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Record of environmental changes in the Triassic–Jurassic boundaryinterval in the Zliechov Basin, Western Carpathians

Jozef Michalík a,⁎, Otília Lintnerová b, Andrzej Gaździcki c, Ján Soták d

a Geological Institute, Slovak Academy of Science, Dúbravská 9, PO Box 106, 84005 Bratislava, Slovakiab Comenius University, Faculty of Science, Mlynská dolina, 842 15, Bratislava, Slovakia

c Institute of Paleobiology, Polish Academy of Sciences, Twarda 51/55, 00-818 Warszawa, Polandd Geological Institute, Slovak Academy of Science, Severná 5, 974 01 Banská Bystrica, Slovakia

Received 1 February 2005; accepted 20 June 2006

Abstract

The Fatra and Kopieniec formations in the Zliechov Basin comprise a record of several environmental crises that may havecontributed to the global Triassic–Jurassic (T–J) boundary events. The diversity of the benthic fauna decreases at the base of the‘transition beds’, the uppermost member of the Fatra Formation. This assemblage comprises many species that do not appear in youngerstrata. In ‘cycle 13’ of the Fatra Formation, stable isotope excursions correspond to significant lithological changes in the sequence. Thenegative δ13C excursion is correlated with a positive δ18O peak and the beds containing the excursion are overlain by a thin layer withunusual lithological and mineral composition. This layer is composed of small calcitized microspheres showing complex alterationduring diagenesis. The origin of this layer, traceable over tens of kilometres, is problematic (impact ejecta, volcanic glass, or alteredaragonitic particles). An increase in terrigenous sediment recorded by sandy beds in the uppermost part of the Fatra Formation indicatesincreased run-off. A sharp lithological boundary between the Fatra and the Kopieniec formations marks a sudden termination ofcarbonate sedimentation, followed by deposition of the ‘boundary clay’. The negative δ13C anomaly recognized in this succession maycorrespond to the isotopic signature at the T–J boundary which has been described from many sections in the world.© 2006 Elsevier B.V. All rights reserved.

Keywords: Western Carpathians; Triassic–Jurassic boundary; Climate change; Eustatic cyclicity; Ramp carbonates; C and O isotopes

1. Introduction

The end of the Triassic and the beginning of theJurassic periods were affected by both climatic changeand biotic turnover, induced by sea-level fluctuationscombined with both poorly known terrestrial (volcanic)

and, perhaps, with extraterrestrial factors (Hallam, 2002;Walkden et al., 2002). These changes resulted in one ofthe five largest Phanerozoicmass extinctions (Stanton andFlügel, 1987; Michalík et al., 1991; Hallam and Wignall,1997, 2000), which took place at the end of Triassic, c.200Ma ago, when up to 80% of marine species were lost.According to Pálfy et al. (2000), the extinction of marineorganisms may have been preceded by the biotic crisis onthe land by as much as several hundred thousand years.Pálfy et al. (2001) and Hesselbo et al. (2002) argue thatCO2 emission by CAMP volcanism and possibly the

Palaeogeography, Palaeoclimatology, Palaeoecology 244 (2007) 71–88www.elsevier.com/locate/palaeo

⁎ Corresponding author. Tel.: +421 7 54773961; fax: +421 7 54777097.E-mail addresses: geolmich@savba.sk (J. Michalík),

lintnerova@fns.uniba.sk (O. Lintnerová), gazdzicki@twarda.pan.pl(A. Gaździcki), sotak@savba.sk (J. Soták).

0031-0182/$ - see front matter © 2006 Elsevier B.V. All rights reserved.doi:10.1016/j.palaeo.2006.06.024

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pysynchronous release of methane from marine clathratemay explain the remarkable negative carbon isotopeexcursion observed in the boundary interval.

The aim of our research is to explain several peculiarphenomena observed in the Triassic–Jurassic (T–J) boun-dary sequence in the Zliechov Basin, which represents themost complete record of this time interval in the WesternCarpathians. These are the eustatic cyclicity of sediments,changes in the fossil assemblages, the occurrence ofspherule layers, slump beds, abrupt cessation of carbonatedeposition, and anomalies in the C and O isotope record.

1.1. Geological setting and stratigraphy of the latestTriassic Fatra Formation

During the latest Triassic, Penninic rifting broke off theAustro-Alpine-Central Carpathian block from the Eur-opean shelf (Kuss, 1983; Michalík, 1993, 1994). Due tofactors such as erosion prior to the start of the Hettangian,tectonic shortening, and poor exposure, the T–J boundaryinterval is not well recorded in Western Carpathiansequences (Fig. 1). Moreover, a paucity of index fossils(i.e. cephalopods and conodonts) hampers a precise bio-stratigraphic determination of the Rhaetian–Hettangianboundary in the studied sections.

On the other hand, the T–J passage beds are wellpreserved in the Fatric domain, which now forms the

Krížna Nappe of the imbricate central Western Carpathianorogenic structure (Michalík, 1980). A shallow marinecarbonate sequence (Fatra Formation) rests transgressivelyon terrestrial Carpathian Keuper deposits in the small pull-apart Zliechov Basin (Michalík, 1974, 1978b). It is over-lainwith a sharp contact bymarine shales of theHettangianKopieniec Formation (Stur, 1859; Goetel, 1917; Gaździckiet al., 1979; Tomašových and Michalík, 2000).

In the 300 km long and 100 km wide basin, ten faciesareas were recognized by Michalík (1973, 1974, 1977).Depositional environments varied from salt marshesthrough carbonate ramp to deeper neritic slope, andwere populated by characteristic benthic associations(Michalík and Jendrejáková, 1978; Michalík, 1978a)dominated by brachiopods (Rhaetina gregaria, seeMichalík, 1975, 1978b, 1980) and/or bivalves (e.g. Pla-cunopsis alpina and Rhaetavicula contorta, Table 1).Corals, sponges, algae and hydrozoans inhabited smallisolated patch-reefs on the edge of carbonate ramp.Storms and sea-level fluctuations combined with slowsubsidence ratio (22–38 mm/ka) hampered formation ofany larger reefal bioconstructions (Michalík, 1982;Tomašových, 2004a). The ramp faced the deeper, dysoxicpart of the basin without benthic fauna (Fig. 2), similarlyas in the Kössen Basin of the Alps (Kuss, 1983).

The thickness of the Fatra Formation is variable, from25 m to a maximum of 116 m in areas of higher

Fig. 1. Geographic and tectonic setting of the four studied T–J sections in the High Tatra Mts.

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Table 1Cyclostratigraphic and biofacies data from thirteen studied sections of the Fatra Formation in the Zliechov Basin

Section 1 2 3 4 5 6 7 8 9 10 11 12 13 14

A. Number of beds in individual cycles in selected sections of the Fatra FormationSuchá Valley ? 3 3 2 4 3 1 1 ErodedLesnianskaValley

1 2 3 5 5 2 4 1 Missing due to erosion

Belianska Valley 4 2 2 3 5 5 5 2 3 missing due to erosionHíreška,Valaská Belá

3 5 4 2 2 2 4 3 2

Zázrivá Valley 1 4 4 4 3 3 3 1 2Štefan ravine ? 2 11 3 2 4 5 2 1Bobrovček 1 2 3 9 4 2 5 5 2 ErodedBystrô 2 3 3 2 2 2 3 4 3 3 1 ErodedDedošova Valley 3 2 2 2 3 2 3 2 2 4 3 1Široký Žlab,Ždiarska Vidla

2 5 3 12 4 3 3 7 2 5 3 3

SlovianskaValley

Not exposed 1 2 2 3 3 2 4 5 2

Furkaska 7 3 9 4 8 8 7 5 2 7 3 2 6 5Kardolína 4 4 8 5 4 6 4 1 1 4 6 5 3 9

B. Dominance of faunal elements in selected sections of the Fatra FormationSuchá Valley ? Por(7) Por(14) Por(10) Por(12) Por(10) Crn(5) Por(36) ErodedLesnianskaValley

Biv(10) Por(8) Biv(5) Por(5) Ant(10) Por(6) Biv(19) Crn(15) Missing due to erosion

Belianska Valley Biv(18) Biv(10) Biv(23) Biv(10) Biv(16) Crn(11) Biv(38) Por(16) Bpd(18 Missing due to erosionHíreška,Valaská Belá

Biv(26) Bpd(13) Biv(12) Por(12) Biv(45) Crn(20) For(16) Biv(36) Biv(23)

Zázrivá Valley Bpd(4) Por(26) Ant(28) Ant(11) Crn(5) Por(6) Bpd(5) Biv(11) Biv(12)Štefan ravine ? Biv(13) Biv(11) Biv(22) Biv(20) Biv(12) Biv(17) Biv(10) Biv(12)Bobrovček ? Biv(8) Biv(6) Biv(13) Biv(20) Biv(25) Biv(22) Biv(42) Biv(18) Biv(5) ErodedBystrô Biv(10) Biv(10) Crn(3.5) Biv(9) Por(10) Biv(13) Biv(37) Biv(16) Biv(28) Crn(9) Crn(20) ErodedDedošova Valley Biv(16) Por(30) Por(42) Ant(56) Gpd(26) Biv(30) Bpd(13) Bpd(22) Crn(10) Ant(20) Biv(22) Biv(40)Široký Žl’ab,Ždiarska Vidla

Biv(10) Bpd(19) Biv(25) Bpd(28) Crn(23) Biv(6.5) Crn(6) Por(9) Crn(2) Crn(6) For(7) Crn(3)

SlovianskaValley

Not exposed Bpd(8) Crn(5) Por(6) Por(8) Por(11) Alg(13) Biv(9) Biv(9) Biv(15)

Furkaska Bpd(19) Por(19) Bpd(12) For(8) Crn(23) Crn(26) Crn(13) Biv(34) Biv(40) Biv(22) Biv(13) Biv(2) Crn(26) Crn(28)Kardolína Biv(10) Biv(25) Biv(15) Biv(15) Ant(18) Crn(5) Crn(21) Crn(21) For(4) Bpd(12) Bpd(9) Ost(9) Biv(8) Crn(13)

A. Number of beds in individual cycles. B. Comparison of biofacies in individual cycles in the sections studied sections, based on percentages (in brackets) of the predominant group of organisms. Alg—algae, Ant—corals, Biv—bivalves, Bpd—brachiopods, Crn—crinoids, For—foraminifers, Gpd—gastropods, Ost—ostracods, Por—sponges.

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pysubsidence. Considerable horizontal and vertical micro-facies variation is typical. Debris of neritic organismsform large accumulations in shallow areas in the westernpart of the basin (Fig. 3). Terrigenous clastic content of thelimestones is generally low,with the exception of sporadicadmixture of clay and silt (Tomašových, 2000). A higherpercentage of clastic quartz is restricted to a delta lobe inthe northern part of the basin (Michalík, 1977, 2003).

Michalík et al. (1977, 1978b) and Gaździcki et al.(1979) divided the Fatra Formation (Figs. 3 and 4) into fiveinformal members. In stratigraphic order, these include 1)the ‘basal beds’, characterized by coquinas with bivalves,gastropods, and rare other stenohaline organisms; 2) the‘lower biostromatic member’ containing more diversifiedassociations of neritic organisms (brachiopods, corals,sponges, crinoids, algae etc.); 3) the ‘barren interval’madeup mainly of dolostones, dolomitized limestones andredeposited clastic limestones that were formed during ashallowing period; 4) the ‘upper biostromatic member’which is similar to the ‘lower biostromatic member’,containing corals, foraminifers and other organisms; and,finally, 5) the ‘transitional beds’ containing several unusuallithological horizons, including micritic ostracodal lime-stones with undulating bedding planes, oolitic limestones,and coquinas with bivalves and foraminifers. Co-occur-

rence of chamosite and goethite ooids in oolitic limestonethat is enriched in chromium and vanadium indicatesmixing of material derived from different environments(Michalík et al., 1982, 2002).

1.2. Stratigraphy of the earliest Jurassic KopieniecFormation

The Kopieniec Formation (defined by Goetel, 1917 inthe Polish part of the Tatra Mts.) consists of a sequence ofbrown claystone with sandstone and limestone intercala-tions. The formation was divided into four informalmembers by Gaździcki et al. (1979): the ‘basal clastics’(also called the ‘Cardinien Sandstein’ by Goetel, 1917),the ‘lower limestone’, the ‘main claystone’ and the ‘upperlimestone’ (also called the Široké Limestone Member byLefeld et al., 1985).

The ‘basal clastic member’ starts with the usually 6 mthick, non-calcareous ‘boundary claystone’. It is overlainby pale quartz sandstone that contains thin muscovite-bearing clay interbeds (Gaździcki et al., 1979). Thedominant heavy mineral in the sandstone is tourmaline(Michalík et al., 1982). The siliciclastic admixturerepresents a feature that is only infrequently observed inthe underlying limestone sequence (Michalík, 1977).

Fig. 2. Palinspastic 3D reconstruction of the depositional basin of Fatra Formation. The basin belonged to a system of ‘Kössen-type’ intra-shelfdepressions dissected by transverse elevations. The length of the block diagram is c. 250 km.

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Fig. 3. Detailed correlation of microfacies, carbonate C and O isotopes, organic carbon (TOC) content and foraminiferan distribution in the Furkaska section. The main microfacies components(micrite, biomorphs, sparite and extraclasts) are divided into two groups, plotted along symmetrical axes. Note the two negative excursions of δ13Ccarb, associated with positive excursions of bothδ18Ocarb and TOC. A third excursion occurs at the base of the Kopieniec Formation. Two important breaks in foraminifer distribution occur close to the negative δ13C excursions.

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Fig. 4. Detailed correlation of microfacies, carbonate C and O isotopes, organic carbon (TOC) content and foraminifer distribution in the Kardolína section. See Fig. 3 for further explanations. Note thecorrelated changes in both the foraminifer distribution and the isotopic ratios.

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Farther to the south, the sandstone is replaced by red huedsiltstone comparable with the ‘Schattwald beds’ of theEastern Alps (Borza et al., 1980).

In strata overlying the Basal Clastic Member, an earlyHettangian age is indicated equivalent to the standardPlanorbis Zone (Rakús, 1993). Up to now, the early

Fig. 5. Characteristic foraminifers of the T–J boundary beds in the Zliechov Basin. 1. Crinoid-foraminifer biopelmicrite (×10) with Ophthalmidiumleischneri (Kristan-Tollmann, 1962). Insets: Involutina liassica (Jones, 1853), a. magnification ×150; b. magnification ×75. Strażyska Valley section,Kopieniec Formation (‘lower limestone’), sample 4, lowermost Hettangian. 2. Foraminifer biosparite with Triasina hantkeniMajzon, 1954. Furkaskasection, Fatra Formation (‘lower biostrome member’), Bed 338, Rhaetian (×10). 3. Coral-crinoid-foraminifer biomicrite (×10) with abundant Glo-mospirella friedli (Kristan-Tollmann, 1962; emend. Brönnimann and Zaninetti, 1970). Inset a. magnification ×80. Lejowa Valley section, FatraFormation, (‘upper biostrome member’), Bed 49, Rhaetian.

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Hettangian Calliphyllum Zone, partially correlated withthe Planorbis Zone, has been reported from three localitiesonly: the mouth of the Sokol Ravine in the Zázrivá Valley(Malá Fatra Mts.), Mt. Babky-Úplazy (Western TatraMts.) and the Rakytov-Tanečnica Hill (Veľká Fatra Mts.).Both at the first and second localities, the ammonitesPsiloceras psilonotum (Quenstedt) and Caloceras cf.torus (d'Orbigny) were found in bioclastic limestone inthe uppermost part of the ‘basal clastic member’, near thebase of the ‘lower limestone member’ of the KopieniecFormation. Middle and Late Hettangian (Liasicus andAngulata Zones) ammonites occur more commonly in themiddle part of the Kopieniec Formation (Rakús, 1975;Gaździcki et al., 1979; Rakús, 1993, 2003).

1.3. Biostratigraphy of the T–J boundary beds

Biostratigraphically important Late Triassic microfos-sils (mainly foraminifers and conodonts, as well aspalynomorphs and coprolites) are rare in the WesternCarpathian neritic carbonate sequences (Gaździcki, 1974,1978; Gaździcki et al., 1979; Gaździcki and Michalík,1980; Błaszyk and Gaździcki, 1982; Gaździcki, 1983;Michalík and Gaździcki, 1983; Fijałkowska and Uchman,1993; Gaździcki et al., 2000). Rapid evolutionary changesof Involutinidae, Ammodiscidae and Ophthalmidiinaewith several evolutionary lineages permit the establish-ment of a relatively detailed foraminiferal zonation. Thesequence of the Rhaetian Glomospirella friedli–Triasinahantkeni assemblage zone and the Hettangian–?Sinemur-ian Ophthalmidium leischneri–Ophthalmidium walfordiassemblage zone (Gaździcki, 1978) is confirmed (Fig. 5).The Glomospirella friedli–Triasina hantkeni zone wascorrelated with both the Choristoceras haueri and Ch.marshi ammonoid zones (Rhaetian), and its extent alsocorresponds to that of the Misikella posthernsteiniconodont zone. The extent of early Early JurassicOphthalmidium leischneri–Ophthalmidium walfordizone may correspond to the Planorbis to Angulata, andpossibly also the Bucklandi standard ammonite zones(Hettangian–?Sinemurian), defining the age of the basalJurassic strata in the Tatra Mts. (Gaździcki, 2003).

2. Subject and methods of study

2.1. Selected sections

Four sections that provide good exposure of thetransition from Rhaetian limestones of the Fatra Forma-tion toHettangian shales of the Kopieniec Formation havebeen selected along the Slovakian and the Polish sides ofthe Tatra Mts. (Figs. 1 and 2): 1) the rocky slope of Mt.

Pálenica above the Kardolína gamekeeper house nearTatranská Kotlina, Belá Tatra Mts. (Fig. 4); 2) a smallravine on the left side of the Lejowa Valley; 3) a gorge on

Table 2C and O isotopes compositions from the T–J boundary beds in theFurkaska and Lejowa sections

Unit Cycle Furkaska section Lejowa section

Bed δ13C δ18O Bed δ13C δ18O

FatraFormation

10 383 2.00 −3.90 33 1.912 −5.24384 1.60 −4.40 40 1.413 −3.93385 1.70 −5.30 42386 1.50 −5.00387 1.50 −4.10

11 388 2.10 −5.20 44 1.78 −3.48388+ 1.60 −3.30 49 2.018 −3.46389 2.00 −4.10390 1.77 −4.18 51 2.16 −3.19391 1.70 −2.00392 1.90 −3.40 55 1.64 −3.87

12 393 1.78 −3.98 59 1.82 −4.39394 1.80 −4.20 60 1.67 −1.67395 0.67 −2.35 61 1.04 −2.08

13 396 1.40 −4.30 62 1.63 −5.12397 1.60 −3.10 64 1.70 −5.83398 2.07 −3.84 65 0.67 −4.73399 1.70 −5.20 66 0.10 −6.09399b 1.38 −3.95 67 −0.46 −4.51400.1 1.13 −4.57400 0.38 −3.89 69 −1.39 −4.02400.1 0.20 −4.30 70400.2 0.89 −4.17 71 −1.35 −0.54400.3 1.03 −4.44

14 401 1.07 −1.12 73401 1.00 0.00402 0.48 −5.20402+ 1.00 −4.60 74403 1.07 −6.71405 1.06 −6.64 75 1.05 −6.67406 1.71 −5.45407 0.90 −5.00408 0.55 −4.95408b 1.33 −5.71408s 0.70 −5.18408v 0.50 −5.67

KopieniecFormation

BoundaryclaystoneCardiniaSandstone

408+ −0.90 −7.20409 1.10 −6.40409+ 1.40 −6.20414 −0.20 −6.00 3.01 0.52 −6.82415 0.30 −5.10416 0.50 −6.20

LowerLimestone

417 0.01 −5.10418 0.50 −5.10419 0.80 −6.90 3.03 0.28 −7.78420 0.50 −5.80422 −0.44 −6.54423 −0.44 −6.87 3.06 −1.25 −8.82

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the slope of Mt. Furkaska in the Juráňova Valley (both inthe Western Tatra Mts., Fig. 3); and 4) an avalanche gullyon Mt. Ždiarska Vidla in the Zadné Meďodoly Valley.These sites were also visited during a field workshop ofIGCP Project 458 (“Triassic and Jurassic BoundaryEvents”) (Michalík, 2003).

2.2. Sequence stratigraphic analysis of sedimentarycycles

The first recognition of sedimentary cycles was basedon macroscopic features represented by the alternation ofdifferent rock types (pure and marly limestone, marl ordolostone beds), by the geometry of lithosomes (changesin bed thickness, correlative unconformities), faciescorrelation (clastic interbeds, sedimentary textures) andfossil content (reefs, biostromes, shell accumulations).Secondly, the analysis was focused on amicrofacies studyof the complete rock sequence using sets of thin sections.Two groups of microfacies elements could be distin-guished. The pelagic bioclasts (tests of plankton, skeletalparticles of nektonic animals along with micrite ofunknown origin), were predominantly accumulatedduring late transgressive and highstand stages of thesea-level (Figs. 3 and 4). On the other hand, the non-pelagic bioclasts (detritus of neritic benthic organisms,terrigenous lithoclasts) were transported basinwardsduring both falling-and lowstand sea-level stages. Thedistinction between the two groups was accentuated bysparry cementation of well-sorted carbonate sediment.The percentage of these components was estimated bymicroscopic analyses (Schäfer, 1969; Soudant, 1972) andevaluated in two-dimensional graphs. These graphs,which represent individual sections, could be fit into apalinspasticmodel (Fig. 2), and their comparison providesinformation on the character, composition and dynamicsof the basin margins that were affected by sea-levelchanges (Michalík, 1993, 1994; Reháková and Michalík,1994; Michalík and Reháková, 1997). Some of thecollected samples (especially those taken from thespherule layers in the Kardolína section) were also studiedusing a Cameca SX100microprobe and a JEOL-JFM 840scanning electron microscope.

2.3. Stable isotope analyses

In total, 48 bulk carbonate samples fromKardolína, 51samples from Furkaska, and 26 samples from the Lejowasection were analyzed for C and O isotope compositions(Figs. 3 and 4, Table 2). Total organic carbon (TOC)content wasmeasured in the first two sections. Rock slicescut from hand samples used for thin sections preparation

were cleaned and crushed. Diagenetic calcite veinletshave been removed under a binocular microscope andselected pieces of the sample were homogenized in anagate mortar. Homogenized samples were used fororganic carbon content and carbon and oxygen isotopeanalyses. Total organic (TOC) and inorganic carboncontent (TIC) wasmeasured on a C-MAT 550. TIC valueswere recalculated to CaCO3 content to assess thecarbonate content of samples selected for isotopeanalyses, which were performed on Finnigan MAT-2mass spectrometers in Prague and Warsaw. The isotopicdata are presented using the standard δ notation in permil(‰), relative to the International Isotopic Standard VPDBwith a precision of 0.01‰.

3. Results

3.1. Eustatic cyclicity

Aweak eustatic signal in comparison with autocyclicmechanisms in Triassic carbonate ramps and platformswith highly variable lithosome geometries makes thesequence analysis difficult (Lakew, 1990; Hofmann et al.,2000). The Fatra Formation sequence, consisting oforganodetrital, oolitic, pelletal limestones, brachiopod-and bivalve coquinas, dark grey marls and laminateddolomicrites, and shows cycles with shallowing upwardspatterns, as indicated by an upsection increase of bioclastsand micrite (Figs. 3 and 4). Fourteen eustatic cycles aredistinguished and correlated over a distance of more than100 km laterally. At the beginning of each cycle, clastsand sparite prevail over other components. Higher up, theproportion of bioclasts and micrite increases. The upper-most part of the cycles consists of micrite or, commonly,dolomicrite. This trend indicates a gradual decrease ofenvironmental energy and defines a shallowing-upwardscycle. The four uppermost cycles, which belong to the‘transitional member’ below the base of the overlyingHettangian Kopieniec Formation, comprise lithologicallydifferent rock types including dark limestonewith nodularsurfaces, ferroan oolitic limestone, and beds containingmicrospheres. The contact with claystones of theKopieniec Formation is sharp (Figs. 3 and 4).

This cyclostratigraphic subdivision has been con-firmed by recently completed quantitative microfaciesanalysis of thirteen selected sections (Michalík, 1977,Tomašových, 2004a,b, Table 1). Fourteen shallowing-upwards cycles of ‘keep-up’ character have beendistinguished in the Fatra Formation sequence. Regularalternation of microfacies elements indicates repetition,hypothetically in 400 ka cycles that correspond to thevariation in orbital eccentricity. The correlation

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highlights the diachronism of the bases of both the Fatraand Kopieniec Formations due to their erosionalcharacter (as first proposed by Gaździcki and Iwanow,1976) resulting in the lowermost cycles (1–3) and theuppermost ones (13–14) are commonly missing.

3.2. Fossil distribution

The succession of benthic assemblages in the FatraFormation has been described byMichalík (1975, 1978a,1982), Roniewicz (1974), Roniewicz and Michalík(1998) and Michalík and Jendrejáková (1978). Renewedcolonization of the marine substrate occurred in eachcycle during sea-level rise and peaked at the maximumflooding. Biogenic production reached maxima in thebiostromal members, in cycles 3 to 5 and 8 to 10. Thesewere separated by a regressive member of the ‘barreninterval’. However, in cycles 11–13, reef builders(sponges, hydrozoans, and corals) and megalodontidsbecame less abundant, and they disappeared in cycle 14.A rich benthic foraminifer fauna is represented byAmmodiscidae, dominated by Glomospirella and Glo-mospira (Fig. 5). From the base of the ‘transitional beds’up to the T–J boundary, the abundance of foraminifersand the sizes of individuals decrease. Three extinctionlevels (at beds 393, 401, and 406) have been observed inthe Furkaska section (Fig. 3), two of them are alsopresent in the Kardolína section (Fig. 4). A renewal of theforaminifer fauna is found in the lower limestonemember of the Kopieniec Formation, dominated by

Ophthalmidium leischeri, nodosariids and involutinids(Involutina liassica, I. turgida, and I. farinacciae).

The distribution of fossil microfauna within the FatraFormation seems to be environmentally controlled. Boththe complex foraminifers (e.g. Triasina hantkeni) anddasycladacean algae are considered to be oligotrophicelements. The foraminifers gradually disappear a fewmetres below the assumed T–J boundary where the testsbecome largest in size and the microsphaeric formsattain the greatest proportion (Bed 55 at Kardolína, andBed 399 at Furkaska). The distribution of dasyclada-ceans shows a similar decline towards the T–Jboundary. The last dasycladacean algae (Placklesiamultipora) occur in Bed 376 of the Furkaska section.Oligotrophic organisms, including corals, disappear inBed 390, where abundant brachiopod shells indicatehigh juvenile mortality. An increase in trophic resourcesis indicated by the common occurrence of cyanobacter-ial nodules of Girvanella (Bed 404), and by intensivemicroborings of skeletal grains by endolithic microbeswhich benefited from increased availability of nutrients(Brasier, 1995). The T–J boundary is marked bypronounced carbonate depletion which might reflectan excess of nutrients such as nitrogen and phosphorus(Hautmann, 2004; Ward et al., 2001). The non-calcar-eous ‘boundary clay’ contains less than 1% CaCO3 andcommon early diagenetic pyrite framboids; it thereforeseems to have been deposited in a nutrient-richenvironment. Drastically impoverished metazoans arerepresented only by rare ostracods and microgastropods.

Fig. 6. Beds 49–57 in the Kardolína section. Marly limestone beds 51 and 52 contain anomalous C and O isotope ratios. Beds 53 and 54 consist ofspherulitic limestone, overlain with slight disconformity by redeposited clastic limestone of Bed 55 that contains reworked spherule-like particles.

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The ostracod fauna, including Ogmoconcha, couldtolerate eutrophic conditions and fluctuating oxygenlevels, thus displaying the opportunistic characteristicsof pioneer species.

3.3. Spherule beds and slumping in cycle 13

Beds 53 and 54 in the Kardolína section (base ofcycle 13, Fig. 8) are rich in small globular particles thatare 0.2 to 0.5 mm in diameter (Fig. 6). The inner part ofeach microsphere is filled by blocky calcite (Fig. 7a–b),

sometimes enclosing enclaves of clayey material andidiomorphic carbonate (probably dolomite) rhombs(Fig. 8b). The outer layer of the microspheres is moreresistant, containing tiny, dispersed quartz crystals andcovered by pallisade calcite grains (Fig. 7a–b). Theoverlying Bed 55 also contains common microspheresbut they are mixed with abraded fragments of marineorganisms (Fig. 7c–d), crinoid ossicles, ooids, androunded lithoclasts (Fig. 7e–f).

Several mineral phases were recognized in the calciticmicrospheres. Their interior contains clusters of

Fig. 7. Photomicrographs of microspheres from the Kardolína section. a. Microspheres from Bed 54 filled by blocky calcite enclosing clayey material.b. Dolomite rhombs forming the last generation of authigenic minerals. c. Microspheres from Bed 55, together with skeletal fragments of benthicorganisms (bivalve shell). d. Calcitized microspheres from Bed 55, together with crinoid ossicles. e–f. Microspheres and rounded lithic fragmentsfrom Bed 54. Cavities filled by sparry calcite. Magnification of all figures ×100.

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phyllosilicates (chlorite, white mica, sericite), feldspar,fluorite, quartz and dolomite. The chlorite and micarepresent alteration products of primary phases (Fig. 7c–d). Fluorite occasionally occurs in more or less regularaggregates in the centre of small microspheres (Fig. 7c).The central part of the hypidiomorphic feldspar crystals of10–20 μm size is more basic (An20), whereas theirmarginal parts are composed of albite. Two quartzgenerations are distinguished. Large (200–300 μm)quartz crystals with surfaces corroded by calcite occurin the limy matrix. Small (10–30 μm), hypidiomorphiccrystals causing the higher acid resistance of rims cover

the crustal part of microspheres (Fig. 7 b, d). They wereprobably formed by remobilization of silica from primarysilicates. Dolomite occurs in the form of irregularlydispersed aggregates and larger rhombs (20–100 μm, Fig.7b). The central part of the crystals is enriched in Fe2+

(siderite) and is depleted in Mg2+ in comparison with theouter part. The dolomite is never substituted by calcite andprobably represents the youngest mineral phase inmicrospheres.

Clasts of the spherulitic limestone were also found inBed 400.1 of the Furkaska section (Fig. 3). This layer isoverlain by an untypical interval that is made up by a

Fig. 8. SEM andmicroprobe images of the microspheres from the Kardolína section. a. Matrix-supported spherulitic limestone from Bed 55, etched byformic acid. b–c. Structure of microspheres, showing partially calcitized interior (blocky calcite grains), rimmed bymore resistant white crust, coveredby pallisade calcite cement; from Bed 55. d. Hypidiomorphic quartz crystals with surfaces corroded by hydrocarbon fluids; from Bed 54. e. Back-scattered electron microprobe images of microspheres from Bed 54. Note enclaves of remaining primary chlorite and sericite (arrow), partiallysubstituted by feldspar (dark grey) and fluorite (white) in the calcitized interior and coated by a crust with scattered quartz crystals. f. Hypidiomorphicplagioclase crystals with oscillatory zoning and surface corroded by hydrocarbon fluids; from Bed 54.

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slumped limestone body, c. 5 m in thickness, with orangeand yellow Fe-rich coatings. Submarine slumping isindicated by erosional channels filled by oolitic limestonein its uppermost part.

3.4. C and O isotope geochemistry

The measured carbon and oxygen isotope ratios andTOC values of carbonate samples are plotted beside thestratigraphical column of the Kardolína and the Furkaskasections where cycles 13 and 14 are preserved (Figs. 3 and4). The carbon and oxygen isotope curves in the latestTriassic part of these two sections are comparable. δ13Cvalues fluctuate in a relatively wide range of −0.9‰ to +2.22‰ at Furkaska and −1.6‰ to +2.25‰ at Kardolína.However, the majority of samples show values between +0.5‰ to +1.7‰ at Furkaska, and +0.7‰ to +2.0‰ atKardolína. The δ13Ccarb fluctuation in both isotope profilesreflect cyclic facies variation and the negative peaks closelyapproximate the extinction levels of the foraminiferanfauna. The δ13Ccarb values from the ‘biostromatic member’of the Fatra Formation are relatively stable in contrast withsamples from the ‘transitional member’, which showgreater variability. Two negative peaks appear in the δ13

Ccarb isotope curve in equivalent levels of the transitionalmember in both sections (Figs. 3 and 4). The first peakappears immediately above the extinction of involutinidforaminifers. The second negative peak is located inspherulitic limestone beds. The positive δ18Ocarb shift justabove this horizon is probably related to the speciallithology of these beds. An increased TOC content from0.5% to 0.7% is recorded in some beds, especially in thespherulitic limestone samples. A third negative δ13C peakoccurs in the topmost carbonate beds, near the supposed T–J boundary level (Figs. 3 and 4).

Preliminary correlation of cycles 13 and 14 relative tothe T–J boundary interval is also supported by the δ13Ccurve from the Lejowa section (Table 2). Here δ13Cdecreases from positive tomarkedly negative (−1.36‰ to−1.04‰) values in beds 69 and 71, similarly to thespherulite horizon of the Furkaska section. Despite themore scarce sampling of the Lejowa section, the δ13Ccurves from the two sections are well correlated.

The oxygen isotope ratios reflect mainly the faciessetting and the diagenetic overprint. The δ18O valuesrange from−2.0‰ to−6.71‰ in the Furkaska successionof shallower ramp setting and from−4.5‰ to−6. 99‰ inthe Kardolína succession that represents a relativelydeeper setting. The δ18O pattern shows higher amplitudefluctuation in the shallower ramp sequence which mayhave been influenced by meteoric water. Less negativevalues (0.0‰ to −0.8‰ at Furkaska and −1.46‰ at

Kardolína) were obtained not only from the spheruliticlimestone but also from the topmost limestone beds of theKardolína section (−2.27‰ to −0.47‰, Fig. 4), whichwere deposited on the slope.

A possible diagenetic overprint in the analyzedsamples was assessed using light microscope and SEMstudies and a correlation analysis of isotopic composi-tions. Recrystallization of the matrix is more common insparitic basal layers of carbonate cycles (Figs. 3 and 4).This observation was taken into consideration for theisotope sample selection. The correlation of δ18O vs.δ13C values in the sample sets is low. However, diageneticrecrystallization cannot be ruled out. Significant dolomi-tization and more positive δ18O values in the Kardolínasection indicate that more saline bottom waters couldaffect the sediments soon after deposition. Typicaldolomite rhombs together with fluorite are visible in thecalcite microspheres which indicate the presence ofevaporitic fluids. The relative position of diageneticphases suggest that the dolomite rhombs may haveformed during a later stage of diagenesis than the outerrim of spherules (Fig. 8).

The δ13C values likely preserve a primary signal andcan be used to determine the T–J boundary in the sections.Facies interpretation, quantitative evaluation of lithology,and the low correlation of δ13C and δ18O values indicatethat the δ18O pattern is also informative but it preferen-tially reflects early diagenetic conditions, which generallymay not have been fundamentally different from theprimary conditions (Veizer et al., 1997). However, apalaeo-temperature interpretation is problematic, due tothe probability of water salinity changes in the course ofthe carbonate ramp evolution.

3.5. Termination of carbonate deposition

The main difference between the Fatra Formation, thatconsists predominantly of limestone, and the overlyingKopieniec Formation, mostly composed of claystone, is intheir carbonate content. Although early Hettangian erosionmay have accentuated the boundary between them, thissignificant change in lithofacies may reflect a majorchange in global climate and in the ocean chemistry.Several sandy beds, indicating enhanced input of terrige-nous quartz (10–42%) and organic (higher plant) debris,were recorded at various levels of the studied sections. Inother sections, termination of carbonate deposition ismarked by oolite beds. In the Furkaska section (Beds 407and 408), mixing of ooids of different genetic types,including chamosite and hematite, is observed.

Nevertheless, a couplet of sandy limestones occurringin cycle 14 is correlatable in several sections and it

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Fig. 9. Lithological correlation of the Kardolína, Furkaska, Lejowa, and Široký žl’ab sections.

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possibly represents the first significant increase insiliciclastic input into the basin (Fig. 9). The amount ofterrigenous material transported from the adjacent landareas into the basin during deposition of the ‘boundaryclaystone’ suggests a sudden climatic shift towards higherprecipitation rate and an increase in the riverine influx.

4. Discussion

A long-termmarine transgression at the end of Triassicwas overprinted by high frequency sea-level changes(Michalík et al., 1976; Michalík, 1980; Michalík et al.,1988, 1991; Michalík, 1993, 1994). The palaeogeo-graphic setting in the earliest Jurassic period wascontrolled by extensive transgression and the reorganiza-tion of sedimentary basins in connection with a rejuvena-tion of tectonic movements (i.e. the early Cimmerianorogeny) and the incipient mid-Atlantic seafloor spread-ing (McRoberts, 1994; Hallam, 2001).Moreover, the end-Triassic volcanogenic CO2 maximum (Hautmann, 2004)resulted in humid conditions, accelerated weathering,enhanced terrestrial input, and interruption of carbonatesedimentation. In a number of regions there is a gap belowthe base of the Early Jurassic sequences. The ZliechovBasin in theWestern Carpathians is one of the areas wherethe sedimentation was almost continuous during the T–Jboundary interval. The T–J boundary beds here contain arecord of several events which may have been caused byglobal changes.

Hesselbo et al. (2002), Pálfy (2003) and Hautmann(2004) interpreted the enriched atmospheric CO2 andrelated temperature elevation at the end of Triassic as aneffect of volcanism. Similarly, volcanic activity wasproposed by Sturesson et al. (2000) as the main agent inorigin of mixed chamosite and haematite ooids. Suchooid types form lens-like concentrations in erosionalchannel fills in the highest bed of the Fatra Formation inthe Furkaska section (Bed 408) which was exploited asiron ore in the 19th century (Kúšik, 1967).

A peculiar limestone layer, composed predominantly(up to 30–60%) of microspheres, occurs somewhatdeeper in the uppermost part of both the Furkaska andKardolína sequences ((Figs. 3 4 and 6)). These beds onlyrarely (in redeposited form only) contain true ooids thatshow concentric structures. They bear resemblance tosmall vitreous particles found in Cretaceous–Paleogenebeds of the Caribbean region (Stinnesbeck et al., 1999;Fouke et al., 2002; Schulte et al., 2003). Themicrospheresfrom our sections might also have been formed by cal-citization of glass droplets originating from a bolide im-pact, such as that postulated to occur immediately beforethe Rhaetian–Hettangian boundary (Walkden et al.,

2002). From this point of view, the occurrence of the Cand O isotope excursions immediately below the base ofthe spherulite bed deserves special attention.

Volcanic explosions could have also produced tiny glassparticles. Effects of distant volcanic activitymay have led toformation of ferruginous crusts, chamosite oolite ores(Sturesson et al., 2000), and elevated siliciclastic input (e.g.the sandy beds in cycle 14, Figs. 3 and 4). The enhancedatmospheric CO2 content in the atmosphere may havecaused a reduction of calcite saturation of sea water and theconcomitant cessation of carbonate deposition (Galli et al.,2005; Huynh and Poulsen, 2005).

Hydrotermal transformation of coated grains isanother alternative to explain the origin of microspheres.In the T–J boundary beds of the Krížna Nappe, Kúšik(1967) found small, transparent, spherical calciticbodies without concentric laminae but enclosed in asiliceous envelope. He regarded them as ‘‘metasomati-cally transformed” ooids and postulated that the crust ofredeposited and hematitised calcite ooids could havereacted with silica and was transformed into chamosite.The observed high δ18O values could thus suggest theeffect of saline solutions that infiltrate the sediment andcause the replacement of aragonite ooids. This hypoth-esis would be supported by the presence of fluorite, thezoning of dolomite and feldspar crystals, and corrosionof crystals of idiomorphic minerals (Fig. 8).

The occurrence of slumping near the T–J boundary isalso of interest. The position of the slumped beds in theFurkaska sectionmight be considered to be connectedwiththe event that produced the spherules. However, a muchthicker (up to 30 m) submarine slump in the Kardolínasection occurs immediately above the ‘boundary clays-tone’ and comprises material from the uppermost beds ofthe underlying Fatra Formation. Therefore, the slumpingcan be simply explained by rapid sedimentation and relatedslope instability at the margin of an elevated delta plain,rather than by an external event. However, slumped bedsoccur at the top of the Rhaetian sequence not only in theCarpathians, but also at Kendelbachgraben in the Alps andin SW Britain (Cotham and Langport Members, Hesselboet al., 2004). Therefore, the interpretation of slopeinstability in different basins during this time may requirefurther study.

The negative δ13C excursion at the T–J boundary incarbonate and organic carbon (McRoberts et al., 1997;Ward et al., 2001), following a similar but earlier excur-sion (Pálfy et al., 2001; Hesselbo et al., 2002; Guex et al.,2003; Galli et al., 2005) indicates the enrichment of thecarbon reservoir in 12C originating from biogenic orbiochemically fractionated carbon. The couple of nega-tive C isotope shifts could have been produced during a

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massive biotic extinction (Holser, 1997; Hallam andWignall, 1997). The negative shift of carbon isotoperatios in carbonates can be also explained by a volcano-genic rise of atmospheric CO2 (Huynh and Poulsen,2005), or by a massive release of methane from oceanicclathrates (e.g. Pálfy et al., 2001). The samples takenfrom the T–J boundary intervals displaying negativeδ13C values also show slightly increased TOC content.Preliminary results indicate a terrigenous origin of theorganic matter. The postulated release of methane fromhydrates stored in bottom sediments can be induced bysea-level fall, warming, or effects of bolide impact orvolcanic activity. The positive δ18O excursion can indi-cate either changes in the temperature (cf. McElwainet al., 1999), water composition, or diagenetic transfor-mation of the sediment. More positive δ18O values indolomitized rocks in these beds can indicate a diageneticoverprint via influx of saline water.

5. Conclusions

The Fatra and Kopieniec Formations in the ZliechovBasin comprise a record of several environmental changeswhich may have been connected with the global T–Jboundary events. The diversity of benthic fauna decreasesat the base of the ‘transition beds’, i.e. the uppermostmember of the Fatra Formation. The corals, foraminifers,and dasycladacean algae comprise forms that never appearagain in younger strata. Eutrophic conditions and highnutrient levels exerted stress on benthic life and favouredan opportunistic fauna, represented by ostracods andmicrogastropods.

In cycle 13, the oxygen and carbon isotope anomaliesare correlated with lithological changes and the extinctionlevels of microfauna. Themost significant C andO isotopeexcursion is followed by thin microsphere-bearing layerswith peculiar lithological composition. The origin of thesmall, spherical, calcitized particles remains uncertain(bolide impact, volcanic eruption, or resorbtion andsubstitution of primary sedimentary particles). However,these layers are traceable across several tens of kilometres(Fig. 9, see arrows).

In the uppermost part of the Fatra Formation, a signi-ficant terrigenous event, expressed by the first majorincrease in riverine influx is recorded. Continental run-offcaused the eutrophication of the basin and resulted in eli-mination of oligotrophic carbonate platform ecosystems.

The boundary between the Fatra and the Kopieniecformations is sharp, marked by a sudden terminationof carbonate sedimentation, followed by the non-car-bonate ‘boundary claystone’ of the Kopieniec Forma-tion. The carbonate carbon isotope anomaly recorded

from these strata is comparable with those previouslydescribed at the T–J boundary from several other sec-tions worldwide.

Acknowledgments

The authors thank those who helped with field workduring 2002–2004, namely the Management andResearch Centre of the Tatra National Park, and post-doctoral students M. Abelova, P. Klepsatel, M. Golej, A.Tomašovych, and J. Čatloš. J. Spišiak provided hisresults of mineralogical study and microprobe analysesof spherule-like particles. I. Holický made scanningelectron microphotographs. A. Svitáčová and K. Mal-kowski provided analyses of the organic carbon.Drawings were made by E. Petrikova. Comments byjournal reviewers P. B. Wignall and J. Haas and editor J.Pálfy substantially improved the manuscript. This studywas supported by VEGA 6026 grant project and by theAPVV 51-011305 project. This is a contribution to theIGCP 458 project.

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