Abstract: This reviewadvancesunderstandingof thepalynostratigraphyof theLateTriassic–EarlyJurassic by correlating the established palynozonations for the northern and southern hemispheres.Previous palynological studies have contributed greatly to our understanding of the biostratigraphy,paleoclimatology and paleogeography of the Upper Triassic. In general, palynology is a good tool forinterregional cross-correlation of marineand non-marine successions because palynomorphs, unlikemost of other fossils, commonly are present in continental and marine environments. Currently,however, biostratigraphical resolution based on Upper Triassic palynomorph assemblages israther low, primarily because of the rarity of successions that are independently dated (i.e. via ammo-noids, conodonts, isotopes, paleomagnetism) to correlate the palynomorph assemblages, but also forother reasons, such as microfloristic provincialism, palaeoenvironmental conditions and differentialpreservation of palynomorph assemblages. During the last few decades many palynological studieshave attempted to integrate and improve the biostratigraphical correlations and paleoclimatologicreconstructions across the Triassic–Jurassic boundary. Several authors have recognized specificmicrofloral assemblages with well-defined and recognizable suites of palynomorphs that enhancethe importance of palynomorphs in the definition of Triassic–Jurassic stages. Comparison of thepalynomorph assemblages from different biostratigraphical stages demonstrates that a changeoccurred in the palynofloral composition of the Tethyan domain between the Carnian and the earliestHettangian that was gradual and without abrupt changes.
Upper Triassic palaeogeography and
palaeoclimatology
The Late Triassic, a geological time interval ofabout 30 million years (Gradstein et al. 2004), wasa period of rapidly changing palaeogeography.During the Triassic, Pangaea, which extendedfrom about 858N to 908S, progressively evolvedfrom maximum continental assembly to the inititia-tion of breakup (Ziegler et al. 1983; Hesselbo 2000;Golonka 2004; Buratti & Cirilli 2007). Beginning inthe Middle Triassic, Pangaea started to divide intothe northern landmass Laurasia (including presentday North America, Europe and Asia) and thesouthern landmass Gondwana (South America,Africa, Antarctica, India and Australia), separatedby the narrow and shallow Tethys seaway. In thesouthern hemisphere, the Cimmerian blocks(present day Turkey, Iran and Tibet) separatedfrom Australia and moved northward, originatingthe back-arc basin of Neotethys, and after the sub-duction of Paleotethys, collided against Eurasia atthe end of the Triassic (Corsin & Stampfli 1977;Sengor 1984; Marcoux et al. 1993; Muttoni et al.1996, 2001; Besse et al. 1998; Gaetani et al. 1998,2005; Dercourt et al. 2000; Moix et al. 2008). Therift between Africa, South America and NorthAmerica was probably already active by the end
of the Triassic, as documented by the Karoosynrift deposits in southeastern Africa (Zerfasset al. 2004; Geiger et al. 2004). This rifting stagetook place contemporaneously with the emplace-ment of the voluminous tholeiitic dikes, sills andflood basalts of the Central Atlantic MagmaticProvince (CAMP) over a total surface area ofabout 10 million km2 in North and South America,Africa and Europe (Marzoli et al. 1999, 2004;Olsen et al. 2003; Verati et al. 2007).
The global Late Triassic climate and thus thefloral distribution were strongly controlled by thesymmetrical distribution of the Pangaean land-masses around the equator and by the presumedabsence of polar ice at high latitude (Frakes &Francis 1988; Parrish 1993; Buratti & Cirilli2007). The latitudinal temperature gradient wasappreciably lower than today, resulting in a moreequable average global temperature with warm tem-perate belts expanded to higher latitudes. Severalauthors have posited a strongly seasonal climatefor both hemispheres influenced by the monsooncirculation and enhanced by the shape of Pangaea(e.g. Hallam 1985; Kutzbach & Gallimore 1989;Simms & Ruffell 1990; Dubiel et al. 1991; Parrish1993). Kent & Olsen (2000) hypothesized a zonalclimate pattern for the Late Triassic, with a narrowequatorial humid zone and an arid belt centred
From: LUCAS, S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 285–314.DOI: 10.1144/SP334.12 0305-8719/10/$15.00 # The Geological Society of London 2010.
around 308, passing to more temperate climates athigher latitude.
Rapid negative excursions in d13C in both car-bonate and sedimentary organic matter have beenobserved in marine Tr–J boundary sections closeto the end-Triassic mass extinction, for example,from Hungary, England, USA, Canada, Spain,Italy, Austria (e.g. Palfy et al. 2001; Hesselboet al. 2002; Guex et al. 2004; Ward et al. 2004;Gomez et al. 2007; Galli et al. 2007; Kurschneret al. 2007; Williford et al. 2007; Van de Schoot-brugge et al. 2008; Gotz et al. 2009; Ruhl et al.2009). Additionally, an apparently synchronousdecrease of the stomatal index of fossil leaves(McElwain et al. 1999) has suggested a disruptionof the global carbon cycle, potentially involvingsome combination of global warming, productivitydecline and methane hydrate release (Palfy et al.2001; Tanner et al. 2004; Ward et al. 2004; Lucas& Tanner 2008; Van de Schootbrugge et al. 2008;Ruckwied & Gotz 2009). It has been suggested bysome authors that there was a global greenhousewarming, estimated at about 38 to 4 8C, and acidicatmospheric pollution resulted from a significantatmospheric loading by CO2 and SO2 related toeruptions of the Central Atlantic Magmatic Pro-vince (CAMP) (Marzoli et al. 1999, 2004; McEl-wain et al. 1999; Hesselbo et al. 2002; Guex et al.2004; Tanner et al. 2004, 2007; Schaltegger et al.2008; Van de Schootbrugge et al. 2008, 2009),although it is yet unresolved if the eruptionsstarted before (Marzoli et al. 2004, 2008) or afterthe Tr–J boundary (Whiteside et al. 2007, 2008).Recent palynological studies (Marzoli et al. 2004;Cirilli et al. 2009), in concert with other data(Jourdan et al. 2009) document that CAMP volcan-ism started before the Tr–J boundary and thereforesupport the hypothesis that the CAMP eruptionshad a causative role in the biotic crisis across theTr–J boundary.
Around the Tr–J boundary evidence of increas-ing humidity and seasonality have been recordedas testified by the large amount of enriched organicmatter and shaly sediments deposited within lowdysoxyc-anoxic basins of the western Tethys (e.g.Cirilli et al. 1999; Berra & Cirilli 1997; Boniset al. 2010). As recorded in the Northern CalcareousAlps, the beginning of black shale depositioncoincides with the onset of the initial negative d13Cexcursion (Kurschner et al. 2007; Ruhl et al. 2009)just before the Tr–J boundary. The climate changeand related anoxic conditions have been well docu-mented on the basis of high resolution palynologicaland geochemical data (Bonis et al. 2010).
The location and shifting positions of the paleo-climatic belts strongly controlled paleofloralcomposition and its geographical distribution.Dolby & Balme (1976) distinguished two different
microfloras during the Carnian, the Onslow micro-flora, a mixture of Gondwanan and European taxa,spread throughout western Europe and northwesternAustralia (308–358S) that is characteristic of conti-nental margins and warm temperate rain-forests,and the Ipswich microflora developed in southernand eastern Australia (extending from 258S inQueensland to more than 758S in Antarctica),which is referable to cool temperate plant commu-nities (Fig. 1). The European taxa of the Onslowmicroflora mainly include the genera Aulisporites,Camerosporites, Duplexisporites, Decussatis-porites (¼ Weylandites), Enzonalasporites, Infer-nopollenites, Minutosaccus, Ovalipollis andSamaropollenites, all of which are absent in theIpswich Microflora (Dolby & Balme 1976; Buratti& Cirilli 2007). In the last three decades, numerousadditional phytogeographic data, such as the recov-ery in the northern hemisphere of the index speciesSamaropollenites speciosus, and other typicalsouthern elements of Onslow microfloral affinity,revealed a broad pattern of Triassic palynofloralprovincialism (e.g. Dolby & Balme 1976; deJersey & McKeller 1981; Fisher & Dunay 1984;Helby et al. 1987; de Jersey & Raine 1990; Cirilli& Eshet 1991; Foster et al. 1994; Buratti & Cirilli2007). The common occurrences of the circum-Mediterranean sporomorphs in the Onslow Micro-flora assemblages support the hypothesis that thedistribution of the Onslow Microflora includes theCarnian miospore assemblages extending fromthe western edge of Tethys to its southeastern shor-elines (e.g. Timor) (Buratti & Cirilli 2007).
The strong palynofloral affinity between theCircum-Mediterranean and Onslow microflorasconfirm the existence of a homogeneous parentplant community. This cosmopolitan communitymost likely grew under the influence of an equableclimatic regime characterised by favourable temp-eratures and humid conditions controlled bymonsoonal circulation. A decrease in macrofloraldiversity is associated with a less pronouncedmicrofloristic provincialism, which in the latestTriassic-earliest Jurassic coincided with themigration events of some plant communities (e.g.Cheirolediaceae) from the Tethyan region tohigher latitudes. The diffusion of Cheirolepidiaceae(Circumpolles producers) probably occurred acrossthe coastal migration pathways created during theCarnian–Norian plate reorganization. The recoveryof Circumpolles taxa from the northern Australianmargin (West Timor) (Martini et al. 2004; Buratti& Cirilli 2007) demonstrates that Cheirolepidiaceaereached the southern hemisphere during the Norian.Their dispersal is thus a very important floral eventthat started in the northern and southern hemi-spheres during the Norian, causing the gradualdecline of Upper Triassic microfloral provincialism,
and, later, the global diffusion of a more homo-geneous Lower Jurassic flora. The dispersion ofthe Classopollis group (Circulina–Classopollis–Corollina complex) may be considered as a post-Carnian global microfloristic event, although itappeared at different times in the northern andsouthern hemispheres.
Palynostratigraphy: limits and
applicability
Palynology has great potential as a tool for inter-regional cross-correlation of marine and non-marinesuccessions because spores and pollens are com-monly present in both continental and marineenvironments. At present, however, the biostratigra-phical resolution based on Upper Triassic palyno-morph assemblages is rather low in comparison totheir potential wide distribution. This low resolutionpartially stems from the rarity of successions,particularly in the Norian, dated independently viabiostratigraphy (i.e. ammonoids or conodonts) orother geochronologic tools (e.g. radio-isotopicdating or paleomagnetic correlation). Currently,few marine successions in the world provide anammonoid-integrated palynostratigraphy. Conse-quently, palynomorph assemblages suffer from a
lack of well-calibrated ranges of the most importanttaxa, which sometimes are diachronous on a regionalscale (Batten & Koppelhus 2002). However, the reallimitation is represented by the classical approachbased only on the concept of LO (Last Occurrence)and FO (First Occurrence) instead of on quantitativemethods (e.g. Brugmann et al. 1994; Bonis et al.2009, among others). Other important factors thatmay limit the extension of palynological correlationsinclude microfloral provincialism, diversity of pala-eoenvironmental conditions and differing grades ofpreservation of palynomorph assemblages. Micro-floral provincialism is directly linked to phytogeo-graphic distribution, which is in turn stronglycontrolled by climate.
Additionally, the striking differences in compo-sition of coeval palynomorph assemblages couldderive from diagenetic overprinting, which maydestroy some sporomorphs, but preserve others.Soon after deposition, during the first early diage-netic stages as well as later during the deep burialdiagenesis, the organic facies can undergo pro-nounced alterations that may drastically reduceand select the amounts and types of sedimentedmicroflora, thereby modifing the original palyno-morph assemblages (Tyson 1995; Batten 2002).Consequently, the use of palynostratigraphicschemes in global scale cross-correlations should
Fig. 1. Distribution of Ipswich and Onslow microfloras (after Buratti & Cirilli 2007).
PALYNOLOGY AND PALYNOSTRATIGRAPHY 287
estimate the possible influence of the above-citedlimiting factors and take into consideration thecompositional variation of the palynomorph assem-blages (cf. concept of phase in Schuurman 1977)rather than the occurrence of single forms.
Carnian palynomorph assemblages of the
northern hemisphere
The Carnian palynological record is distingui-shed by the blooming of Circumpolloid generasuch as Camerosporites secatus, Duplicisporites
granulatus, Paracirculina scurrilis, Praecirculinagranifer, and the monosaccate pollen Patinasporitesdensus (Fig. 2). Bisaccate pollen grains are scarce incomparison to the older Ladinian palynomorphassemblages, although some taxa continue to besignificant for palaeofloristic reconstructions (i.e.Samaropollenites speciosus, and the protosaccateOvalipollis pseudoalatus (Dolby & Balme 1976;Visscher & Krystyn 1978; Visscher & Van derZwan 1981; Cirilli & Eshet 1991; Brugman et al.1994; Hochuli & Frank 2000; Warrington 2002;Buratti & Carrillat 2002; Buratti & Cirilli 2007;Traverse 2008).
Callialasporites dampieri
Camerosporites secatus
Cerebropollenites macroverrucosus
Cerebropollenites thiergartii
Classopollis meyerianus
Classopollis murphyae
Classopollis torosus
Duplicisporites granulatus
Enzonalasporites vigens
Granuloperculatipollis rudis
Heliosporites reissingeri
"Lueckisporites" cf. L. singhii
Lunatisporites rhaeticus
Paracirculina quadruplicis
Partitisporites novimundanus
Patinasporites densus
Pinuspollenites minimus
Porcellispora longdonensis
Pseudoenzonalasporites summus
Retitriletes semimuris
Rhaetipollis germanicus
Ricciisporites tuberculatus
Samaropollenites speciosus
Trachysporites fuscus
Tsugaepollenites pseudomassulae
Vallasporites ignacii
Carnianearly late
Norianearly late
Ladin. Rhaetianlate early
Hett.
Fig. 2. Stratigraphic ranges of selected sporomorphs of the northern hemisphere.
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In the European domain, the Carnian stage,including the Ladinian–Carnian boundary, waspalynologically defined on the base of the occur-rence of Ovalipollis pseudoalatus in assemblageswith numerous distinctive species such as Came-rosporites secatus, Duplicisporites granulatus,Ellipsovelatisporites plicatus, Enzonalasporitesvigens, Infernopollenites spp. and Triadispora spp.(phase 1 of Schuurman, 1977, 1979) (Figs 2–4).
In Central Europe (Poland), Orłowka-Zvolinska(1983; 1985) distinguished five ammonoid-controlled palynomorph assemblage Zones cover-ing the Upper Triassic. The oldest Heliosaccusdimorphus assemblage can be correlated with thelate Ladinian phases of western Europe, succeededstratigraphically by the Carnian Ovalipollis–Tria-dispora and Toroisporis–Camarozonosporites–Aulisporites assemblages and by the uppermostCarnian–Norian (which includes also the Rhaetian)Classopollis assemblage, marked by the presence ofRiccisporites at the top.
Roghi (2004) provided a palynological charac-terization of the Julian–Tuvalian boundary in anammonoid-controlled section of the Julian Alps(Italy) based on the FAD of significant sporomorphs(e.g. Duplicisporites continuus, Pseudoenzonalas-porites summus in the Austrotrachyceras austria-cum Zone, Ricciisporites tuberculatus in theTropites dilleri Zone and Granuloperculatipollisrudis in the Tropites subbullatus Zone), within theDuplicisporites continuus assemblage spanning theJulian–Tuvalian interval (Figs 2–4).
The Carnian successions from the CanadianArctic (Fisher & Bujak 1975; Fisher 1979) and thewestern Barents Sea (Hochuli et al. 1989; Mørket al. 1992) are comparable with those from conti-nental Europe, except for some differences (War-rington 2002). The Circumpolles (e.g. Classopollisgroup) are rarer in the Barents Sea than in Europe,and some taxa such as Camerosporites and Duplicis-porites seem to have their LAD in the lower Carnian.Both differences could reflect a colder climate athigher latitudes or, alternatively, poorer independentstratigraphic control (Warrington 2002).
In North America, palynologically-dated non-marine Upper Triassic successions crop out in thesouthwestern and eastern USA, for example, theChinle Formation in Arizona and New Mexico(Dunay & Fisher 1979; Litwin et al. 1991; Cornet1993), the Dockum Group in Texas (Fisher &Dunay 1984; Cornet 1993) and the ChathamGroup of North Carolina (Cornet & Olsen 1985;Litwin & Ash 1993). Litwin et al. (1991) proposedthree miospore assemblage Zones for the ChinleFormation in the southwestern USA (Fig. 5) corre-lated with analogous assemblages of the NewarkSupergroup in the eastern USA (Dunay & Fisher1974; Cornet & Traverse 1975; Manspeizer &
Cousminer 1988; Robbins et al. 1988). Zone I,marked by abundant Brodispora striata, Equisetos-porites chinleanus, Lagenella martini, Minutosac-cus crenulatus, Samaropollenites speciosus and theLAD of Lunatisporites aff. L. noviaulensis, wasassigned to the late Carnian. Zone II contains theFADs of Camarozonosporites rudis, Enzonalaspor-ites vigens, Heliosaccus dimorphus, Ovalipollisovalis, Pseudoenzonalasporites summus and othertaxa (e.g. Alisporites spp., Cycadopites stonei,Guthoerlisporites cancellosus), and the LADs ofBrodispora striata, Camerosporites secatus, Equise-tosporites chinleanus and Lagenella martini. Thisassemblage, which has close similarities with Euro-pean ones, was dated as Tuvalian. Zone III, foundedon the FADs of Camerosporites verrucosus and Kyr-tomisporis spp., was dated as early Norian because ofthe absence of significant Carnian taxa (e.g. Brodis-pora striata and Camerosporites secatus) and by thepresence of Pseudoenzonalasporites summus.
The Camerosporites secatus phase
Starting from the definition of Schuurman’s phase I,Visscher & Krystyn (1978) introduced the ‘Camer-osporites secatus phase’, characterized by a rapiddiversification of the Circumpolloid genera (e.g.Camerosporites secatus, Duplicisporites granulatusand Praecirculina granifer), associated with a groupof monosaccate spores (Enzonalasporites vigens,Patinasporites densus, Pseudoenzonalasporitessummus and Vallasporites ignacii) and with thebisaccate Samaropollenites speciosus. This phasewas introduced to provide a practical palynologicaltool for world-wide correlation because its charac-teristic assemblage was recorded in many palaeo-floristic provinces: in the Alpine Triassic ofEurope (Visscher & Brugman 1981), Italy (Vander Eem 1983; Broglio Loriga et al. 1999; Roghi2004; Mietto et al. 2007), southern Albania (Cirilli& Montanari 1994), Israel (Cirilli & Eshet 1991),Africa (Bourmouche et al. 1996), Tunisia (Mehdiet al. 2009), USA (Fisher & Dunay 1984; Litwin& Ash 1993), Arctic Canada (Fisher & Bujak1975), Australia (Dolby & Balme 1976) andTimor (Martini et al. 2000). Some of the aboverecords come from independently dated (ammonoidand/or conodonts) sections, for example, in Sicily,within the Tropites subbullatus Zone or the‘Anatropites-Bereich’, dated as Tuvalian (Visscher& Krystyn 1978); in the Western Dolomites (Italy)within the Daxatina canadensis Subzone (Vander Eem 1983; Broglio Loriga et al. 1999; Miettoet al. 2007) (Figs 3–4); in Austria from theTrachyceras aon Zone (Cordevolian), the Halobiarugosa/Carnites floridus Zone (Julian) and theTropites subbullatus Zone (Tuvalian) (Dunay &Fisher 1978).
PALYNOLOGY AND PALYNOSTRATIGRAPHY 289
Fig. 3. Schematic comparison of the main palynozonations across the Late Triassic and Early Jurassic of thenorthern hemisphere. Data from: (1) Kurschner et al. 2007; (2) Krystyn et al. 2007a, b; (3) Broglio Loriga et al.1999; Mietto et al. 2007; (4) Roghi 2004; (5) Van der Eem 1983; (6) Schuurman 1977, 1979; (7) Morbey 1975, 1978;(8) Lund 1977; (9) Orbell 1973. FO: first occurrence; LO: last occurrence. Z (Zone); Sz (Subzone); Palynozones from
S. CIRILLI290
The Camerosporites secatus phase was originallyconsidered as an exclusively Carnian palynologicalevent and later extended into the Ladinian (Visscher& Brugman 1981; Van der Eem 1983; Besems 1981,1983; Brugman et al. 1994). However, the long rangedistribution of this phase (upper Ladinian to Carnian)made it unsuitable for a detailed biostratigraphicalsubdivision as many authors introduced furthersubdivisions. As already suggested by Visscher &Krystyn (1978), an upper Ladinian trend may be dis-cerned from the occurrence of Camerosporitessecatus with Echinitosporites iliacoides, Retisulcitesperforates, Lunatisporites spp., and, at least in theAlpine–Mediterranean part of Europe, with the‘northern’ element Staurosaccites quadrifidus,while an exclusively Carnian age may be definedon the gradual diversification of circumsulcateform genera and monosaccates belonging to theEnzonalasporites–Patinasporites–Pseudoenzona-lasporites–Vallasporites group (Warrington 2002).The FAD of Patinasporites densus marks the base
of the early Carnian (cf. the vigens–densus phaseof Van der Eem 1983) (Figs 2–3), which is com-monly associated with the first appearance ofVallasporites ignacii at the base of the Carnian(Visscher & Brugman 1981; Van der Eem 1983;Fisher & Dunay 1984; Blendinger 1988; Hochuliet al 1989; Broglio Loriga et al. 1999; Hochuli &Frank 2000; Warrington 2002; Roghi 2004; Miettoet al. 2007), while the appearance of Paracirculinaquadruplicis and Samaropollenites speciosus dif-ferentiates the youngest Carnian (Tuvalian) Camer-osporites secatus phase (Visscher & Krystyn 1978;Visscher et al. 1980; Cirilli & Eshet 1991; Cirilli &Montanari 1994; Warrington 2002).
The palynological Ladinian–Carnian
boundary
In Italy (Prati di Stuores section, western Dolo-mites), Van Der Eem (1983) established three
Fig. 4. Main palynozones and palynological assemblages across the Ladinian–Carnian boundary. Z, Zone; Sz,Subzone; FO, first occurrence; LO, last occurrence.
PALYNOLOGY AND PALYNOSTRATIGRAPHY 291
ammonoid-controlled palynomorph assemblagesacross the Ladinian–Carnian transition (Figs 3–4),which in ascending order are: the secatus–dimor-phus and secatus-vigens phases (late Ladinian,Longobardian), the vigens–densus phase (earlyCarnian) and the densus-maljawkinae phase(Julian). The same section (Prati di Stuores) is acandidate GSSP section for the Ladinian–Carnianboundary, herein based on the FAD of the ammoniteDaxatina canadensis (Broglio Loriga 1999; Miettoet al. 2007). The miospore assemblages for thissection, correlated with ammonites and other fossilgroups, show a compositional variation across theboundary distinguished by the FAD of Concentri-cisporites cf. C. bianulatus, Duplicisporites granu-latus, Enzonalasporites vigens, Gordonisporafossulata, Kyrtomisporis ervii, and ‘Lueckisporites’cf. singhii in the upper part of the regoledanusSubzone, and by the LAD of Nevesisporites vallatusat the putative boundary, with additional elementshaving a wider Ladinian-early Carnian range (e.g.Reticulatisporites dolomiticus, Partitisporites novi-mundanus). Aulisporites cf. A. astigmosus, Camer-osporites secatus, Duplicisporites verrucosus,Patinasporites densus and Vallasporites ignacii,have their first occurrence in the lower part of thecanadensis Subzone, while Camerosporites pseudo-verrucatus, Samaropollenites speciosus and Wey-landites magmus first occur in the uppermost partof this Subzone. According to this new palyno-morph zonation, the base of the vigens-densusphase in the Prati di Stuores/Stuores Wiesensection is lowered by at least 130 metres withrespect to Van der Eem (1983).
The Norian–Rhaetian miospore
assemblages in the northern hemisphere
In the Boreal realm, the Norian and Rhaetian paly-nomorph assemblages are documented in severalsections, some of which, especially those of Rhae-tian age, are independently dated: the SvalbardArchipelago, Arctic Canada, North Sea andBarents Sea (Fisher & Bujak 1975; Lund 1977,2003; Fisher 1979; Pedersen & Lund 1980;Guy-Ohlson 1981; Smith 1982; Hochuli et al.1989; Dybkjær 1991; Batten & Koppelhus 2002;Koppelhus & Batten 2002; Lindstrom & Erlstrom2006). They share common taxa such as Camarozo-nosporites spp., Chasmatosporites spp., Cingulizo-nates rhaeticus, Kyrtomisporis spp., Ovalipollisspp., Rhaetipollis germanicus, Triancoraesporitesreticulatus, Zebrasporites spp., and Granulopercu-latipollis rudis, Limbosporites lundbladii, Quadrae-culina anellaeformis, Ricciisporites tuberculatus,the latter four taxa dominating the Rhaetianassemblages (Fig. 2). Lund (1977) established four
palynozones for the Danish Basin across the LateTriassic and Early Jurassic (Fig. 3), although thereare some local variations in the relative abundancesof taxa (Dybkjær 1988, 1991): (1) the Norian–earlyRhaetian Classopollis–Enzonalasporites and theearly Rhaetian Ricciisporites–ConbaculatisporitesZones, characterized by the presence of Ricci-isporites tuberculatus and by the absence of olderforms such as Enzonalasporites vigens; (2) themiddle Rhaetian Rhaetipollis–Limbosporites Zone,marked by the occurrence of Limbosporites lund-bladii and Rhaetipollis germanicus; (3) the lateRhaetian Ricciisporites–Polypodiisporites Zone,defined by the occurrence of Polypodiisporitespolymicroforatus (¼ Convolutispora microrugu-lata) and Semiretisporis spp., rare Rhaetipollis ger-manicus and abundant trilete ornamented spores.This last Zone correlates with the Classopollis–Ricciisporites Zone of Dybkjær (1991), introducedas an alternative Zone because of the sporadicoccurrence of Polypodiisporites polymicroforatusand Semiretisporis in the Danish Basin; and (4)the Hettangian Pinuspollenites–TrachysporitesZone is marked by the common presence ofPinuspollenites minimus, by trilete spores such asConbaculatisporites spp., Trachysporites spp. andUvaesporites reissingerii and by the sporadic pres-ence of Cingulizonates rhaeticus and Limbosporiteslundbladii.
Koppelhus & Batten (2002) proposed a detailedpalynozonation across the Late Triassic-Early Juras-sic valid for northwestern Europe. This includes, inascending order: (1) The Norian–early RhaetianClassopollis–Enzonalasporites Zone subdividedinto the Norian–early Rhaetian Classopollis–Porcellispora Subzone and the early RhaetianGranuloperculatipollis rudis and Enzonalaspor-ites–Conbaculatisporites Subzones; (2) the early tomiddle Rhaetian Ricciisporites–Conbaculatispor-ites Zone; (3) the middle Rhaetian Rhaetipollis–Limbosporites Zone; (4) the late RhaetianRicciisporites–Polypodiisporites (Lund 1977) andClassopollis–Ricciisporites Zones (Dybkjær 1991);and (5) the Hettangian Pinuspollenites–Trachyspor-ites Zone, below the FAD of Cerebropollenitesmacroverrucosus, which is herein considered asSinemurian–Pliensbachian (cf. Cerebropollenitesmacroverrucosus Zone of Dybkjær 1991). Themiddle to late Rhaetian assemblages from theBoreal realm are overall typified by the occurrenceand occasional abundance of some significant taxasuch as Cingulizonates rhaeticus, Classopollis spp.,Kraeuselisporites (¼ Heliosporites) reissingeri,Limbosporites lundbladii, Lunatisporites rhaeticus,Ovalipollis spp., Rhaetipollis germanicus, Triancor-aesporites reticulates and Semiretisporis gothae.
Homogeneous Norian–Rhaetian miosporeassemblages have been recorded in several regions
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of Europe, for example, England, Germanic andAlpine domains (e.g. Morbey & Dunay 1978;Schuurman 1979; Visscher et al. 1980; Visscher &Brugman 1981; Smith 1982; Brenner 1986; Batten& Koppelhus 2002; Warrington 2002; Hounslowet al. 2004), Circum-Mediterranean area (e.g.Brugman & Visscher 1988; Conway et al. 1990;Cirilli et al. 1993, 1994; Barron et al. 2006;Gomez et al. 2007) and Western Ciscaucasia (e.g.Yaroshenko 2007) and documented from indepen-dently dated sections in Europe (e.g. NorthernCalcareous Alps, Austria: Morbey & Neves 1974;Morbey 1975; Mostler et al. 1978; Krystyn &Kurschner 2005; Krystyn et al. 2007a, b; Kurschneret al. 2007; Bonis et al. 2009). However, the Norianpalynological assemblages still need to be indepen-dently dated because of the scarcity of Norian con-tinental successions that might be correlated tocoeval marine beds. At present, two candidateGSSP sections have been proposed for the base ofthe Norian: Black Bear Ridge (Williston Lake,B.C. Canada) (Orchard 2007) and Pizzo Mondello(Sicily, Italy) (Nicora et al. 2007) but both lackpalynological assemblages to correlate withother proxies.
The available palynozonations show that, at leastin Europe, the Norian palynomorph assemblagesreveal transitional features, characterised mainlyby the progressive LADs of Carnian taxa and bythe progressive appearance of species that becomedominant in the Rhaetian assemblages. The mostsignificant Carnian forms that still persist in the
early Norian are Camerosporites spp., Duplicispor-ites spp., Ellipsovelatisporites rugosus, Enzonalas-porites spp., Paracirculina quadruplicis spp.,Patinasporites densus, Pseudoenzonalasporitessummus and Vallasporites ignacii, together withsome bisaccates such as Triadispora spp. and Infer-nopollenites parvus. Most of these taxa (e.g.Enzonalasporites vigens, Patinasporites densus,Vallasporites ignacii) seem to disappear in theupper part of the Norian or, locally, in the earlyRhaetian (Morbey 1975; Schuurman 1977, 1979;Visscher et al. 1980; Fisher & Dunay 1981;Visscher & Brugman 1981) (Fig. 2). The gradualevolution from Norian to Rhaetian assemblagesis underlined by the progressive dominance ofgenera belonging to the Circumpolles Group(Classopollis, Geopollis, Gliscopollis, Granu-loperculatipollis) in association with Ovalipollis,Rhaetipollis and Ricciisporites. A suite of additionaldiversified forms that also may continue in theRhaetian are Acanthotriletes varius, Camarozonos-porites spp., Chasmatosporites spp., Densosporitesfissus, Heliosporites reissingeri, Kyrtomisporisspp., Limbosporites lundbladii, Lycopodiaciditesrugulatus, Perinosporites thuringiacus, Quadraecu-lina anellaeformis, Triancoraesporites spp., Uvae-sporites argenteaeformis and Zebrasporites spp.(the TL, LR and TK Zones of Morbey 1975; phaseIII and IV of Schuurman 1977) (Figs 3, 6). Thelate Rhaetian assemblages are marked by the FADof Retitriletes semimuris (phase IV of Schuurman1977; Orbell 1973; Visscher et al. 1980; Visscher
Fig. 5. Miospore assemblage zones proposed for the Chinle Formation in the southwestern USA (modified by Litwinet al. 1991). FO, first occurrence; LO, last occurrence.
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& Brugman 1981) and by the LADs of a group oftaxa, such as Granuloperculatipollis rudis, Luna-tisporites rhaeticus, Ovalipollis pseudoalatus,Protohaploxypinus sp. cf. P. microcorpus, Tsugae-pollenites pseudomassulae and of a group of triletespores. Tsugaepollenites pseudomassulae is acommon form in the European Rhaetian assem-blages (Fig. 2), because it has been frequentlyrecorded in England (Hounslow et al. 2004; War-rington et al. 2008), Austria (Morbey & Neves1974; Hillebrandt et al. 2007; Kurschner et al.2007; Bonis et al. 2009a), southern France (Pey-bernes et al. 1988; Frechengues et al. 1993), Italy(Cirilli et al. 1994; Jadoul et al. 1994; Galli et al.2007), Spain (Baudelot & Taugourdeau–Lantz1986; Vachard et al. 1990; Calvet et al. 1993; Fre-chengues et al. 1993; Barron et al. 2006) and,more seldom, in Norian strata (e.g. Mallorca inBoutet et al. 1982).
Recent palynological data from key Austriansections, such as the Steinbergkogel section nearHallstatt, and the Kuhjoch section near Hinter-riss in Tyrol, which have been proposed as candidateGSSPs for the Norian–Rhaetian boundary (Krystyn& Kurschner 2005; Krystyn et al. 2007a, b) andthe Rhaetian–Hettangian boundary (Kurschneret al. 2007; Hillebrandt et al. 2007), respectively,fit well with the TL and LJ concurrent-range palyno-zones established by Morbey & Neves (1974) withinthe Karparthian facies of Kossen beds in the Ken-delbach section (Figs 3, 6, 7). The integration of
palynological, ammonoid, conodont and isotopedata from the candidate GSSP section (Krystyn &Kurschner 2005; Krystyn et al. 2007a, b) confirmsthe transitional nature of the Norian–Rhaetianpalynological boundary. The authors distinguishedtwo palynological Zones, the lowermost containingtypical older ‘Carnian’ elements (Ellipsovelatis-porites rugosus, Enzonalasporites vigens, Partitis-porites spp., Patinasporites toralis, Triadisporaspp., Vallasporites ignacii), and the upper typifiedby the FADs of taxa such as Chasmatosporitessp., Limbosporites lundbladii and Quadraeculinaanellaeformis. Additional taxa such as Classopollismeyerianus, Classopollis torosus, Granulopercula-tipollis rudis, Ovalipollis pseudoalatus, Rhaetipollisgermanicus, Ricciisporites tuberculatus, and Tsu-gaepollenites pseudomassulae, range throughoutthe two palynological Zones. An acme of dinoflagel-late cysts (Heibergella, Noricysta, Rhaetogonyau-lax) has been also recorded at the transition of thetwo palynozones, which also approximately coincidewith the Cycloceltites–Vandaites ammonoid event(Figs 3 & 6).
On the whole, the uppermost Triassic assem-blages of the Boreal realm (e.g. Lund 1977; Pedersen& Lund 1980; Hochuli et al. 1989; Dybkjær 1991;Batten & Koppelhus 2002; Lindstrom & Erlstrom2006) show affinity with those found in continentalEurope with some exceptions, such as Limbosporiteslundbladii, which is more common in the northernassemblages, Quadraeculina anellaeformis, which
Fig. 6. Main palynozones and palynological assemblages across the Norian–Rhaetian boundary. Z, Zone; Sz, Subzone;FO, first occurrence; LO, last occurrence. Palynozones from Morbey 1975: TR, Classopollis torosus–Granuloperculatipollis rudis; TL, Riccisporites tuberculatus–Hystrichosphaeridium langi; RG, Rhaetogonyaulaxrhaetica–Rhaetipollis germanicus; Rk, Rhombodella kendelbachia; LL, Carnisporites lecythus–Zebrasporiteslaevigatus.
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seems to appear earlier in the Boreal realm than inthe European domain, and the Circumpolles group(Classopollis spp.), which has a wider distributionin central and southern Europe (Hochuli et al.1989). Other taxa, such as Tsugaepollenites pseudo-massulae, are not common in the Boreal realm,except for the British Rhaetian (Orbell 1973;Warrington et al. 2008).
In contrast, the composition of Norian and Rhae-tian miospore assemblages from the Middle East(e.g. east–central Iran) differs considerably fromthat of the representative European assemblages.The Norian microflora shows more affinity withthe Gondwanan (e.g. Eastern Australia and NewZealand), while the Rhaetian assemblages, althoughmaintaining a strong affinity with the microflora ofthe southern hemisphere, show a gradual enrichmentin cosmopolitan species found in both the northernand southern hemispheres (Cirilli et al. 2005).
The Rhaetipollis germanicus assemblage zone
A considerable number of palynological studies inthe Rhaetian reference sections of the Boreal andEuropean realms point to Rhaetipollis germanicusas a valuable biostratigraphical marker to typifythe Rhaetian stage (Figs 2–3) (e.g. Rhaetipollisgermanicus assemblage Zone of Orbell 1973;
However, the FAD of Rhaetipollis germanicus isyet uncertain, although the ammonoid-dated Norianpalynomorph assemblages of Svalbard providedconstraints for a Norian first appearance of thistaxon (Smith 1982). Recent data from Austria docu-ments its lowest entrance in the middle Rhaetian,although the entry level of Rhaetipollis germanicusremains unknown ‘because of the carbonaceousnature (facies) of the deposits. . .underlying theRhaetian deposits (Kossen or Zlambach beds)’(Krystyn et al. 2007a, p.193; Krystyn et al. 2007b,p.165). At the St. Audrie’s Bay section, Rhaetipollisgermanicus makes its first occurrence in the LateAlaunian to Early Sevatian Twyning Mudstone Fm.(Hounslow et al. 2004); in the German Keuper itslowermost occurrence has been recorded in thelower Rhaetian (¼ Sevatian 2 of the Alpinedomain) (Lund 1977, 2003; Schulz & Heunish 2005).
The LAD of Rhaetipollis germanicus has beendated as late Rhaetian (Orbell 1973; Achilles1981; Brenner 1986) as shown in Austria wherethe Rhaetipollis germanicus Assemblage Zoneranges into the upper part of the pre-planorbisbeds (Morbey 1975; Kurschner et al. 2007),although it was also rarely recorded in the oldest
Fig. 7. Schematic correlation of two GSSP candidate sections based on palynological and geochemical proxies acrossthe Triassic–Jurassic boundary. LAD, Last appearance datum. Data from: Hounslow et al. 2004; Hillebrandt et al. 2007;Kurschner et al. 2007; Warrington et al. 2008; FO, first occurrence; LO, last occurrence.
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Jurassic strata (Warrington 2002; Traverse 2008)(Figs 2, 3, 6). Phase III of Schuurman (1979),regarded as transitional between phase II (Norian)and phase IV (late Rhaetian), registers the rapiddecrease and disappearance of Rhaetipollis germa-nicus. In the Southern Alps (Italy), Rhaetipollis ger-manicus disappears just before the blooming ofKrauselisporites reissingeri (Galli et al. 2007).Also, in Northern Europe (North Sea basin, EastGreenland and Scania) the last occurrence ofRhaetipollis germanicus has been dated as ‘middleRhaetian’ (Lund 1977; Pedersen & Lund 1980;Guy-Ohlson 1981). The diachronous first and lastoccurrences of this taxon, as well as those of othertaxa, may be related to the climate-controlledprovincialsm of the parent flora (cf. palynologicalsection in Krystyn et al. 2007a, b).
The palynological Triassic–Jurassic
boundary
One of the main obstacles to describing and con-straining the age of the palynomorph assemblagesat the Tr–J boundary and at the base of Hettangianis the lack, until now, of a formally accepted GSSPsection for this time interval. Many sections havebeen proposed during the last decade, some ofthem including palynostratigraphic zonations;examples include the St. Audrie’s Bay–DonifordBay section, England (Warrington et al. 1994,2008) and the Kuhijoch section, in the Tyrol ofAustria (Hillebrandt et al. 2007). Recently,members of the International Subcommission onTriassic Stratigraphy (ISTS) and the InternationalSubcommission on Jurassic Stratigraphy (ISJS)voted to recommend the Kuhjoch section, Karwen-del Mountains, northern Calcareous Alps (Austria)(Hillebrandt et al. 2007) as the GSSP section andthe Ferguson Hill Section, New York Canyon,Nevada (USA) as the ASSP (Guex et al. 2006;Lucas et al. 2007). The two proposals were formallyapproved by the ISJS in August 2008 (during the33rd International Geological Meeting in Oslo)and are presently waiting for the approval of theInternational Commission on Stratigraphy (ICS).In the proposed GSSP section the Tr–J boundarylies 5.8 metres above the top of the KossenFormation, within the level bearing the FO ofPsiloceras cf. P. spelae, which Guex consideredas the primary marker defining this boundary. Themicrofloral record across the boundary, which ischronologically constrained by carbon-isotope stra-tigraphy and marine (ammonoid and conodont)biostratigraphy, is distinguished by a few notablepalynostratigraphic events that correlate with thoseof the Tiefengraben section in the eastern part ofthe Eiberg basin (Kurschner et al. 2007) (Figs 2,3, 6). Four palynomorph assemblage Zones have
been proposed. These are, in ascending order: (1)the Rhaetipollis–Limbosporites Zone (RL Zone),from the uppermost part of the Kossen Formation,is dominated by Classopollis meyerianus and minorClassopollis torosus, with Limbosporites lundbladii,Ovalipollis pseudoalatus, Rhaetipollis germanicusand Ricciisporites tuberculatus. Marine palyno-morphs such as dinoflagellate cysts (Dapcodiniumpriscum, Rhaetogonyaulax rhaetica), and acritarchs(e.g. Cymatiosphaera polypartita, Micrhystridiumsp.) are also present; (2) The Rhaetipollis–Porcel-lispora Zone (RPo Zone) from the SchattwaldBeds is distinguished by the increase in Calamos-pora tener, Classopollis torosus, Classopollismeyerianus, Convolutispora microrugulata andDeltoidospora spp. in the lower–middle part, bythe decrease in Classopollis spp. and the disappear-ance of Rhaetipollis germanicus in the upper part(lower part of the Schattwald Beds); (3) TheTrachysporites-Porcellispora (TPo Zone) occursbetween the Schattwald Beds and the beds con-taining Psiloceras cf. P. spelae, where Ovalipollispseudoalatus has its highest appearance, Classo-pollis torosus temporarily disappears, and Helios-porites reissingeri, Porcellispora longdonensisand Trachysporites fuscus are abundant; (4) TheTrachysporites–Heliosporites (TH Zone) liesabove the FO of Psiloceras cf. P. spelae, where Clas-sopollis torosus reappears, Cerebropollenites thier-gartii makes its first occurrence, and Heliosporitesreissingeri and Porcellispora longdonensis becomedominant in the lower-middle part and declinein the upper part; and (5) The Trachysporites–Pinuspollenites Zone (TPi Zone), Jurassic in age,is identified by the appearance of Pinuspollenitesminimus associated with abundant Classopollisspp., Heliosporites reissingeri and Trachysporitesfuscus. It is noteworthy that the first occurrence ofCerebropollenites thiergartii is close to the Psilo-ceras cf. P. spelae horizon, within the lower part ofthe main negative d13C excursion (Kurschner et al.2007; Ruhl et al. 2009). On the basis of these con-straints and by correlation with other previouslyestablished palynological Zones, Cerebropollenitesthiergartii, at least in the European domain, seemsto be the first and the only morphologically distinct,post-Triassic species suitable for interregional corre-lations between terrestrial and marine sections(Kurschner et al. 2007).
Recent data from additional key sections of theEiberg Basin provided a detailed palynologicalzonation (Bonis et al. 2009) which is equivalent,except for minor differences, with the palynomorphassemblage zones from the Tiefengraben. The zona-tion scheme for the Eiberg Basin correlates wellwith those of Morbey (1975) and Schuurman(1977, 1979) for the Alpine realm (Figs 3, 6). TheRhaetipollis–Limbosporites Zone correlates to the
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MI Subzone of Morbey (1975) and to phase IIIof Schuurman (1977, 1979), the assemblages ofwhich contain Classopollis spp., Cingulizonatesrhaeticus, Limbosporites lundbladii, Ovalipollispseudoalatus, Rhaetipollis germanicus and Ricciis-porites tuberculatus. The Rhaetipollis–Porcellis-pora and the Trachysporites–Porcellispora Zonesare the equivalent of the FG Subzone of Morbey,including the transition from the Rhaetian to theHettangian Stage, defined by the LAD of Densos-porites fissus, Retitriletes gracilis, Triancorae-sporites ancorae, and by abundant Classopollistorosus, Heliosporites reissingeri and Ricciisporitestuberculatus. The Rhaetipollis–Porcellispora andthe Trachysporites–Porcellispora Zones corre-spond to phase IV of Schuurman (1977, 1979),which is distinguished by the rapid disappearanceof Ovalipollis, Rhaetipollis and taeniate bisaccatepollen and by abundant Heliosporites reissingeri.The Trachysporites–Heliosporites and the Trachy-sporites–Pinuspollenites Zones fit Jurassic phaseV of Schuurman (1979), dominated by Classopollisspp. and distinguished by the lack of typical Rhae-tian forms except for some taxa (e.g. Heliosporites,Retitriletes).
The established carbon-isotope curve, as well aspalynostratigraphic and other biostratigraphicalevents of the Austrian section, also correlate withthe section at St Audrie’s Bay on the Somersetcoast of England (Fig. 7), previously a GSSP candi-date section (Warrington et al. 1994, 2008) andothers nearby in Somerset and South Wales(Orbell 1973; Warrington 1974; Hounslow et al.2004), where the Tr–J boundary has been conven-tionally placed at the lowest occurrence of theHettangian ammonite Psiloceras planorbis. Palyno-logical data integrated with other fossils and proxiessuch as magnetostratigraphy and chemostratigraphy(Warrington 2002; Hounslow et al. 2004; Warring-ton et al. 2008), show miospore assemblages fromthe Triassic Penarth Group dominated by Circum-polles (Classopollis, Geopollis, Gliscopollis, Gran-uloperculatipollis), Ricciisporites tuberculatus, andminor Ovalipollis pseudoalatus, Rhaetipollis ger-manicus, which have their LADs at the top mostof the Penarth Group, below the FAD of Psiloceras.Ricciisporites tuberculatus ranges into the basal partof the Lias Group, and Kraeuselisporites reissingeriand Quadraeculina anellaeformis range from thePenarth Group into beds that contain Hettangianammonites, while the Triassic-type palynomorphsrapidly disappear. There are no other significantfirst occurrences that can be specifically correlatedwith the Tr–J boundary (Warrington 1974, 2002),except for Cerebropollenites thiergartii and Cere-bropollenites macroverrucosus, whose occurrencesin the Somerset succession are in the HettangianPsiloceras planorbis beds, as in other European
sections (Clement-Westerhof et al. 1974; Morbey1975, 1978; Van Erve 1977; Visscher et al. 1980;Fisher & Dunay 1981; Kurschner et al. 2007).
Additionally, the palynological Zones of theAustrian section match those of central and southernEuropean sections where similar assemblages havebeen recorded across the Tr–J boundary in Spain(Baudelot & Taugourdeau-Lantz 1986; Vachardet al. 1990; Calvet et al. 1993; Frechengues et al.1993; Barron et al. 2006; Gomez et al. 2007), inthe Lusitanian Basin and France (Adloff et al.1974; Adloff & Doubinguer 1982; Grignac &Taugourdeau-Lantz 1982; Peybernes et al. 1988;Frechengues et al. 1993), in Italy (Cirilli et al.1993, 1994; Galli et al. 2007), and in the WesternCarpathians and Hungary (Gotz et al. 2009;Ruckwied & Gotz 2009).
The Austrian Zones also correlate well withthe palynozonations of Northern Europe andScandinavia (Lund 1977, 2003) (Fig. 3). The TriassicRhaetipollis–Limbosporites and Rhaetipollis–Porcellispora Zones compare with the Rhaetipol-lis–Limbosporites Zone of Lund (1977); the Trachy-sporites–Porcellispora Zone shows a relationshipwith the Ricciisporites–Polypodiisporites Zone ofLund (1977), primarily in the absence of Rhaetipollisand the abundance of Convolutispora microrugulata(¼ Polypodiisporites polymicroforatus in Lund1977). The Early Jurassic Trachysporites–Helios-porites and Trachysporites–Pinuspollenites Zonescompare with the Pinuspollenites–TrachysporitesZone of Lund (1977).
In most of the European domain, the Tr–Jboundary seems to be characterized by only aminor extinction/turnover of the terrestrial macro-flora and microflora against a background of moregradual change (e.g. Orbell 1973; Schuurman1979; Pedersen & Lund 1980; Fisher & Dunay1981; Knoll 1984; Ash 1986; Hallam & Wignall1997; Hallam 2002; Tanner et al. 2004; Lucas &Tanner 2004, 2007, 2008; Kurschner et al. 2007).A floral change has been documented at the Tr–Jboundary in northwestern Europe, for example,Greenland (McElwain et al. 1999, 2007; McElwain& Punyasena 2007), Germany and Sweden (Van deSchootbrugge et al. 2009). In the latter two sites Vande Schootbrugge et al. (2009) record a floral changeat the base of the Triletes Beds, where typicalRhaetian pollen, such as Ricciisporites tuberculatusand Rhaetipollis germanicus, decrease until theydisappear. They are replaced in the earliest Hettan-gian by assemblages dominated by pollen fromCheirolepidiaceae and Taxodiaceae conifers(e.g. Classopollis spp., Perinopollenites elatoides,Pinuspollenites minimus) and bisaccates fromcorystosperm seed ferns (e.g. Alisporites spp.).accompanied by an acme of lycopodiophyte spores(e.g. Kraeuselisporites reissingeri).
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In the Danish part of the Danish Basin, the Tr–Jtransition is associated with an abundance of thecheirolepidacean conifer pollen Classopollis spp.,and follows a mass-occurrence of dinoflagellatecysts (Dybkjær 1988, 1991; Lindstrom & Erlstrom2006). Variable abundances of Classopollis-typepollen, commonly associated with hot and/or aridclimate conditions (Vakhrameev 1991; McElwainet al. 2007), are documented in numerous sectionsspanning the Tr–J transition, such as: GreatBritain (Warrington & Ivimey-Cook 1990; Houn-slow et al. 2004), central and southern Europe (seeKurschner et al. 2007 for updated references),north Africa (Marzoli et al. 2004), and also in thesouthern hemisphere, such as in Australia (Back-house & Balme 2002; Grice et al. 2005), Timor,Indonesia (Martini et al. 2000, 2004) and Antarctica(Shu et al. 2000).
In the Eastern North America basins (i.e. NewarkSupergroup basins) the palynological Tr–J bound-ary has been placed since the 1970s in the strataimmediately beneath the lowest CAMP basaltflow (e.g. the Orange Mt. Basalt in the Newarkbasin), to coincide with a palynological turnoverevent that is characterized by a low diversity palyno-morph assemblage recording a significant loss ofLate Triassic taxa (Fowell & Olsen 1993; Fowellet al. 1994; Fowell & Traverse 1995; Olsen et al.2002; Whiteside et al. 2007). This rapid turnoverhas been primarly defined on the LADs of Ovalipol-lis ovalis, Patinasporites densus, Vallasporitesignacii, on the increase in the Classopollis typepollens such as Classopollis meyerianus, Classopol-lis murphyae, Classopollis torosus, and on a bloom-ing of trilete fern spores (fern spike) (e.g. Cornet &Traverse 1975; Cornet & Olsen 1985; Fowell &Olsen 1993; Fowell et al. 1994; Fowell & Traverse1995; Whiteside et al. 2007). Additionally, thisputative Tr–J boundary is in close stratigraphicproximity to a magnetic reversal (chron E23r:Olsen et al. 2002), a negative carbon isotope excur-sion (Olsen et al. 2002), and is also characterized (inthe Newark and Fundy basins) by a moderate Iranomaly (Olsen et al. 2002). However, the abovepalynostratigraphic criteria for placing the Tr–Jboundary have been strongly critiqued (Gradsteinet al. 1994; van Veen 1995; Lucas & Tanner 2007,2008; Cirilli et al. 2009), primarly because theLAD of Patinasporites densus in the northern andsouthern hemispheres is documented as lateNorian to early Rhaetian (Sevatian–Rhaetianboundary). The large distribution of this taxon indifferent palaeoclimate belts (Buratti & Cirilli2007) would exclude the possibility of floral pro-vincialism (Fowell & Olsen 1993) controlling adiachronous distribution of the parent plant inNorth American and Tethyan domains. Addition-ally, the increase in the Classopollis type pollens
is, as seen above, a widespread palynological eventspanning the Tr–J transition in the northern andsouthern hemispheres and not only Jurassic in age.Recent data from the Fundy Basin of Nova Scotia,Canada, document the presence of Triassic sporo-morphs in the Scots Bay Member, which overliesthe North Mountain Basalt, and therefore above thelast occurrence of Patinasporites densus (Cirilliet al. 2009). In the Blomidon Formation, at PartridgeIsland section, no significant palynological declineor turnover has been observed below the NorthMountain Basalt that can be attributed to a massextinction event. Furthermore, although a true fernspike is absent at this site, a level enriched infern spores (e.g. Converrucosisporites cameronii,Dictyophyllidites harrisii and Dictyophyllidites sp.)is recorded 15 cm below the putative Tr–J bound-ary, but not above the boundary as in the NewarkBasin (Fowell & Olsen 1993; Fowell et al. 1994;Whiteside et al. 2007), suggesting the relative abun-dance of fern spores may record a short-term episodeof climate change and atmospheric acidification(i.e. related to CAMP eruptions) rather than aglobal event of recolonization after mass exctinction(Cirilli et al. 2009; Van de Schootbrugge et al. 2009).
Upper Triassic miospore assemblages
of the southern hemisphere
As a consequence of Late Triassic provincialism,different palynological zonations, mostly based onintegrated dinoflagellate and spore-pollen palynos-tratigraphy, have been provided for eastern andwestern Australia (Helby et al. 1987; Brenner1992; Backhouse & Balme 2002; Grice et al.2005), Queensland (de Jersey 1975), Victoria Land,Antarctica (Kyle & Schopf 1982), and NewZealand (de Jersey & Raine 1990) (Fig. 8). Theassemblages from New Zealand and easternAustralia, apart from a limited occurrence in theGalilee Basin of Queensland, can be assigned to theIpswich Microflora, whereas the assemblages ofnorthwestern Australia and Timor belong to theOnslow Microflora (Foster et al. 1994; Martiniet al. 2000, 2004; Buratti & Cirilli 2007). Correlationbetween these palynological zonal schemes and thestandard geological timescale has been based onscattered data of marine biota, primarly fora-miniferans, and sparse occurrences of ammonites,conodonts and/or dinoflagellate cysts (Helby et al.1987; Nicoll & Foster 1994, 1998).
Western Australia
In western Australia three important Late Triassicpalynological Zones have been proposed to corre-late with conodont biozones (Dolby & Balme
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1976; Helby et al. 1987; Nicoll & Foster 1998;Backouse & Balme 2002; Grice et al. 2005)(Fig. 8): (1) the mostly Carnian to earliest NorianSamaropollenites speciosus Oppel Zone, markedby a significant decline of Enzonalasporites vigensand by an increase of Falcisporites australis; (2)the Norian Minutosaccus crenulatus Oppel Zone,defined by a considerable decline of the Tethyantaxa such as Enzonalasporites vigens, Patinaspor-ites densus and Samaropollenites speciosus; and(3) the Rhaetian, probably extending into Hettan-gian, Ashmoripollis reducta Oppel Zone, character-ized by the FAD of Zebrasporites, by the decline ofbisaccates such as Falcisporites spp., and by thedominance of Classopollis spp. at the top. Thetransition from the Ashmoripollis reducta to Classo-pollis torosus Zones characterizes the Tr–J bound-ary, defined by the last occurrence of Triassicspores (Backhouse & Balme 2002; Grice et al. 2005).
Eastern Australia and Indian Peninsula
A different Late Triassic palynozonation has beenestablished for eastern Australia based on two longrange Zones (Fig. 8): the Craterisporites rotundusOppel Zone, assigned to the Carnian (de Jersey1975) and recognized by the FAD of Craterisporites
rotundus at its base, by the dominance of Duplexis-porites problematicus (¼ Striatella seebergensis)and Falcisporites australis and by the FAD ofPolycingulatisporites crenulatus at the top; andthe Polycingulatisporites crenulatus Oppel Zone,defined by the FAD of Polycingulatisporites crenu-latus at the base, by a rapid decline in Falcisporitesaustralis, by the LAD of Playfordiaspora velataand by a rapid increase in Classopollis torosus atthe top. This Zone has been variously assigned tothe Rhaetian–Hettangian by de Jersey (1975) andto the Norian extending to Hettangian by Helbyet al. (1987).
The Polycingulatisporites crenulatus Zone ofEastern Australia has been correlated with theArcuatipollenites tethyensis Zone of the IndianPeninsula (Tripathi 2000), where the UpperTriassic–Lower Jurassic assemblages of the Dubra-jpur Formation commonly contain Araucariacites,Callialasporites, Classopollis, Dictyophyllidites,Infernopollenites, Minutosaccus, Nevesisporitesand Staurosaccites.
New Zealand and Tanzania
The palynological zonations of eastern Australiamay be reasonably compared with New Zealand
Fig. 8. Schematic comparison of the main palynozonations across the Late Triassic and Early Jurassic in the southernhemisphere. (1) de Jersey & Raine 1990; (2) de Jersey 1975; de Jersey & Raine 1990; (3) Dolby & Balme 1976; Helbyet al. 1987; de Jersey & Raine 1990; Nicoll & Foster 1998; Backhouse & Balme 2002; Grice et al. 2005.
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miospore zonations (cf. Murihiku Supergroup: deJersey & Raine 1990), considering the distributionof the Late Triassic floral and microfloral provinci-alism (Fig. 1). Four palynological Zones and oneSubzone have been established from the Late Trias-sic (Oretian, Otamitan, Warepan, Otapirian) to theHettangian (Aratauran) interval (de Jersey &Raine 1990), correlated by invertebrate faunaswith the international stratigraphic stages (Fig. 8).In ascending order they are: (1) the long-ranginglate Ladinian to early Norian Annulispora follicu-losa Zone, defined by the FAD of Annulispora folli-culosa and, in its upper part by the FAD ofAnnulispora microannulata, which identifies thelate Carnian–early Norian Annulispora microannu-lata Subzone. This Subzone contains other FADssuch as Alisporites warepanus, Aratrisporites flexi-bilis, Rogalkaisporites spp., Stereisporites antiqua-sporites and Striatella seebergensis in theuppermost part of the Annulispora microannulataSubzone (early Norian); (2) The early to lateNorian Polycingulatisporites crenulatus Zone,marked by the FAD of Polycingulatisporites crenu-latus and by the LAD of Lundbladispora denmeadiand Aratrisporites flexibilis. Polycingulatisporitesmooniensis enters in the lower part of the Zone inthe assemblage with Camarozonosporites rudisand Retitriletes austroclavatidites (de Jersey &Raine 1990); (3) The Rhaetian Foveosporites more-tonensis Zone is defined by the acme of Densoispor-ites psilatus, by the LAD of Aratrisporites spp., bythe scarce presence of Perinopollenites elatoidesand by the presence of Rugaletes awakinoensis inassociation with Classopollis sp. cf. C. chateaunovi.Limbosporites antiquus is present in the upper partof the Zone; and (4) The Hettangian Retitriletesaustroclavatidites Zone distinguished by the FADof Retitriletesaustroclavatidites, by the LAD ofAratrisporites spp., by the acme of Densoisporitespsilatus and by abundant Classopollis spp. n theupper part of the Zone. Taxa such as Annulisporafolliculosa and Annulispora microannulata arecommonly present as accessory forms also in themiddle Carnian to early Norian of eastern Australia(Craterisporites rotundus Zone: de Jersey 1975),and in the Carnian to early Norian of westernAustralia (Dolby & Balme 1976). Polycingulatis-porites mooniensis is another cosmopolitan form ofthe southern hemisphere that is present as an acces-sory taxon in the middle Norian of eastern Australia(Queensland), New Zealand (Stevens 1981) andNew Caledonia (de Jersey & Grant-Mackie 1989).
An Ipswich microfloral affinity has also beenrecognized in Tanzania where the Rhaetian micro-floral assemblages of the Mkuju Formation shareseveral taxa with the microfloras from NewZealand, southern and Eastern Australia andMalagasy such as Duplexisporites problematicus
The palynomorph assemblages of Argentina arecomparable with the Ipswich microflora (Zavattieri& Batten 1996), and have close similarities withthose of eastern Australia, South Africa, Antarcticaand other Gondwanan domains. The oldest palyno-logical records are from the Cacheuta Formation inMendoza Province (Orlando 1954; Jain 1968), andfrom various Triassic basins of western Argentinastratigraphically correlated through palynostratigra-phy (Yrigoyen & Stover 1969). A Middle Triassic toCarnian microflora reported from the Cuyo Basin iscomposed of a large number of spores (e.g. Aratris-porites compositus, Auritulinasporites scanicus,Cadargasporites cuyanensis, Calamospora spp.,Camerosporites verrucosus, Dictyophylliditesmortonii, Guthoerlisporites cancellosus, Punc-tatosporites spp., Rugulatisporites spp.), abundantbisaccate (e.g. Alisporites spp., Falcisporitesnuthallensis, Klausipollenites schaubergeri, Minu-tosaccus acutus, Platysaccus spp., Protodiploxypi-nus spp., Triadispora spp.) and other gymnospermpollen (e.g. Accinctisporites spp., Cycadopitesspp.). The Carnian–earliest Norian has been docu-mented from the Los Rastros and IschigulastoFormations in the Ischigualasto–Villa UnixnBasin, San Juan–La Rioja provinces, on the basisof palynomorph assemblages containing Alisporitesspp., Cacheutasporites wielandii, Cadargasporitescuyanensis, Convolutispora microrugulata,Dictyophyllidites mortonii, Klausipollenites spp.,Monosulcites spp., Osmundacidites spp., Platysac-cus spp., Punctatosporites walkomii, Rugulatispor-ites neuquenensis, Spheripollenites spp. andVesicaspora spp., among others (Zavattieri &Batten 1996; Ottone et al. 2005). In the North Pata-gonian Basin, Rıo Negro Province, the continentalsuccession of the Cerro Puntudo Formation yieldeda rich palynofloral assemblage showing a Norian toRhaetian transitional composition. Spores such asAnnulispora folliculosa, Apiculatisporis lentus, Ara-trisporites spp., Craterisporites rotundus, Neorais-trickia densata, Osmundacidites spp., Striatellaseebergensis, pollen such as Alisporites spp., Indu-siisporites parvisaccatus, Klausipollenites deci-piens, Laricoidites intragranulosus, Triadisporaspp., Voltziaceaesporites heteromorphus and a fewClassopollis in the upper part of the succession,co-occur with the oldest Middle Triassic–earliestLate Triassic taxa. The youngest Triassic Argenti-nian assemblages (Norian–early Rhaetian), havebeen recorded from the Chihuido Formation at the
Llantenes locality, in the Malargxe Basin, southernMendoza Province, which yielded a microfloramainly composed of Baculatisporites comaumensis,Dictyophyllidites mortonii, Guthoerlisporites can-cellosus, Leptolepidites spp., Neoraistrickia taylorii,Polypodiisporites ipsviciensis, and pollen such asAlisporites spp. and Classopollis spp.
Conclusions
The compositional changes of miosporeassemblagesfrom the Late Triassic to Early Jurassic appearprimarily to be latitudinally, and therefore largelyclimatically controlled. Whereas palynostratigraphyprovides reliable results on regional stratigraphic
Fig. 12. Selected Norian and Rhaetian sporomorphs from the northern hemisphere. (1–4, 7–11, 13, 17) Classopollistorosus Reissinger 1950; (5) Classopollis murphyae (Cornet & Traverse 1975) Traverse 2008; (6, 12, 14–16, 19)Classopollis meyerianus (Klaus 1960) de Jersey 1973; (18) Granuloperculatipollis rudis Venkatachala & Goczan 1964emend. Morbey 1975.
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correlations, global correlations are complicated bymany factors, such as differences in the compositionof coeval assemblages, and by discrepancies in thetaxonomy of sporomorphs. Nevertheless, althoughlong distance correlations are commonly limited byclimatic and environmental factors, palynologymay be very useful in solving regional correlations.This is particularly evident for the Europeandomain, where the palynostratigraphy is well docu-mented and the available data provide a detailedregional picture of the Late Triassic and Early
Jurassic European microfloral composition. Further-more, microfloral associations provide useful sup-plementary information concerning the distributionof parent flora and the pattern of floral provinces.In the future, multidisciplinary studies based onthe integration of different biostratigraphical andgeochronological tools (e.g. magnetostratigraphy,chemostratigraphy) should better constrain the agesof palynomorph assemblages and improve the poten-tiality of palynology for world-wide correlationsamong marine and non-marine sections.
Fig. 13. Late Triassic sporomorphs from the southern hemisphere and Iran. (1, 2) Camerosporites secatus Leschik 1956emend. Scheuring, 1978; (3) Patinasporites densus Leschik 1956 emend. Scheuring, 1970; (4) Vallasporites ignaciiLeschik 1956; (5, 6) Enzonalasporites vigens Leschik 1956; (7, 8) Partitisporites novimundanus Leschik 1956; (9, 10)Samaropollenites speciosus Goubin 1965; (11) ‘Lueckisporites’ cf. L. singhii Balme 1970; (12) Brodispora striataClarke 1965; (13) Perinopollenites elatoides Couper 1958; (14) Ephedripites primus Klaus 1963; (15) Trachysporitesfuscus Nilsson 1958; (16) Polycingulatisporites mooniensis de Jersey & Paten 1964.
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Taxonomic note
Systematic nomenclature of the Corollina–Circulina–Classopollis complex: Traverse (2004,2008) proposed, by means of the procedures outlinedin the International Code of Botanical Nomenclature(Greuter et al. 2000), to conserve the morphogenericname Classopollis against Corollina and Circulina,which are commonly used for the same fossilpollen forms. This proposal has been accepted bythe International Botanical Congress in 2005. There-fore, in the present contribution the generic nameClassopollis will be used to replace Corollina andCirculina for the species: Corollina meyeriana(¼ Gliscopollis meyeriana), Corollina torosa,Corollina murphyi, Corollina sp. cf. C. chateaunovi(cf. also Cornet & Waanders 2006). Figures 9–13illustrate representative Late Triassic sporomorphs.
First of all many thanks to Spencer Lucas as editor of thisbook. I gratefully acknowledge Nicoletta Buratti (PerugiaUniversity, Italy) for having made available her palynolo-gical collection and for her comments and suggestions,Roberto Rettori (Perugia University) for his constructivediscussions on Triassic biostratigraphy and LawrenceTanner (Le Moyne College, Syracuse, N.Y.) for hisfriendly encouragement and for reviewing the manuscript.I am also deeply grateful to the other reviewers,W. M. Kurschner (Utrecht University, NL), J. Ogg(Purdue University, Indiana, USA) and J. B. Riding(British Geological Survey, UK) for their helpful com-ments and suggestions that surely improved the qualityof this article. Part of this study was financially supportedby the Italian Minister of Scientific Research (PRIN05–07,S. Cirilli).
Appendix I
Species list of cited taxa
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Annulispora microannulata de Jersey 1962Apiculatisporis lentus Playford 1982
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