isotope fluc-tuation (Isozaki et al., 2007a,b), (3) trend change in Sr-isotope ratio(Veizer et al., 1999; Kani et al., 2008), (4) sea-level change from along-term regression to a transgression (Haq and Schutter, 2008),and (6) pattern change in geomagnetic polarity (Gradstein et al.,2004; Steiner, 2006).

The most unique feature in the above-listed geologic phenom-ena related to the G-LB event is the sharp pattern shift in geomag-netic polarity called the Illawarra Reversal (Fig. 8). This remarkablechange in geomagnetic field reversal pattern took place around the

nterval is characterized by double-phased plume winter-related events; i.e. two setsthe G-LB and the P-TB. Major geological phenomena occurred twice, uniquely

t around the Wordian–Capitanian boundary (Guadalupian) when a superplume was


Wordian–Capitanian boundary (ca. 265 Ma), marking a sharp tran-sition from the long-lasting Kiaman Reverse Superchron (Late Car-boniferous to Middle Permian) to the Permian–Triassic MixedSuperchron (Late Permian to Late Triassic) (e.g., Jin et al., 2000;Gradstein et al., 2004; Steiner, 2006). The Illawarra Reversal wasrecently confirmed also in the mid-Panthalassan paleo-atoll car-bonates in Japan (Kirschvink and Isozaki, 2007; Fig. 5). Such aremarkable change in the behavior of the Earth’s magnetic field

Fig. 8. Geomagnetically unique periods in the Phanerozoic (left) and the Permian magnmodified from Gradstein et al. (2004) and Courtillot and Olson (2007). Note the clear conPermian) and the Permo-Triassic Mixed Superchron (late Middle Permian to Late Triassicmode to a frequently oscillating one. In order to trigger such instability of dipole, a rapiboundary (D00 layer). The episodic launching of a mantle superplume coupled with downdisturbance (Fig. 9A). The Mid-Cretaceous Normal Superchron and the Ordovician Moyamong geomagnetic intensity, influx of galactic cosmic radiation, and surface climate of

provides direct evidence for the appearance of a major thermalinstability at the core-mantle boundary that may have been relatedto the probably synchronous launching of a superplume (Isozaki,2009). When the stability of Earth’s dipole field declined, a de-crease in geomagnetic intensity could have been triggered by alter-ation of molten iron circulation patterns in the outer core, driven inturn by events at the core-mantle boundary (e.g., Tarduno et al.,2006; Courtillot and Olson, 2007).

etostratigraphy punctuated abruptly by the Illawarra Reversal at ca. 265 Ma (right)trast between the Kiaman Reverse Superchron (Late Carboniferous to early Middle

). The Illawarra Reversal recoded a sharp change in geodynamo from a stable dipoled disturbance of thermal conditions is needed immediately above the core-mantle-welling of a megalith (subducted slabs) is the most likely cause of such a thermal

ero Reverse Superchron provide another good examples to check the possible linkthe Earth.


Through consideration of the superplume-launching event inthe late Middle Permian, Isozaki (2009) further extended the

Fig. 9. Schematic diagrams of the integrated ‘‘plume winter” scenario. The episodic activcause of the double-phased mass extinction at the G-LB and P-TB. (A) The Illawarra ReversCapitanian (late Middle Permian) (ca. 265 Ma), a dramatic change took place at the coroceanic slabs beneath Pangea) from the mantle transition zone (410–660 km deep). Tappearance of thermal instability on the core’s surface that drove a major mode change ointensity of the field, thus increased the flux of galactic cosmic rays to the atmosphere. ByKamura cooling event (Figs. 5 and 7) that resulted in the selective decline/extinction of tthe expansion of O2 minimum zone in mid-depth. It is noteworthy that the sea level reachplume winter condition (modified from Isozaki, 2007b): The breakup of the supercontinbase of a supercontinental lithosphere nearly 5 million years after the Illawarra Reversadeep), a mantle superplume branched into plural secondary plumes due to the densitymillion years of the Paleozoic–Mesozoic transition, two episodes of ‘‘plume winter” occusecondary plumes that branched from the main superplume in two steps. Through the(volcanic cooling, toxic gas emission, pouring acid rain, and sunlight-blocking dust/aerososhutdown of sunlight and global cooling by the dust/aerosol screen (plume winter) appglobal scale is critical in maintaining bio-diversity. A short-time darkness and coldness cThe following flood basalt volcanism drove the surface environment into global warmingthe stratification of the superocean Panthalassa. Although caused by the similar supevolcanic cooling, whereas the P-TB event lacked the former process. The lowest biodivershort between the two episodes for retrieving the full biodiversity back to the Paleozoic

‘‘plume winter” scenario to connect the unique event in historyof the geodynamo and geomagnetic field with the concurrent

ity of a mantle superplume beneath the supercontinent Pangea was the most likelyal and the Kamura event (modified from Isozaki, 2009): Around the beginning of thee/mantle boundary (D00 layer) by the drop of a relatively cold megalith (subductedhe prominent geomagnetic event called Illawarra Reversal (Fig. 8) recorded thef the geodynamo in the outer core. The unstable dipole likely lowered geomagneticthe enhanced cloud coverage, the Earth’s albedo increased to trigger the Capitanian

he unique tropical fauna and the unusually high oceanic productivity coupled withed the lowest of the Phanerozoic and the superanoxia started around this time. (B) Aent Pangea was intimately linked to the impingement of a superplume-head at thel, i.e., around the G-LB (ca. 260 Ma). In the mantle transition zone (ca. 410–660 kmcontrast of rocks related to mineral phase transitions along depth. During the 20rred intermittently, i.e., at the G-LB and P-TB. These were triggered by two sets of

violent magmatism of felsic and partly kimberlitic nature, various volcanic hazardsls) led the Earth’s surface to a ‘‘plume winter” condition. Besides everything else, theear to have been the most significant because the cessation of photosynthesis on aan halt photosynthesis on a global scale and disorganize the pre-existing food web.

(plume summer), delayed post-extinction recovery in biodiversity, and intensifiedrplume-related processes, the G-LB event involved both geomagnetic cooling andsity was recorded across the P-TB event probably because the time interval was too

plateau level (Fig. 1A).


changes in the biosphere around the G-LB (Fig. 9A). For launching ahuge mass from the deep mantle, the fall-down of a mass of sub-ducted slabs (megalith) is necessary from the upper mantle tothe D00 layer above the core-mantle boundary, and this providesthe most compelling mechanism for the destabilization of the geo-magnetic dipole. Such an event would place a huge, relatively coolmass directly onto the bottom of mantle, driving both upwelling ofa warm superplume through volume displacement and rapid prop-agation of thermal instabilities from the D00 layer into the outercore. In fact, this is the one and only practical way to modify themode of geodynamo rapidly and drastically. Thus the IllawarraReversal recorded in surface rocks provides a likely prime markerfor the launching of a superplume from the core-mantle boundaryaround 265 Ma.

Destabilization of the geomagnetic dipole may increase the like-lihood of frequent polarity changes and may also weaken fieldintensity, as often observed in short-term fluctuations in the Qua-ternary. Assuming a constant galactic cosmic ray flux, increasingcosmic ray penetration and bombardment of the Earth’s atmo-sphere would track with decreasing geomagnetic field intensityas the Earth’s field provides the primary armor deflecting awayhigh-energy cosmic radiation (Tinsley, 2000). Although the funda-mental processes of cloud formation are still under scrutiny (e.g.,Haigh, 2007), an increase in high-energy cosmic ray flux may en-hance both ionization of atmospheric molecules (aerosols) andresulting formation of cloud nuclei; i.e., the more cosmic radiation,the more cloud coverage over the Earth (e.g., Dickinson, 1975;Svensmark, 2007). Along this vein of reasoning, increased cloudcover leads to increased albedo, decreased sunlight (heat) flux tothe Earth’s surface, and decreased global surface temperatures.

If this were the case for the Illawarra Reversal in the Guadalu-pian, a geodynamo-driven decrease in surface temperature couldreasonably explain the Capitanian Kamura cooling event coupledwith the lowest sea level in the Phanerozoic (Figs. 2, 5, 7 and 8A;Isozaki, 2009). Global cooling may have accelerated ocean mixing,providing an increased nutrient supply to the surface ocean todrive the Kamura high-productivity event. Increased surface pro-ductivity, in turn, may have expanded the mid-depth oxygen-min-imum zone of the oceans and knocked the oxic, early Capitaniandeep-sea into anoxia. The onset of the superanoxia was likely re-lated not to a global warming by plume-generated volcanism butto the Capitanian Kamura cool event, whereas the developeddeep-sea anoxia was maintained for a while under the followingglobal warming across the G-LB. In this regard, the drawdown ofatmospheric CO2 recorded in the d13Ccarb was not a likely causeof the cooling as originally proposed (Isozaki et al., 2007a) butprobably one of the consequences of the geomagnetic cooling.

Around 260 Ma, nearly 5 m.y. after the putative superplumelaunching from the core-mantle boundary, the first magmatismderived from secondary plumes appeared on the surface, leadingto the first ‘‘plume winter” (Fig. 9B). This event likely acceleratedglobal cooling to complete the G-LB extinction. Against the cur-rently popular interpretations of global warming and relevantextinction both at the G-LB and P-TB, the significance of coolingdriven by mantle plume activity is emphasized in this article.The two extinction-bearing boundary events likely shared ‘‘volca-nic cooling” induced by secondary plumes; however, the P-TBevent did not experience ‘‘geomagnetic cooling”. This made themajor difference between the two events and highlighted theuniqueness of the G-LB event or the Capitanian interval.

At the end of discussion, other geologic periods with long-termstable geomagnetism are checked from the similar viewpoints(Fig. 8). The mid-Cretaceous Normal Superchron (ca. 118–83 Ma)demonstrates another such link between geomagnetic fieldstability and climate, but going in the opposite direction.The Earth’s atmosphere during the Long Normal Superchron, a

geomagnetically stable interval, would have experienced a low cos-mic ray flux and thus higher anticipated surface mean temperatures.This is in agreement with the long-term warm climate for the mid-Cretaceous coupled with a high sea level, thermally stagnant oceans,and resulting episodic oceanic anoxia (e.g., Larson, 1991). Detailedmagnetostratigraphy of the Ordovician is still unknown; however,the ‘‘relatively stable” period called Burskan reversed bias polarityinterval (particularly its later part; Algeo, 1996) or Moyero ReverseSuperchron (ca. 490-460 Ma; Pavlov and Gallet, 2005) appears con-cordant with the long-lasting warm period that allowed the rapiddiversification of metazoans (e.g., Servais et al., 2009).

In contrast, the mid-Carboniferous to mid-Permian Kiaman Re-verse Superchron (ca. 310–265 Ma) chronologically corresponds toa much colder period, including the Gondwana glaciations(Gradstein et al., 2004). This case apparently contradicts the geo-magnetic field stability/climate model. Shaviv and Veizer (2003)resolved this contradiction, however, by explaining two indepen-dent variables for controlling the cosmic ray flux, i.e. the intensityof radiation source(s) and the strength of magnetic shield by Sunand Earth. Cosmic ray flux into the Earth’s atmosphere dependsnot only on geomagnetic and helio-magnetic field intensity butalso on the number, size, and proximity of galactic cosmic raysources. Due to the proximity between our solar system and asupernova-ridden spiral arm of the Milky Way galaxy, the flux ofcosmic radiation increased in the mid-Carboniferous to EarlyPermian interval (including the Kiaman Superchron), independentof changes in the geo- and helio-magnetic field. The period of thisoscillation is estimated ca. 143 ± 10 m.y. (Shaviv and Veizer, 2003)that is much longer than that for the heliomagnetism and muchshorter than that of superplume-induced geomagnetic rhythm,but concordant with the well-known long-term cycle of green-house-icehouse of the Phanerozoic. As to the two remarkablesnowball Earth events in the Paleo- and Neo-proterozoic, i.e., the2.3–2.2 Ga and 0.7–0.6 Ga events, Veizer (2005), Svensmark(2007), and Maruyama et al. (submitted) further speculated forthe cosmoclimatological link between the major cooling event inhistory and increased galactic cosmic radiation in terms of the dis-tance from the neighboring galaxies.

7. Conclusion

Recent studies in Japan on the accreted ancient mid-oceanicrocks have provided vital information on the unique double-phased global event that occurred in the lost Panthalassasuperocean during the Paleozoic–Mesozoic transitional interval.In addition to the conventional data from the peri-Pangean shallowmarine settings, all the new information from Panthalassa con-firmed that the era-boundary event between the Paleozoic andMesozoic was not a simple short-term episode, like the end-Creta-ceous case with one large meteorite impact and a resultant massextinction. Instead, the event was composed of much longer-ranged complicated geologic/geophysical phenomena that tookplace in various parts of the Earth from its core to the surface. Inparticular, the double-phased nature of the extinction and relevantenvironmental changes at the G-LB and the P-TB was clearly dem-onstrated in a global context, and this requires at least two inde-pendent causes that were chronologically separated for at least8 m.y. from each other. Evidence from field geology suggests thata superplume activity in mantle most likely played the main rolein extinction-related environmental changes at the G-LB and P-TB. The double-phased felsic (plus kimberlitic) volcanism,double-phased environmental turmoil, and double-phased extinc-tion (triple–double) during the Paleozoic–Mesozoic transitioninterval, are by and large summarized into two repeated sets oftrigger, process, and consequence that were essentially driven by


double-phased activity of secondary plumes branched from a mainsuperplume. The plume-derived violent volcanism in two stepslikely caused ‘‘plume winter” twice at the G-LB and P-TB to drivethe extinctions. The G-LB event (or Capitanian event) was distinctfrom the P-TB event, however, as associated with the lowest sea le-vel, the lowest Sr-isotope ratio, onset of superanoxia, and theprominent change in geomagnetic stability.

The integrated version of the ‘‘plume winter” scenario appearspromising because this interpretation connects various geologicalphenomena occurred in multiple spheres of the Earth, from its coreto surface, and uniquely in the Paleozoic–Mesozoic transitioninterval, and because it likely explains all concordantly in one pic-ture. Many parts of this ‘‘grand-unified explanation” presentedhere, however, are obviously too immature to be validated imme-diately, therefore, this challenging new interpretation needs to bechecked and polished through serious criticisms from many quar-ters. Nevertheless, this explanation may revive general interests inseeking possible interactions between terrestrial and extraterres-trial phenomena, in a completely different way from the widelyknown bolide impact scenario. The viewpoint and methods usedfor analyzing the Paleozoic–Mesozoic transitional interval mayprovide a template for further application to paleoenvironmentalstudies on the pre-200 Ma deep past, i.e., the Paleozoic and Pre-cambrian time. In particular, re-evaluation from alternative per-spectives is needed for major events in the history of biosphereand life evolution during the Precambrian.

Acknowledgments

The author thanks Hodaka Kawahata, Kayo Minoshima, HidekiWada, Masaaki Musashi, Tomomi Kani, Susumu Nohda, HisayoshiIgo, Ayano Ota, Akihisa Kasuya, Yao Jianxin, Ji Zhangsheng, MasakiOgawa, Shigenori Maruyama, Joseph L. Kirschvink, Isaac Hilburn,and Paul G. Falkowski for their valuable co-operation and discus-sion in the field and laboratory. Adam Woods and Paul B. Wignallprovided constructive review comments. Brian F. Windley cor-rected the English. This research was supported by a Grant-in-Aid of the Japan Society of Promoting Science (Nos. 16204040,20224012).

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