GEOLOGICA CARPATHICA, 53, 6, BRATISLAVA, DECEMBER 2002
351 — 367
LOWER TRIASSIC SHALLOW MARINE SUCCESSION
IN THE BÜKK MOUNTAINS, NE HUNGARY
and PÁL PELIKÁN
Geological Research Group of the Hungarian Academy of Sciences, Pf. 120, 1518 Budapest, Hungary; email@example.com
Hungarian Geological Institute, Stefánia út 14, 1443 Budapest, Hungary; firstname.lastname@example.org
(Manuscript received December 7, 2001; accepted in revised form June 18, 2002)
Abstract: This paper presents the Lower Triassic sequence in the Bükk Mountains (NE Hungary). On the basis of the
revised macrofossil-collections and previous conodont studies, a biostratigraphic correlation and a chronostratigraphic
subdivision of the formations and members is given. Hindeodus parvus (Kozur et Pjatakova), Isarcicella isarcica
(Hukriede), Claraia aurita (Hauer), Eumorphotis gr. multiformis (Bittner), Eumorphotis cf. hinnitidea (Bittner), Costatoria
subrotunda (Bittner), Eumorphotis kittli (Bittner), Tirolites illyricus Mojsisovics, Tirolites seminudus Mojsisovics,
Costatoria costata (Zenker), Rectocornuspira kalhori Brönnimann, Zaninetti et Bozorgnia, Cyclogyra? mahajeri
Brönnimann, Zaninetti et Bozorgnia, and Meandrospira pusilla (Ho) are the main index fossils. Two sections, which
represent the entire Lower Triassic succession, were studied from a sedimentological point of view. These are the
Gerennavár section for the Gerennavár Limestone Formation, and the Lillafüred section for the Ablakoskővölgy Forma-
tion, respectively. The sedimentological data and facies interpretations are summarized in facies models. During Early
Triassic time the depositional area of the Bükk Mountains were situated in the epeiric shelf of the Western Tethys. The
model of the Gerennavár Limestone Formation reflects the high-energy, tide- and wave-dominated shallow shelf, char-
acterized by a protected—stabilized muddy sand flat and a high-energy sand belt. The model of the siliciclastic, lower part
of the Ablakoskővölgy Formation indicates a mixed, dominantly siliciclastic shallow shelf with deposition in the coastal,
shoreface and transitional zones. The model of the carbonate, upper part of the Ablakoskővölgy Formation refers a
storm-controlled shelf with facies representing the whole spectrum from the peritidal—shallow shoals to the low-energy
deeper subtidal zone below the storm wave-base. Lithologic and facies comparison of the Lower Triassic succession of
the Bükk Mountains to other sequences of the Western Tethyan depositional area reveals many differences, and fewer
similarities, which suggests local controls on depositions, that is locally different terrigenous siliciclastic input, different
subsidence rate, antecedent topography, and dominance of the local physical regime.
Key words: Lower Triassic, NE Hungary, Bükk Mountains, stratigraphy, facies interpretation, facies model, fossils.
Reinvestigations of the Mesozoic succession in the Bükk
Mountains have given rise to much controversy for the last
two decades. It seems that the only exception is the Lower
Triassic part since the basic stratigraphic subdivision by
Schréter (1935, 1953, 1954) and Balogh (1964) is commonly
accepted. The lithostratigraphic subdivision of Lower Triassic
deposits was established by Balogh (1980) and was refined by
Pelikán (1985a, 1995) as a result of a mapping program of the
Hungarian Geological Institute starting in 1979.
The aim of this paper is to summarize the present-day
knowledge of the Lower Triassic sequence in the Bükk
Mountains and to present a general outline of its stratigraphy
and sedimentology. A review of the fossil collections of the
Hungarian Geological Institute provides a possibility for com-
parison of the succession to the Western Tethyan biozonation,
and for the chronostratigraphic subdivision. Sedimentological
studies of the Gerennavár and Lillafüred section have resulted
in more detailed facies interpretations and a better understand-
ing of the facies successions of the Lower Triassic formations
in the Bükk Mountains.
The Bükk Mountains are situated in Northern Hungary,
south of the Inner Western Carpathians. It is a part of the
Bükk Composite Unit (Bükkia Composite Terrane) which be-
longs to the Pelso Megaunit (Pelsonia Composite Terrane)
(Kovács et al. 2000) (Fig. 1). However, Paleozoic—Mesozoic
sequences were deposited in the northwestern neighbourhood
of the Inner Dinarides according to the paleogeographical re-
construction by Protić et al. (2000), and Filipović et al. (in
press). The block of Paleozoic—Mesozoic sequences of the
Bükk Mountains was displaced northeastward to its recent posi-
tion along a large-scale, regional dextral lateral fault-zone
(Mid-Hungarian Lineament) during the Tertiary (summary in
Fodor & Csontos 1998).
The succession of the region comprising the Bükk Moun-
tains is composed of intensely deformed, anchimetamorphic
Mesozoic rocks, surrounded by a non-metamorphic Paleo-
gene-Neogene cover (Árkai 1973, 1983; Csontos 1999). The
Lower Triassic formations are exposed only in the Northern
anticline (details in Less et al. in press). Because of low-grade
metamorphism and strong ductile deformation of the sedi-
352 HIPS and PELIKÁN
mentary rocks, the original macroscopic depositional struc-
tures are only scarcely visible, which makes facies interpreta-
tions more difficult.
The Lower Triassic sequence (Fig. 2) is more completely
represented by the Bálvány section (lower part of the Geren-
navár Limestone), the Gerennavár section (Gerennavár Lime-
stone Formation, Fig. 3), and an overturned but continuous
section at Lillafüred (Ablakoskővölgy Formation, Fig. 4). The
two latter sections were studied in detail from a sedimentolog-
ical point of view and were amplified by observations from
Fossil assemblages and stratigraphy
A review of the available fossils in the collection of the
Hungarian Geological Institute is presented here. Most of the
macrofossils were collected by Legányi and Schréter
(Schréter 1935, 1953, 1954) and Balogh (1964). They have
collected and described relatively rich faunal assemblages
from a few locations. Especially, the surroundings of
Bogdány-tető (at the non-metamorphosed northern limb of the
Northern anticline) yielded ‘numerous’ fossils. According to
Less et al. (in press), the Ablakoskővölgy Formation crops out
there, however, the original stratigraphic position of the mem-
bers was disturbed by complicated tectonic displacements.
Otherwise, there are only scanty data on the occurrence of the
marker species. Based on the descriptions of Schréter (1935,
1953, 1954) and Balogh (1964), and the results of mapping
(Less et al. in press), the major part of the index fossils could
be assigned to lithostratigraphic units, established by Balogh
(1980) and Pelikán (1985a, 1995). However, precise positions
of the markers within the formations or members could not al-
ways be reconstructed. Systematic collection of macrofossils
from the Scythian part has not been carried out since
Schréter’s and Balogh’s collections. The preservation of the
fossils is generally not good.
Foraminifers were determined by Bérczi-Makk (Bérczi-
Makk 1986, 1987; Bérczi-Makk et al. 1995; Pelikán 1995)
and Oravecz-Scheffer (Pelikán 1985a). Kozur (1985, 1988)
studied the conodont and ostracod assemblages near the Per-
Biostratigraphic zonation of the Western Tethyan sections
is mostly based on molluscs (Krystyn 1974; Broglio Loriga et
al. 1983; Neri & Posenato 1985; Broglio Loriga & Mirabella
1986; Broglio Loriga & Posenato 1986; Broglio Loriga et al.
1990; Posenato 1992) and conodonts (Perri 1991; Yin 1993;
Yin et al. 1996; Zhang et al. 1996; Orchard 2001), and partly
foraminifers (Salaj et al. 1983; Broglio Loriga et al. 1990), but
it seems that some taxa of the foraminifers are only facies-in-
dicators (see Hips 1996). However, as in many parts of the
Western Tethyan domain, within the Lower Triassic the
boundary between the Induan and Olenekian stages cannot be
determined due to the simple, low-diversity fauna. A more de-
tailed chronostratigraphic subdivision of the formations and
their members in the Bükk Mountains can be made only on
the basis of the three-fold subdivision (Griesbachian, Namma-
lian, Spathian), as in other Western Tethyan sequences. The
paper does not deal with the problem of the Permian/Triassic
Gerennavár Limestone Formation
The underlying ‘Leptodus member’ of the Upper Permian
Nagyvisnyó Limestone Formation is composed of black, thin-
bedded limestones rich in micro- and macrofossils (Fülöp
1994). The lowermost beds of the Gerennavár Limestone For-
mation indicate a sharp lithological change.
Lithostratigraphic definition. The Gerennavár Limestone
Formation consists of predominantly grey, thick-bedded oo-
lites, and subordinately of thin-bedded mudstones. The low-
er part of the formation is subdivided into two characteristic
horizons. The ‘basal bedset’ of the Gerennavár Limestone
Formation (Figs. 2—3) is approximately 1 m in thickness,
and consists of dark grey clayey marls, followed by sand-
stones, calcareous sandstones, sandy clayey marls, and in a
few sections sandy dolomites. It is overlain by the ‘transi-
tional bedset’ ca. 6—7 m in thickness, which is made up of
thin- to thick-bedded, dark grey, fine laminated limestones.
Above them thin- to thick-bedded mudstones are recorded,
and upwards thick-bedded oolite limestones form the bulk of
the formation ca. 100—130 m in thickness (Figs. 2—3). Local-
ly coarse crystalline, late diagenetic dolomite lenses occur.
Fossils. From the ‘basal bedset’, Csontos-Kis (in Pelikán &
Csontos-Kis 1990) reported the occurrence of bivalves and
Fig. 1. A – Schematic terrane map of the Circum Pannonian Region
(Kovács et al. 2000); 1 – Flysch Belt, 2 – Klippen Belt, 3 – North-
ern Calcareous Alps, 4 – Early Alpine unit related to European con-
tinental margin, 5 – Early Alpine shelf sequences related to the
Apulian (Southern Alps and Outer Dinarides) continental margin,
6 – Ophiolites of the Penninic Ocean, 7 – Ophiolites of the Vardar
Ocean, 8 – Major strike-slip zones; B – Position of the Northern an-
ticline in the structure of the Bükk Mountains (after Csontos 2000).
LOWER TRIASSIC SHALLOW MARINE SUCCESSION 353
brachiopods, but some of the species are not valid, because
the species proposed by Schréter have never been published.
This rich Upper Permian (Changhsingian) macrofauna is char-
acterized by the abundance of Pectinoid bivalves, while bra-
chiopods and gastropods are rarer.
From the ‘transitional bedset’ only microfossils of Chan-
ghsingian age were found, that is foraminifers: Earlandia
dunningtoni (Elliott), E. tintinniformis (Mišik), E. deformis
Bérczi-Makk, Neotuberitina reitlingerae (Mikl. Maklay),
Globivalvulina graeca Reichel, Geinitzia sp., Ammodiscus
sp., Pachyphloia sp. (Bérczi-Makk 1986, 1987; Bérczi-Makk
et al. 1995), Agathammina pusilla (Geinitz), Paraglobival-
vulina mira Reitlinger (Kozur 1988), ostracods: Indivisia
buekkensis Kozur, Goranella sp., Judahella bogschi Kozur;
holothuridea: Theelia dzulfaensis Mostler et Rahimi-Yazd;
and conodont: Ellisonia transita Kozur et Mostler (Pelikán
1985b; Kozur 1985, 1988; pers. commun. in Fülöp 1994).
Kozur (1988) reported the occurrence of the conodont
Hindeodus latidentatus (Kozur; Mostler et Rahimi-Yazd),
which is identical with H. praeparvus Kozur (Kozur 1996),
from the ‘transitional bedset’. Hindeodus parvus (Kozur et
Pjatakova), which is regarded as the index fossil of the basal
Triassic (Orchard 2001), is not reported near the lower bound-
ary of the Gerennavár Limestone Formation. The ‘basal’ and
the ‘transitional bedsets’ contain Permian and longer-life fau-
Above a conodont-free interval the first appearance of the
conodont Hindeodus parvus (Kozur et Pjatakova) with an ad-
vanced form was recorded from ca. 15 m (in the core drilling
Mályinka-8, Pelikán 1985b), and 20 m (Kozur 1988) above
the lower boundary of the formation. Still higher, approxi-
mately from the middle part of the formation, the conodonts
Isarcicella isarcica (Huckriede), Hindeodus parvus (Kozur et
Pjatakova), and Ellisonia aequabilis Staesche, and the ostra-
cod Hollinella tingi (Patte) were encountered (Kozur in Pe-
likán 1985a; Kozur 1985, 1988). Kozur determined Calli-
cythere postiangulata Wei, Liuzhinia parva Wei, Liuzhinia
sp., Bairdia sp., Polycope sp. from the lower half of the for-
mation (in Pelikán 1985b). From the upper half of the forma-
tion, Cyclogyra? mahajeri Brönnimann, Zaninetti et Bozo-
rgnia, Spirorbis phlyctaena Brönnimann et Zaninetti were
also reported (Oravecz-Scheffer in Pelikán 1985a). The bi-
valve Claraia clarai (Emmrich) was mentioned by Schréter
(1953) from the Bálvány section, but only one dubious speci-
men is present in the collection of the Hungarian Geological
Institute. Moreover Claraia aurita (Hauer) and C. gr. aurita
Fig. 2. Stratigraphic subdivision, and biostratigraphic markers of the Lower Triassic sequence in the Bükk Mountains. For comparison
conodont biostratigraphy (only in the upper Changhsingian and Griesbachian) and mollusc biozonation in the Western Tethys are also in-
dicated. Vertical subdivision is time proportional (radiometric data of the Permian/Triassic boundary is the proposal of the Permian/Tri-
assic Boundary Working Group 1999, based on Zhang et al. 1992; Claoué-Long et al. 1991; Renne et al. 1995; Bowring et al. 1998; Met-
calfe et al. 1999, two other data are compiled by Gradstein et al. 1994).
354 HIPS and PELIKÁN
(Fig. 5) occur in the upper part of the formation and in the
transitional beds towards the overlying unit (Schréter 1935,
Biostratigraphic correlation and chronostratigraphy. Cono-
donts from the ‘transitional bedset’ refer to the H. latidenta-
tus—C. meishanensis Zone (see Zhang et al. 1996). Between
the ‘transitional bedset’ and the first occurrence of the ad-
vanced form of Hindeodus parvus, which does not occur be-
fore the Isarcica Zone (Kozur pers. commun.), there is no
marker in the Bükk Mountains. On the basis of the conodont
data, at least a ca. 55—60 m thick mudstone—oolite succession
within the lower half of the formation, upwards from the ap-
Fig. 4. Generalized log of the Ablakoskővölgy Formation in the
Lillafüred section. Legend on Fig. 3. SB: a supposed sequence
Fig. 3. Generalized log of the Gerennavár Limestone Formation in
the Gerennavár section (after Péró 1983). 1 – ‘Basal bedset’; 2 –
pearance of Hindeodus parvus could be correlated to the Isar-
cica Zone. The lower boundary of the C. wangi-griesbachi
Subzone, which is a part of the Claraia Zone and defined in
the Dolomites (Broglio Loriga et al. 1983), coincides roughly
with the appearance of the Isarcicella isarcica (Kozur 1985;
Broglio Loriga et al. 1990). The uppermost part of the forma-
tion yields abundant Claraia gr. aurita allowing us to corre-
late these beds with the lower part of the C. aurita Subzone.
On the basis of the biostratigraphic results, the lowermost
part (‘basal’ and ‘transitional bedsets’) of the formation repre-
sents the uppermost Permian (Changhsingian). The Permian/
Triassic boundary could not be marked at present, further de-
tailed studies are necessary for delineation of the boundary.
The appearance of the advanced form of Hindeodus parvus
and Isarcicella isarcica marks the beginning of the Upper
Griesbachian, and that of the Claraia aurita assigns the begin-
ning of the Dienerian (Nakazawa 1977; Broglio Loriga et al.
1990). There are several hints, which suggest that the forma-
tion comprises the entire Griesbachian and reaches up to the
lower part of Dienerian (lower part of the Nammalian). These
indirect pieces of evidence are 1) the sedimentology, that is
continuous transitions are recorded between the different
lithologies and facies, 2) the biostratigraphy, that is the am-
biguous specimen of ?Claraia clarai (Emmrich) might record
the C. clarai Subzone. Moreover, the Griesbachian/Nammali-
an boundary is detectable in the uppermost part of the forma-
tion, with the appearance of Claraia aurita.
LOWER TRIASSIC SHALLOW MARINE SUCCESSION 355
This formation is ca. 300 m thick, and composed by alterna-
tion of limestones, dolomites, and sandstones—shales. It can be
divided into four members (Figs. 2 and 4), each displaying a
typical lithological and facies development.
The Ablakoskővölgy Sandstone Member is defined by alter-
nation of red and green sandstones, siltstones, and shales 40—
100 m in thickness, but thin or thick intercalated limestone
beds generally occur. There are gradual lithological changes
both from the underlying formation, and towards the overly-
ing member. The lowermost part of the member yields rich
Claraia gr. aurita (Hauer) assemblages. Eumorphotis gr. mul-
tiformis (Bittner), which is incorporated within marly fine-
crystalline limestones, and Costatoria subrotunda (Bittner)
‘Pseudomonotis’ lóczyi Bittner (Fig. 7), which is preserved
within brownish sandstones—siltstones, are reviewed from the
collections of Schréter (1935) and Balogh (1964). The lower
part of this member corresponds to the uppermost Claraia
Zone, that is the upper C. aurita Subzone on the basis of oc-
currence of the index fossil. Otherwise, this member contains
index fossils of the lower part of the Eumorphotis Zone, that
is the E. multiformis and E. hinnitidea Subzones (see Broglio
Loriga & Posenato 1986; Broglio Loriga et al. 1990). Accord-
ing to the biostratigraphic data this member is Nammalian in
age ranging from the lower Dienerian to the top of the Smith-
ian (see Broglio Loriga et al. 1990). The C. aurita Subzone is
attributed to the Dienerian, and the E. multiformis—E. hinniti-
dea Subzones together have been indirectly referred to the
Nammalian, because they are located below the Tirolites cas-
sianus Beds (Broglio Loriga et al. 1990).
The Lillafüred Limestone Member is up to 150 m in thick-
ness and composed of grey limestones dissected by greenish,
yellowish grey marl intercalations. Schréter (1935, 1954) and
Balogh (1964) mentioned Natiria costata (Münster) (Fig. 8A)
from many outcrops from the area of the Northern anticline.
According to the revision of the fauna, the occurrence of Eu-
morphotis kittli (Bittner) shows that (at least the lower part of)
the member belongs to the E. kittli Subzone (see Broglio
Loriga et al. 1990), which suggests Spathian age for this
member. Schréter (1935) collected Tirolites specimens from
the Bogdány-tető location. He described the fossil assemblag-
Fig. 5. Specimens of Claraia gr. aurita from the upper part of the Gerennavár Limestone Formation; left valves, A) and B) from
Nekézseny, Bikkfolyás Valley (51 and 45 in the collection of the Hungarian Geological Institute (cHGI further on)), C) near Bánkút
(cHGI 204). Scale bars have mm subdivision.
Fig. 6. Costatoria subrotunda (Bittner) from Ablakoskővölgy Sand-
stone Member. A – right valve, Mályinka, southwest of Bogdány-
tető 424 (cHGI 74); B – left valve, Mályinka, from the topmost
bed south of Bogdány-tető 424 (cHGI 75). Scale bars have mm
Fig. 7. ‘Pseudomonotis’ lóczyi Bittner, left valve, from Ablakos-
kővölgy Sandstone Member, Mályinka, east of Bogdány-tető
(cHGI 68). M: 1.5
es partly from brownish limestones (identified as Lillafüred
Limestone) and partly from marls with thin limestone beds
(identified as Savósvölgy Marl), both with ‘Tirolites cassian-
us (Quenstedt)’. But specimens in the collection are not differ-
entiated according to two lithologies (to two lithostratigraphic
units), since he recognized and mapped only one ‘Campil’
Member. Thus, after many years these two Tirolites groups
could not be separated. It is supposed that a pile of fragments,
and a couple of whole specimens, Tirolites illyricus Mojsiso-
vics (Fig. 9A) were collected from the Lillafüred Limestone
Member. None of them could be identified as Tirolites cas-
356 HIPS and PELIKÁN
sianus (Quenstedt) in accordance with the taxonomical study
by Posenato (1992).
The Savósvölgy Marl Member consists predominantly of
grey, greenish grey shales, clayey marls, and marls, 40—150 m
in thickness. Thin micritic limestone intercalations are very
characteristic for the entire member. Sand-rich deposits are re-
corded from the western part of the Northern anticline. A rela-
tively rich gastropod fauna is reported by Schréter (1935) and
Balogh (1964) including Natiria costata (Münster), Naticella
Fig. 9. Tirolites from the Lillafüred Limestone and/or Savósvölgy
Marl Members; A – Tirolites illyricus Mojsisovics, Mályinka, east
of Bogdány-tető 424 (cHGI 93); B – Tirolites seminudus Mojsiso-
vics, Mályinka, west of Bogdány-tető 424 (cHGI 82). Scale bars
have mm subdivision.
Fig. 10. Microfacies of crinoidal tempestite layer in the Újmassa
Limestone Member, Lillafüred section – packstones with crinoids
(C), and co-occurrence of Meandrospira pusilla (Ho) (white ar-
rows) and Cyclogyra? mahajeri Brönnimann, Zaninetti et Bo-
zorgnia (black arrows). Sample: 84.
subtilistriata (Frech), and ‘Turbo’ rectecostatus Hauer. Cos-
tatoria costata (Zenker) was found from the thin-bedded
limestones—marls (Schréter 1935; Balogh 1964). In the collec-
tion of the Hungarian Geological Institute, Tirolites seminu-
dus Mojsisovics (Fig. 9B), and Dinarites sp.? are supposed to
have been collected from this member from the Bogdány-tető
location (see Schréter 1935). According to the study by
Posenato (1992) these above-mentioned species represent the
early and middle stages of the tirolitids phylogenetic trend be-
longing to the T. cassianus Zone proposed by Krystyn (1974).
On the basis of the occurrence of Tirolites and Costatoria cos-
tata, this member also belongs to the middle part of the
The Újmassa Limestone Member is composed of dark grey
platy and nodular, strongly bioturbated limestones with marl
and shale intercalations or flasers. In its upper half, the bio-
clastic limestones often alternate with laminated mudstones.
Dolomite lenses, or layers also occur. Its maximum thickness
is ca. 60 m. Rectocornuspira kalhori Brönnimann, Zaninetti et
Bozorgnia and Cyclogyra? mahajeri Brönnimann, Zaninetti
et Bozorgnia were recently recognized in the Bükk Mountains
(Fig. 10). On the basis of the occurrence of Costatoria costata
(Zenker), Meandrospira pusilla (Ho), this member corre-
sponds to the upper part of the C. costata Zone (Broglio
Loriga et al. 1990), and its age is upper Spathian (uppermost
Olenekian). There is a gradual lithological and facies transi-
tion towards the overlying Hámor Dolomite Formation, which
is regarded to be Anisian in age. However, similarly to many
other Western Tethyan sections, there is no evidence for the
Scythian/Anisian boundary in the Bükk Mountains.
Glomospira sinensis Ho, Glomospirella shengi Ho, and
Meandrospira pusilla (Ho) are the most frequent foraminifers
in the carbonate part of the Ablakoskővölgy Formation (Pe-
likán 1995). They refer to Spathian age. Schréter (1935, 1954)
and Balogh (1964) collected many Alpine faunal elements
from outcrops hardly identifiable at present. Most of them
have no chronostratigraphic value within the Lower Triassic,
like Unionites canalensis (Catullo) (Fig. 8B), U. fassaensis
(Wissmann), Bakevellia sp., ‘Pecten’ sp., Neoschizodus laevi-
Fig. 8. A – Natiria costata (Münster) from the Lillafüred section
(cHGI 152); B – Unionites canalensis (Cat.) left valve, from the
Lillafüred section (cHGI 160). Scale bars have mm subdivision.
LOWER TRIASSIC SHALLOW MARINE SUCCESSION 357
Gerennavár Limestone Formation
The lowermost part of the formation (ca. 7—8 m) is different
from the bulk of the formation, especially the lowermost 1 m
(Fig. 3). Clayey marls with fine siliciclastic limestone layers
and sandstones of the ‘basal bedset’ were deposited with a
sharp lithological change from the underlying carbonate suc-
cession. The ‘transitional bedset’ is composed of thin-bedded
dark grey limestones (ca. 6—7 m) characterized by alternations
of microbial lamination and fine even lamination, with silt
and sand-sized biodetritus sometimes normally graded (Figs.
11 and 12).
The predominant part of the formation consists of two types
of limestone: 1) thin- to thick-bedded, dark grey laminated or
massive finely crystalline limestones (mudstones, occasional-
ly wackestones); 2) thick-bedded well-sorted oolites (Fig. 13),
or pure bioclastic grainstones, which are mainly crinoidal
grainstones (Figs. 14—15). The wackestones are composed of
recrystallized microsparitic matrix with fragments of ostra-
cods and bivalves. The thick-bedded oolites are massive, and
there is no visible structure inside the beds. The lack of cross-
bedding could be explained by well-sorting of the ooid sand,
and absence of micrite.
The lower part of the succession is composed of mud-
stones—wackestones, whereas oolite grainstones first appear
ca. 20 m above the top of the ‘transitional bedset’. The oolite
beds become predominant upsection. There is a short, gradual
transition with increasing siliciclastic interlayers towards the
The oolite beds contain ooids, which display different de-
grees of diagenetic alterations. The following ooid-types can
be recognized on the basis of their microfabrics. The first type
contains typical coated grains with multiple tangential con-
centric laminae (Fig. 13.1). This type represents the initial
depositional form. However, the nuclei of these ooids are gen-
erally already neomorphosed. The second type of rounded
grains are fully micritized, thus, they are regarded as peloids
(Fig. 13.2). Grains of the third type, which are recrystallized
into microspars, have rounded surfaces, but no internal struc-
tures (Fig. 13.3). They preserve remnants of micrite enve-
lopes, or micritized patches. Dolomitized grains, the fourth
type, display only the hollows of the ooids (Fig. 16). In some
layers and lenses sand-sized grains have a characteristic yel-
lowish color. The coloration could be due to the weathering of
dolomitized ooids and bioclasts originally most probably
composed of aragonite and/or high Mg-calcite. Otherwise, the
other types of ooid (types of 1—3) are typically grey.
The above-described types represent alteration stages of the
diagenesis. In the first phase micritization took place, and re-
sulted in rounded peloids. The micritization is regarded as a
very early, syndepositional diagenetic processes accom-
plished by micro-organisms (Bathurst 1971; Reid & Macinty-
re 2000). In the second phase most probably the micritized
grains and the nuclei of the ooids were recrystallized. The
triggering processes are interpreted as aggrading neomor-
phism which took place during burial diagenesis (see Bathurst
Dolomitization occurred in the burial diagenetic realm and
represents the third phase of the diagenetic alterations. The
burial diagenesis is indicated by the petrography – the coarse
crystalline baroque dolomite (Fig. 16) was formed by fabric
selective and fabric destructive replacive dolomitization.
Additionally, the pseudo-oolite type (cortoid) of the grain-
stones is very common consisting of rounded bioclasts, most-
ly echinoderm and mollusc fragments (0.2—10 mm), whereas
each grain is coated by a thin micritic envelope (Fig. 15).
These are not strictly coated grains. These micritic crusts rep-
resent alteration of the grain surfaces by boring micro-organ-
isms. In the crinoidal limestones degrading neomorphism, as a
result of burial diagenesis, is recorded (Fig. 15).
Facies interpretation. The lithological change at the base of
the formation most probably also reflects the fundamental en-
Fig. 11. Photomicrograph of fine laminated mudstones from the peritidal facies of ‘transitional bedset’ of Gerennavár Limestone Forma-
tion, Gerennavár section. Note gradation in sample A), and very thin lamination, which most likely refers to the microbial micritization in
sample B). Samples: A – G-2/1295, B – G-2a/2c.
358 HIPS and PELIKÁN
vironmental change during the latest Permian. Thick laminites
of the ‘transitional bedset’ were probably deposited in a re-
stricted, peritidal environment. Crinkled, discontinuous, thin
stromatolite lamination refers to microbial colonization. Grad-
ed planar laminae indicate deposition from waning-flow, or
storm currents. Weak bioturbation is most likely due to the
changeable environmental conditions. Summarizing, the
‘transitional bedset’ suggests a highly stressed environment.
Formation of ooids took place in a shallow subtidal zone,
where the energy of water was sufficiently high to keep the
sand grains in permanent movement. Studies of modern envi-
ronments (summary e.g. in Halley et al. 1983; Tucker &
Wright 1990), have shown that oolite sand bodies are associ-
ated with bank, platform or shelf margins, which are either
tide-dominated or windward open, storm- and tide-dominated.
The recent analogies suggest that the Lower Triassic oolite
shoals were initiated on topographic highs whereas skeletal
sand bodies were deposited at first, since bioclastic grainstone
layers underlie the oolite succession.
Thick successions of oolite beds suggest amalgamation of
the shoal lobes in the major part of the Gerennavár Limestone
Formation. In sand belts, grains in the shield zone are com-
monly reworked to wide fans, which migrate towards the
‘platform interior’ during storms (e.g. Lilly Bank, see Ball
Fig. 12. Thin, irregular lamination in microbial mat, peritidal fa-
cies in the upper part of the ‘transitional bedset’ of Gerennavár
Limestone Formation, Bálvány section.
Fig. 13. Typical microfacies of oolite grainstones from the Geren-
navár Limestone Formation with the types of ooids: 1 – ooid with
recrystallized nucleus and a number of concentric, thin micrite
crusts, 2 – micritized ooid, 3 – entirely recrystallized ooid with
only one micrite envelope. Note the elongation of ooids according
to the shear stress. Sample: G-1/19.
Fig. 15. Photomicrograph of packstones with recrystallized, round-
ed crinoid fragments which preserved their thin micrite envelopes,
Gerennavár Limestone Formation, Gerennavár section, sample G-1/
15. Note the result of degrading neomorphism (burial diagenesis).
This means that the crinoid fragments are composed of not only one
but a few spars.
Fig. 14. Photomicrograph of bioclastic grainstones with mollusc
shell fragments, Gerennavár Limestone Formation, Gerennavár
section, sample G-2a/10a.
LOWER TRIASSIC SHALLOW MARINE SUCCESSION 359
1967). Another example of oolite formation is the Joulter’s
Cays, where the ooids were deposited in lobe-shape fans de-
veloped at the ends of tidal channels, through which the ooids
were transported from the shoal, and at bank spillovers (Harris
1979). The oolite deposits of the Gerennavár Limestone For-
mation could be compared to the above mentioned oolite sand
bodies. Thin-bedded mudstone interlayers within the thick-
bedded oolite grainstones are interpreted as deposits of major
storms similar to those described from the Bahamas by Shinn
et al. (1993), and by Major et al. (1996). The cortoid contain-
ing beds may have formed shallower flanks of the oolite
shoal, where microbes altered the grain surfaces during calm
The interlayering mudstones—wackestones in the lower and
upper part of the formation most probably represent a low-en-
ergy subtidal facies. The facies of the beds, where scattered
ooids embedded within micrite matrix in the upper part of the
formation, could be interpreted as shallow subtidal stabilized
sand flat bankward to the shoal. Its mixed muddy peloidal—
ooidic sand deposits and facies could be compared to the
modern Joulter’s area, in the Bahamas (see Harris 1979).
The poor fossil record of this formation could be mainly the
result of the depopulation effect of the Permian/Triassic ex-
tinction, and partly due to the disadvantageous conditions for
the population in the niche of a high-energy mobile sand belt.
Otherwise, the preservation potential must also have been low
in this environment because of the disintegration and frag-
mentation of skeletons. This could explain the missing marker
of the Claraia clarai Subzone.
In summary, the following vertical facies changes are re-
corded in the succession of the formation: a restricted very
shallow marine environment at the base of the formation (per-
itidal), which evolves upwards into a low-energy shallow sub-
tidal (lagoon), and then into a high-energy intertidal—shallow
subtidal (oolite shoal) environment. At the upper part, the
mixing of the muddy and peloidal—oolitic sand refers to de-
creasing water energy (stabilized sand flat).
Ablakoskővölgy Sandstone Member
Fine siliciclastics prevail with intercalations of carbonate
layers (Fig. 4). The following litho-types are observed in the
siliciclastic bedsets (not in stratigraphic order): 1) reddish
brown shales; 2) alternating greyish green, rarely brown, par-
allel laminated clayey or carbonate siltstones and silty marls;
3) alternating greyish green sand-streaked siltstones, and par-
allel, or sometimes cross-laminated thin-bedded, mica-rich
sandstones. The mature sandstones contain predominantly
detrital quartz grains and additionally detrital feldspar up to
m in size. Clay minerals are illite, and rarely kaolinite.
Because of low-grade metamorphism and strong schistosity,
the original sedimentary structures of this member are hardly
The thick-bedded limestones (wackestones—packstones)
contain red ooids or fragments of ooids and detrital quartz
grains. This lithofacies is characterized by reddish staining
due to the high Fe-content. The thin-bedded grey or beige
limestones are typically mudstones, generally recrystallized
microspar with much less siliciclastics, and often with juve-
nile pelecypod shells. The laminated or thin-bedded mud-
stones typically alternate with greenish grey calcareous silt-
An upward thickening and coarsening stacking pattern
characterizes the major part of this member. However, an up-
ward fining trend and darkening of colors becomes prevalent
in the upper part and the transitional interval towards the over-
lying carbonate member. With upwardly increasing carbonate
content this member continuously evolves into the overlying
Facies interpretation. Sediment deposition most likely took
place mainly in the coastal, shoreface and transitional zones of
a shallow shelf. In the siliciclastic succession, the oolite grain-
stones are intercalated as distinct thick beds, which refers to
Fig. 16. Photomicrograph of dolomitized ooids in microspar matrix. Late diagenetic baroque dolomite partly, or fully replaced the ooids.
Note zonation and curved crystal surface of the baroque dolomite (e.g. B – lower left), and the preserved round surface of the host grains
(B – two uppers). Gerennavár Limestone Formation, Gerennavár section, samples: A – G-1/5, B – G-1/22b.
360 HIPS and PELIKÁN
tern of the basal transitional and lower parts may refer to a rel-
ative sea-level rise. The crinoidal wackestones—packstones
represent storm events between fairweather and storm wave-
base which are generally rare in the lower part and become
more frequent upwards. Thin lenses composed of washed
crinoidal fragments suggest patchily distributed thin storm ve-
Shoals consisting of washed and well-rounded crinoidal
and ooid sand represent deposition around the fairweather
wave-base. Their position in the succession reflects highstand
progradation. Mudstones above the oolites were probably de-
posited in a shallow peritidal zone and subjected to dolomiti-
zation afterwards. Thin uneven, laminated dolocrete crusts in-
side the dolomite body most probably indicate subaerial
exposures. Early diagenetic near surface dolomitization by
relative sea-level drop is indicated by pervasive finely crystal-
line stratiform dolomites and the associated erosional surface
on top of the succession. In the upper part of the member, a
the reworking of ooids most probably by storms. They were
either fed by temporal, patchily distributed shoals, or by thin
tidal delta lobes. However, the in situ cross-bedded deposits
of the oolite shoal or tidal delta were not recognized.
Lillafüred Limestone Member
This member can be subdivided into three parts (Fig. 4).
The lower part shows a gradual transition from the underlying
member with upwardly increasing carbonate content. The
bulk of the lower part is composed of dark grey, thin-bedded,
partly thick-bedded finely crystalline limestones (mudstones—
wackestones) and marly bedsets. The latter are either dark
grey calcareous marls or greenish brown clayey siltstones oc-
casionally with thin platy mudstones. Lenses, or thin- to
thick-bedded bioclastic limestone beds, generally crinoidal
wackestones—packstones, are more abundant upwards. A
thickening-upward trend up to the middle part of the member
The middle part of the member consists of upward thicken-
ing and coarsening bedsets. In the basal part dark grey lami-
nated or thin-bedded mudstones alternate with thick-bedded
crinoidal limestones separated by marl layers. Thick-bedded,
well rounded and well sorted crinoidal—oolite limestones
(grainstones—packstones) occur upsection (Fig. 17). The se-
quence is terminated by a dolomite cap consisting of massive,
finely crystalline beds with occasionally preserved thin irreg-
ular lamination (dolocrete crusts). Its top is truncated by an ero-
Above the erosional surface, dark grey crinoidal limestones
with intraclasts and peloids occur which form the base of the
upper part of the member (Fig. 18). Upsection, upward fining
and thinning bioclastic, crinoidal limestones (packstones),
and mudstones with an upwardly increasing number of marl
intercalations are characteristic. Crinoidal limestones (pack-
stones) occasionally exhibit ripples on the bedding surfaces
Facies interpretation. The lower part of the member was
deposited in the deeper subtidal zone, mostly below the storm
wave-base indicated by the fine deposits and the dominant
thin bedding. The thinning- and fining-upward stacking pat-
Fig. 18. Sharp surface (at the head of hammer) truncating the mas-
sive, light grey dolomite succession, peritidal facies, is overlain by
dark grey intraclastic crinoidal limestones, shoal facies, in about the
mid-part of the Lillafüred Limestone Member, Lillafüred section
(overturned section but the photograph is rotated into the original
depositional position). Hammer is 33 cm long.
Fig. 17. Photomicrograph of crinoidal—oolite limestones, shoal fa-
cies, from the mid part of Lillafüred Limestone Member, Lil-
lafüred section. Sample: 55.
Fig. 19. Ripples in crinoidal limestones, shoals facies, in upper
part of the Lillafüred Limestone Member, Lillafüred section. Coin
is 2.5 cm in diameter.
LOWER TRIASSIC SHALLOW MARINE SUCCESSION 361
thinning- and fining-upward succession of crinoidal lime-
stones and mudstones—marls recorded an upwardly deepening
trend from shallow subtidal below the storm wave-base. This
is most likely due to a relative sea-level rise.
In summary, the vertical stacking pattern of the facies in
this member exhibit a deepening-, followed by a shallowing-,
and then a deepening-upward trend. Considering the thickness
of the deepening—shallowing-upward succession (ca. 75 m) in
the lower part of the member, this trend most likely suggests
third-order relative sea-level changes. A sequence boundary is
recognized on top of the shallowing-upward cycle, that is on
top of the finely crystalline massive dolomites, approximately
in the middle part of this member (Fig. 4).
Savósvölgy Marl Member
This member is characterized by monotonous grey, green-
ish grey shales, clayey marls, and marls (containing gastro-
pods and ammonites) with frequent, thin platy mudstone in-
tercalations (Fig. 4). Sandier development was reported in the
western part of the Northern anticline.
Facies interpretation. Lack of wave- and storm-induced
structures, and the relatively rich ammonite fauna in the suc-
cession suggest that the deposition took place in an open shelf
environment, most probably below the storm-wave base. It is
an open question what was the control of terrestrial siliciclas-
Újmassa Limestone Member
This member (Fig. 4) is characterized by platy, thin-bedded
(1—4 cm), dark grey nodular limestones, mudstones—wacke-
stones—packstones (Fig. 10). Their bedding surfaces are often
covered by clayey marl flasers (Fig. 20). The intensity of bio-
turbation varies from minor individual burrows to strong bio-
turbation, and related nodular structure. The burrow-mottled
pattern resembles Thalassinoides-type burrows. Beds with
only minor bioturbation preserved their original depositional
structures, for example normal gradation in thin skeletal lens-
es or layers, or faint planar lamination becoming more com-
mon upwards. Dolomite lenses, or layers appear at the transi-
tion zone to the overlying formation.
Facies interpretation. The dark colour of the rocks and faint
planar lamination refer to partial restriction near the bottom.
Lack of wave structures and common graded bedding indi-
cates that this succession was composed of a series of distal
tempestites deposited within a muddy sequence below the
Facies model and correlation of sedimentary units
The Alpine sedimentary history of the Dinarides—Alps—Car-
pathians started in the middle Permian with terrigenous clas-
tics of alluvial plains. First marine ingressions reached the
southeastern domains (according to the recent co-ordinates) of
the Western Tethys, and cyclic sabkha, and then subtidal
open-marine ramp carbonates of ‘Bellerophon-type’ forma-
tions were deposited in the late Permian (e.g. Nagyvisnyó
Limestone Formation in the Bükk Mountains). A renewed
long-term sea-level rise shifted the coastline further west—
northwest, and the coastal zone of the shallow sea reached
the western part of the Northern Calcareous Alps in the late
Scythian (Gwinner 1971; Tollmann 1976; Mostler & Ross-
ner 1984; Broglio Loriga et al. 1990; Krainer 1993; Rüffer
& Zühlke 1995).
The distribution of carbonate and siliciclastic sediments
temporally and laterally changed on the extensive shallow
seafloor of the Western Tethys. On a large-scale a general lat-
eral trend can be recognized in the decreasing amount and
grain size of siliciclastics from the coastal areas towards the
shelf interior. However, the temporal distribution exhibited
different trends in different facies domains, but generally the
siliciclastic influx intensified during the Dienerian and Smith-
ian in many depositional areas.
According to the Paleozoic-Mesozoic paleogeographic re-
construction the depositional area of the Bükk Mountains was
in the northwestern neighbourhood of the Inner Dinarides,
next to that of the Jadar Block (Yugoslavia), and the Sana—
Una Unit (Bosnia-Herzegovina) (Protić et al. 2000). During
the Early Triassic very similar sequences were deposited in all
three units, which is reflected in very uniform lithological and
facies development of time equivalent formations. These three
units together composed a uniform facies domain during the
In the depositional area of the Bükk Mountains, siliciclastic
sedimentation prevailed during the Nammalian, otherwise
carbonates deposited. The disposition of sedimentary environ-
ments was controlled by the position of the fairweather and
storm wave-base. On the basis of the vertical hierarchy of
lithofacies, three characteristic parts can be distinguished.
They are represented by the Gerennavár Limestone, and the
siliciclastic and carbonate parts of the Ablakoskővölgy For-
mation. On the basis of grain types such as ooids, and early
diagenetic processes (dolomitization) the sediments may have
been generally deposited under arid—semi-arid climate. Con-
trols of the periodic siliciclastic input might have been either
Fig. 20. Dark grey platy, nodular limestones, restricted outer ramp
facies of the Újmassa Limestone Member, Lillafüred section
(overturned section but the photograph is rotated into the original
depositional position). Coin is 2.5 cm in diameter.
362 HIPS and PELIKÁN
the climate (temporal humid periods) or tectonic uplifting of
The global absence of reef-builders is a general characteris-
tic of the Lower Triassic sequences (Heckel 1974; James
1984; Flügel 1982). The faunal assemblage is generally char-
acterized by cosmopolitan, simple, opportunistic forms,
which belong to the heterozoan association (a term of James
1997), and reflects a mass extinction aftermath (Schubert &
Bottjer 1995). They indicate that stressful environmental con-
ditions persisted during Scythian time.
Gerennavár Limestone Formation
Peritidal microbial and fine evenly laminated mudstones,
mudstones—wackestones of a shallow subtidal lagoon envi-
ronment, and bioclastic-oolite limestones of a high-energy
subtidal zone characterize the Gerennavár Limestone. The oo-
ids were formed in the high-energy, permanently agitated
surge-zone and transported into shallower (stabilized sand
flat) environments. The ooid-dominated ramps in general
probably reflect low rates of biological carbonate production
in shallow-water environments, which were widespread in pe-
riods when shallow-water framework reefs were globally ab-
sent or scarce as during the Mississippian and Jurassic
(Wright & Faulkner 1990; Burchette & Wright 1992).
The depositional interval of the Gerennavár Limestone For-
mation can be correlated to that of the Tesero Horizon,
Mazzin Member, Andraz Horizon, and about the lower half of
the Siusi Member in the Dolomites (see Broglio Loriga et al.
1983). The differences are more obvious than the similarities
in comparison of the two successions. The uniform develop-
ment of mudstones and ooidic grainstones is a characteristic
feature of the succession of the Gerennavár Limestone in the
Bükk Mountains, whereas alternations of variable limestones,
early diagenetic dolomites, marls, and fine siliciclastics, make
up the succession in the Dolomites. The lithology of the latter
one is more marly and rich in detrital clay and silt.
Similarities can hardly be found in the lower units of the
two areas, which overlie the Bellerophon-type upper Permian
carbonates, that is siliciclastics of the ‘basal bedset’ and fine
evenly and microbially laminated mudstones (Bükk Moun-
tains) and oolite grainstones—packstones of the Tesero Hori-
zon, or offshore mudstones in the Cadore, Comelico, and
Southwest Carnia areas (Dolomites).
Mostly muddy lithology characterizes the time equivalent
succeeding parts of the sequences, that is in the Gerennavár
Limestone above the ‘transitional bedset’, up to the first ap-
pearance of ooids (ca. 20 m in thickness)(see Fig. 3), which
can be correlated with the lower part of the Mazzin Member.
However, the sedimentary structures and their facies are dif-
ferent, since deposition occurred in low-energy environments,
in a lagoon in the area of the Bükk Mountains, but on a deeper
subtidal shelf in the area of Dolomites. The upward succeed-
ing time equivalent part of oolite limestones of the Geren-
navár Limestone in the Isarcia Zone (ca. 45 m in thickness)
and the upper part of the Mazzin Member are also rather dif-
The facies developments of the overlying part in the se-
quences display great differences, as well. High-energy oolite
shoal and stabilized sand flat is interpreted for the major part
of the Gerennavár Limestone. Because of the highly monoto-
nous development, trends were not recognized in the oolite
sequence. This is partly due to the fact, that the system had a
potential to accrete vertically and therefore the shallow marine
environment could have persisted for a relatively longer time.
There is no analogy for the Andraz Horizon for correlation in
the sequence of the Bükk Mountains. However, the lower unit
of the Siusi Member consists partly of oolites, but their colour
and bedding is rather different. Lithology and facies develop-
ment of the lower half of the Siusi Member is also quite differ-
ent from the equivalent part of the Gerennavár Limestone For-
mation, except for the Trento and Valsugana area. In these
paleohighs, the oolitic-bioclastic bodies several metres in
thickness occur at various levels within the Siusi Member.
order depositional sequence has been recognized
within the uppermost Bellerophon Formation, the Tesero Ho-
rizon and the Mazzin Member, and the self margin wedge and
the transgressive systems tract of another one in the Andraz
Horizon and the lower part of the Siusi Member in the Dolo-
mites (Broglio Loriga et al. 1983; Neri 1991; De Zanche et al.
1993). Facies changes within the lower and upper part of the
Gerennavár Limestone Formation mark a shift between low-
energy and high-energy environments, with no significant
change in depositional depth. As a consequence the individual
depositional sequences of the Dolomites cannot be correlated,
which is most likely due to local controls dominated on sedi-
mentation in the depositional area of the Bükk Mountains.
The south-easterly zones of the Northern Calcareous Alps
and Inner Western Carpathians were reached by the sea dur-
ing the latest Permian—earliest Scythian. The deposition on
the shallow shelf, for example in the tidal flat, coastal zone
and shallow shoreface zone, is controlled by the strong detri-
tal siliciclastic input. Evaporite facies near the Permian/Trias-
sic boundary, that is the Haselgebirge Formation in the North-
ern Calcareous Alps, Perkupa Evaporite Formation in the
Silica Nappe, Inner Western Carpathians, and the succeeding
siliciclastics, that is the Werfen Schist, and Bódvaszilas Sand-
stone Formations are rather different from the time equivalent
Gerennavár Limestone Formations in the Bükk Mountains
(see Tollmann 1976; Mostler & Rossner 1984; Kovács 1992;
Ablakoskővölgy Sandstone Member
During the deposition of the Ablakoskővölgy Formation
the terrigenous clastic input increased in two periods, in the
Nammalian and mid-Spathian. Sediment supply was drastical-
ly changed during the Dienerian (in the Claraia aurita Zone) in
the depositional area of the Bükk Mountains. Various marine
sediments, clastics and carbonates, characterized many West-
ern Tethyan depositional areas during the Nammalian, as in
the upper part of the Siusi Member, Gastropod Oolite Mem-
ber, and Campil Member in the Dolomites (Broglio Loriga et
al. 1983; Broglio Loriga et al. 1990). The terrigenous input
achieved a peak during Smithian defined as the ‘Campil
event’ in the Dolomites (Italy) and in the Transdanubian Mid-
Mountains (Hungary) (Broglio Loriga et al. 1990). In the dep-
ositional area of the Bükk Mountains the temporal distribu-
LOWER TRIASSIC SHALLOW MARINE SUCCESSION 363
tion of siliciclastic supply was rather steady. Siliciclastic de-
posits unambiguously predominate over the carbonates, al-
though limestone intercalations are recorded in the entire suc-
cession as well as in the Smithian (Ablakoskővölgy Sand-
stone Member). In the Outer Dinarides (see Aljinović 1995),
Northern Calcareous Alps (see Mostler & Rossner 1984) and
Inner Western Carpathians (see Bystrický 1964, 1973; Salaj et
al. 1983; Hips 1996, 1998) the clastic input was continuously
significant and inhibited carbonate production in the course of
the early Scythian time.
The similarity all over the Western Tethyan depositional
area is that the intense siliciclastic input seems to be terminat-
ed at the beginning of the Spathian (see Bystrický 1964, 1973;
Salaj et al. 1983; Herak et al. 1983; Sćavničar & Šušnjara
1983; Mostler & Rossner 1984; Broglio Loriga et al. 1990;
Michalík 1994; Aljinović 1995; Hips 1996, 1998). By late
Scythian—earliest Anisian time the fine terrigenous compo-
nents almost completely disappeared and consequently car-
bonates and evaporites were deposited as in the Hámor Dolo-
mite Formation of the Bükk Mountains.
The time equivalent sedimentary units of the Ablakos-
kővölgy Sandstone Member are the upper part of the Siusi
Member, Gastropod Oolite Member, and Campil Member in
the Dolomites. They have generally similar lithological and
facies developments with the consideration that the red oolite
beds are distributed throughout the entire Ablakoskővölgy
Sandstone Member, and they do not compose a distinct mem-
ber as in the Dolomites (Gastropod Oolite). Depositional se-
quences recognized in the Dolomites, that is the highstand sys-
tems tract in the upper Siusi Member and lower Gastropod
Oolite, and a transgressive systems tract and a highstand sys-
tems tract, a part of the succeeding sequence in the majority of
Gastropod Oolite and Campil Members were not recorded in
the Bükk Mountains.
From Lillafüred Limestone to Újmassa Limestone Member
The carbonate stage of the Ablakoskővölgy Formation from
the Spathian clearly indicates predominantly lower energy
depositional environments, whereas finely crystalline lime-
stones and marls prevail in the deeper ramp setting. However,
while distal storm layers are relatively abundant in intervals,
deposits of the high-energy shallower environment as bioclas-
tic beds and thicker beds of partly dolomitized oolite—crinoi-
dal ‘shoals—amalgamated storms sheets’ appear only in a rel-
ative short interval in the middle part of the sequence. The
lithological and facies characters of the Lillafüred Limestone
correlate well to the Val Badia Member except for its basal
peritidal horizon and uppermost cross-bedded sandstone
In the lower part of the Lillafüred Limestone, the vertical
facies arrangement may refer to a relative sea-level rise. The
striking transgression at the beginning of the Spathian with
the appearance of Tirolites sp. was recorded not only in the
whole Western Tethys, that is in the Southern Alps (Broglio
Loriga et al. 1990), Balaton Highland (Budai & Haas 1997),
Drau Range (Krainer 1987, 1993), Northern Calcareous
Alps (Mostler & Rossner 1984), Carpathians (Patrulius et al.
1971; Hips 1998), Dinarides (Herak et al. 1983; Aljinović
1995), but also from the Peri-Tethyan areas (Bleahu et al.
1994), and in the western USA (Schubert & Bottjer 1995)
and Western Canada Sedimentary Basin (Davies 1997).
The stacking pattern of the Lillafüred Limestone Formation,
except for its lower part, displays a regressive—transgressive
trend with a sequence boundary at the top of the finely crys-
talline dolomites approximately in the middle part of this
member. Because of the lack of detailed biostratigraphy in the
sequence of the Bükk Mountains, reliable correlation to the
depositional sequences of the Dolomites (De Zanche et al.
1993; Rüffer & Zühlke 1995) is not possible.
The time equivalent formations in the uppermost part of the
Scythian sequences exhibit major differences in terms of li-
thology and sedimentology between the Bükk Mountains and
the Dolomites. Only a few similar features are recognized.
Small-scale sequences consist of the Cencenighe Member in
the Dolomites, with open shelf muddy basal part and tide-con-
trolled oolitic-bioclastic calcarenites, and siltstones—marls of
supratidal mud flat (Broglio Loriga et al. 1983, 1990). Where-
as no small-scale sequences are recognized within the deep
ramp mudstones—marls of the Savósvölgy Marl Member in
the Bükk Mountains. On the one hand, the open shelf muddy
basal parts of the small-scale sequences of the Cencenighe
Member display a similar lithological character to the mud-
stones—marls of Savósvölgy Marl Member. On the other hand,
the rest of the small-scale sequences of the Cencenighe Mem-
ber are markedly different in every respect from the succes-
sion of the Savósvölgy Marl. Moreover, no correspondence
between the peritidal varicoloured fine siliciclastics—dolo-
mites of San Lucano Member (Dolomites) and the deep ramp
distal tempestites—mudstones of the Újmassa Limestone
Member (Bükk Mountains). Nevertheless, the Újmassa Lime-
stone closely resemble the time equivalent Szinpetri Lime-
stone Formation in the Silica Nappe of the Inner Western Car-
pathians (see Hips 1998; Koša & Janočko 1999) since both
have the same bio- and lithofacies. According to Michalík
(1994) the strong restriction of sea-water circulation between
the extremely shallow seas in the depositional area of the
Western Carpathians caused hyper-salinity and low oxygen
regime during the Spathian/Anisian boundary time interval. A
similar situation most likely occurred in the area of the Bükk
Epeiric shelf and ramp models
On the basis of the comparison of time equivalent facies of
the Dinaridic—Alpine—Carpathian depositional areas of West-
ern Tethys (the above-mentioned examples) two similar facies
models can be obtained for two intervals of the Early Triassic.
They are on the one hand an epeiric shelf model (see Pratt &
James 1986; Wright & Burchette 1996) for the Griesbachian—
Nammalian, and on the other hand a carbonate ramp model
(see Ahr 1973; Read 1985; Burchette & Wright 1992) for the
The Western Tethyan depositional area was initiated as a
ramp during the middle-late Permian, but evolved into an
epeiric shelf during the Early Scythian, when the sea flooded
364 HIPS and PELIKÁN
more territories on the wide continental shelf (see Mostler &
Rossner 1984; Krainer 1993; Rüffer & Zühlke 1995). As a re-
sult, most probably very shallow sea-water covered the exten-
sive shelf area in the course of the Griesbachian and Namma-
lian. Storms and winds strongly influenced the circulation and
deposition (see e.g. Aljinović 1995), but tidal activity was
also recognized locally (e.g. in the formation of longer life oo-
lite shoals, Bükk Mountains).
Although there are lithological variations in the characters
of the sequences of the different depositional areas, which are
partly carbonate, or pure siliciclastic in periods, but mainly
mixed carbonate—siliciclastic, the successions are generally
represented by either dominantly shallow subtidal facies, or
shallowing-upward cycles from shallow subtidal to intertidal,
and supratidal deposits. These cycles are well documented for
example in the Outer Dinarides (Sćavničar & Šušnjara 1983;
Aljinović 1995), and in the Dolomites (Broglio Loriga et al.
1983). These cycles appear to represent deposition on and
around scattered, ephemeral, low-relief shoals whose shore-
line prograded into the surrounding shallow subtidal areas to
generate areally restricted shallowing-upward small-scale cy-
cles. This type of epeiric platform model was described from
the Ordovician by Pratt & James (1986).
Subsidence over such a large area was most probably not
uniform. The different subsidence rate could be an explana-
tion for the difference of time equivalent formations, for ex-
ample between the Mazzin Member (Dolomites) and the oo-
lites of Gerennavár Limestone. Additionally, the different
local physical regime was most likely the other control factor,
for example during mid-Griesbachian the low-energy envi-
ronment in the facies domain of the Dolomites and high-ener-
gy, tide- and wave-dominated environment in the facies do-
main of the Bükk Mountains.
A prominent feature of the sequences is a characteristic
deepening facies trend as a result of a relative sea-level rise at
the beginning of the Spathian. As a consequence a deeper
subtidal environment mostly below or around the storm wave-
base was the site of deposition in the marine settings and fur-
ther zones were reached by the sea. Generally this event was
coupled with increase of carbonate content in the sediments
and decrease of the grain size of detrital clastics. However, the
northwestern margin of the basin is characterized by stronger
terrigenous influx (Mostler & Rossner 1984). Different de-
grees of storm redeposition are widely recognized in the se-
quences, thus a major part of the shelf was storm-controlled
(e.g. Broglio Loriga et al. 1990; Aljinović 1995; Hips 1998).
Generally in the Western Tethyan sequences, appearance of
oolite bedsets above a deeper shelf facies reflects a shallow-
ing-upward trend in the successions. These oolite shoals
framed a marine basin extending to the northwest with normal
salinity (model by Mostler & Rossner 1984). During the late
Scythian, the sedimentation continued in the inner ramp zone
with carbonates and partly evaporites in the depositional areas
of the Dolomites, Northern Calcareous Alps, and Transdanu-
bian Mid-Mountains of Hungary, whereas thin-bedded car-
bonates deposited in the deep ramp zone around or below the
storm wave-base in the depositional areas of Bükk Mountains,
Silica Nappe in Hungary and Slovakia, and part of the Outer
Dinarides. Most probably the facies differentiation was con-
trolled by the different subsidence. Appearance of the basinal
red nodular limestones in the Hellenides (Chios section, Gaeta-
ni et al. 1992) indicated the early stage of a rifting (Mostler &
Rossner 1984). As a consequence of the accelerated subsid-
ence in some part of the Western Tethyan depositional area a
gentle slope was generated.
Biostratigraphic correlation of the sequence in the Bükk
Mountains, as in other Western Tethyan sections, is mostly
based on molluscs, bivalves and partly on conodonts and fora-
minifers, and from the Spathian on ammonites. Several index
fossils have been identified from the Bükk Mountains: Hind-
eodus parvus (Kozur et Pjatakova), Isarcicella isarcica
(Huckriede), Claraia aurita (Hauer), Eumorphotis gr. multi-
formis (Bittner), Eumorphotis cf. hinnitidea (Bittner), Costa-
toria subrotunda (Bittner), Eumorphotis kittli (Bittner), Tiro-
lites illyricus Mojsisovics, Tirolites seminudus Mojsisovics,
Costatoria costata (Zenker), Rectocornuspira kalhori Brönni-
mann, Zaninetti et Bozorgnia, Cyclogyra? mahajeri Brönni-
mann, Zaninetti et Bozorgnia, and Meandrospira pusilla
On the basis of the revised fossil-collections of Schréter
(1935) and Balogh (1964), and conodont studies by Kozur
(1985, 1989, 1996) the Lower Triassic sequence of the Bükk
Mountains was subdivided on the basis of the biostratigraphic
zonation proposed in the Dolomites by Broglio Loriga et al.
(1983, 1990), and Yin et al. (1996), Zhang et al. (1996). The
generally scanty occurrence of index fossils allowed only a
basic comparison to the biozonation proposed in the Western
Tethyan areas. Using this zonation a chronostratigraphic sub-
division of the formations and members is assigned. The Ge-
rennavár Limestone Formation contains the Isarcica conodont
Zone (with advanced form of Hindeodus parvus) and C. aurita
Subzone, (C. clarai Subzone is ambiguous because of scarce
specimens), thus this formation is Griesbachian and lower Di-
enerian in age. In the Ablakoskővölgy Formation the Ablako-
skővölgy Sandstone Member contains fossils of the C. aurita
Subzone, and Eumorphotis gr. multiformis—E. hinnitidea Sub-
zones (with Costatoria subrotunda), thus its age is Nammali-
an. In the Lillafüred Limestone and Savósvölgy Marl Mem-
bers, ammonites of the Tirolites cassianus Zone are
recognized, thus they are Spathian in age. The Újmassa Lime-
stone Member contains Costatoria costata, Rectocornuspira
kalhori—Cyclogyra? mahajeri and Meandrospira pusilla,
which refer to the upper Spathian age.
The Lower Triassic succession can be subdivided into three
stages. The first stage is represented by the Gerennavár Lime-
stone Formation and reflected a) partly, in the lowermost part
in low-energy peritidal and shallow subtidal lagoon facies,
and b) predominantly in a typical high-energy, tide- and
wave-dominated sand belt and shallow subtidal sand flat. The
second stage includes the siliciclastic, lower part of the Abla-
koskővölgy Formation. Since the terrigenous clastic supply
increased in this period this is a mixed, but predominantly si-
liciclastic shallow shelf with sedimentation in the coastal,
shoreface and transitional zones. The third stage is represent-
LOWER TRIASSIC SHALLOW MARINE SUCCESSION 365
ed by the carbonate, upper part of the Ablakoskővölgy Forma-
tion, and reflected in a storm-controlled sedimentation in the
whole spectrum of environments from peritidal to deeper sub-
tidal below the storm wave-base.
Although, the sedimentation of the ‘Werfen Formation’ in
the Western Tethyan area occurred generally on a gently slop-
ing shallow, epeiric shelf, and partly on a ramp at the end of
the Scythian, the deposits and the facies can be very different
in the time equivalent sequences. It suggests the importance
of local control factors in sedimentation, that is terrigenous si-
liciclastic input, different subsidence rate, antecedent topogra-
phy, and dominance of local physical regime. The Lower Tri-
assic succession in the Bükk Mountains, in term of
sedimentology, displays many lithological and facies differ-
ences relative to that of the Dolomites. Nevertheless, close
similarities can only be documented in short intervals between
sequences of different paleogeographical position.
Acknowledgments: The first author thanks Gy. Less for his
guidance in the field. We appreciate the encouragement from
J. Haas and S. Kovács by means of which the first Hungarian
version of this review was written. The paper benefited from
critical comments of the GC reviewers, R. Posenato, J.
Michalík and an anonymous reviewer. We are grateful to M.
Pellérdyné and Cs. Péró for the photographs. The fieldwork
was partly sponsored by the Hungarian Scientific Research
Fund (OTKA) No. T 037966 and T 037595. We are indebted
to the Bükk National Park for supporting our studies in the
Ahr W.M. 1973: The carbonate ramp – an alternative to the shelf
model. Gulf Coast Association of Geological Societies Trans-
actions, 23, 221—225.
Aljinović D. 1995: Storm influenced shelf sedimentation – an ex-
ample from the Lower Triassic (Scythian) siliciclastic and car-
bonate succession near Knin (Southern Croatia and Western
Bosnia and Herzegovina). Geol. Croatica 48, 1, 17—32.
Árkai P. 1973: Pumpellyite-prehnite-quartz facies Alpine metamor-
phism in the Middle Triassic volcanogenic-sedimentary se-
quence of the Bükk Mountains, Northeast Hungary. Acta Geol.
Hung. 17, 1—3, 67—83.
Árkai P. 1983: Very low- and low-grade Alpine regional metamor-
phism of the Paleozoic and Mesozoic formations of the Bükki-
um, NE Hungary. Acta Geol. Hung. 26, 1—2, 83—101.
Ball M.M. 1967: Carbonate sands of Florida and the Bahamas. J.
Sed. Petrology 32, 2, 556—591.
Balogh K. 1964: Die geologischen Bildungen des Bükk-Gebirges.
Ann. Inst. Geol. Publ. Hung. 48, 2, 245—553.
Balogh K. 1980: Correlation of the Hungarian Triassic. Acta Geol.
Hung. 24, 1, 3—48.
Bathurst R. 1971: Carbonate sediments and their diagenesis. Devel-
opments in Sedimentology, 12, Elsevier, Amsterdam, 1—620.
Bérczi-Makk A. 1986: Microfauna study of the Gerennavár section.
Manuscript, Hungarian Geological Institute, Budapest (in
Bérczi-Makk A. 1987: Earlandia (Foraminifera) species from the Per-
mian—Triassic boundary in N Hungary. Ann. Inst. Geol. Publ.
Hung. 1985, 215—226 (in Hungarian, English summary).
Bérczi-Makk A., Csontos L. & Pelikán P. 1995: Data on the (upper
Permian) Foraminifer fauna of the Nagyvisnyó Limestone For-
mation from borehole Mályinka-8 (Northern Hungary). Acta
Geol. Hung. 38, 3, 85—250.
Bleahu M., Mantea G., Bordea S., Panin ., tefănescu M., Sikić
K., Haas J., Kovács S., Péró Cs., Bérczi-Makk A., Konrád
Gy., Nagy E., Rálisch-Felgenhauer E. & Török Á. 1994: Tri-
assic facies types, evolution and paleogeographic relations of
the Tisza Megaunit. Acta Geol. Hung. 37, 3—4, 187—234.
Bowring S.A., Errwin D.H., Jin Y., Martin M.W., Davidek K. &
Wang W. 1988: U/Pb zircon geochronology and tempo of the
end-Permian mass extinction. Science 280, 1039—1045.
Broglio Loriga C., Góczán F., Haas J., Lenner K., Neri C., Oravecz-
Scheffer A., Posenato R., Szabó, I. & Tóth-Makk Á. 1990: The
Lower Triassic sequences of the Dolomites (Italy) and Trans-
danubian Mid-Mountains (Hungary) and their correlation.
Mem. Sci. Geol. 42, 41—103.
Broglio Loriga C., Masetti D. & Neri C. 1983: The Werfen Forma-
tion (Scythian) in the western Dolomites: sedimentology and
biostratigraphy. Riv. Ital. Paleont. 88, 1982, 4, 501—598 (in
Italian, English summary).
Broglio Loriga C. & Mirabella S. 1986: The genus Eumorphotis
Bittner 1901 in the Scythian biostratigraphy, Werfen Forma-
tion (Dolomites). Mem. Sci. Geol. 38, 241—281 (in Italian, En-
Broglio Loriga C. & Posenato R. 1986: Costatoria (Costatoria?)
subrotunda (Bittner 1901) a Smithian (Lower Triassic) marker
from Tethys. Riv. Ital. Paleont. Stratigr. 92, 2, 189—200.
Budai T. & Haas J. 1997: Triassic sequence stratigraphy of the
Balaton Highland, Hungary. Acta Geol. Hung. 40, 3, 307—335.
Burchette T.P. & Wright V.P. 1992: Carbonate ramp depositional
systems. In: Sellwood B.W. (Ed.): Ramps and Reefs. Sed. Ge-
ology 79, 3—57.
Bystrický J. 1964: Stratigraphie und Dasycladaceen des Gebirges
Slovenský Kras. Ústr. ústav. geol., redakcia. Bratislava, 21—
35, 161—165 (in Slovak, German summary).
Bystrický J. 1973: Triassic of the West Carpathian Mts. X
gress of the Carpatho-Balkan Geological Association, Guide-
book to Excursion D. Xth Congress of the Carpatho-Balkan
Geological Association, 5—117.
Claoué-Long J.C., Zhang Z. & Du S. 1991: The age of the Permain—
Triassic boundary. Earth Planet. Sci. Lett. 105, 182—190.
Csontos L. 1999: Structural outline of the Bükk (N Hungary). Föld.
Közl. 129, 4, 611—651 (in Hungarian, English summary).
Csontos L. 2000: Stratigraphic reevaluation of the Bükk Mts (N.
Hungary). Föld. Közl. 130, 1, 95—131 (in Hungarian, English
Davies G.R. 1997: The Triassic of the Western Canada Sedimentary
Basin: tectonic and stratigraphic framework, paleogeography,
paleoclimate and biota. Bull. Canada Petrol. Geol. 45, 4, 434—
De Zanche V., Gianolla P., Mietto P., Siorpaes C. & Vail P. 1993:
Triassic sequence stratigraphy in the Dolomites (Italy). Mem.
Sci. Geol. 45, 1—27.
Filipović I., Jovanović D., Sudar M., Pelikán P., Kovács S., Less
Gy. & Hips K. (in press): Comparison of the Variscan—Early
Alpine evolution of the Jadar Block (NW Serbia) and “Bükki-
um” (NE Hungary) terranes; some paleogeographic implica-
tions. Slovak Geol. Mag.
Flügel E. 1982: Evolution of Triassic reefs: Current concepts and
problems. Facies 6, 297—328.
Fodor L. & Csontos L. 1998: Structural geological research in Hun-
gary: a review. Föld. Közl. 128, 1, 123—143 (in Hungarian, En-
Fülöp J. 1994: Geology of Hungary, Paleozoic II. Akadémiai Kiadó,
Budapest, 222—229 (in Hungarian).
Gaetani M., Jacobshagen V., Nicora A., Kauffmann G., Tselepidis
366 HIPS and PELIKÁN
V., Fantini Sestini N., Mertmann D. & Skourtsis-Coroneou V.
1992: The Early—Middle Triassic boundary at Chios (Greece).
Riv. Ital. Paleont. Stratigr. 98, 2, 181—204.
Gradstein F.M., Agterberg F.P., Ogg J.G., Hardenbol J., van Veen
P., Thierry J. & Huang Z. 1994: A Mesozoic time scale. J.
Geophys. Res. 99, B12, 24,051—24,074.
Gwinner M.P. 1971: Geologie der Alpen. E. Schweizerbart’sche
Verlagsbuchhand, Stuttgart, 1—477.
Halley R.B., Harris P.M. & Hine A.C. 1983: Bank margin environ-
ment. In: Schole P.A., Bebout D.G. & Moore C.H. (Eds): Car-
bonate depositional environments. AAPG, Tulsa, Oklahoma,
Harris P.M. 1979: Facies anatomy and diagenezis of a Bahamian
ooid shoal. Sedimenta 7, 1—163.
Heckel P.H. 1974: Carbonate buildups in the geological record: a
review. In: Laporte L.F. (Ed.): Reefs in Time and Space. Soc.
Econ. Paleontol. Mineral., Spec. Publ. 18, 90—154.
Herak M., Sćavničar & Šušnjara A., Đur anović Ž., Krystyn L. &
Gruber B. 1983: The Lower Triassic of Muc – Proposal for a
standard section of the European Upper Scythian. Schr. Erd-
wissenschaft. Komm. 5, 93—106.
Hips K. 1996: The biostratigraphic significance of the Cyclogyra?—
Rectocornuspira association (Foraminifera, Lower Triassic):
Data from the Aggtelek Mountains (Northeastern Hungary).
Neu. Jb. Geol. Paläont. Mh. 7, 439—451.
Hips K. 1998: Lower Triassic storm-dominated ramp sequence in
northern Hungary: an example of evolution from homoclinal
through distally steepened ramp to Middle Triassic flat-topped
platform. In: Wright V.P. & Burchette T.P. (Eds.): Carbonate
ramps. Geol. Soc. London, Spec. Publ. 149, 315—338.
James N.P. 1984: Reefs. In: Walker R.G. (Ed.): Facies Models.
Geosci. Canad., Repr. Ser. 1, 229—244.
James N.P. 1997: The cool-water carbonate depositional realm. In:
James N.P. & Clarke J.A.D. (Eds.): Cool-water carbonates.
Soc. Sed. Geol. Spec. Publ. Tulsa, Oklahoma, USA, 56, 1—20.
Koša E. & Janočko J. 1999: Storm-dominated mixed siliciclastic-
carbonate “Szin” ramp (Gl’ac Unit of the Silicicum Superunit,
Inner Western Carpathians): implication for Lower Triassic
eustacy. Slovak Geol. Mag. 5, 3, 201—212.
Kovács S. 1992: Tethys “western ends” during the Late Paleozoic
and Triassic and their possible genetic relationships. Acta
Geol. Hung. 35, 4, 329—369.
Kovács S., Szederkényi T., Haas J., Buda Gy., Császár G. & Nagy-
marosi A. 2000: Tectonostratigraphic terranes in the pre-Neo-
gene basement of the Hungarian part of the Pannonian area.
Acta Geol. Hung. 43, 3, 225—328.
Kozur H. 1985: Biostratigraphic evaluation of the Upper Paleozoic
conodonts, ostracods and Holothurian sclerites of the Bükk
Mts. Part II: Upper Paleozoic ostracods. Acta Geol. Hung. 28,
Kozur H. 1988: The Permian of Hungary. Z. Geol. Wiss. 16, 11—12,
Kozur H. 1996: The Conodonts Hindeodus, Isarcicella and Sweeto-
hindeodus in the Uppermost Permian and Lowermost Triassic.
Geol. Croatica 49, 1, 81—115.
Krainer K. 1987: Zusammensetzung und fazielle Entwicklung des
Alpinen Buntsandsteins und der Werfener Schichten im westli-
chen Drauzug (Karten/Osttirol). Jb. Geol. Bundesanst. 130, 1,
Krainer K. 1993: The Alpine Buntsandstein Formation of the Drau
Range (Eastern Alps, Austria): Transition from fluvial to shal-
low marine facies. In: Lucas S.G. & Morales M. (Eds.): The
nonmarine Triassic. New Mexico Museum of Natural History &
Science Bull. 3, 267—275.
Krystyn L. 1974: Die Tirolites-fauna (Ammonoidea) der untertrias-
sischen Werfener Schichten Europas und ihre stratigraphische
Bedeutung. Sitzberg. Öster. Akad. Wiss. Matem.-naturw. Kl.
Abt. I 193, 1—3, 29—50.
Less Gy. (Ed.), Gulácsi Z., Kovács S., Pelikán P., Pentelényi L. &
Sásdi L. (in press): Geological map of the Bükk Mts., 1:50,000.
Hungarian Geological Institute, Budapest.
Major R.P., Bebout D.G. & Harris P.M. 1996: Recent evolution of a
Bahamian ooid shoal: Effects of Hurricane Andrew. GSA Bul-
letin 108, 2, 168—180.
Metcalfe I., Nicoll R.S., Black L.P., Mundil R., Renne P., Jagodzin-
ski E.A. & Wang C. 1999: Isotope geochronology of the Per-
mian-Triassic boundary and mass extinction in South China.
In: Yin H. & Tong J. (Eds.): Pangea and Paleozoic-Mesozoic
transition. China University of Geosciences Press, Wuhan,
Michalík J. 1994: Notes on the paleogeography and paleotectonics
of the Western Carpathian area during the Mesozoic. Mitt. Ös-
terr. Geol. Gesell. 86, 101—110.
Mostler H. & Rossner R. 1984: Mikrofazies und Palökologie der
höheren Werfener Schichten (Untertrias) der Nördlichen Kal-
kalpen. Facies 10, 87—144.
Nakazawa K. 1977: On Claraia of Kashmir and Iran. J. Petrol. Soc.
India 20, 191—204.
Neri C. 1991: Sequence stratigraphy of the Early Triassic Werfen
Formation (Dolomites, Northern Italy). Dolomieu Conference
on Carbonate Platforms and Dolomization, Abstract volume,
Neri C. & Posenato P. 1985: New biostratigraphical data on upper-
most Werfen Formation of Western Dolomites (Trento, Italy).
Geol. Paläont. Mitt. Innsbruck 14, 3, 83—107.
Orchard M. 2001: Executive notes, from the Chair. Albertiana
Patrulius D., Bleahu M., Popescu I. & Bordea S. 1971: The Triassic
Formations of the Apuseni Mountains and of the East Car-
pathian Bend. Guidebook to Excursion of the 2
loquium Carpatho-Balkan Association, September 1971,
Geological Institute, Bucharest, 5—51.
Pelikán P. 1985a: Geological key-profile, Gerenna-vár, Szilvás-
várad, Bükk Mountains, Gerennavár Limestone Formation.
Geological key-sections in Hungary, 6. Hungarian Geological
Institute, Budapest, 5.
Pelikán P. 1985b: Evaluation report of the Mályinka-8 core drilling.
Manuscript, Hungarian Geological Institute, Budapest, 1—50
Pelikán P. 1995: Bükk Unit, Gerennavár Limestone Formation,
Ablakoskővölgy Formation. In: Barabás-Stuhl Á., Bérczi-
Makk A., Budai T., Csillag G., Dosztály L. & Haas J. et al.
(Eds.): Lithostratigraphical units in Hungary Triassic. Hungar-
ian Geological Institute, Budapest, 105—111 (in Hungarian).
Pelikán P. & Csontos-Kis K. 1990: Geological key-profile, N
Bálvány, Nagyvisnyó, Bükk Mountains, Nagyvisnyó Lime-
stone Formation, Gerennavár Limestone Formation. Geologi-
cal key-sections in Hungary, 136. Hungarian Geological
Institute, Budapest, 5.
Péró Cs. 1983: Evaluation report of the sample collection in the Ge-
rennavár-2 section. Manuscript, Hungarian Geological Insti-
tute, Budapest, 1—13 (in Hungarian).
Perri M.C. 1991: Conodont biostraigraphy of the Werfen Formation
(Lower Triassic), Southern Alps, Italy. Boll. Soc. Paleont. Ital.
30, 1, 23—46.
Posenato R. 1992: Tirolites (Ammonoidea) from the Dolomites, Ba-
kony and Dalmatia: Taxonomy and biostratigraphy. Eclogae
Geol. Helv. 85, 3, 893—929.
Pratt B.R. & James N.P. 1986: The St George Group (Lower Or-
dovivian) of western Newfoundland: Tidal flat island model
for carbonate sedimentation in shallow epeiric seas. Sedimen-
tology 33, 313—343.
LOWER TRIASSIC SHALLOW MARINE SUCCESSION 367
Protić L., Filipović I., Pelikán P., Jovanović D., Kovács S., Sudar
M., Hips K., Less Gy. & Cvijić R. 2000: Correlation of the Car-
boniferous, Permian and Triassic sequences of the Jadar Block,
Sana-Una and “Bükkium” Terranes. In: Karamata S. & Janko-
vić S. (Eds.): Proceedings of the International Symposium Ge-
ology and Metallogeny of the Dinarides and Vardar Zone.
Acad. Scien. Arts. Rep. Srpska, Collection and Monographs, 1,
Dept. Nat., Mat. and Techn. 1, 61—69.
Read F. 1985: Carbonate platform facies models. Amer. Assoc.
Petrol. Geol. Bull. 69, 1—21.
Reid P. & Macintyre I. 2000: Microboring versus recrystallization:
further insight into the micritization process. J. Sed. Res. 70, 1,
Renne P.R., Zhang Z., Richards M.A., Black M.T. & Basu A.R.
1995: Synchrony and causal relations between Permian-Trias-
sic boundary crisis and Siberian flood volcanism. Science 269,
Rüffer Th. & Zühlke R. 1995: Sequence stratigraphy and sea-level
changes in the Early to Middle Triassic of the Alps: A global
comparison. In: Haq B.U. (Ed.): Sequence stratigraphy and
depositional response to eustatic, tectonic and climatic forcing.
Kluwer Academic Press, Dordrecht, 161—207.
Salaj J., Borza K. & Samuel O. 1983: Triassic foraminifers of the
West Carpathians. Geologický ústav Dionýza Štúra, Bratislava,
Sćavničar B. & Šušnjara A. 1983: The geologic column of the Low-
er Triassic at Muć (Southern Croatia). Acta Geol. 13, 1, 1—25.
Schréter Z. 1935: Über die Triasbildungen des Bükk-Gebirges.
Földt. Közl. 65, 4—6, 90—105 (in Hungarian, German sum-
Schréter Z. 1953: Recherches géologiques aux environs de
Nagyvisnyó. Ann. Inst. Geol. Publ. Hung. 1951, 155—167 (in
Hungarian, French summary).
Schréter Z. 1954: Geologic and hydrogeologic conditions of the old
mass of Bükk Mountains. Hidr. Közl. 34, 7—8, 287—294, 369—
381 (in Hungarian).
Schubert J.K. & Bottjer D.J. 1995: Aftermath of the Permian—Trias-
sic mass extinction event: Paleoecology of Lower Triassic car-
bonates in the western USA. Palaeogeogr., Palaeoclimat.,
Palaeoecol. 116, 1—39.
Shinn E.A., Steinen R.P., Dill R.F. & Major R.P. 1993: Lime-mud
layers in high-energy tidal channels, a record of hurricane de-
posits. Geology 21, 603—606.
Tollmann A. 1976: Analyse des klassischten Nordalpinen Mesozoi-
kums, Stratigraphie, Fauna und Fazies der Nördlichen Kalkal-
pen. Franz Deuticke, Wien, 50—65.
Tucker M. & Wright V.P. 1990: Carbonate Sedimentology. Black-
well Science, Oxford, 1—482.
Wright V.P. & Burchette T.P. 1996: Shallow-water carbonate envi-
ronments. In: Reading H.G. (Ed.): Sedimentary environments
and facies. Blackwell Scientific Publications, Oxford, 325—394.
Wright V.P. & Faulkner T.J. 1990: Sediment dynamics of Early
Carboniferous ramps: a proposal. Geol. J. 25, 139—144.
Yin H. 1993: A proposal for the global stratotype section and point
(GSSP) of the Permian-Triassic boundary. Albertiana 11, 4—30.
Yin H., Zhang K., Wu S. & Peng Y. 1996: Global correlation and
definition of the Permian-Triassic boundary. In: Yin H. (Ed.):
The Palaeozoic-Mesozoic boundary. Candidates of the global
stratotype section and point of the Permian-Triassic boundary.
Dedicated to the 30
International Geological Congress. Chi-
na University of Geosciences Press, Wuhan, 3—28.
Zhang Z., Claoué-Long J. C., Ma G. & Du S. 1992: Age determina-
tion of the Permian-Triassic boundary at Meishan, Changxing,
Zhejiang Province. Geol. Rev. 38, 4, 272—281 (in Chinese, En-
Zhang K., Ding M., Lai X. & Liu J. 1996: Conodont sequences of
the Permian-Triassic boundary strata at Meishan section, South
China. In: Yin H. (Ed.): The Palaeozoic-Mesozoic boundary.
Candidates of the global stratotype section and point of the
Permian-Triassic boundary. Dedicated to the 30
Geological Congress. China University of Geosciences Press,