Global and Planetary Change - Universiteit Utrechtforth/publications/deLeeuw_2013.pdfA. de Leeuw et al. / Global and Planetary Change 103 (2013) 82–98 83 environments were characterized

Global and Planetary Change 103 (2013) 82–98

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Global and Planetary Change

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Paleomagnetic and chronostratigraphic constraints on the Middle to Late Mioceneevolution of the Transylvanian Basin (Romania): Implications for Central Paratethysstratigraphy and emplacement of the Tisza–Dacia plate

Arjan de Leeuw a,⁎, Sorin Filipescu b, Liviu Maţenco c, Wout Krijgsman a, Klaudia Kuiper c, Marius Stoica d

a Paleomagnetic Laboratory ‘Fort Hoofddijk’, Utrecht University, Budapestlaan 4, 3584 CD, Utrecht, The Netherlandsb Department of Geology, ‘Babeş-Bolyai’ University, Kogălniceanu St. 1, 400084 Cluj-Napoca, Romaniac Faculty of Earth and Life Sciences, VU University Amsterdam, De Boelelaan 1085, 1081 HV Amsterdam, The Netherlandsd Department of Geology, Faculty of Geology and Geophysics, University of Bucharest, Bălcescu Bd. 1, Bucharest, 010041, Romania

⁎ Corresponding author. Fax: +31 30 253 1677.E-mail addresses: [email protected] (A. de Leeuw

(S. Filipescu), [email protected] (L. Maţenco), kr(W. Krijgsman), [email protected] (K. Kuiper), m

0921-8181/$ – see front matter © 2012 Elsevier B.V. Alldoi:10.1016/j.gloplacha.2012.04.008

a b s t r a c t
a r t i c l e i n f o
Article history:Received 6 May 2011Revised 31 March 2012Accepted 12 April 2012Available online 15 May 2012

Keywords:MiocenechronostratigraphyParatethysTransylvanian Basinsedimentary history40Ar/39Ar datingpaleomagnetismmagnetostratigraphytectonic rotationenvironmental change

From theOligocene onwards, the complex tectonic evolution of the Africa–Eurasia collision zone led to paleogeo-graphic and biogeographic differentiation of the Mediterranean and Paratethys, two almost land-locked seas, inthe area formerly occupied by the western Tethys Ocean. Episodic isolation of the basins triggered strong faunalendemism leading to the introduction of regional stratigraphic stages for the Paratethys. Chronostratigraphiccontrol on the Paratethys stages remains rudimentary compared to the cyclostratigraphically constrained Med-iterranean stages. This lack of chronostratigraphic control restricts the insight in the timing of geodynamic, cli-matic, and paleobiogeographic events and thereby hinders the identification of their causes and effects. In thispaper, we here derive better age constraints on the Badenian, Sarmatian and Pannonian Central Paratethys re-gional stages through integrated 40Ar/39Ar,magnetostratigraphic, and biostratigraphic research in the Transylva-nian Basin. The obtained results help to clarify the regionsMiddle Miocene geodynamic and paleobiogeographicevolution. Six new 40Ar/39Ar ages were determined for tuffs intercalating with the generally deep marine basininfill. Together with data from previous studies, there is now a total of 9 radio-isotopically dated horizons in thebasin. These were traced along seismic lines into a synthetic seismic stratigraphic column in the basin center andserve as first order tie-points to the astronomically tuned Neogene timescale (ATNTS). Paleomagnetically inves-tigated sections were treated similarly and their polarity in general corroborates the 40Ar/39Ar results. The inte-grated radio-isotopic and magnetostratigraphic results provide an improved high-resolution time-frame for thesedimentary infill of the Transylvanian Basin. Early Badenian deepwater sedimentation is characterized by accu-mulation of theDej Tuff Complex in response to a period of intensive volcanism, the onset ofwhich is constrainedbetween the first occurrence (FO) of Orbulina suturalis at 14.56 Ma and 14.38±0.06 Ma. During the subsequentBadenian Salinity Crisis (BSC) up to 300 m of salt accumulate in the basin center. The faunal turnover thatmarksthe Badenian–Sarmatian Boundary is dated at 12.80±0.05 Ma. A second phase of intense volcanism occurs at12.4 Ma and leads to deposition of themiddle Sarmatian tuff complex (Ghiriş, Hădăreni, Turda and Câmpia Turziituffs). Rates of sediment accumulation strongly diminish in the basin center at the onset of the Pannonian stagecoincident with an approximately 20° CW tectonic rotation of the Tisza–Dacia plate. Concurrent enhanced upliftin the Eastern a'nd Southern Carpathians leads to the isolation of the Central Paratethys and triggers the transi-tion from marine to freshwater conditions. An additional Pannonian to post-Pannonian 6° of CW rotation is re-lated to the creation of antiform geometries in the Eastern Carpathians which are notably larger in the norththan in the south. An 8.4 Ma age is determined for the uppermost Pannonian sediments preserved in the centralpart of the Transylvanian Basin. Two sections belonging tomiddle Pannonian Zone D, and the lower part of ZoneE (Subzone E1) are found to cover the 10.6–9.9 Ma time-interval.

© 2012 Elsevier B.V. All rights reserved.

), [email protected]@[email protected] (M. Stoica).

rights reserved.

1. Introduction

During the Cenozoic, collision of the northward moving African platewith Eurasia led to the Alpine Orogeny. Along the course of this process,the vast water mass of the Tethys Ocean disintegrated and gave birth tothe Mediterranean and Paratethys Seas (e.g. Seneš, 1973; Rögl, 1999)(Fig. 1). The complex tectonic evolutionof theAfrica–Eurasia convergence

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Paratethys

Africa

Eurasia

Mediterranean

W C E

Fig. 1. Middle Miocene paleogeographic map for the Africa–Eurasia collision zone after(Rögl, 1999). Northern Europe after Meulenkamp and Sissingh (2003). Red frame indicatesthe Transylvanian Basin. W, C, and E indicate the Western, Central and Eastern Paratethysrespectively.

83A. de Leeuw et al. / Global and Planetary Change 103 (2013) 82–98

zone, with its multitude of colliding micro-continents (Csontos et al.,1992; Stamfli and Kozur, 2006; Schmid et al., 2008; Ustaszewski et al.,2008), induced a similarly intricate regional paleogeographic historywith frequently changing seaways and land-bridges (Rögl, 1999;Harzhauser et al., 2002; Popov et al., 2004, 2006; Harzhauser andPiller, 2007).

Gradual growth of the Alpine-Carpathian-Dinaridic orogenic sys-tem induced progressive restriction of the Western, Central and East-ern Paratethys (Fig. 1) Repeated isolation led to large changes inwater chemistry (Matyas et al., 1996; Peryt, 2006; Harzhauser et al.,2007; Vasiliev et al., 2010b) and inflicted periods of severe endemism(e.g. Steininger et al., 1988; Magyar et al., 1999a; Harzhauser et al.,2002, 2003; Kovač et al., 2007). This geodynamically controlled paleo-geographic and biogeographic differentiation causedmajor difficulties instratigraphic correlation between the different parts of the Paratethysand the Mediterranean (Piller and Harzhauser, 2005) and thus led tothe establishment of regional chronostratigraphic scales for the CentralParatethys summarized in the series “Chronostratigraphie und Neo-stratotypen” (Cicha et al., 1967; Steininger et al., 1971; Papp et al.,1973, 1974; Báldi and Seneš, 1975; Papp et al., 1978, 1985; Stevanovicet al., 1990). A high resolution chronologic framework was establishedover the past few decades for the Mediterranean stages, on which theglobal timescale currently relies (Lourens et al., 2004a). In comparison,chronologic control on most of the Paratethys regional stages remainsrudimentary, although significant progress has been achieved in thelast decade (e.g. Krijgsman et al., 2010 and references therein; Lireret al., 2009; Paulissen et al., 2011). In particular, precise dating of theMiddle-Late Miocene sediments of the Central Paratethys (Geary et al.,2000;Magyar et al., 2007; Vasiliev et al., 2010a), remains an outstandingproblem.

In order to overcomepart of this problemwe conducted an integrat-ed biostratigraphic, magnetostratigraphic and geochronologic study inthe Transylvanian Basin. This basin was part of the Central Paratethysduring the Miocene and it exposes a thick pile of siliciclastic depositssuitable for paleomagnetic investigations. A number of intercalatingvolcano-sedimentary products (mostly tuffs) provide key stratigraphichorizons potentially datable with the 40Ar/39Ar method. The continuityof sedimentation, reasonable exposures due to tectonic exhumation ofthe basin towards the end of the Miocene, and a significantly loweramount of tectonic disturbance and related structural complicationsand unconformities (Sztanó et al., 2005; Horváth et al., 2006; Krézseket al., 2010) in comparison with other parts of the Central Paratethys,make the Transylvanian Basin a favorable study location. The limitedsize of the Transylvanian Basin and its mature stage of exploration(Krézsek et al., 2010) provide the opportunity to correlate outcropsover large distances based on subsurface imagery.

The Transylvanian Basin also provides an optimal location to ex-amine the intriguing post-20 Ma rotation associated with the in-vasion of the Tisza–Dacia plate into the Carpathian embayment,because of its largely undeformed Middle-Late Miocene sedimentarysuccession. A precise quantification of the translation mechanics ofTisza–Dacia around the Moesian indenter, driven by the roll-back ofa slab attached to the European continent (e.g. Balla, 1987; Royden,1988; Ustaszewski et al., 2008), is hampered by contrasting amountsof rotation derived by the paleomagnetic, paleogeographic and struc-tural geology studies, which vary from 10° to almost 70° clockwise ro-tation after the Paleogene (Csontos et al., 1992; Fügenschuh andSchmid, 2005; van Hinsbergen et al., 2008).

We aim to derive better age constraints on the Middle to LateMiocene regional stages (Badenian, Sarmatian, and Pannonian) ofthe Central Paratethys through integrated magnetostratigraphic,radio-isotopic and biostratigraphic research in Transylvania. Takingadvantage of the resulting chronostratigraphic and paleomagnetic data,the amount aswell as the partitioning in time of theMiocene rotation ofTisza–Dacia can be unraveled. These data provide a more detailed in-sight in the mechanism of emplacement of the Tisza–Dacia block,supported by recent advances in the understanding of the exhumationand kinematics of the Carpathians.

2. Geological setting

The Transylvanian Basin is a 200 km long and 250 km wide semi-isolated back-arc basin (Krézsek et al., 2010) situated in the eastern partof the Central Paratethys and bounded by the Eastern Carpathians, South-ern Carpathians and Apuseni Mountains (Fig. 2). The Middle to LateMiocene subsidence and subsequent exhumation of the TransylvanianBasin are both closely related to, and concurrent with, major episodes ofdeformation in the Carpathians in conjunction with the subduction ofthe Eastern European underneath the Tisza–Dacia plate and the collisionthat followed (Matenco et al., 2010).

During the Early Miocene deep marine conditions were restrictedto the northern part of the basin while continental settings prevailedelsewhere. A regional transgression during the Middle Mioceneestablished relatively deep marine settings throughout the basinwith mixed carbonate-siliciclastic platforms near the margins andsiliciclastic environments in deeper areas (Fig. 3, Filipescu andGîrbacea, 1997; Krézsek et al., 2010). At this time, the Central Para-tethys was well connected with the Mediterranean (Fig. 1) and therewas a great similarity in their paleontological record. The onset ofcalc-alkaline volcanism related to subduction processes in the exteriorCarpathians induced acidic volcanism that led to widespread tuff depo-sition in the Central Paratethys (Pécskay et al., 1995; Seghedi et al.,2004; Pécskay et al., 2006). In the Transylvanian Basin an up to 50 mthick tuff complex accumulated (Fig. 3, Seghedi and Szakács, 1991).This so called Dej Tuff thus constitutes amajor lithostratigraphicmarkerwithin the up to 100 m thick lower Badenian sedimentary package.

Severe restriction of the Central Paratethys at 13.82 Ma (Peryt,2006; de Leeuw et al., 2010) induced hypersaline conditions. Theseled to the deposition of around 300 m of salt in the deeper parts ofthe Transylvanian Basin and gypsum along the western margin(Krézsek and Bally, 2006).

In the late Badenian, the Central Paratethys re-connected with theEastern Paratethys that occupied the current Black Sea and CaspianSea regions. Hypersaline conditions disappeared even though there-united Paratethys remained semi-isolated (Popov et al., 2004;Kovač et al., 2007). At this time, subsidence rates in the TransylvanianBasin increased tremendously and a, locally over 3 km thick, upperBadenian to Pannonian siliciclastic sedimentary infill accumulated(Krézsek and Filipescu, 2005). The lowermost post-salt successiononlaps the evaporites in the northern, western and southern partsof the basin and thickens from a few 100 m in the west to morethan 2000 m in the southeast (Krézsek and Bally, 2006). Depositional

MesozoicophyolitesDaciaTisza

0

5

10

km

0

5

10

km

20 40 60 80 100 120 140 160km EW

Pn

Eastern Carpathians

Apuseni Mountains

Sm

Bn

Odorhei

Reghin

Sibiu

Aiud

Blaj

Alba Iulia

10 km

Basement

Eocene (Ec)

E. Oligocene (EO)

L. Oligocene (LO)

Early Miocene (EM)

Badenian (Bn)

Sarmatian (Sm)

Pannonian (Pn)

Neogene Volcanics

Quartenary

Fault

Badenian halite

Sampling locality

City

Legend

v

Sm

Pn

Bn

EM

Ec

Sm

Bn

Ec

EM

diapir

diapir

diapir

BnSm

Pn

B

C

EOLO

A

N50 km

Southern Carpathians

Eas

tern

Car

pat

hia

ns

Cluj

ApuseniMountains

Figure 1b

Figure 7

Figure 1c

Turda

Viforoasa

Câmpia Turzii

Apahida

Dej

Gherla

Câmpia Turzii 1Câmpia Turzii 2Turda

Oarba D Oarba Beta

Oarba A, B

Fig. 2. Geological map (A) and east–west cross section of the Transylvanian Basin (B). Panel C is an enlarged geological map of the area indicated in panel A.

84 A. de Leeuw et al. / Global and Planetary Change 103 (2013) 82–98

environments were characterized by clay-dominated deep marine fans(Krézsek and Filipescu, 2005). During the late Badenian and Sarmatian,the basin's depocenter was located at its current eastern margin andthe corresponding part of the infill thus thins out towards the west(Krézsek et al., 2010).

The Sarmatian part of the basin infill comprises siliciclastics, suchas marls and sandstones, subordinate conglomerates and evaporites.These demonstrate a higher variety of sedimentary environments.Sea-level changes in combination with incipient tectonics in the

adjacent mountain chains triggered alternation between normal ma-rine, brackish, and lacustrine conditions (Krézsek et al., 2010).

At the advent of the Pannonian, the Central Paratethys became fullyisolated from the marine realm, and turned into a lake (Magyar et al.,1999a). The Pannonian sediments of the Transylvanian Basin compriselacustrine fans and lowstand deltas with sand dominated facies in themore proximal-, and marl dominated facies in the central, more distalpart of the basin (Krézsek et al., 2010). The regressive trend was inter-rupted by a phase of regional uplift towards the end of the Pannonian

Ammonia acmeOstracods

rare foraminifera (Miliammina), ostracodsPorosononion aragviensis / Streptochilus

Dogielina sarmatica

Elphidium reginumVaridentella reussi

Articulinasarmatica

Anomalinoides dividens

Tenuitellinata

benthic foraminifera diversification

benthic foraminifera, mysids

benthic foraminifera diversification

benthic foraminifera (miliolids)

rare benthic (brackish) foraminifera

Anomalinoides acme, rare planktonics

Velapertina

barren

Globoturborotalita druryi / Globigerinopsis grilli

Orbulina suturalis / Globorotalia bykovae

Praeorbulina glomerosaStreptochilus pristinum

planktonics and rare benthicsbenthic foraminifera, rare planktonics

Bogdanowiczia acme, planktonic foraminifera

planktonic and deep bethic foraminifera

planktonic foraminifera, radiolarians, pteropods

rare benthic foraminifera

planktonic and benthic foraminifera

planktonic and rare benthic foraminifera

planktonic foraminifera

Pan

noni

anea

rlyS

arm

atia

nea

rlyla

teea

rlyla

tem

iddl

eB

aden

ian

Stages

W/SW evaporites

middle fan

outer fan

shelf

outer fanmiddle fan

outer fan

middle fan

fandelta

delta

Gypsum Salt Limestone Shale / MarlShale & sandstone

Sandstone Conglomerate Hiatus

Shallow DeepLithology and sedimentologyForaminifera biozones Micropaleontological assemblages

Dej Tuff Complex

Apahida Tuff

Hâdâreni Tuff

Oarba Tuff

Fig. 3. Sedimentary facies and biozonation of the Transylvanian Basin after Krézsek and Filipescu (2005), and Krézsek and Bally (2006). Biozonation after Popescu (2000), Krézsekand Filipescu (2005) and Filipescu and Silye (2008).


that exhumed the basin to its present-day altitude of between 300 and600 m (Matenco et al., 2010).

Volcanic activity was common in the neighbouring EasternCarpathians and Apuseni Mountains during the Badenian, Sarmatianand Pannonian (Szakács and Seghedi, 1995; Seghedi et al., 2004). Itgenerated a large number of volcaniclastics in the sedimentary se-quence of the Transylvanian Basin. These among others include theApahida Tuff, Hădăreni Tuff, Ghiriş Tuff, and Oarba Tuff.

3. Stratigraphic framework and sampling approach

Our research relies on the stratigraphic framework of Krézsek andFilipescu (2005), and Krézsek and Bally (2006). Sections were inves-tigated using an integrated stratigraphic approach taking multiplesamples for both biostratigraphic and paleomagnetic analysis at thesame stratigraphic level. We focused on sections that include tuffsand tuffites and sampled them for 40Ar/39Ar dating, and biostratigraphy.We use the Astronomically Tuned Neogene Timescale (Lourens et al.,2004a) with adaptations by Hüsing et al. (2010) and supplementedwith short term fluctuations in the geomagnetic field identified byKrijgsman and Kent (2004) as a global chronostratigraphic referenceframework.

In order to retrieve the relative stratigraphic relation, all sections(Fig. 2) were traced along seismic reflectors to a composite strati-graphic column (Fig. 4) near the basin's depocenter where the Middleto Upper Miocene succession is as complete as possible (Fig. 2). Foreach of the sections, an uncertainty in stratigraphic position was esti-mated based on the proximity to available seismic lines, clarity ofline-to-line reflector correlation and the presence of potential sedi-mentary architectures affecting precision, such as clinoforms withlower reflectivity. The synthetic seismic section in Fig. 4 is verticallyscaled in seconds two-way travel time. A depth conversion of all ele-ments projected into this line was performed using average intervalvelocities derived from neighboring well logs (sonic or density logs,see De Broucker et al., 1998; Krézsek et al., 2010).

4. 40Ar/39Ar dating of key stratigraphic horizons

There are several tuffs in the Middle to Late Miocene sedimentaryinfill of the Transylvanian Basin, discussed in detail by Bedelean et al.(1991). Some of these, e.g. the Bor a-Apahida, Hădăreni, Ghiriş,Sărmă el and Bazna tuffs, extend over a wide area. Their strong acous-tic impedance contrast facilitates correlation across seismic lines(Ionescu, 1994) and they represent excellent regional stratigraphicmarkers. Several of these tuff levels were sampled for 40Ar/39Ar dat-ing in order to acquire absolute ages for the corresponding strati-graphic horizons.

4.1. Methods

The volcanic ashes were processed at the Department of IsotopeGeochemistry (VU University Amsterdam). Bulk samples were crushed,disintegrated in a calgon solution, washed and sieved over a set ofsieves between 63 and 250 μm. The residue was subjected to standardheavy liquid as well as magnetic mineral separation techniques fork-feldspar. Except for the sample from Apahida, all samples containedample k-feldspar for 40Ar/39Ar dating. The k-feldspar separates wereleached with a 1:5 HF solution in an ultrasonic bath for 5 min. Aftercareful handpicking, theywere loaded in a 10 mm IDquartz vial togeth-er with Fish Canyon Tuff (FC-2) and Drachenfels (Dra-1, f250-500 andDra-2, f>500) sanidine that served as flux monitors. The vial was irra-diated in the Oregon State University TRIGA reactor in the cadmiumshielded CLICIT facility for 10 h.

Upon return in the laboratory, mineral separates were split into atleast 9 duplicate fractions and loaded in a Cu-tray. The loaded Cu-traywas pre-heated to ~200 °C under vacuum using a heating stage and aheat lamp to remove undesirable atmospheric argon. The tray wasthen placed in the sample house and the system (extraction line+sample house) was degassed overnight at ~150 °C. Incrementalheating was performed with a Synrad CO2 laser in combination witha Raylase scanhead as a beam delivery and beam diffuser system.After purification the resulting gas was analyzed with a Mass Analyzer

0.0

1.0

2.0

3.0

0.0

1.0

2.0

3.0

sTW

TsT

WT

76 78 80 82 84 86 88 90 92 km 96

Vh2-Bs

Vh1

Pn

Bn

Vh2-Bs

Vh1

Pn

Bn

Salt

Viforoasa

Oarba D

Oarba A

Oarba Beta

Turda

Câmpia Turzii

Gherla

Apahida

Dej Tuff, Dej

Câmpia Turzii Tuff

Oarba Tuff

Fig. 4. Synthetic seismic section near the depocenter of the Transylvanian Basin. For location see Fig. 2 and Beldean et al. (2010). Surface outcrops have been traced along seismicreflectors from their respective position in the basin into this seismic section in order to determine their position in the stratigraphic succession. The section is vertically scaled inseconds two-way travel time. Vertical brackets indicate the uncertainty in stratigraphic position. For a detailed explanation we refer to the main text.


Products LTD 215-50 noble gas mass spectrometer. Beam intensitieswere measured in a peak-jumping mode in 0.5 mass intervals overthe mass range 40–35.5 on 2 a Balzers 217 secondary electronmultipli-er. System blanks were measured every three to four steps. Mass dis-crimination was monitored by frequent analysis (~every 10 h) ofaliquots of air. The irradiation parameter J for each unknownwas deter-mined by interpolation using a second-order polynomial fitting be-tween the individually measured standards.

All age calculations use the decay constants of Steiger and Jäger(1977). Steps with less than 1% radiogenic argon were immediatelydiscarded. The age for the Fish Canyon Tuff sanidine flux monitorused in age calculations is 28.201±0.03 Ma (Kuiper et al., 2008).The age for the Drachenfels sanidine flux monitor is 25.42±0.03 Ma(Kuiper et al. in preparation). Correction factors for neutron interfer-ence reactions are 2.64±0.017×10−4 for (36Ar/37Ar)Ca, 6.73±0.037×10−4 for (39Ar/37Ar)Ca, 1.211±0.003×10−2 for (38Ar/39Ar)Kand 8.6±0.7×10−4 for (40Ar/39Ar)K. Errors are quoted at the 1σlevel and include the analytical error and the error in J. All relevantanalytical data as well as error determination can be found in the on-line supplementary material.

4.2. Results

The calculated weighted mean ages, peaks in probability densitydistribution, isochron ages and several other parameters for each ofthe investigated ashes are listed in (Table 1). The full analytical dataare available in the supplementary material. The MSWD was smaller

than the statistical T ratio at the 2 sigma level for all reported weight-ed mean ages. Isochron ages are concordant with the weighted meanages and the trapped 40Ar/36Ar component is in all cases atmospheric.The calculated weighted mean ages are also concordant with thepeaks in probability density distribution and we interpret them to re-flect the respective tuff's crystallization age.

The lower Badenian part of the basin infill is dominated byvolcaniclastic deposits of the Dej Tuff complex (Fig. 2.). We dated twosamples of the Dej Tuff taken from outcrops near Ciceu Giurgeşti andDej. For these samples a larger number of repetitive experiments hadto be performed, since both contained a large reworked component.This resulted in abnormally high ages for part of the experiments thattherefore had to be rejected. For each of the samples at least 7 experi-ments provided consistent, homogenous and stratigraphically realisticages (Fig. 5). These were selected to calculate their weighted meanages. The resulting statistically equivalent weighted mean ages of thesamples of the Dej Tuff taken at Ciceu Giurgeşti and Dej are 14.38±0.06 Ma and 14.37±0.06 Ma (Table 1).

Several tuff levels occur within a small stratigraphic distance inthe middle Sarmatian part of the basin infill. The lowermost of theseis the Hădăreni Tuff, which is recognized as a seismic reflectorthroughout the larger part of the basin (Fig. 4). It crops out in a classiclocality at the base of the Brâul Alb (‘White Belt’) hill near Hădăreniwhere it was sampled. At the top of the same hill, stratigraphicallyabout 100 m higher, there is another tuff level. This level, known asthe andesitic Ghiriş Tuff (Mârza and Mészáros, 1991), also crops outat a site just west of Ghirişu Român where it was sampled. Betweenthe Apuseni Mountains and Hădăreni, where the E–W flowing Arieş

Table 140Ar/39Ar ages for several tuffs of the Transylvanian Basin. MSWD is Mean Square Weighted Deviates, N is the total number of repetitions in the single fusion experiments and n isthe amount included when calculating the weighted mean and inverse isochron age. %40Ar(r) is the radiogenic amount of 40Ar. Errors are given at 95% confidence level. MSWD,40Ar* (%), and Inverse isochron intercept were determined based on the experiments selected for calculation of the weighted mean or isochron age. Ages in bold are consideredto most accurately represent the respective tuffs crystallization age.

Sample LaboratoryID

Location Method Weightedmean age(Ma)

n/N MSWD Peakprobabilitydistribution(Ma)

%40Ar(r) Mineral Cystalsize (μm)

Inverseisochronage

InverseisochronIntercept

IsochronMSWD

Reference

Dej AL62,VU78-25

N47.159060,E023.882390

Multipletotal fusion

14.37±0.06 7/13

1.33 14.34 95 k-feldspar

200–250 14.29±0.13

328±48 1.46 This study

Ciceu Giurgeşti AL58,VU78-24

N47.241480,E024.032890


14.38±0.06 8/27

0.91 14.41 90 k-feldspar

180–200 14.31±0.10


CâmpiaTurzii

AL49,VU79-19

N46.565760,E023.890187


12.35±0.04 14/15

0.52 12.35 99 k-feldspar

250–500 12.36±0.04


Hădăreni AL59,VU78-22

N46.471167,E023.985811


12.37±0.05 9/13

0.13 12.37 88 k-feldspar

250–500 12.37±0.05


GhirişuRomân

AL56,VU78-26

N46.799330,E023.953930


12.38±0.04 8/9 0.75 12.39 98 k-feldspar

200–250 12.38±0.05


Turda AL52,VU78-17

N46.566940,E023.822020


12.37±0.04 10/10

1.4 12.39 98 k-feldspar

200–250 12.36±0.05


Oarba OM-C,VU59-5

N46.457052,E024.288622


11.62±0.04 11/13

1.71 86 Botite 300–500 Vasiliev etal., 2010a


River has cut outcrops in the south-directed hillsides, the diapirs ofthewestern salt diapir alignment (Krézsek and Bally, 2006) cause tiltingand multiple repetition of upper Badenian to middle Sarmatian strata.Two additional middle Sarmatian tuff levels were sampled in outcropsnear Turda and Câmpia Turzii respectively (Fig. 2). The 40Ar/39Ar exper-iments for the Turda, Hădăreni, Câmpia Turzii and Ghiriş tuffs providedhomogenous age populations (Fig. 5). The peaks in the probability den-sity distributions are concordant with the calculated weighted meanages andwe interpret them to reflect the respective tuff's crystallizationage. The Turda, Hădăreni, Câmpia Turzii andGhiriş tuffs are thus respec-tively 12.37±0.04 Ma, 12.37±0.05 Ma, 12.35±0.04 Ma, and 12.38±0.04 Ma old (Table 1).

5. Resulting time frame and sedimentation rates

The performed 40Ar/39Ar measurements thus provided ages for sixstratigraphic horizons. The number of horizons with a radio-isotopically

single experimentwith analytical uncertainty

probability density distribution

10 12 14 160

0.002

0.004

0.006

0.008

0.01

11 13 15

Oarba-C

Dej

Turda

Câmpia Turzii

Age (Ma)

Pro

babi

lity

Fig. 5. Probability density diagrams for the multiple total fusion experiments for theinvestigated tuffs.

determined age can be augmentedwith the base and top of the evaporitesequence for which a 13.8 Ma and 13.36 Ma 40Ar/39Ar age was recentlyestablished (de Leeuw et al., 2010; de Leeuw et al., unpublished results;de Leeuw, 2011). Vasiliev et al. (2010a) furthermore determined an11.62±0.04 Ma age for Oarba Tuff that is located in the uppermostSarmatian part of the basin infill (Fig. 4). These radio-isotopic ages pro-vide a first order timeframe for the Middle to Late Miocene infill of theTransylvanian Basin that enables the calculation of average sedimentationrates.

At the onset of the Badenian the Paratethys was well connectedwith the Mediterranean and endemism does not hamper biostrati-graphic correlation to the global timescale. The base of the Badeniancoincides with the Early-Middle Miocene boundary (Papp et al., 1978)which is dated at 15.97 Ma (Lourens et al., 2004a). The 13.8 Ma onsetof evaporite deposition provides a minimum age for the top of thelower Badenian. Whereas the early Badenian thus lasted for over2 Ma, not more than 100 m of sediments accumulated (Krézsek andFilipescu, 2005). Rates of deposition were consequently rather low(maximum of 0.05 m/kyr).

After the Badenian salinity crisis (BSC), rates of deposition in-crease tremendously. The upper Badenian to upper Sarmatian partof the basin infill is around 2.6 km thick in the basin centre (Fig. 6).The 13.36 Ma termination of the BSC provides a maximum age forits base and the 11.62 Ma old Oarba Tuff provides an approximateage for its top. Based on these two absolute ages an average sedimen-tation rate of 1.5 m/kyr can be calculated.

Our synthetic seismic section (Fig. 4) demonstrates that theHădăreni, Ghiriş, Turda and Câmpia Turzii tuffs are all distributed inan interval with a thickness lower than 0.1 sTWT, which is around130 m given the average interval velocity of Sarmatian sediments.We collectively call these tuffs the ‘middle Sarmatian tuff complex’(MSTC) since they demonstrably co-occur in a relatively limited strat-igraphic interval and according to our 40Ar/39Ar data follow uponeach other in a short time-interval.

Based on our composite stratigraphic column for the late Badenianand Sarmatian (Fig. 6), there are about 1.44 km of sediments betweenthe top of the salt (13.36 Ma) and the base of the MSTC (12.36 Ma).The respective upper Badenian and lower Sarmatian sedimentswere thus deposited at a rate of approximately 1.44 m/kyr. Theupper Sarmatian sediments overlying the MSTC have a stratigraphicthickness near 1.16 km. Using the 11.62 Ma age for the Oarba Tuffas tie-point this package was deposited at a comparable rate of

uppe

r B

aden

ian

mid

dle

Bad

.S

arm

atia

n

Sectionname

Tracedposition

Magnetostratigraphy

A. dividens Zone

Tenuitellinata Zone

Velapertina Zone

A. dividens Zone

Badenian - Sarmatian boundary

A. dividens Zone

2.0

1.0

1 .5

2.0

1.0

1 .5

3.0

2.5

4.0

3.5

3.0

DepthParatethysstages

unce

rtai

nty

Oarba tuff (11.62±0.06 Ma)

Câmpia Turzii 2Hâdâreni Tuff (12.37±0.04 Ma)

Oarba Beta

Turda*

Câmpia Turzii*

Gherla

Apahida

Turda Tuff

Câmpia Turzii Tuff

sTwTkm

13.0

12.0

13.5

12.5

11.5

ATNTSTime (Ma)Polarity

(12.35±0.04 Ma)

50

110

5

20

150

10

(12.37±0.04 Ma)

40

top salt (13.36 Ma)

Badenian Salinity Crisis evaporites C5ABn

C5AAr

C5AAn

C5Ar.3r

C5Ar.2n

C5Ar.2r

C5Ar.1n

C5Ar.1r

C5An.2n

C5An.1r

C5An.1n

C5r.3r

C5r.2n

C5r.2r

Fig. 6. Biostratigraphy, magnetostratigraphy and 40Ar/39Ar ages for the investigated upper Badenian to Sarmatian sections and tuffs, and their consequential correlation to theATNTS (Krijgsman and Kent, 2004; Lourens et al., 2004a; Hüsing et al., 2010). Stratigraphic column and position of the sections therein based on the synthetic section in Fig. 4.Scaling and first order correlation of the stratigraphic succession to the ATNTS is based on the radio-isotopic ages for the top of the evaporites and the Oarba Tuff. Due the scaleimposed by the high sedimentation rate detailed lithological changes in the depicted individual sections would be indiscernible. Detailed bio-, litho-, and magnetostratigraphicinformation on each of the outcrops is provided in the supplementary material.


around 1.56 m/kyr. This indicates that from the end of the BadenianSalinity Crisis to the final part of the Sarmatian deposition in theTransylvanian Basin was very rapid and occurred at a fairly constantrate. Lower sedimentation rates are obviously to be expected nearthe western basin margin and in the vicinity of rising diapirs thatstarted their upward movement already in the late Badenian(Krézsek and Filipescu, 2005).

It should be noted that the calculated Badenian, Sarmatian and Pan-nonian sedimentation rates do not permit direct quantitative conclu-sions on basin evolution without a proper subsidence history analysistaking into account factors such as paleo‐bathymetry, compaction and

sea-level variations. However, these sedimentation rates provide afirst order qualitative indication of the sedimentary influx associatedwith processes such as subsidence or increases in input from the sourceareas.

6. Paleomagnetism

Sections were sampled for paleomagnetic investigation since theirmagnetostratigraphic polarity pattern can provide additional con-straints for the correlation to the global timescale and the directionaldata can be used to unravel the amount as well as the partitioning in


time of the Miocene rotation of the Transylvanian Basin (see Fig. 11and Table 2 for locations and results).

We will discuss the magnetic polarity pattern of the investigatedsections, taking the composite stratigraphic column (Figs. 6, 8) andthe stratigraphic position of the sampled sections inferred from it asa starting point. In order to place the magnetostratigraphic patternof the investigated sections into the composite stratigraphic columnat the right scale, a scaling factor has to be determined. Accordingto the geological map (Raileanu et al., 1968) and our own observa-tions the stratigraphic distance between the Badenian–Sarmatianboundary and the Câmpia Turzii Tuff is around 400 m in the areawhere the respective sections were sampled. In the basin center, thecorresponding interval is around 500 m and thus 1.25 times asthick. The upper Badenian to middle Sarmatian sections were scaledaccordingly in the composite stratigraphic column. The stratigraphicdistance between the Oarba Beta and A outcrop is, based on fielddata, approximately 530 m. In the composite stratigraphic column,the same stratigraphic interval is nearly 800 m, or 1.5 times as thick.The magnetostratigraphic pattern of Oarba Beta was thus scaled ac-cordingly, which indicates that its earlier correlation to the ATNTSby Vasiliev et al. (2010a) does not agree with the correlation to ourcomposite stratigraphic column. Its limited stratigraphic extent incombination with its position in the stratigraphy precludes the occur-rence of two reversals (Figs. 6, 8).

Approximate scaling factors for the Pannonian sections can also bederived from the sequence of outcrops at Oarba de Mure . The strati-graphic distance between Oarba A and D is, based on field data,around 120 m. In the composite stratigraphic column, the same inter-val covers about 215 m and is thus 1.8 times as thick. All Pannoniansections are scaled accordingly.

6.1. Methods

All samples were collected with a hand-held electric drill. The ori-entation of the standard paleomagnetic cores and corresponding bed-ding planes were measured by means of a magnetic compass, andcorrected for the local magnetic declination. In the laboratory, ther-mal as well as alternating field demagnetization techniques were ap-plied to isolate the characteristic remanent magnetization (ChRM).

PIG04

N

up/W

FA05

N

up/WCT2 120

N

up/W

CT16

N

up/W

TH TH T

TH TH A

marl marl m

marl mtuffite

0 ˚C

200 ˚C

325 ˚C

300 ˚C

200 ˚C

20 ˚C

100 ˚C

200 ˚C

315 ˚C150 ˚C

260 ˚C470 ˚C

Fig. 7. Bedding plane corrected characteristic demagnetization diagrams for paleomagneticare marked ‘th’ while those for alternating field demagnetized samples are marked ‘af’. SectApahida; PIG, Pâglisa; CT, Câmpia Turzii; VIF, Viforoasa.

The natural remanent magnetization (NRM) of the samples was mea-sured after each demagnetization step on a 2G Enterprises DC Squidcryogenic magnetometer (noise level 3•10–12 Am2). Heating oc-curred in a laboratory-built, magnetically shielded furnace employing10–30 °C temperature increments. AF demagnetization was accom-plished by a laboratory-built automated measuring device applying5–20 mT increments up to 100 mT by means of an AF coil interfacedwith the magnetometer. The presence of iron sulfides in the studiedsamples was anticipated. In order to overcome the problem ofgyroremanence during alternating field demagnetization the specifi-cally designed per component demagnetization scheme of Dankersand Zijderveld (1981) was applied. In addition, small (2–5 mT) fieldsteps were taken in the 20–40 mT range. The ChRM was identifiedthrough assessment of decay-curves and vector end-point diagrams(Zijderveld, 1967). ChRM directions were calculated by principalcomponent analysis (Kirschvink, 1980) and are based on at leastfour consecutive temperature or field steps.

6.2. Results

Several characteristic demagnetization diagrams, of mostly marlsand clays, are depicted in Fig. 7. Their NRM intensity ranges between1 and 1000*10−4 A/m. A viscous overprint is generally removed at100 °C or 15 mT respectively. During progressive stepwise thermaldemagnetization two and sometimes three components can beestablished. A first, low temperature component appears between100 and 200 °C. The corresponding directions are generally of normalpolarity and close to the present day field direction. A notable excep-tion is the Câmpia Turzii section, where the 100–200 °C componentbears both normal and reversed polarities. A second component ap-pears between 200 and 300 °C. For the Badenian and Sarmatian sec-tions, the corresponding directions are clearly different from thepresent-day field direction and indicative of a significant counter-clockwise rotation. Approximately 90% of the NRM had been removedat 300 °C in samples from the Făgăra , Gherla and Câmpia Turzii sec-tions. For these localities we interpret the 200–300 °C component asthe ChRM. For samples of the oimuş and Viforoasa sections, a highertemperature (300–360 °C) component could be established. At 360 °Cgenerally 99% of their NRM had been removed. We thus interpret the

AP35

N

up/W

CT21

N

up/W

VIF12N

up/W

TU33

N

up/W

H TH

F TH

arl marl

arl marl

15 mT

40 mT

80 ˚C

180˚C

285˚C

315˚C

0 mT

25 mT

100 mT

100 mT

100 ˚C260 ˚C

280 ˚C

360 ˚C

samples from the Transylvanian Basin. Diagrams for thermally demagnetized samplesions are indicated with abbreviations: CT2, Câmpia Turzii 2; FA, Făgăraş; TU, Turda; AP,

Fig. 8. Lithology, biostratigraphy, magnetostratigraphy, 40Ar/39Ar ages, and consequential correlation to the ATNTS for the investigated upper Karpatian to lower Badenian sections.1. and 5. GPTS according to Lourens et al. (2004a), Hüsing et al. (2010), and Krijgsman and Kent (2004). 2. Biozones according to Krézsek and Filipescu (2005). Ages of bioeventsaccording to Abdul Aziz et al. (2008) and Lourens et al. (2004b). 3. Central Paratethys regional stages. 4. Confidence interval for the acquired 40Ar/39Ar ages for the Dej Tuff.


300–360 °C as the ChRM for the Viforoasa and oimuş sections. In al-ternating field demagnetization diagrams of marly samples generallyonly a single component appears. The corresponding directions areestablished in the 20–40 mT interval. Above 40 mT the onset ofgyroremanence often distorts the demagnetization diagrams.

The samples taken from the Pâglişa section consisted of volcanictuff or tuffite, and displayed a different demagnetization behavior.Their NRM intensity was generally around 1000*10−4 A/m. Twocomponents were identified during thermal demagnetization. A firstcomponent appears between 100 and 260 °C. The corresponding

directions are close to the present day field direction and we interpretthis component as a present-day field overprint. A second componentis isolated between 280 and 480 °C. Although also of exclusively nor-mal polarity, the corresponding directions differ significantly fromthe present-day field direction. We interpret the 280–480 °C compo-nent as the ChRM of the Pâglişa section.

The ChRM temperature interval and gyroremanent behavior uponAF demagnetization indicate that the magnetic carrier in most sam-ples from the Transylvanian Basin is an iron sulfide, most likelygreigite. Iron sulfides are very commonly the main magnetic carrier

Fig. 9. Biostratigraphy, magnetostratigraphy, and 40Ar/39Ar ages for the investigated upper Sarmatian and Pannonian sections, and their consequential correlation to the ATNTS(Krijgsman and Kent, 2004; Lourens et al., 2004a; Hüsing et al., 2010). Stratigraphic column and position of the sections therein based on the synthetic section in Fig. 4. Scalingand first order correlation of the stratigraphic succession to the ATNTS is based on the 40Ar/39Ar for the Oarba Tuff (Vasiliev et al., 2010a) and correlation of the reversal in theŞoimuş section to the top of C5n.2n.


in Middle to Late Miocene of the Paratethys (Vasiliev et al., 2010a).Correlation of the iron sulfide based magnetic polarity patterns ofboreholes in the Pannonian Basin to the GPTS (Sacchi et al., 1997;Juhász et al., 1999; Magyar et al., 1999b; Sacchi and Horváth, 2002;Sacchi and Müller, 2004; Magyar et al., 2007) is often intricate. Thereliability of greigite as a magnetic carrier has moreover frequentlybeen questioned, because greigite is thermodynamically metastableand the timing of NRM acquisition by greigite is not well constraineddue to its diagenetic formation (Vasiliev et al., 2008). The straightfor-ward correlation of the magnetostratigraphy of greigite bearing sectionsin the Carpathian Foredeep (Vasiliev et al., 2007), and Mediterranean(Hüsing et al., 2009), however, shows that greigite may preserve itsNRM for geological times and that reliable magnetostratigraphic resultscan be obtained if only a dedicated greigite specific demagnetizationapproach is taken.

6.3. Correlation to the timescale

The polarity of the sampled sections in general corresponds to thepolarity predicted for the respective stratigraphic interval by thetime-scale correlation based on 40Ar/39Ar data and thus provides ad-ditional support (Fig. 6). The Făgăra outcrop, which has a reversedpolarity, correlates to chron C5Ar.3r. The marls underlying the tuff ex-posed at Apahida were deposited during chron C5Ar.2n, and the re-versed polarity interval exposed at Gherla correlates to the base ofchron C5Ar.1r. At the top of the Turda section the magnetic polaritychanges from normal to reverse. This reversal correlates to theC5An.2n to C5An.1r transition taking the acquired 12.37±0.04 Maradio-isotopic age for the Turda Tuff into account.

The magnetostratigraphic pattern of the Câmpia Turzii, however,requires additional explanation. The 150 m thick Câmpia Turzii sec-tion covers ~0.2 Ma. Its polarity is dominantly reversed, but a largenumber of single samples have a normal polarity. The demagnetiza-tion diagrams on which the magnetostratigraphic pattern is basedare of high quality for both normal and reversed samples. They do,on the other hand, all display multiple and sometimes even up tofour components. These can be of normal or reversed polarity inde-pendent of their demagnetization temperature or field interval. We

consider it most likely that the primary magnetic component of theCâmpia Turzii section is reversed, but demagnetization behavior isso complex that reliable magnetostratigraphic information is hard toextract.

For the uppermost Sarmatian to Pannonian part of the basin infillwe do not have 40Ar/39Ar dates that can guide first order estimates onsedimentation rates and correlation to the ATNTS. We thus have torely on the magnetostratigraphic information acquired (Fig. 9). Forthe correlation of the magnetostratigraphic pattern of the Oarba sec-tions we follow (Vasiliev et al., 2010a), except for Oarba Beta. The po-larity of the newly sampled Pannonian sections and sites is generallynormal. We thus correlate the Târnăveni, Viforoasa, and oimuş sec-tions to the long normal Chron C5n.2n (Fig. 9). The short reversedintervals of the Viforoasa section might represent small-term fluctua-tions in the geomagnetic field (Krijgsman and Kent, 2004). The up-permost part of oimuş, which bears a reversed polarity, in our viewcorrelates to C5n.1r. This correlation implies that in the Pannonianthe sedimentation rate is much lower than during the upper Badenianto Sarmatian time-interval and amounts to 0.36 m/kyr (Fig. 10). Thisis in agreement with seismic interpretations that demonstrate a grad-ual change from subsidence to uplift during the Pannonian stage(Matenco et al., 2010). Our synthetic seismic stratigraphy showsthat the Sighi oara and Mediaş sections are slightly younger thanoimuş. They have been correlated to the timescale based on the cal-culated sedimentation rate for the Pannonian part of the basin infill(Fig. 9).

Along the course of our study it has become clear that the quality ofdemagnetization results for and the preservation of NRM in the sedi-ments of the Transylvanian Basin is highly variable. Sections for whichdemagnetization resultswere of low qualitywere immediately excludedfrom further analysis. The polarity of most sections with high quality de-magnetization diagrams corresponds to the one expected based on cor-relation to the GPTS (Lourens et al., 2004a) according to our 40Ar/39Arbased chronostratigraphic framework. This consistency provides inde-pendent support for the reliability of the acquired paleomagnetic results.The intricate polarity pattern of the Câmpia Turzii section, however,points out that magnetostratigraphic interpretation in iron-sulfide dom-inated sediments remains delicate and great care has to be taken.

Fig. 10. Sedimentation rates for the upper Badenian to Sarmatian, and Pannonian part of the basin infill. It should be noted that these do not allow direct quantitative conclusions onbasin evolution without a proper subsidence history analysis taking into account factors such as paleo‐bathymetry, compaction and sea-level variations. However, these sedimentationrates provide a first order qualitative indication of the sedimentary influx associated with processes such as subsidence or increases in input from the source areas.


7. Biostratigraphy

7.1. Methods

All sections were investigated in detail for micropaleontologicalanalyses primarily focusing on foraminifera and ostracods. Sedimentsamples of approximately 1 kg were processed by standard micropa-leontological methods, sieved on 63 μmmesh and hand-picked underthe microscope. A detailed description of the micropaleontological as-semblage characteristic for each section or part thereof is provided inthe supplementary material.

Each section is attributed to one or more particular biozones(Fig. 3). The biostratigraphic zonation of the Badenian and Sarmatianwith foraminiferal assemblages is based on Popescu (2000), Filipescuand Silye (2008), and Beldean et al. (2010). The biostratigraphic zo-nation for the Pannonian is based on ostracods and follows (Jiřiček,1975, 1985), who referred to the standard Pannonian Biozones ofPapp (1951) and Papp et al. (1985).

During the early Badenian the Paratethys and Mediterranean werewell connected (Rögl, 1999; Popov et al., 2004) and biostratigraphicevents can be correlated. Endemism hampers direct correlation ofthe middle Badenian, Sarmatian and Pannonian sections. Moreover,relatively deep water environments characterized the larger part ofthe Transylvanian Basin at that time. Most of the specifically designedbiostratigraphic zonations for the Middle Miocene of the Central Par-atethys rely on shallow water benthic taxa (e.g. Görög, 1992). Theyare therefore not always suitable for the deep-water sediments ofthe Transylvanian Basin (Filipescu and Silye, 2008). This sometimescomplicates attribution of the sampled sections to a specific biozone.Fitting the ostracod assemblages from the Transylvanian Basin in thechronostratigraphic framework of the Pannonian stage is rather diffi-cult because of pertinent differences regarding the Pannonian ostra-cod biozonation between authors (Pokorný, 1944; Brestenská, 1961;Krstić, 1973; Jiřiček, 1975, 1985; Krstić, 1985; Rundić, 1997, 2006).

Assignment of the late Badenian to Pannonian sections to a specificbiozone is for all these reasons often tentative.

7.2. Results

Three successive biozones can be identified in the Ciceu Giurgeştisection (Fig. 8). The assemblage of the lowermost part of the sectionis dominated by biserial planktonic foraminifera of the Early MioceneStreptochilus pristinum Biozone (Beldean et al., 2010). The second partof the outcrop, which starts with a conglomerate bed with clasts up to15 cm in diameter, contains foraminifera of the early BadenianPraeorbulina glomerosa Biozone. Samples from the remaining part ofthe section belong to the Orbulina suturalis Biozone.

Nannoplankton data by Mészáros and Şuraru (1991) indicate thatthe Dej Tuff in the Pâglişa area pertains to the NN5 calcareous nanno-plankton zone. The foraminiferal assemblage identified at Pâglişa be-longs to the interval characterized by the S. pristinum to O. suturalisbiozones.

The Făgăraş outcrop exposes sands and marls that belong to themiddle to upper part of the late Badenian Velapertina Biozone. SEMinvestigations of samples from the Făgăraş 2 outcrop revealed fre-quent microperforate globigerinids belonging to the genera Tenuitellaand Tenuitellinata assigned to the top Badenian Tenuitellinata Biozone(Filipescu and Silye, 2008).

The faunal assemblage of the marls of the Apahida section is domi-nated by small-sized foraminifera which are mostly in their juvenilestage. Specimens of Tenuitellids occur together with Anomalinoidesdividens. The former are very common planktonics in the late Badenianand rare in the lowermost Sarmatian, whereas a bloom of latter marksthe lowermost Sarmatian. This situation suggests a transitional zonepositioned probably just below the Badenian–Sarmatian boundary. Asimilar situation occurs at Unguraş where the micropaleontologicalassemblage included several Tenuitellinata specimens and one speci-men of A. dividens. The Gherla section clearly belongs to the lower

DexpPDF

Pn

Bd+Sm

-20 -10 0 10 20 30 40 50Declination (°)

16

15

14

13

12

11

10

9

8A

ge (

Ma)

Fig. 11. PDF, present day field direction; Dexp, declination expected based on the10 Ma reference pole for Europe (Torsvik et al., 2008). Observed paleomagnetic decli-nation for each section and uncertainty therein according to results in Table 2. Sectionages follow the correlation in Figs. 6, 8 and 9. Average directions for the Badenian toSarmatian, and Pannonian time intervals are indicated with shaded rectangles. Theirwidth corresponds to the confidence interval. The difference between the expected(Dexp) and observed declination reveals the amount and direction of rotation.


transgressive part of the Sarmatian corresponding to the A. dividensBiozone. Dilution of the assemblage in combination with reworking,and a biozonation based on benthic foraminiferawhich strongly dependon the local environment (see Piller et al., 2007) makes biozonation ofthe Hădăreni, Ghiriş, Câmpia Turzii 2 and Turda sections difficult. Thediluted micropaleontological assemblage from the Hădăreni outcropcontains a few specimens of small rotaliids (Elphidium spp., Nonionspp.) and miliolids (Quinqueloculina spp., Sinuloculina spp., Varidentellaspp.). These characterize the lower half of the Sarmatian, but the exactbiozone remains unclear. The Câmpia Turzii section exposes 150 m ofclays, marls, and sandstones that belong to the A. dividens Biozone.The Câmpia Turzii 2 section, situated stratigraphically 250 m abovethe Câmpia Turzii section, exposes some 30m ofmarls and subordinate

Table 2Paleomagnetic data and rotation results. Section: name of paleomagnetic sampling locality. Agelongitude of sampling locality. Observed direction: inclination (I) and declination (D) of meancalculated applying fisher statistics to the observed directions, (A95) the 95% confidence circle(n) the number of sites used to calculate mean direction, (pol) the polarity of the site. Referen(2008). Rotation: vertical axis rotation with 95% confidence limit (positive indicates clockwisflattening are derived from observed direction minus expected direction at locality calculated

Section Age Site location Observed direction R

Lat Long Is Ds α95 k Α95 n Pol A

(Ma) (Ma) (°N) (°E) (°) (°) (°) (°N) (°E) (°) (

Mediaş 9.6±0.1 46.17 24.36 54.4 19.0 4.1 141 4.9 10 n 1Sighişoara 10.0±0.1 46.24 24.81 58.4 10.8 6.8 59 8.9 9 n 1oimuş 10.2±0.1 46.36 24.96 51.4 6.3 3.6 78 3.6 21 n 1Viforoasa 10.4±0.1 46.43 24.81 53.2 10.4 2.7 84 2.9 34 n 1Târnâveni 10.6±0.1 46.34 24.30 41.0 14.2 7.8 27 6.9 14 n 1Oarba D 11.0±0.1 46.46 24.28 66.8 6.9 3.9 114 5.9 13 n 1Oarba B 11.2±0.1 46.46 24.29 58.5 355.3 7.2 61 9.2 8 n 1Average Pn 46.25 24.50 53.4 10.1 2.0 47 2.1 106 n 1Turda 12.3±0.1 46.57 23.83 46.1 41.2 4.5 42 4.8 25 n 1Apahida 12.3±0.1 46.82 23.76 52.6 35.1 4.9 155 5.9 7 n 1Unguraş 12.7±0.1 47.09 24.09 57.2 0.3 5.8 70 7.7 10 n 1Fâgâras 12.9±0.1 45.87 25.00 51.8 24.8 12.0 51 13.0 4 r 1Pâglişa 14.3±0.2 47.01 23.66 54.3 20.1 5.3 68 6.5 12 n 1Average Bd+Sm 46.25 24.50 51.5 30.0 3.3 33 4.0 57 n 1

sandstones with on top of it an 8 m thick tuff bed, named the CâmpiaTurzii Tuff here. Biostratigraphic samples from this section contain fre-quent A. sarmatica in combination with rare A. dividens and most likelybelongs to the Articulina Zone (middle-upper Volhynian), in whichAnomalinoides is rare but still occurs. The 150 m of marls and sand-stones overlying the Turda Tuff might belong to the Elphidium reginumBiozone.

The biostratigraphy of five successive upper Sarmatian and lowerPannonian outcrops at Oarba de Mureş was studied in detail bySztanó et al. (2005), Vasiliev et al. (2010a), and Filipescu et al.(2011). Vasiliev et al. (2010a) attribute the Oarba Beta section tothe Porosonion aragviensis Biozone. Renewed inspection of the bio-stratigraphic samples, however, pointed out that the whole sectioncontains species of Elphidium, but no specimens of Dogielina sarmaticawere encountered. This demonstrates that it is difficult to apply thezonations established in shallow marine settings to the deeperareas. We thus refrain from exact biozone determination for theOarba Beta section. The larger part of outcrop A belongs to theupper Sarmatian Porosononion aragviensis Biozone. Samples fromthe top part of outcrop A contain typical Pannonian deep-water ostra-cods and thus Filipescu et al. (2011) place the Sarmatian–Pannonianboundary 2.3 m below the top of the section. This is in good agree-ment with the results of Sütő and Szegő (2008), who found massivenumbers of Mecsekia ultima 3.4 m below the top and the indextaxon for the base of the Pannonian (Spiniferites bentonii pannonicus)1.4 m below the top. Outcrops B, C and D are Pannonian in age and con-tain a rich ostracod association specific for low salinities. Sztanó et al.(2005) identified several mollusks indicative of the Lymnocardiumpraeponticum Zone of the early Pannonian in the lower part of outcrop B.

The micropaleontological assemblage of the Viforoasa and Şoimuşsections is dominated by Pannonian ostracods. The larger part of the os-tracod assemblages can be attributed to Zone 5 (Amplocypris abscissaZone) and the lower part of Zone 6 (Hemicytheria croatica Zone) of thebasal part of the upper Pannonian (lower Serbian) as defined by Krstić(1973, 1985) and mentioned by Rundić (1997, 2006). They can also becorrelated with Zone D and the first part of Zone E (Subzone E1)according to Jiřiček (1975, 1985), who referred to the standardPannonian biozones of (Papp, 1951). The lower part of Viforoasa sectionseems to be older than the Şoimuş section andmight correlate to the up-permost Slavonian (Propontoniella candeo Zone) (Krstić, 1973, 1985).The Târnăveni, Mediaş and Sighişoara sectionswere exclusively sampledfor rotation research and only few biostratigraphic samples were taken

: age of the section according to our chronostratigraphic results. Site location: latitude andpaleomagnetic direction in tilt corrected coordinates. (α95) is the 95% confidence circlecalculated applying fisher statistics to corresponding VGPs, (k) the precision parameter,ce pole: age, latitude and longitude, and A95 (95% confidence) according to Torsvik et al.e rotation). Flattening: flattening of inclination with 95% confidence limit (rotation andfrom reference pole).

eference pole p Expected I Expected D Rotation Flattening

ge Lat Long A95 Ix±δΙ Dx±δΔ R±ΔΡ F±ΔΦ

°) (°) (°) (°) (°) (°) (°) (°) (°)

0 87.2 125.0 2.5 44.4 63.9±2.0 3.9±3.6 15.1±6.3 9.5±3.60 87.2 125.0 2.5 44.3 64.0±2.0 3.9±3.6 6.9±10.8 5.6±5.70 87.2 125.0 2.5 44.2 64.1±2.0 4.0±3.6 2.3±5.4 12.7±3.30 87.2 125.0 2.5 44.1 64.1±2.0 4.0±3.6 6.4±4.6 10.9±2.70 87.2 125.0 2.5 44.3 64.0±2.0 3.9±3.6 10.3±8.8 23.0±6.40 87.2 125.0 2.5 44.1 64.1±2.0 4.0±3.6 2.9±8.5 −2.7±3.50 87.2 125.0 2.5 44.1 64.1±2.0 4.0±3.6 −8.7±11.5 5.6±6.00 87.2 125.0 2.5 44.3 64.0±2.0 3.9±3.6 6.2±3.9 10.6±2.20 87.2 125.0 2.5 44.0 64.2±2.0 4.0±3.6 37.2±5.9 18.1±3.90 87.2 125.0 2.5 43.8 64.4±2.0 4.0±3.6 31.1±7.1 11.8±4.20 87.2 125.0 2.5 43.5 64.6±1.9 4.0±3.6 −3.7±9.1 7.4±4.90 87.2 125.0 2.5 44.7 63.7±2.0 3.9±3.6 20.9±16.0 11.9±9.70 87.2 125.0 2.5 43.6 64.5±1.9 4.0±3.6 16.1±7.8 10.2±4.50 87.2 125.0 2.5 44.3 64.0±2.0 3.9±3.6 26.1±5.1 12.5±3.1


froma single stratigraphic level in each of the quarries. The faunal assem-blage indicates a Pannonian age.

8. A new chronostratigraphy for the Middle to Upper Miocene ofthe Transylvanian Basin

Our new radio-isotopic, magnetostratigraphic and biostratigraphicresults allow, in combination with the high resolution seismic correla-tion of surface outcrops, the establishment of a new chronostratigraphyfor the Middle to Upper Miocene of the Transylvanian Basin.

8.1. Late Karpatian and early Badenian: the S. pristinum, P. glomerosaand O. suturalis biozones and deposition of the Dej Tuff complex

The Ciceu Giurgeşti, Dej and Pâglisa outcrops (Fig. 8) reflect the earlyBadenian evolution of the last megasequence deposited in the Transyl-vanian Basin. The lower part of the Ciceu Giurgeşti, Dej, and Pâglişa sec-tions comprise the sediments that accumulated before the onset ofvolcanic activity related to the Dej Tuff complex. The presence of thelate Karpatian and two Badenian biozones in an interval of only 24 mcorroborates extremely low sedimentation rates. The transition fromKarpatian to Badenian (i.e. the Early to Middle Miocene transition)occurs after 7 m (measured from the base of the section) and before11.5 m where P. glomerosa is first registered. In the Mediterranean theFO of P. glomerosa was dated to occur at 15.97 Ma (Lourens et al.,2004b). A second biostratigraphically correlatable horizon (Fig. 8) isthe FO of O. suturalis, indicative for the start of the O. suturalis Biozone.It was dated to occur at 14.56 Ma in the Western Mediterranean(Abdul Aziz et al., 2008).

The 14.38±0.06 Ma 40Ar/39Ar age for the Dej Tuff from CiceuGiurge ti provides the third tie-point to the global timescale. Deposi-tion of this tuff slightly postdates the FO of O. suturalis (Fig. 8) andprovides an independent minimum age constraint on the first occur-rence of O. suturalis in the Central Paratethys. Since this dated tuff ho-rizon is one of the basal horizons of the Dej Tuff complex, the acquiredage should be indicative of the onset of related volcanism. The statis-tically indistinguishable 14.37±0.06 Ma age determined for the Dejtuff at Dej provides additional support for this conclusion. This agecontrasts strongly with the much older age of between 15.4 and15 Ma previously attributed to the Dej Tuff based on K-Ar and fissiontrack measurements (Szakács et al., 2000). Since the potassium-argonmethod is restricted to whole rock measurements, the strongreworked component that our experiments demonstrate to be pre-sent in the Dej Tuff, might have induced seemingly higher ages.Since multiple single fusion 40Ar/39Ar experiments are performedon only a small selected fraction of minerals in the rock, detectionof, and correction for reworking is possible. We thus conclude thatthe here calculated 14.37±0.06 Ma and 14.38±0.06 Ma crystalliza-tion ages provide more reliable ages for the Dej Tuff.

The part of the tuff complex sampled at Pâglişa also pertains to theO. suturalis Biozone and is exclusively of normal polarity. It was thusdeposited during chron C5ADn. Our results imply that accumulationof the over 80 m thick volcaniclastic complex took much shorterthan previously thought. Tuffs and tuffites belonging to the Dej Tuffcomplex disappear from the sedimentary record before the onset ofthe BSC at 13.82 Ma. Tuff deposition can thus have lasted a maximumof 0.6 Ma.

8.2. Late Badenian and Sarmatian biostratigraphy

The base of Velapertina Biozone corresponds to the end of the sa-linity crisis and is consequently 13.36 Ma. The boundary betweenthe Velapertina and Tenuitellinata Biozones is located in the Făgăraş2 section. Our results also provide new age constraints on theBadenian–Sarmatian boundary and the Velapertina s.l., Tenuitellinataand A. dividens Biozones. The upper Badenian Făgăraş 2 outcrop,

described by Filipescu and Silye (2008), falls into the latest BadenianTenuitellinata Biozone. The micropaleontological assemblage of themarls of the Apahida Tuff reflects the very top part of the same bio-zone and suggests it is just below the base of the Sarmatian. TheGherla section belongs to the lower Sarmatian Anomalinoides Biozoneand therefore provides an upper limit for the position of theBadenian–Sarmatian boundary, which is consequently positioned be-tween 2.73 and 2.93 km, and is dated between 12.70 and 12.83 Ma(Fig. 6). The observed micropaleontological assemblages suggestthat the boundary is stratigraphically closer to the Apahida than tothe Gherla section and is thus likely close to 12.8 Ma old (the ApahidaTuff is traditionally placed at the boundary — Vancea, 1960). This agefor the transition from Badenian to Sarmatian in the TransylvanianBasin is very similar to the age attributed to this transition along thewestern margin of the Central Paratethys by Harzhauser and Piller(2004) on the basis of sequence stratigraphic arguments, and ingood agreement with the down-well magnetostratigraphic resultsof Paulissen et al. (2011) for the Vienna Basin. It is remarkably youn-ger than the 13.32 Ma age found by Lirer et al. (2009) through cyclo-stratigraphic calibration of well-data in the same basin. Many wells inthe Vienna Basin, including the Spannberg-21 studied by Paulissenet al. (2011) are, however, known to have a hiatus spanning the inter-val of the Badenian–Sarmatian boundary. A hiatus in the Eichorn-1well cannot be excluded (Lirer et al., 2009), which may explain thediscrepancy between our results and those of Lirer et al. (2009).

The Gherla and Câmpia Turzii 1 outcrops belong to the A. dividensBiozone (Fig. 6). The transition between the A. dividens, Varidentellareussi, and A. sarmatica biozones is hard to pinpoint since it occursgradually. A. dividens continues to be present in the A. sarmatica Bio-zone although it becomes less abundant. Some taxa characteristicfor the Articulina Zone moreover already occur at the top of the A.dividens zone. However, at the level of the Câmpia Turzii 2 andHădăreni outcrops the assemblage becomes more diverse than inthe underlying strata. The abundance of Anomalinoides decreasesand the assemblage is enriched in specimens of miliolids (Articulina,Varidentella, Sinuloculina) and rotaliids (Elphidium spp.), most likelyin response to environmental changes in the basin. This suggests itis adequate to place the zone-boundary at this level. The A. dividensBiozone should, in our opinion, therefore end at the stratigraphiclevel of the Hădăreni Tuff. The Câmpia Turzii Tuff is the first tufflevel above the Câmpia Turzii 1 section, which pertains to the A.dividens Biozone. The faunal assemblage is, in analogy with theHădăreni Tuff, enriched at the level of the Câmpia Turzii Tuff. Thissuggests that the Hădăreni and Câmpia Turzii Tuffs represent a singletuff level that coincides with the top of the A. dividens Biozone. Thetop of the Anomalinoides is consequently around 12.4 Ma old. Differ-entiation between biozones in the investigated deep water sectionsbecomes more problematic above this level and we thus refrainfrom age estimates for the A. sarmatica, Dogielina sarmatica and thebase of the P. aragviensis biozones, which were separated using ben-thic taxa, which strongly depend on the local environment.

8.3. The Sarmatian–Pannonian Boundary

In the Oarba de Mure section, situated near the centre of the Tran-sylvanian Basin, the Sarmatian–Pannonian boundary is located 46 mabove the 11.62±0.04 Ma old Oarba Tuff. If the average rate of deposi-tion of 1.5 m/kyr calculated for the upper Badenian to Sarmatian infillof the basin were applicable to the Oarba de Mure A section, then theSarmatian–Pannonian boundary would be around 11.59 Ma old. Thiswould be in good correspondence with the results of Paulissen et al.(2011) and ter Borgh et al. (2013-this volume), whose mag-netostratigraphic correlations show that the Sarmatian–Pannonianboundary is also around 11.6 Ma old in the Vienna Basin and the south-ern Pannonian Basin respectively. Based on calibration of the mag-netostratigraphic pattern of outcrop A–D of the Oarba de Mure


section to the ATNTS, Vasiliev et al. (2010a) have nevertheless deter-mined an age of 11.3±0.1 Ma for the Sarmatian–Pannonian boundaryin the Transylvanian Basin. This implies significantly lower sedimenta-tion rates for themiddle and upper part of the Oarba deMure A outcrop.

Biostratigraphic evidence for a younger Sarmatian–Pannonianboundary in Transylvania is unequivocal. The basal Pannonian ostra-cod zone A is missing at Oarba de Mure (Filipescu et al., 2011),suggesting that the marine conditions characteristic for the Sarmatianstage persisted longer in the Transylvanian Basin than in thePannonian Basin and inhibited colonization by the fresh‐water ostra-cods that characterize the base of the Pannonian there. This wouldprovide additional support for a younger (i.e. 11.3 Ma) age for theSarmatian–Pannonian boundary in Transylvania, were it not that Dino-flagellate and mollusc-records (Sztanó et al., 2005; Sütő and Szegő,2008) characteristic for the basal part of the Pannonian are present.

We here choose to adopt the magnetostratigraphic correlation ofthe Oarba de Mure sections by Vasiliev et al. (2010a) and attributean age of 11.3 Ma to the Sarmatian–Pannonian boundary. This age im-plies that the Transylvanian Basin became isolated from the EasternParatethys 0.3 Ma after the Pannonian Basin became isolated, as dis-cussed in detail by ter Borgh et al. (2013-this volume).

8.4. The Pannonian

Marked differences exist regarding the Pannonian ostracod bio-zonation between different authors (Pokorný, 1944; Brestenská, 1961;Krstić, 1973; Jiřiček, 1975, 1985; Krstić, 1985; Rundić, 1997, 2006).More-over, ostracod assemblages are commonly environment sensitive. Pan-nonian ostracod associations from the central part of the TransylvanianBasin, including the Viforoasa and Şoimuş sections, reflect a more deep,basinal and possibly fresher water environment (Krézsek and Filipescu,2005) than in the other parts of Pannonian Basin where the ostracodbiozones were defined. The micropaleontological assemblage of theViforoasa and Şoimuş sections can, however, be correlatedwith theMid-dle Pannonian Zone D and the lower part of Zone E (Subzone E1)according to Jiřiček (1975, 1985). According to ourmagnetostratigraphiccorrelation of these outcrops and their position in the synthetic seismicstratigraphic column (Fig. 9), they correspond to the time-interval be-tween 10.6 and 9.9 Ma. This is in good agreement with results byHarzhauser et al. (2004) who attribute an age of 10.56 Ma to the baseof Zone D. These authors do, on the other hand, attribute an age of10.36 Ma to the top of Zone E1. Adoption of a 10.36 Ma maximum agefor the Şoimuş section would lead to an even larger difference in sedi-mentation rates for the Sarmatian and Pannonian intervals in the Tran-sylvanian Basin. Moreover, the polarity of the uppermost part of thissection warrants correlation to a reversed Chron. We therefore prefercorrelation of the reversal in the Şoimuş section to the transition fromC5n.2n to C5n.1r and infer an age of 10.0 Ma for this level. This is ingood agreement with the age of faunal assemblages similar to those ofthe Şoimuş and Viforoasa sections in the magnetostratigraphically cali-brated Beočin section in the southern Pannonian Basin (ter Borgh et al.,2013-this volume). Based on extrapolation of the sedimentation ratethat results from this correlation to the Pannonian deposits overlyingthe sampled sections results in an around 8.4 Ma age for the uppermostPannonian deposits in the depocenter of the Transylvanian Basin.

9. The rotation of the Transylvanian Basin during the Miocene

The paleomagnetic directions established in ourmagnetostratigraphicresearch canbeused to calculate thehorizontal plane rotation of the Tran-sylvanian Basin during and since the Middle Miocene and thus provideinsight in the tectonic evolution of the Tisza–Dacia terrain. The data qual-ity requirements for paleomagnetic rotation analysis are higher than formagnetostratigraphic studies since not only the polarity but also thedirection of the ChRM needs to be firmly established. The Gherla, CâmpiaTurzii and Oarba Beta sections are discarded from further analyses,

because the corresponding demagnetization diagrams are either crypticor of insufficient quality to determine a paleomagnetic direction reliableenough for rotation analysis. For each of the remaining sections andsites the average declination is subsequently calculated using Fisher sta-tistics on the VGP distribution and applying the Vandamme cutoff(Vandamme, 1994) to discard outliers (Table 2, Fig. 11).

The net rotation per site was calculated and the observed averagedeclination was with the direction of the expected magnetic field,based on the location of the paleomagnetic pole for Eurasia with acorresponding age (Torsvik et al., 2008). After inspection of theresults, sites were grouped according to age. It shows that Badenianand Sarmatian sites have a significantly larger rotation than Pan-nonian sites. The absence of major faults in the Middle Miocene infillof the Transylvanian Basin, except for those that resulted from localup-doming due to salt movements, favors the treatment of thebasin as a single rigid block. All Badenian and Sarmatian sites wereconsequently grouped together and the same was done for Pan-nonian sites. An average rotation per time interval was subsequentlycalculated (Table 2).

The average rotation of the Badenian and Sarmatian sites is 26.1±5.1°. The average rotation of the Pannonian sites is 6.2±3.9°. Thisindicates the Transylvanian Basin rotated approximately 20° CW(clockwise) from the late Sarmatian to early Pannonian and another6° CW since the Pannonian.

10. Comparison with structural geological observations

A number of structural geological features can be observed thatcorroborate CW rotation of the basin during the Middle to Upper Mio-cene and fulfill its mechanical requirements. Coincident with thetransition from Sarmatian to Pannonian in the Transylvanian Basin(11.3 Ma, Vasiliev et al., 2010a) the collision of Tisza–Dacia with theEastern-European craton evidently accelerated in the Eastern Car-pathians and Transylvanian Basin. Exhumation data demonstratethat parts of the Eastern Carpathians underwent enhanced upliftaround 12–11 Ma, coeval with the last significant event of nappe stac-king (Sanders et al., 1999; Merten et al., 2010; Necea, 2010). Alongthe eastern margin of the Transylvanian Basin, near the contact withthe Eastern Carpathians, successive uplifts on the order of 300 moccur in the latest Sarmatian and earliest Pannonian (Matenco et al.,2010). These were associated with strain partitioning along theboundary between the Tisza–Dacia and ALCAPA continental blocks,where 25–40 km of sinistral strike-slip was recorded along the DragoVodă–Bogdan Vodă (DVBV) fault system (Tischler et al., 2007;Gröger et al., 2008). This sinistral movement has been accommodat-ed by differential shortening at the contact between the Eastern Car-pathians and the European lower plate (Tischler et al., 2007) andnecessitates a concurrent clockwise rotation of Tisza–Dacia with re-spect to the European foreland on the order of 10°. In the meantime,final docking of the South Carpathians against the Moesian plate oc-curred. Amounts of S-ward vergent shortening are variable alongthe contact between them and range from zero at its westernmostlimit to around 35–50 km near the SE Carpathians ( tefănescu,1988). This asymmetry in shortening also implies around 15° ofclockwise rotation of Tisza–Dacia with respect to stable Europe.The around 10° of difference in rotation estimates between structur-al geological observations and our paleomagnetic data might be ac-commodated by further strain partitioning at the contact betweenTisza–Dacia and stable Europe (see Fügenschuh and Schmid, 2005)or reflect uncertainties in either of the observations.

The collision that started in the late Sarmatian continued until ~8–9 Ma although its effects in terms of exhumation and kinematics werereduced (Gröger et al., 2008; Merten et al., 2010). Deformation migrat-ed towards the interior of the thrust wedge forming large antiformalgeometries. These were notably larger in the northern part of the East-ern Carpathians than in the SE Carpathians (Matenco et al., 2010). This


difference in exhumation can be related to a larger amount of shorten-ing in the Eastern Carpathian areas adjacent to the northern part of theTransylvania Basin and therefore explain the observed 6° Pannonian topost-Pannonian rotations.

Our results are in agreementwith the near 30° CWrotation recordedby sediments directly underlying the Sarmatian strata of the SouthernCarpathians (Dupont-Nivet et al., 2005). Sediments younger than6.2 Ma are, according to these authors, on the other hand, not rotated.These findings demonstrate that Tisza–Dacia (including the Transylva-nian Basin and surrounding orogens) rotated as a rigid block aroundthe Moesian indenter during the Middle to Late Miocene, without therequirement of internal differential strain partitioning structures.

The observed CW rotation of the Transylvanian Basin is driven bythe eastward roll-back of the European slab during its subductionbelow the Eastern Carpathians (Royden, 1988) and provides newconstraints on the concurrent emplacement of the Tisza–Daciaupper plate into the Carpathian embayment.

11. Conclusions

Our new radio-isotopic, magnetostratigraphic and biostratigraph-ic results provide a new chronostratigraphy for the Middle to UpperMiocene of the Transylvanian Basin when stratigraphically calibratedto a synthetic seismic section in the basin center through subsurfacetracing. This improves insight in the basin's evolution and moreoverprovides new time constraints on the Middle Miocene CentralParatethys regional stages. The magnitude and timing of paleomagnet-ically determined tectonic rotations furthermore give new constraintson the emplacement of the Tisza–Dacia plate into the Carpathianembayment.

In the early Badenian (16 to 13.82 Ma), when marine connectionsbetween the Central Paratethys and the Mediterranean still existed, upto 100 m of relatively deep water sediments accumulated under a slightextensional regime (Krézsek et al., 2010). Intense volcanism led to thedeposition of the Dej Tuff complex, an important lithostratigraphicmarker for the lower Badenian. The onset of Dej Tuff volcanism is con-strained between the first occurrence (FO) of O. suturalis at 14.56 Ma(Abdul Aziz et al., 2008) and 14.38±0.06 Ma, based on new 40Ar/39Arages for two of the lowermost tuff levels. Tuff accumulation ceased be-fore the onset of the Badenian Salinity Crisis (BSC) at 13.82 Ma (deLeeuw et al., 2010). The BSCwas triggered by a glacio-eustatic restrictionof the connection between the Central Paratethys and theMediterranean(de Leeuw et al., 2010) and lead to the accumulation of up to 300 m ofsalt in the central part of the Transylvanian Basin. The establishment ofa marine connection to the Eastern Paratethys at the onset of the lateBadenian ends salt deposition restores normal marine conditions. Thestrongly endemic nature of late Badenian faunal assemblages suggestthe marine connection to the Mediterranean ceased to exist (see Rögl,1999; Piller et al., 2007). Accumulation rates strongly increased in com-parison with the early Badenian, and a marked W–E thickening of theupper Badenian in the basin suggests this relates to an increase in tecton-ic activity in the Eastern Carpathians. The transition from Badenian toSarmatian is marked by a strong faunal turnover (Harzhauser andPiller, 2007). Most foraminifera, apart from some Tenuitella andTenuitellinata (see Filipescu and Silye, 2008), and over 500 gastropodspecies disappear. We date the Badenian–Sarmatian Boundary in theTransylvanian Basin at 12.80±0.05 Ma. The upper Badenian conse-quently lasts from 13.36 to 12.80±0.05 Ma. At 12.4 Ma, a second pulseof intense volcanic activity occurred, leading to deposition of the Ghiriş,Hădăreni, Turda and Câmpia Turzii tuffs that are collectively called themiddle Sarmatian tuff complex. The tuffs accumulated in a short timeas indicated by their statistically indistinguishable ages of 12.38±0.04 Ma, 12.37±0.05 Ma, 12.37±0.04 Ma, and 12.35±0.04 Ma. The in-terval between the onset of the Sarmatian and themiddle Sarmatian tuffcomplex belongs to the Anomalinoides dividens Biozone which accord-ingly covers the 12.80±0.05 Ma to 12.4 Ma time interval. The basin is

struck by a second major environmental change with a heavy impacton its biota at the Sarmatian–Pannonian boundary, in Transylvaniadated at 11.3 Ma by Vasiliev et al. (2010a). The complete disappearanceof foraminifera from the faunal record and a subsequent proliferation ofostracodsmark a transition to freshwater conditions registered through-out the whole Central Paratethys. This was triggered by tectonic uplift ofthe Eastern and Southern Carpathians which isolated the Central Para-tethys from the Eastern Paratethys (Krézsek et al., 2010). Concomitantenhanced uplift in the Carpathians is corroborated by the timing of pa-leomagnetically determined tectonic rotation of the Tisza–Dacia block.The average rotation of sampled Badenian and Sarmatian sites is 26.1±5.1°. The average rotation of the Pannonian sites is 6.2±3.9°. This indi-cates the Transylvanian Basin rotated approximately 20° CW from thelate Sarmatian to early Pannonian. This rotation was accommodatedsinestral strike slip along theDrago Vodă–BogdanVodă (DVBV) fault sys-tem and differential shortening observed in the Eastern and SouthernCarpathians. Magnetostratigraphic correlation of the investigatedPannonian sections, which belong to middle Pannonian Zone D, and thelower part of Zone E (Subzone E1) of Jiřiček (1975, 1985), to the ATNTS(Lourens et al., 2004a) indicates that these cover the 10.6–9.9 Ma time-interval. This suggests an 8.4 Ma age for the uppermost Pannoniandeposits in the depocenter of the Transylvanian Basin. Our new chrono-stratigraphic results confirm that during the late Badenian and Sarmatian,coincident with intensive nappe stacking in the neighboring EastCarpathians (Matenco et al., 2007), sediment accumulation rates in thecenter of the Transylvanian Basinweremuch higher than during the Pan-nonian, when a gradual change from subsidence to uplift occurred(Matenco et al., 2010). In the Eastern Carpathians, deformation at thistime migrated towards the interior of the thrust wedge. The resultingantiformal geometries were significantly larger in the NE Carpathiansthan in the SE Carpathians (Matenco et al., 2010) which can explain theobserved 6° Pannonian to post-Pannonian rotation.

Supplementary data to this article can be found online at http://dx.doi.org/10.1016/j.gloplacha.2012.04.008.

Acknowledgments

We would like to express great appreciation for the commitment ofAndrei Briceag, ArnoudSlootman, and JaapVerbaaswhohave contributedto the conceptual and practical advancement of this research. This studywould not have been possible without the benevolence of multiplebrick factories that have provided access to their quarries. We thank Mr.Vasile for his hospitality, Mr. Moldovan for logistic support, Roel vanElsas for assistance with mineral separation, Jan Wijbrans, GuillaumeDupont-Nivet, Douwe van Hinsbergen for discussion, and Cor Langereisand four anonymous reviewers for critically reviewing the manuscript.This study was supported by the Netherlands Research Centre for Inte-grated Solid Earth Sciences (ISES) and by the Netherlands Organizationfor Scientific Research (NWO/ALW).

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