Geomorphology 134 (2011) 470–482

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Geomorphology

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Wind erosion under cold climate: A Pleistocene periglacial mega-yardang system inCentral Europe (Western Pannonian Basin, Hungary)

Krisztina Sebe a,⁎, Gábor Csillag b, Zsófia Ruszkiczay-Rüdiger c, László Fodor b, Edit Thamó-Bozsó b,Pál Müller b, Régis Braucher d

a University of Pécs, Department of Geology, 7624 Pécs, Ifjúság ú. 6., Hungaryb Geological Institute of Hungary, 1143 Budapest, Stefánia út 14., Hungaryc Eötvös University, Institute of Geography and Earth Sciences, Department of Physical Geography, 1117 Budapest, Pázmány P. sétány 1/c, Hungaryd CEREGE, UMR6635, CNRS, Aix Marseille Universités, BP80 13545, Aix en Provence, France

⁎ Corresponding author. Tel.: +36 304690717; fax: +E-mail addresses: krisztina.sebe@gmail.com (K. Sebe

rzsofi@ludens.elte.hu (Z. Ruszkiczay-Rüdiger), fodor@mbozso@mafi.hu (E. Thamó-Bozsó), muller.paal@gmail.cobraucher@cerege.fr (R. Braucher).

0169-555X/$ – see front matter © 2011 Elsevier B.V. Adoi:10.1016/j.geomorph.2011.08.003

a b s t r a c t

a r t i c l e i n f o

Article history:Received 19 May 2011Received in revised form 1 August 2011Accepted 2 August 2011Available online 7 August 2011

Keywords:YardangPannonian BasinPeriglacialClimatePleistoceneWind erosion

Mega-yardangs are known to form under extremely arid, warm or in some instancesmoderate climates. Here wereport on themega-yardangfield of the Pannonian Basin (Hungary),whichwas uniquely formedunder periglacial,probably semi-arid conditions, and discuss its implications on factors controlling yardang formation in general.The yardang field is composed of advanced, streamlined forms and of less aerodynamic ridges separated by well-developed corridors. The length of the largest features exceeds 60 km, their height varies around 150 m. Togetherwith ventifacts, deflation hollows and wind-blown sand areas, the yardang field is part of an integrated eolianfeature system driven by northerly to northwesterly winds. Dating of ventifact surfaces with in situ producedcosmogenic nuclides places the onset of significant wind erosion to the Early Pleistocene (1.5 Ma). OSL-datedyoung loess, fromwhich well-developed yardangs were sculpted, and wind-blown sands show that eolian actioncontinued until the early Holocene. The long temporal range of wind erosion compared to average yardangformation rates and the rarity of advanced yardang forms indicate intermittent and relatively short, most probablymillennial-scale deflation periods. Since the Early Holocene the system has been inactive, and demonstrates thatyardangs can persist in a recognizable form for several millennia even under less arid conditions.The main control on yardang formation appears to be the existence of strong unidirectional winds, in thePannonian Basin facilitated mainly by the channeling effect of topographic lows in the surrounding mountainchains. Yardangs can form even under semi-arid climates; however, the rarity of these occurrences suggests thatthis requires a special association of environmental conditions. Yardang formation is facilitated by homogeneouslithology, namely by widespread weakly consolidated, sandy sediments. The distribution of yardang fields iscontrolled by vertical neotectonic movements, which provided uplifted areas to be sculptured and topographicobstacles that sheltered downwind sediments. Faulting can be excluded as a controlling factor in mega-yardangformation, nor is the role of jointing supported.

36 72501531.), csillag@mafi.hu (G. Csillag),afi.hu (L. Fodor),m (P. Müller),

ll rights reserved.

© 2011 Elsevier B.V. All rights reserved.

1. Introduction

Mega-yardangs are among the largest geomorphic features pro-duced by wind erosion. Themaximum length of these long and narrow,often streamlined ridges can reach several tens of kilometers, theirwidth more than 1 km, their height 100 m, with the width of valleysseparating them reaching as much as 2 km (Goudie, 2007; Laity, 2009).Yardangs are present in all continents on Earth except for Antarctica,though in Europe only a small, partly relictfield ofmeso-scale forms has

been reported from Spain (Gutiérrez-Elorza et al., 2002). Besides theseveral unresolvedquestions in connectionwith their formation (see forexample Laity, 2009), their study is also important from the aspect ofplanetary geomorphology, since yardangs have been identified in largenumbers onMars and probably onVenus aswell (Goudie, 2007; de Silvaet al., 2010).Most yardangs havebeen reported to formunderhot desertclimates, while a few fields are located in Andean plateau deserts withmoderate temperatures (Inbar and Risso, 2001; Bailey et al., 2007; deSilva et al., 2010); no publications on yardang formation under coldclimates are known to us. Fossil—Permian/Triassic andMiddle Jurassic—large-scale yardangs have been described in the literature in only a fewcases (Tewes and Loope, 1992; Jones and Blakey, 1993); most of theknown yardang fields are active (e.g. Halimov and Fezer, 1989) or theirformation has stopped only recently, probably during the late Holocene(e.g. Vincent and Kattan, 2006). They are considered to be short-livedfeatures under less arid climates, since vegetation quickly stops wind

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erosion and fluvial erosion renders the eolian forms unrecognizable(McCauley et al., 1977a; Laity, 2009).

The Pannonian Basin, surrounded by the Carpathians, the Alps andthe Dinarides, is the largest intermountain basin of Europe (Fig. 1). Themajority of its area is covered by the Latest Miocene clastic sequence ofLake Pannon, a remnant of Central Paratethys, and by Quaternaryterrestrial sediments including loess. The topography of the westernpart of the basin is dominated by a fan-shaped system of rectilinearvalleys and ridges (Fig. 2). The formation mechanism of the valleysystem has been debated for more than a century. The main theoriesexplaining this peculiar pattern have been the following: 1) the locationand strike of valleys was determined by faults and most of the materialfrom the valleys was removed by wind erosion (e.g. Lóczy, 1913;Cholnoky, 1929; Jámbor, 1992, 2002); 2) the dominant agent in theformation of valleys and ridges was tectonics, and their appearancewasonly slightly modified by eolian and fluvial processes (Gerner, 1992;Magyari et al., 2005); and 3) the location of valleys was controlled byfaults and fluvial erosion incised them (e.g. Erdélyi, 1961, 1962; Szilárd,1967; Marosi, 1970). Our research over the past few years revised thequestion and showed that the system of valleys and ridges is actually aseries of yardangs and corridors. The arrangement, location andorientation of ridges and valleys was determined by wind erosion, andother processes played only a minor role by modifying their shapes(Fodor et al., 2005; Ruszkiczay-Rüdiger et al., 2007, 2009; Csillag et al.,2010). Geomorphological and chronological data indicate yardangformation mostly during the cold stages of the Pleistocene, whichwere characterized by a periglacial climate in the Pannonian Basin. Ourstudy area is therefore the first mega-yardang system known to besculptured under cold-climate conditions.

The aim of this article is twofold. First we describe the ridge-and-valley topography in the western Pannonian Basin and present lines ofevidence proving its eolian erosional origin. After this we briefly

Fig. 1. The Pannonian Basin (PB) in Central Europe

describe the Pleistocene climatic and surface conditions of thePannonian Basin. We then discuss the implications of these resultsconcerning yardang formation in general. The main criteria widelysupposed to be necessary for yardang formation are an extremely aridclimate with scattered or no vegetation; a unimodal wind regime or abimodal one with a strongly dominant direction; in addition, relativelyhomogeneous lithology and jointing are also considered important(McCauley et al., 1977a; Goudie, 2007; Laity, 2009). We try to answertwoquestions in connectionwith eachof these criteria: 1)was the givenfactor active during the development of the eolian landscape of thePannonian Basin, and 2) is it essential for yardang formation in general?We also examine the preservation potential of yardang fields.

2. Geomorphological and geological setting

The Pannonian Basin is located in eastern Central Europe and isencircled by 2000–3000 m high mountain ranges (Fig. 1). About twothirds of the basin surface is occupied by plains lying between 70 and150 m a.s.l, while the rest is comprised of hills and low mountains,with peak elevations around 250 and 900 m a.s.l., respectively. Thebasin has never been glaciated, ice caps and glaciers formed only onthe surrounding mountain ranges. The ridge and valley systemoccupies the western half of the basin (Fig. 2), its extent is roughly280 km in the NE–SW and 180 km in the NW–SE directions. Locatedbetween the latitudes 46 and 48° N, this is the northernmost reportedoccurrence of yardangs.

From the aspect of wind erosion, surface geology can be subdividedinto two types. 1) The consolidated rocks of the NE–SW-strikingTransdanubian Range: Paleo-Mesozoic rocks, Paleogene–Neogenesediments, and lavas and pyroclastics of the Miocene–Quaternarycontinental basalt volcanism. These rocks bear widespread marks ofsmall-scale wind abrasion (Budai et al., 1999, 2008; Csillag et al., 2010).

. Ice sheet limits from Svendsen et al. (2004).

Fig. 2. Shaded relief model of the western part of the Pannonian Basin showing the fan-shaped system of rectilinear valleys and ridges. LHP: Little Hungarian Plain; MF: Mezőföld;BS: Belső-Somogy; M: Marcali Ridge; T: Tapolca Basin; K: Kál Basin; MG: Mór Graben.

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2)Thepoorly consolidated, thickUpperMiocene sediments—dominantlysands and silts—of Lake Pannon, a remnant of the Central Paratethys, anda variably thin Plio-Quaternary fluvial sequence also dominated by sand.These siliciclastics are highly erodible and provide ample sources ofcorrasive material for wind abrasion. Mega-yardangs formed in theuplifted parts of these deposits.

3. Eolian landforms

3.1. Mega-yardangs and corridors

The dominant geomorphic features on poorly consolidatedsediments in the Western Pannonian Basin are the narrow, verylong, straight valleys (corridors). Their maximum length exceeds80 km, their bottom width varies between 650 and 2500 m withtypical widths around 600–700 m, which classifies them among thelargest features of this type on Earth. The corridors are usually flat-bottomed or gently U-shaped. In areas undergoing subsidence—e.g. Wand SW of Lake Balaton (Fodor et al., 2005)—they have partially beenfilled up by fluvial and lacustrine sediments and, thus, are wider now,reaching a width of as much as 4 km. One of the longest valleys isoccupied by an important river of the region (river Zala) due totectonically induced river capture (Gábris and Mari, 2007), but mostvalleys lack an established drainage pattern or have only underfitstreams, and possess one or more valley-floor divides. Late Pleisto-cene to Holocene denudation processes have caused headward

erosion of streams; however, in the majority of the corridors thesegments near the valley-floor divides are remnants of the originalhorizontal, smooth wind-carved surface. At some sites wind-blownsands occur within the corridors. Though the width of some ridges(e.g. Marcali Ridge, Fig. 2) is evidently controlled by the size of anupwind obstacle, the wavelength of the ridge-and-valley system incross-section is remarkably uniform, with a modal value of valleyspacing around 7 kmmeasured in the central part of the system, southof Lake Balaton. This is close to the spacing of large yardang corridorsmeasured by Brookes (2001). The valleys and the enclosed ridges arenot exactly parallel, instead they converge toward a 60–70-km widearea at the transition of the Eastern Alps and theWestern Carpathians.

The long ridges separated by the described corridors have twomorphological types (Fig. 3).

1) Features with a classical, streamlined shape, definitely emergingon all sides from the surroundings.

2) Features with a streamlined shape on the windward side andlaterally bounded by corridors. The delimitation of these land-forms is uncertain at their lee ends. These are the largest featuresof all.

Thedistributionof the ridge types is not random.Type1 ridgesdirectlyface the northerly–northwesterly winds and are located in the westernthird of the field and in theNW foreland of the Transdanubian Range. Thenortheastern two thirds of the yardang field are located downwind of theTransdanubian Range and are dominated by type 2 ridges.

Fig. 3. Relationship of mega-yardang types and other eolian landforms with topographic obstacles. The numbers mark the prows of mega-yardangs and indicate their type asdescribed in the text.

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Ridges were flat-topped originally, though small troughs parallel tothe yardang axeswere carved into someof themby thewind. The ridgeshave beendissected to variable extent by subsequentfluvial erosion andshaped bymass movements. Ridge flanks have convex–concave slopes,with their lower parts covered by slope sediments up to 10 m thickshowing obvious inactivity of the valleys as wind channels.

Yardangs can be recognized primarily by their shape and by theirparallel arrangement infleets facing thedominantwind (McCauley et al.,1977a; Laity, 2009). They are elongate, characteristically streamlinedridges often likened to inverted ship hulls: their prow is blunt-ended,they are usually widest in their first third and have tapering tails. Theytend to occur on plains, separated by U-shaped or flat-bottomed troughsor—in amore advanced stageof evolution—scattered individually.Mega-yardangs are distinguished frommicro- andmeso-yardangs by their sizeand by their tendency to have a less aerodynamic shape (Goudie, 2007;Laity, 2009).We can state that the studied linear ridges and valleys of thePannonian Basin fulfill these requirements and can be interpreted asyardangs and yardang troughs/corridors, respectively. The overall shapeof both meso- and mega-scale ridges and of valleys corresponds to thedefinition, in spite of their subsequent modification and dissection byfluvial processes and mass movements. More importantly, they form acoherent system, with a consistent alignment of ridges and roughlyuniform size and wavelength in cross-section, characteristic of yardangfields (Mainguet, 1970; Laity, 2009).

3.1.1. Origin of the linear valley system— alternativesHere we discuss some agents which have been cited as alternative

explanations of the landforms.

3.1.1.1. Tectonic origin. A strong tectonic control on NW–SE-trendingridges and valleys has been a widespread idea in Hungarian literature.The major “argument” for tectonic control was, in most cases, thelinearity of landforms. However, with the help of industrial seismic

reflection profiles the Miocene and partly the neotectonic Plio-Quaternary fault pattern was imaged in detail (Fodor et al., 1999;Ruszkiczay-Rüdiger et al., 2007; Bada et al., 2010, respectively). Thesestudies show that faultswith displacements larger than the resolution ofthe data (e.g. N20 m) have little in commonwith linear N–S to NW–SE-trending valleys. Seismic data from the eastern basin of Lake Balatonallow a step even further: no structures of N20 cm displacement can berecognized in thenorth-westerly extension of valleys (Badaet al., 2010).In some cases, there is a rough parallelism between imaged faults andlandforms, like in the Mór Graben and Tapolca Basin, which act as windchannels (Fig. 3). However, detailed mapping demonstrated that thefaults are curved, branching, and show stepovers, and do not matchvalley margins exactly (Fodor et al., 2008).

The presence of small-displacement faults (e.g. with displace-ments b20 m) cannot be ruled out by seismic reflection profiles, butcan be checked using outcrop-scale fault-slip data and detailedmapping. Surface structural measurements reported numerous faultswith variable orientations, part of them parallel with NW–SE-strikinglandforms (Magyari et al., 2005; Bada et al., 2010). The variability offracture directions excludes their origin from one homogeneous stressfield, i.e. from one deformation phase. One should either supposechanges of the stress field during the Quaternary or attribute some ofthe fractures to non-tectonic processes, like gravitational sliding,decompressional fracturing along steep slopes etc. An additionalproblem is related to fracture scaling: faults with b20 m displacementshould follow displacement-length scaling laws (Walsh andWatterson,1992) and, thus, should not be as longas theNW-trending linear valleys.These faults might have controlled the formation of linear landformsonly locally but not along their entire lengths.

The landform-controlling role of joints or of fractures withunknown displacement is more problematic; they cannot be analyzedby fault-slip data, and do not necessarily obey the scaling laws. On theother hand, outcrop-scale joint data show similar patterns as small-

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scale faults, thus these fractures seem to have the same problems ininterpretation (stress field, gravitational origin etc.).

Some valleys south of Lake Balaton were interpreted to haveformed along small folds (Magyari et al., 2005). The supposed foldswere observed locally, and thus they can only control very short valleysegments. In some cases (e.g. Mór Graben) mapping and dip data areclearly against the presence of folds parallel to valleys.

All these remarks show that the role of large faults can be ruledout, and the possible role of other structures (small faults, joints,folds) might have exerted only a very limited and localized control onyardang formation.

3.1.1.2. Fluvial origin. A fluvial origin of the valleys is problematic forseveral reasons. Their morphology does not show the characteristicfeatures of fluvial valleys either in map view or in cross-section. Longtracks of flat valley-floors near the drainage divides lack significantalluvial sediments and signs of fluvial erosion, and in several cases areoverlain by wind-blown sand. Besides, no fossil fluvial network isknown that would have been able to carve the valleys. For example,south-bound waters of the Balaton Highland—a potential source area—are presently drained by a NE–SW-oriented stream (Eger Creek).Maars next to themodern stream, which erupted on valley floor (Auer

Fig. 4. Meso-yardangs. a) A row of meso-yardangs in a fault-controlled wind channel acrossediments (near Bakonysárkány; location shown on Fig. 4a). Wind came from the NW.

et al., 2007), indicate that a similarly oriented valley existed here asearly as the Pliocene.

3.2. Meso-yardangs

Around the field of mega-yardangs as well as between the large-scale landforms similarly sculptured and oriented but smaller troughsand ridges can be observed in a zone of varying width. Although theirsize classifies them into the mega-yardang category of Goudie (2007)and Laity (2009), in this paper we refer to them as meso-scale forms.

Meso-yardangs are typical features on Tertiary and Quaternaryweakly consolidated sediments of the study area. They are usuallyadvanced, streamlined forms, with lengths often above 1 km (Fig. 4).Their length:width ratio is usually around 4:1–5:1.The largest of them(Látó-hegy in the Mór Graben, Fig. 4a) is 2.9 km long, nearly 600 mwide and 40 m high at its front. Meso-yardangs are most frequent onLate Pleistocene geomorphic levels but also occur on older surfaces.

The Mezőföld region E of Lake Balaton (Fig. 2), which is mostlycoveredby thick loess, is dominated bymore subdued landforms. Ridgesand troughshavemuch lower relief thanmega-yardangs;however, theyare similarly long and straight. Their linearity, orientation andparallelism indicates that they are wind-carved features as well.

s the Transdanubian Range. L: Látó-hegy. b) Meso-yardang carved of Oligocene fluvial

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3.3. Ventifacts

Ventifacts occur all over the area of the fan-shaped valley systemand directly upwind of it where appropriate lithologies are present(Figs. 5 and 6). They are absent elsewhere, except for a fewoccurrences in the NE corner of the Pannonian Basin. Ventifacts areof diverse materials: quartzite, sandstone, limestone, dolomite, flint,basalt, silicified wood and calcretes (Jámbor, 1992; Csillag et al.,2010). They dominantly occur as pebbles within or on the surface ofalluvial fans, fluvial terraces or pediment sheet-wash deposits, butlocally extensive bedrock surfaces have been wind-abraded as well.The latter are typical on strongly cemented Lake Pannon sandstonesbut also occur on Permian sandstones and Tertiary basalts (Csillag etal., 2010; Ruszkiczay-Rüdiger et al., in press). These bedrock ventifactsbear abundant wind striations, which—especially flutes—are consid-ered to be the most reliable indicators of air flow direction (Knight,2008). The 11 sites that were suitable for measuring flute directionsare located along a 120 km-long line roughly perpendicular to theorientation of the mega-yardangs; all of them referred to NNW to NWwinds (Fig. 3).

3.4. Deflation hollows

Wind-carved depressions are numerous (Budai et al., 2008; Fodoret al., 2008; Csillag et al., 2010) and occur in various topographicalpositions. The largest ones are several km long, partly filled by shallowlakes and are located on the lee side of the Transdanubian Range withtheir long axis parallel to the mountains (Fig. 3). Their formation wasprobably initiated by the topographic obstacle formed by thismountain range as described in the model of Mainguet (1972). Thesmallest blowout hollows are only a few tens of meters in diameter,occur on hummocky pediment surfaces and were produced by thedeflation of loose sand from between patches capped by protectivehard calcretes.

3.5. Wind-blown sands

Wind-blown sands are widely distributed (Fig. 6). In the NE andNW part of the basin, they are only superficially reworked fluvialsediments. At some sites near the Transdanubian Range (e.g. Császár,see Section 3.7.2) several m-high dunes accumulated. In Belső-Somogy (Fig. 2) eolian sand is up to 10 m thick (Marosi, 1970). Thelargest and thickest accumulation lies in the middle of the basinbetween the rivers Danube and Tisza, downwind of the fan-shapedvalley system. Here the total thickness of wind-blown sandsintercalated with fluvial deposits can reach up to 150 m and theyoverlie Lower Pleistocene fluvial deposits (Miháltz, 1953; Molnár,1961, 1965). Based on the downwind grain size decrease and/orthickness increase, these wind-blown sands were originated from theNW (Cholnoky, 1940; Rónai, 1985), while those in Belső-Somogy fromthe north (Marosi, 1970); Borsy (1977) inferred NNW to NNE winddirections from parabolic dune orientations in NE Hungary. The

Fig. 5. Ventifacts. (A) Bedrock ventifact: silica-cemented Upper Miocene sandstone. (B)WindSE foreland of Vértes Hills. Coin is 2 cm in diameter.

superposition of several generations of dunes and deflationallandforms (e.g. Marosi, 1970) indicates variable climatic conditions.

3.6. The mega-yardangs in the eolian system of the Pannonian Basin

Wind directions inferred from dune orientations are in goodagreement with the orientation of yardangs, corridors and windstriations, suggesting that these features are all parts of the sameeolian system. Because of the limits of dating techniques, the temporalcorrelation of deflation and sand dune accumulation is problematic.Considering the sand fraction, the Pleistocene Pannonian Basin was anearly complete eolian sedimentary system from source to sink(Fig. 6), but probably the largest part of dust-sized material alsomoved within the basin.

Eolian landforms indicate that the majority of geomorphicallyeffective winds entered the Pannonian Basin between the EasternAlps and the Western Carpathians. The restricted distribution ofventifacts and wind-blown sands in the eastern half of the basin andparabolic dune orientations point to another entrance of lesserimportance within the Northeastern Carpathians. We suppose thatdeflation has been most intense in the Little Hungarian Plain, wherefunneled winds enter the basin (similarly to the case described byWashington et al., 2006). The marginal areas of the plain are wind-beveled lowlands which host advanced forms of yardangs (Fig. 3).Fluvial sediments have been entrained and transported by the wind,as shown by Upper Pleistocene wind-reworked fluvial (“fluvio-eolian”) sands and by new ventifact finds on the plain.

The uplifting region of the Transdanubian Range and the belt ofmega-yardangs have been areas of both deflation and sand transport.From themountain range, thick UpperMiocene sandy sediments havebeen removedmostly by thewind from hilltops and small depressions(e.g. Kál Basin) (Csillag et al., 2010; Ruszkiczay-Rüdiger et al., inpress). A large amount of material is missing from the inter-yardangcorridors as well. The range as a topographic obstacle also initiated theformation of deflation hollows on the lee side.

From the eolian sand fields, Belső-Somogy (Fig. 2) has long beenidentified as a deposition area linked to the inter-yardang corridors(Lóczy, 1913; Cholnoky, 1929). However, the material of the largestdune field between the modern Danube and Tisza rivers (Fig. 6),according to previous authors, originates from the alluvial fan of theformerly NW–SE trending Danube, mostly based on mineralogicalinvestigations (e.g. Molnár, 1961). In contrast, we propose that UpperMiocene–Quaternary sediments west of the Danubewere carried hereby northwesterly winds and provided a considerable part, possiblythe bulk of the accumulated eolian sand. This view is supported by thelocation of the dune field downwind of the yardangs. In addition, inthe Early–Middle Pleistocene the Danube could not serve as asediment source under a northwesterly wind regime because it raneast of its present position (Gábris and Nádor, 2007). However, fromthe LateWürmian, when the river occupied its modern course (Gábrisand Nádor, 2007), the two sediment sources could become mixed.Even then, the Danube, whichmust have been frozen inwinter—in the

flutes on Permian sandstone. (C) Ventifacted pebbles in a Pleistocene alluvial fan in the

Fig. 6. The periglacial eolian feature system of the Pannonian Basin. Ny: Nyírség. White arrows indicate winds entering the basin. The ventifact distribution data set is grounded onthe summary of Jámbor (1992), extended by numerous own field data and a few literature references.

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main deflation season—did not act as an obstacle against eolian sandtransport from the west.

Without discussing dust transport, we must note that during thewindy intervals of the Pleistocene, the Pannonian Basin may haveacted as a major dust source, similarly to several global dust hotspotsoccurring in topographic depressions (Washington et al., 2006). Thedust could originate from the fine fractions of the Upper Miocene andyounger (fluvial) sediments but also from the mechanical weathering(grinding) of the sand fraction of these deposits during transport, in away supposed by Smith et al. (1991) and Enzel et al. (2010).

3.7. Chronological data

In order to establish an absolute chronology ofwind erosion events,wind-polished surfaceswere datedusing in situ produced cosmogenic10Be in the CEREGE-CNRS laboratory in France (Ruszkiczay-Rüdigeret al., in press), while OSL dating was undertaken at the GeologicalInstitute of Hungary (with gamma spectrometry measurementsperformed at the Eötvös Loránd Geophysical Institute of Hungary)on wind-blown sands and other sediments which had connection towind activity.

3.7.1. Exposure age dating of wind-polished rock surfaces using in situproduced 10Be

Ventifacts were sampled in the Kál and Tapolca Basins (Figs. 2 and3), where denudation is considered to have happened mostly viadeflation. These are erosion resistant, silica-cemented remnants of theUpper Miocene Lake Pannon sediments, which are otherwise easilyaffected by erosion. They are composed mainly of quartz and are,therefore, ideal for cosmogenic 10Be studies.

Recently it has been demonstrated that it is possible to determineboth denudation rate and exposure age when using depth profiles andconsidering all particles (neutrons and muons) involved in theproduction of cosmogenic nuclides (Siame et al., 2004; Braucher etal., 2009; Hein et al., 2009). Using the erosion rate-corrected exposureages of these rock surfaces lying at various elevations—constrained bymodeling the changes of the 10Be concentration with depth—it ispossible to quantify the regional denudation rate of the surroundingloose Upper Miocene deposits.

3.7.1.1. Exposure ages. Depth profiles on hard rock provided 10Beexposure age—denudation rate pairs as shown in Table 1 (Ruszkiczay-Rüdiger et al., in press). Accordingly, the 200 m level in the basins wasexposed 1.56±88 Ma ago, the 170 m horizon 865±65 ky ago, andeolian denudation reached the 140 m level 287±23 ka ago. At 175 melevation (Kővágóörs), 10Be measurements occurred along a loosesand depth profile with the aim of constraining directly thedenudation rate of this material, and control our calculations usingthe age and elevation of the hard rock depth profiles. Fast denudationof the loose sand did not allow the determination of the exposure timeat this location; the reported age reflects only a steady state ofcosmogenic nuclide accumulation and decay.

A first conclusion on exposure ages derived from the depth profilesdemonstrates that deflation was active and strong already ca. 1–1.5 Ma ago. No wind erosion of similar age has been reported fromEurope so far.

3.7.1.2. Denudation rates. Measurements of 10Be concentrations alongdepth profiles of exposed wind-facetted rocks allowed us to derive a

Table 1Cosmogenic 10Be exposure age—denudation rate pairs for the sampled depth profiles (for more details and measurement technique refer to Ruszkiczay-Rüdiger et al., in press).

Location Latitude (DD) Longitude (DD) Elevation (m) Lithology 10Be exposure time (y) 10Be denudation rate(m/My)

Kelemen-kő 46.8906 17.5560 198 Sandstone 1560,534 ±88,442 3.44 ±0.19Kővágőörs 46.8642 17.6217 175 Sand 257,293 ±46,871 56.07 ±11.95Kőmagas 46.8648 17.4978 171 Sandstone 864,813 ±65,191 3.84 ±0.29Salföld 46.8323 17.5630 141 Sandstone 286,799 ±23,121 0.00 ±0.00

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local denudation rate of 3.44–3.84 m/My. This moderate rate explainswhy the silica-cemented rocks form local highs.

Secondly, a regional denudation rate was calculated using the 10Beexposure times and the elevations of these wind-polished hard rocksurfaces, which are interpreted as remnants of fossil geomorphichorizons. Accordingly, regional denudation in the Tapolca and KálBasins varies between 40 and 80 m/My, which is in harmonywith the10Be denudation rate of 56.1±12.0 m/My calculated directly from theloose sand depth profile at Kővágóörs (Ruszkiczay-Rüdiger et al., inpress).

3.7.2. OSL ages of sediments related to wind activityTo date periods of strong wind activity sediments for OSL dating

were collected from four outcrops in the forelands of the Vértes Hills(Fig. 4a). OSL dating was performed on quartz grains extracted fromthe 160 to 200 μm fraction of the sediments by the standard methodof sample preparation. Luminescence measurements were carried outusing a Risø TL/OSL DA-15C/D automatic reader with a calibrated90Sr/90Y beta source (more details on the measurement techniqueare in Thamó-Bozsó et al., 2010a). Equivalent doses of the sampleswere measured on small aliquots, and they showed a more or lessasymmetric distribution. Therefore, based on the statistical charac-teristics of the distributions (Bailey and Arnold, 2006; Arnold et al.,2007), the OSL ages were calculated by the central and minimum agemodels of Galbraith et al. (1999). Considering the previouslypublished age data and the geomorphology of the area, the centralages of our samples represent the true depositional ages inmost cases.

The central age of the dune sands with 5 m-high foresets at theCsászár sand pit (sample ‘Cs’) is 9–8 (±1) ka, though their minimumage, 6–4 (±1) ka is also acceptable (Thamó-Bozsó et al., 2010a), i.e. itis not contradictory to the climatic and landscape evolution history ofthe area. Sand dunes with similar OSL ages were also reported fromDunavarsány (Ujházy et al., 2003), located about 70 km to ESE. Theseages indicate strong wind activity in the early Holocene, which wasable to accumulate large dunes.

Two loess sections were dated, because some yardangs werecarved from young loess. Among themeso-yardangs which formed onloess, the Kazal Hill at Lovasberény is the most characteristic, with aplanar top, very straight shape, and remarkable size: it is ~1 km long,100–140 m wide at the base, and 20 m high. The loess on the Triassicdolomite at Csákberény (Orondpuszta dolomite quarry; sample ‘Csb-2’) turned out to be 14±2 ka old, while the loess at Lovasberény(sample ‘Lb’) was deposited about 30 ka ago. These ages are in goodagreement with earlier results of luminescence dating of loess (17–15±2 ka) and radiocarbon ages of charcoal in the loess (30–15 ka) atLovasberény (Pécsi, 1982; Mészáros, 2007). Loess sequences andloess-like sediments with similar 14–30 ka old ages were also dated atnumerous sites in the Pannonian Basin including both the plains andthe hilly regions (e.g. Wintle and Packman, 1988; Frechen et al., 1997;Novothny et al., 2002, 2009; Thamó-Bozsó et al., 2010b).

Alluvial fan sediments at the Csákberény gravel pit (sample ‘Csb-1’) were dated because this fan prograded into a former ephemerallake, the basin of which had been formed by deflation (Budai et al.,2008; Csillag et al., 2010). These sediments were deposited 42–40(±4) ka ago (Thamó-Bozsó et al., 2010a). This shows that there was a

strong deflation period in the area more than 42±4 ka ago, duringwhich the lake basin was excavated.

These OSL data indicate strong eolian activity during the cold anddry climates of the last glacial and also during the early Holocene. Thelast phase of yardang formationmust have post-dated the 15 ka age ofyoung loess at Lovasberény (Budai et al., 2008). Luminescence agedata of the key loess sections in Hungary (Paks, Mende, Basaharc)indicate some discontinuities. Themain time gaps are between 55 and35 ka and from 25 to 20 ka (Frechen et al., 1997). These discontinu-ities were caused either by the absence of loess accumulation and/orerosion or deflation of the loess.

Based on TL, IRSL, OSL, and radiocarbon age data of eolian sands inHungary (Borsy, 1991; Ujházy et al., 2003; Nyári et al., 2007; Novothnyet al., 2010; Thamó-Bozsó et al., 2010a) and other data (e.g. Gábris,2003; Gábris and Nádor, 2007) from the Late Glacial and Holocene, wecan infer climatically forced eolian sand movement between 27 and7.2 ka and possibly between 6.5 and 5.6 ka in the dry periods. Sandmovement also occurred repeatedly during historical times from theSubboreal up to the 19th century (e.g. Horusitzky, 1901; Gábris, 2003;Nyári et al., 2007; Novothny et al., 2010), but this was due to humanimpact and is, thus, not related to yardang formation.

3.7.3. Timing of yardang formationYardang formation rates in weakly consolidated sediments can be

considerably high and are usually around a few m/ka. Some examplesare: N4 m/ka in sand (Mojave Desert; Clarke et al., 1996); N2 m/ka inlacustrine and swamp deposits (Kharga, Egypt; Goudie, 2007); 1.6–4 m/ka in diatomites (Bodélé, Chad; Washington et al., 2006); ~5 m/ka in lacustrine silts (Lop Nor; Halimov and Fezer, 1989); and 20 m/kain lacustrine fine sediments (Lop Nor; McCauley et al., 1977b). Thevalues suppose constant denudation rates and, thus, are obviouslyminimum rates in most cases. Counting with a characteristic relativeheight of 150 m for large yardangs in the Pannonian Basin and evenwith a high denudation rate of 10 m/ka, which is readily acceptablefor the easily eroded well-sorted sands that dominate the region,gives a formation time of 15,000 years. Together with the highexposure ages, this implies that the landforms were sculpted in anintermittent way, in millennial-scale intervals as longest. The fact thatthe yardangs represent the initial phases of the landscape evolutionmodel both of Brookes (2001) and of Halimov and Fezer (1989) alsorefers to short, unsustained deflation periods. Windy periods like thishave been reported from loess at several sites in the Pannonian Basin(e.g. Frechen et al., 1997; Marković et al., 2008) and from pelagicsediments and maar lake deposits (Rea, 1994; Dietrich and Seelos,2010). Though deflation started in the Early Pleistocene, it wasprobably most intense during the Late Pleistocene, as shown byslightly increased regional denudation rates (Ruszkiczay-Rüdiger etal., in press), by climatic considerations and by the young age of largedeflation hollows.

4. Climate and surface cover during Pleistocene glacials in thePannonian Basin

During the Pleistocene glaciations the Pannonian Basin was ice-free, with a cold and dry periglacial climate indicated by widespreadloess sequences and the contained paleontological record. Mean

478 K. Sebe et al. / Geomorphology 134 (2011) 470–482

temperatures were characterized by a clearly detectable north–southgradient, with a growing abundance of thermophilous taxa towardsthe south (Sümegi and Krolopp, 2002) andmean annual temperaturesas low as −3 °C (Kordos and Ringer, 1991). In accord, permafrostconditions varied as well, from discontinuous permafrost in thenorthern part of the basin to a lack of cryogenic features in the south(Pécsi, 1961; van Vliet-Lanoë et al., 2004; Marković et al., 2008).Reconstructed precipitations were 300–500 mm/year with a pro-nounced winter drying (van Huissteden and Pollard, 2003), or “250–500 mm lower than today” [i.e. 550–850 mm] (Járainé Komlódi,2003).

Snow cover must have been thin or absent (possibly blown away),shown by permafrost phenomena requiring uncovered ground fortheir formation (e.g. sand wedges) (Pécsi, 1961; Kovács et al., 2007)and by the widespread presence of small ventifacts hardly emergingabove the ground.

The dominant vegetation during the glaciations was cold steppe orsteppe tundra, with a varying amount—usually patches—of arborealvegetation, shown by pollen, plant macrofossil (charcoal), molluscand mammal data (e.g. Willis et al., 2000; Sümegi and Krolopp, 2002;Járainé Komlódi, 2003; Gábris and Nádor, 2007). The parabolic dunefields (Nyírség, Belső-Somogy; Figs. 2 and 6) with their uniformlyoriented dunes are indicative of open tundra, in contrast withparabolic dunes in the boreal forest zone, which are frequentlymultilobate and have varied orientations due to polymodal windsinfluenced by the arboreal vegetation (Filion and Morisset, 1983).

5. Discussion: controls on yardang formation

5.1. Wind regime

5.1.1. Wind directionsFor the wider periglacial belt of Europe, climate reconstructions

indicated almost exclusivelywesterlywinds for various cold intervals ofthe Late Pleistocene (Poser, 1950; Huijzer and Isarin, 1997; Isarin et al.,1997; van Huissteden and Pollard, 2003), though northerly winds havealso been proposed based on the downwind sorting of eolian sediments(Huijzer and Isarin, 1997). Fromgeochemical investigationsBuggle et al.(2008) suggested northerly/northwesterly katabatic winds for theperiglacial area east of the Carpathians. For the Pannonian Basin, thereconstruction of Poser (1950) based on geomorphic evidence and theLGM calculation of van Huissteden and Pollard (2003) roughly agreewith our landform-indicatedwind directions, while other models showwesterly, even southwesterly circulation for the area (Isarin et al., 1997;the MIS3 cold model of van Huissteden and Pollard (2003)).

The surface wind system of the Pannonian Basin indicated bylandforms shows remarkable stability over a time-scale of at least tensof thousands of years. Air flow directions indicated by Pleistocene–early Holocene landforms—including wind striations on all datedsurfaces—all agree and are in accordance with modern wind di-rections (Mersich et al., 2003). Prevailing modern wind directions areremarkably constant throughout the year as well (Mersich et al.,2003), in contrast with other parts of Central and Western Europe(Poser, 1950).

This stability can be explained by the topographic control of thesurrounding mountains on low-level winds: where local air flow isstrongly controlled by the topography, wind directions tend to bestable even if the regional circulation changes (e.g. Sharp, 1949;Griffiths et al., 2009). The Carpathians and Alps allowed winds toenter the basin through two low segments; modern zonal westerliesstill enter the basin through the aforementioned two “gaps” (Mersichet al., 2003). Cold Pleistocene winds moved closer to the ground,consequently, they were more prone to be influenced by thetopography. The deflecting effect of the Carpathians was demonstrat-ed outside the mountains by the orientation of loess ridges (gredas)(Rozycki, 1967). The topographic arrangement is very similar in a

yardang field in the Qaidam Basin in Central Asia (roughly between38°30′N 92°7′E and 39°N 92°45′E), where yardangs and corridorsform two fan-shaped systems downwind of a larger and a smaller passin the mountains surrounding the basin. The control of regionaltopography on winds is evident in many other yardang fields as wellwhich do not lie in a basin position; these include some of the mostspectacular fields like the one between the Tibesti and Ennedi Mts. inAfrica (McCauley et al., 1977a; Washington et al., 2006; Laity, 2009).

5.1.2. Wind strengthModern wind speeds in the study area are moderate: mean annual

wind speeds are around 3 m/s, while winds above 6 m/s, which couldbe able to entrain unvegetated sand, occur with a frequency of 7–8%.Winds above 15 m/s occur as unsustained gusts (Mersich et al., 2003).Sufficient wind speeds in the Pleistocene Pannonian Basin could begenerated by an interplay of climatic and topographic effects.

Pleistocene glaciations were characterized by stronger circulationand higher cyclonic activity (Isarin et al., 1997; Christiansen, 2004)and, thus, geomorphically more effective winds than today, even inareas directly not affected by the ice sheet (Vincent and Kattan, 2006;Washington et al., 2006; Enzel et al., 2010). In the periglacial belt, theincreased pressure and temperature gradients created by the ice massfurther enhanced air circulation (Isarin et al., 1997; van Huisstedenand Pollard, 2003; Seppälä, 2004).

Topographic factors can increase the frequency, the speed andpersistence of winds through several factors (Liu et al., 2000;Washington et al., 2006; Trachte et al., 2010): 1) dry and cold airmasses can be produced in elevated areas, 2) they flow down overrelatively steep slopes (downslope flow), and 3) become channeledby gaps in barrier ranges or by narrow valleys bounded by mountains(gap flow). Flow lines diverge at the basinward exit of the gap orvalley, and wind strength diminishes away from the pass. In the caseof the Pannonian Basin, the Scandinavian ice sheet acted as a producerof cold air masses, but also as elevated relief: model simulations forthe ice sheet indicated ice thicknesses of up to 2700 m for the LGMand values approaching 2000 m also for 60 ka and 90 ka ago(Svendsen et al., 2004), with the reconstructed geometry of the icesheet characterized by steep margins comparable to the modernAntarctic ice. The fanning of flow lines after the gaps is shown bylandform orientations and by the boundaries of the distribution areasof eolian features, while the downwind decrease in air speed by thesouth(east)ward growing importance of loess and by the fact thatmost advanced yardang forms and the majority of ventifacts arelocated in the north.

Some or all of these topographic factors are present at modernyardang fields which lie in a basin location. The largest yardang field ofthe Earth in the Qaidam Basin is surrounded by high mountains andreceives an extremely dry, foehn-type wind through the passes of the4000–5000 m high Altyn Tag from the Tarim Basin (Halimov andFezer, 1989). The foehn winds are strong enough to interrupt cyclonicactivity in the summer. The Lut Desert is also surrounded bymountains of more than 3000 m relative height and is sculpted byvery strongwinter winds (McCauley et al., 1977a; Liu et al., 2000). Thetopographic situation is similar in the case of the yardangs of the RioIca Valley in Peru, where local winds are superposed on the tradewinds channeled by a tectonic trough (McCauley et al., 1977a), andaround Lop Nur (Laity, 2009).

5.2. Aridity of the climate

Yardangs have been reported from regions with precipitation wellabove the 50 mm/year limit proposed for yardang formation. Thehighest one was 400 mm/yr for the Ebro Basin, Spain, where smallyardangs are actively forming (Gutiérrez-Elorza et al., 2002). Thesesuggest that hyper-aridity may not be indispensable for yardangformation. The case of the Pannonian Basin supports this idea: no

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records of hyper-aridity during the Quaternary have been found sofar. However, in cold settings aridity controls wind erosion in a morecomplex way than in warm climates, and the various interlinkedmechanisms can reinforce each other in enhancing wind erosion.

Model calculations for the northern periglacial belt showed thatdeflationwas not controlled bywind speeds but by surface conditions,primarily plant cover (Sima et al., 2009). Vegetation growth isrestricted by cold and dry winds, by a reduced or lacking snow coverand by the limited availability of water due to frost. The sparse plantcover of the Pannonian Basin has also been explained by the stronglydecreased precipitation (Járainé Komlódi, 2003). If once broken up,the plant cover recovers very slowly and a few periods with strongwinds can keep the ground open and deflation active for decades(Seppälä, 2004). This is in accord with the calculations of Sima et al.(2009), which showed that in the northern periglacial belt only a fewgusty days per year were responsible for the geomorphological effect(there dust deflation) of eolian activity. If snow cover is missing,interstitial ice can leave by sublimation and deflation can be effective.In the modern subarctic zone, deflation is extremely active close toand just outside of the northern boreal forest limit, which coincideswith the zones of sporadic and discontinuous permafrost (Seppälä,2004). The above characteristics weremost probably all present in thePannonian Basin during the Pleistocene cold stages. In addition, afterthe formation of incipient corridors, winds channeled in them mayhave locally inhibited the regeneration of plant cover even in periodswithout extremely strong winds.

5.3. Homogeneous material

The hypothesis that mega-yardangs tend to occur in relativelyhomogeneous rocks without complex structures (Goudie, 2007) issupported by the Pannonian Basin. The Upper Miocene sedimentsmaking up the majority of the basin surface and being the material ofmega-yardangs show strong lateral and also considerable verticalhomogeneity, with a thickness reaching several hundred meters.Numerous meso-yardangs were formed on Oligocene fluvial andlittoral sediments and on loess as well. In addition to their spatialuniformity, the Upper Miocene lacustrine sediments are characteris-tically well-sorted and dominantly sandy, which makes themespecially sensitive to deflation (Seppälä, 2004). Rock types differentfrom the dominant sand–silt lithology fail to host yardangs. Due to thelarge extent of sand-dominated rocks, abrasive material was easilyavailable everywhere.

Within the large area of relatively homogeneous lithology, thelocation of wind erosion features is not uniform. Instead, it wascontrolled by vertical tectonic movements: yardangs and corridorsformedwhere themoderately consolidated sediments were located inan elevated position. In many places in the world, the final depth ofcorridors and the shape of their bottoms are controlled by underlyingmore resistant strata (McCauley et al., 1977a). In the Pannonian Basin,uplift rate—in itself or indirectly, by affecting groundwater levels—seems to be responsible for valley depth, since the corridors areentrenched only into the upper portion of the thick Upper Miocenesequence. Neotectonic vertical movements influenced the preserva-tion of these sediments and yardang formation by creating hard-rocktopographic barriers perpendicular to the wind. In the lee of theuplifted blocks of the Transdanubian Range, long yardangs—typicallytype 2, i.e. less streamlined ridges—or at least well-developed yardangtroughs stretch downwind, similarly to the system described byMainguet (1972). In contrast, fault-controlled wind channels and thearea in their lee are heavily deflated and are followed by wind-blownsand deposits downwind (Figs. 3 and 6). The westernmost part of themega-yardang field is located downwind of thick Pleistocene gravelsheets, which sheltered the underlying sediments andmay have actedas a low, nevertheless continuous obstacle to wind erosion.

5.4. Joints, faults

Laity (2009) notes that the role of jointing and fluvial erosion inthe initial channelization of wind is poorly understood. OurPannonian case study contributes to this general question andseems to give partial answers.

The authors who interpreted the linear valleys in the PannonianBasin as wind channels have all considered their orientation andlocation to be fault-controlled (Cholnoky, 1929; Jámbor, 1992, 2002);however, the role of faults controlling yardang orientation could bedisproved (see Section 3.1.1). Our studies, in accord with other recentpapers (e.g. de Silva et al., 2010), demonstrated that jointing does notbasically control yardang orientation. In addition to opposingobservations listed in Section 3.1.1, jointing can be ruled out becausethe areal extent of mega-yardang fields is usually too large to becharacterized by a uniform stress field and uniform fractureorientation. Ridge-and-swale topography tends to have a relativelyconstant wavelength in cross-section for any given group of yardangs(Mainguet, 1970), and this is valid for the Pannonian Basin as well.Although fault arrays may consist of more or less evenly spaced,parallel faults at a small scale, this pattern is highly improbable acrossareas tens or rather hundreds of km in extent. The same is true for thestream network: regular, parallel drainage patterns are usually localphenomena and the topography of thousands of km2 cannot originatefrom them.

There is a general belief that faults or joints (fractures in general)have the ability to facilitate or even to initiate eolian incision. Although itmight be true for cemented rocks, this role is questionable in poorlyconsolidated sediments. Preliminary observations in the PannonianBasin are opposing: most fractures develop as deformation bands intheir early stage; these few mm- to few cm-thick bands are somewhatmore consolidated (“cemented”) andwater-saturated thansurroundingundeformed sediments and thus probably more resistant to erosion.

6. Persistency of mega-yardangs

The mega-yardangs of the Pannonian Basin have been stronglymodified by various erosional processes: drainage networks incisedinto them and the steep yardang slopes have been reshaped bygullying, slope wash and mass movements. The corridors haveexperienced accumulation of fluvial or slope sediments. This strongmodification under the moister interglacial/-stadial climates is due tothe only weakly resistant material, large size and relatively old age ofthe ridges. In spite of all these, the yardangs retained their generalform and are readily identified. At present, themega-yardang field liesunder a moderate continental climate, with a mean annual precip-itation of 550–850 mm decreasing from W to E.

Yardangs are considered to be unable to persist long in arecognizable form under less arid climates (McCauley et al., 1977a;Laity, 2009), though the latter author mentions the Ebro Basin fossilyardangs (Gutiérrez-Elorza et al., 2002) as an exception. The EbroBasin now receives a precipitation of 400 mm/year, but data as towhen yardang formation finished are lacking.

Thoughatpresentwedonotpossessexact data on the timingofwinderosion in the Pannonian Basin, dating results and geomorphologicalconsiderations (Section 3.7) suggest that the formation of the yardangfield might have started in the Early Pleistocene, but their sculpturinghappened mostly during the Middle to Late Pleistocene and up to theEarlyHolocene. Thismeans that the features haveexisted for at least tensof thousands of years, but possibly even for an estimated 1–1.5 Ma.Yardangs were actively sculptured during relatively short, repeatedintervals separated by long periods—supposedly up to tens of thousandsof years—of quiescence, and they were able to “survive” these periodsincluding the largest part of the Holocene.

There are numerous records illustrating that eolian landforms areable to survive changes in the wind regime and also in other climatic

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elements, most importantly precipitation (Thorson and Schile, 1995;Gutiérrez-Elorza et al., 2002; Parsons and Abrahams, 2009; Pye andTsoar, 2009). If Pleistocene dune fields are preserved—even if vegetated—up to today, then it seems reasonable to suppose that yardangs, thematerial of which is at least weakly consolidated, are able to retain theirshape through at least several millennia. In the case of mega-yardangs,preservation is facilitated by the large dimensions as well.

7. Conclusions

The peculiar fan-shaped ridge-and-swale system of the westernPannonian Basin is the northernmost reported mega-yardang field andtheonly one reported to have formedunder a cold—periglacial—climate.The system of rectilinear valleys and ridges was originally formed ascorridors and yardangs, and they experienced only minor subsequentmodifications by fluvial and mass wasting processes. Both faults andfluvial erosion can be ruled out as initial formative factors. Corridors andmega-yardangs, together with meso-yardangs, ventifacts, deflationhollows and wind-blown sand areas, form an integrated eolianlandform system driven by northerly to northwesterly winds. Winderosionwasmost intense near the passes in the surroundingmountains,especially near the wide gap between the Alps and the Carpathians;mega-yardangs were located in the erosional/transportation zone,while wind strength diminished and accumulation increased towardsthe S–SE. The yardang field was sculptured dominantly from wide-spread Upper Miocene lacustrine sediments in uplifting areas, whileeolian sand deflated from the wind channels was deposited downwindof them, mostly in areas of subsidence.

The system is at present inactive; its last activity can be dated tothe Early Holocene. Dating of ventifact surfaces with in situ producedcosmogenic 10Be places the onset of significant wind erosion to theEarly Pleistocene, while the fact that well-developed yardangs werealso sculpted from the youngest loess (dated by OSL) shows that thelast, Dryas cold stage and possibly the warm and dry Holocene Borealhad important effects on the evolution of the eolian features. The longtemporal range of wind erosion compared to average yardangformation rates and the rarity of advanced forms among the winderosional features indicate intermittent and relatively short, mostprobably millennial-scale deflation periods, which compare well withperiods of strong winds recorded in loess sequences and with theabrupt North Atlantic climate changes.

The case of the Pannonian Basin has important implications withregard to the most critical factors in yardang development in general.

1. The main control on yardang formation in the Pannonian Basin andprobably in all yardang fields is the existence of strong unidirec-tional winds: they determine the location (spacing) and orienta-tion of wind-carved features. In the Pannonian Basin, stable winddirections were maintained on a time-scale of at least 1.5 Ma bythe channeling effect of the topographic lows (passes) in thesurrounding mountain chains; this situation seems to be commonin basin-located yardang fields. Sufficient wind speedwas providedby the enhanced atmospheric circulation of the Pleistocene glacialscomplemented by topographical effects, namely channeling anddownslope flow.

2. Yardangs can form not only in hyper-arid but even in semi-aridclimates, as shown by the case of the Pannonian Basin and of someother sites (e.g. Spain, Andes). However, the rarity of occurrencessuggests that this requires a special association of environmentalconditions. In the study area, special conditions were related tothe periglacial climate: a location in the most deflation-sensitiveperiglacial zone (boreal forest–tundra boundary, with scatteredpermafrost), zonal precipitation decrease during the glacial phases,and strong ice sheet-related winds. After the formation of initialyardang corridors, vegetation-free rock surfaces could have locally

been maintained in them by channeled winds even during steppe-dominated periods.

3. As in other regions, homogeneous lithology—here erosion-pronesandy–silty deposits—facilitated yardang formation. Yardangsdeveloped where these sediments were in a tectonically upliftedposition, and often downwind of topographic barriers perpendic-ular to the air flow.

4. Jointing, together with the fluvial network, can be ruled out as aprerequisite of mega-yardang formation and as a regional controlon yardang orientation, but their role is probably similar in the caseof meso-scale features as well. This can be explained partly by thesize of the yardang fields and partly by the regularity of the formsless characteristic of fracture-dominated areas or of drainagepatterns.

Acknowledgements

Our research was supported by the Hungarian Scientific ResearchFund (OTKA) projects NK60455, K062478 and F042799, by the BolyaiJános Scholarship of the Hungarian Academy of Sciences, by the MarieCurie Fellowship, by the Research Scholarship of the French Embassyof Hungary, the French-Hungarian Balaton-Tét Project (FR-32/2007)and by the EGT/Norwegian FinancingMechanism and ZoltánMagyaryPublic Foundation of Higher Education, Hungary. The AMS measure-ments performed at the ASTER AMS national facility (CEREGE, Aix enProvence) are supported by the INSU/CNRS, the French Ministry ofResearch and Higher Education, IRD and CEA. The authors thank twoanonymous reviewers and Andrew Plater for their helpful commentsand corrections on the manuscript.

References

Arnold, L.J., Bailey, R.M., Tucker, G.E., 2007. Statistical treatment of fluvial dosedistributions from southern Colorado arroyo deposits. Quaternary Geochronology2, 162–167.

Auer, A., Martin, U., Németh, K., 2007. The Fekete-hegy (Balaton Highland Hungary)“soft-substrate” and “hard-substrate” maar volcanoes in an aligned volcaniccomplex—implications for vent geometry, subsurface stratigraphy and thepaleoenvironmental setting. Journal of Volcanology and Geothermal Research159, 225–245.

Bada, G., Szafián, P., Vincze, O., Tóth, T., Fodor, L., Spiess, V., Horváth, F., 2010. Theneotectonic habitat of the eastern part of Lake Balaton and its broader environs:inferences from high resolution seismic profiling. Bulletin of the HungarianGeological Society 140, 367–390 (in Hungarian, with English abstract).

Bailey, R.M., Arnold, L.J., 2006. Statistical modelling of single grain quartz Dedistributions and an assessment of procedures for estimating burial dose.Quaternary Science Reviews 25, 2475–2502.

Bailey, J.E., Self, S., Wooller, L.K., Mouginis-Mark, P.J., 2007. Discrimination of fluvial andeolian features on large ignimbrite sheets around La Pacana Caldera, Chile, usingLandsat and SRTM-derived DEM. Remote Sensing of Environment 108, 24–41.

Borsy, Z., 1977. Evolution of relief forms in Hungarian wind-blown sand areas. FöldrajziKözlemények 25, 3–12.

Borsy, Z., 1991. Blown sand territories in Hungary. Zeitschrift für Geomorphologie N.F.Supplement-Band 90, 1–14.

Braucher, R., Del Castillo, P., Siame, L., Hidy, A.J., Bourles, D.L., 2009. Determination ofboth exposure time and denudation rate from an in situ-produced 10Be depthprofile: a mathematical proof of uniqueness. Model sensitivity and applications tonatural cases. Quaternary Geochronology 4, 56–64.

Brookes, I.A., 2001. Aeolian erosional lineations in the Libyan Desert, Dakhla Region,Egypt. Geomorphology 39, 189–209.

Budai T., Császár G., Csillag G. Dudko A., Koloszár L., Majoros Gy., 1999. Geology of theBalaton Highland. Explanation to the Geological Map of the Balaton Highland, 1:50000. Occasional Papers of the Geological Institute of Hungary 197, Budapest, 257pp.

Budai T., Fodor L., Császár G., Csillag G., Gál N., Kercsmár Zs., Kordos L., Pálfalvi S.,Selmeczi I., 2008. Geology of the Vértes Hills. Explanatory Book to the GeologicalMap of the Vértes Hills (1:50 000). Geological Institute of Hungary, Budapest,368 pp.

Buggle, B., Glaser, B., Zöller, L., Hambach, U., Marković, S., Glaser, I., Gerasimenko, N.,2008. Geochemical characterization and origin of Southeastern and EasternEuropean loesses (Serbia, Romania, Ukraine). Quaternary Science Reviews 27,1058–1075.

Cholnoky, J., 1929. Magyarország földrajza. A Föld és élete 6. Franklin Társulat,Budapest. 530pp. (in Hungarian).

Cholnoky, J., 1940. A futóhomok elterjedése. Bulletin of the Hungarian GeologicalSociety 70, 258–294 (in Hungarian).

481K. Sebe et al. / Geomorphology 134 (2011) 470–482

Christiansen, H.H., 2004.Windpolished boulders and bedrock in the Scottish Highlands:evidence and implications of Late Devensian wind activity. Boreas 33, 82–94.

Clarke, M.L., Wintle, A.G., Lancaster, N., 1996. Infra-red stimulated luminescence datingof sands from the Cronese basins, Mojave Desert. Geomorphology 17, 199–205.

Csillag, G., Fodor, L., Sebe, K., Müller, P., Ruszkiczay-Rüdiger, Zs., Thamóné Bozsó, E.,Bada, G., 2010. The role of wind erosion in the surface development ofTransdanubia during the Quarternary. Bulletin of the Hungarian Geological Society140, 445–468 (in Hungarian, with English abstract).

de Silva, S.L., Bailey, J.E., Mandt, K.E., Viramonte, J.M., 2010. Yardangs in terrestrialignimbrites: synergistic remote and field observations on Earth with applicationsto Mars. Planetary and Space Science 58, 459–471.

Dietrich, S., Seelos, K., 2010. The reconstruction of easterlywind directions for the Eifel region(Central Europe) during the period 40.3–12.9 ka BP. Climate of the Past 6, 145–154.

Enzel, Y., Amit, R., Crouvi, O., Porat, N., 2010. Abrasion-derived sediments underintensified winds at the latest Pleistocene leading edge of the advancing Sinai–Negev erg. Quaternary Research 74, 121–131.

Erdélyi, M., 1961. Külső-Somogy vízföldtana 1. Hidrológiai Közlöny 41, 445–458 (inHungarian).

Erdélyi,M., 1962. Külső-Somogy vízföldtana 2. Hidrológiai Közlöny 42, 56–65 (inHungarian).Filion, L., Morisset, P., 1983. Eolian landforms along the eastern coast of Hudson Bay,

Northern Québec. Collection Nordicana 47, 73–94.Fodor, L., Csontos, L., Bada, G., Györfi, I., Benkovics, L., 1999. Tertiary tectonic evolution

of the Pannonian basin system and neighbouring orogens: a new synthesis ofpaleostress data. In: Durand, B., Jolivet, L., Horváth, F., Séranne, M. (Eds.), TheMediterranean Basins: Tertiary Extension Within the Alpine Orogen: GeologicalSociety, London, Special Publications, 156, pp. 295–334.

Fodor, L., Bada, G., Csillag, G., Horváth, E., Ruszkiczay-Rüdiger, Zs, Palotás, K., Síkhegyi,F., Timár, G., Cloetingh, S., Horváth, F., 2005. An outline of neotectonic structuresand morphotectonics of the western and central Pannonian Basin. Tectonophysics410, 15–41.

Fodor L., Csillag G., Lantos Z., Budai T., Selmeczi I., 2008. Geological map of the VértesHills, 1:50000. Geological Intitute of Hungary, Budapest.

Frechen, M., Horváth, E., Gábris, Gy, 1997. Geochronology of Middle and UpperPleistocene loess sections in Hungary. Quaternary Research 48, 291–312.

Gábris, Gy, 2003. The periods of the history of the Earth for the last 30 thousand yearsand the periods of the movement of aeolian sand. Földrajzi Közlemények 127, 1–14(in Hungarian with English abstract).

Gábris, Gy., Mari, L., 2007. The Pleistocene beheading of the Zala River (West Hungary).Földrajzi Értesítő 56, 39–50.

Gábris, Gy., Nádor, A., 2007. Long-term fluvial archives in Hungary: response of theDanube and Tisza rivers to tectonic movements and climatic changes during theQuaternary: a review and new synthesis. Quaternary Science Reviews 26,2758–2782.

Galbraith, R.F., Roberts, R.G., Laslett, G.M., Yoshida, H., Olley, J.M., 1999. Optical dating ofsingle and multiple grains of quartz from Jinmium rock shelter, northern Australia:Part I. Experimental design and statistical models. Archaeometry 41, 339–364.

Gerner, P., 1992. Recent stress field in Transdanubia (Western Hungary). Bulletin of theHungarian Geological Society 122, 89–105.

Goudie, A.S., 2007. Mega-yardangs: a global analysis. Geography Compass 1, 65–81.Griffiths, P.G., Webb, R.H., Fisher, M., Muth, A., 2009. Plants and ventifacts delineate late

Holocene wind vectors in the Coachella Valley, USA. Aeolian Research 1, 63–73.Gutiérrez-Elorza, M., Desir, G., Gutiérrez-Santolalla, F., 2002. Yardangs in the semiarid

central sector of the Ebro Depression (NE Spain). Geomorphology 44, 155–170.Halimov, V.M., Fezer, F., 1989. Eight yardang types in Central Asia. Zeitschrift für

Geomorphologie N.F. 33, 205–217.Hein, A.S., Hulton, N.R.J., Dunai, T.J., Schnabel, C., Kaplan, M.R., Naylor, M., Xu, S., 2009.

Middle Pleistocene glaciation in Patagonia dated by cosmogenic-nuclide measure-ments on outwash gravels. Earth and Planetary Science Letters 286, 184–197.

Horusitzky, H., 1901. A bábolnai állami ménesbirtok agrogeológiai viszonyai. A MagyarKirályi Földtani Intézet Évkönyve 13, 170–202 (in Hungarian).

Huijzer, A.S., Isarin, R.F.B., 1997. The reconstruction of past climates using multi-proxyevidence: an example of the Weichselian Pleniglacial in Northwest and CentralEurope. Quaternary Science Reviews 16, 513–533.

Inbar, M., Risso, C., 2001. Holocene yardangs in volcanic terrains in the Southern Andes,Argentina. Earth Surface Processes and Landforms 26, 657–666.

Isarin, R.F.B., Renssen, H., Koster, E.A., 1997. Surface wind climate during the YoungerDryas in Europe as inferred from aeolian records and model simulations.Palaeogeography, Palaeoclimatology, Palaeoecology 134, 127–148.

Jámbor, Á., 1992. Pleistocene ventifact occurrences in Hungary. Acta GeologicaHungarica 35, 407–436.

Jámbor, Á., 2002. Occurrences and importance of the Pleistocene ventifacts in Hungary.Bulletin of the Hungarian Geological Society 132, 101–116.

Járainé Komlódi, M., 2003. A Kárpát-medence növényzetének kialakulása. In: Láng, I.,Bedő, Z., és Csete, L. (Eds.), Növény, állat, élőhely. : Magyar Tudománytár, Vol. 3.Hungarian Academy of Sciences Centre for Social Studies—Kossuth Press, Budapest,pp. 39–65 (in Hungarian).

Jones, L.S., Blakey, R.C., 1993. Erosional remnants and adjacent unconformities along aneolian–marine boundary of the Page Sandstone and Carmel Formation, MiddleJurassic, South-Central Utah. Journal of Sedimentary Petrology 63, 852–859.

Knight, J., 2008. The environmental significance of ventifacts: a critical review. Earth-Science Reviews 86, 89–105.

Kordos, L., Ringer, Á., 1991. Climate- and archeostratigraphic correlation of the Hungarianupper Pleistocene Arvicolidae stratigraphy. Annual Report of the Geological Instituteof Hungary from 1989, pp. 523–534 (in Hungarian, with English abstract).

Kovács, J., Fábián, Sz.Á., Schweitzer, F., Varga, G., 2007. A relict sand-wedge polygon sitein north-central Hungary. Permafrost and Periglacial Processes 18, 379–384.

Laity, J.E., 2009. Landforms, landscapes, and processes of aeolian erosion. In: Parsons,A.J., Abrahams, A.D. (Eds.), Geomorphology of Desert Environments. Springer,pp. 597–627.

Liu, M., Westphal, D.L., Holt, T.R., Xu, Q., 2000. Numerical simulation of a low-level jetover complex terrain in Southern Iran. Monthly Weather Review 128, 1309–1327.

Lóczy, L.sen., 1913. A Balaton környékének geologiai képződményei és ezeknek vidékekszerinti telepedése. Balaton Committee of the Hungarian Geographical Society,Budapest. 617p p. (in Hungarian and German).

Magyari, Á., Musitz, B., Csontos, L., van Vliet-Lanoë, B., 2005. Quaternary neotectonics ofthe Somogy Hills, Hungary (part I): evidence from field observations. Tectono-physics 410, 43–62.

Mainguet, M., 1970. Un étonnant paysage: les cannelures gréseuses du Bembéché (N. duTchad). Essai d'explication géomorphologique. Annales de Géographie 79, 58–66.

Mainguet, M., 1972. Le Modelé Des Grès: Problèmes Généraux. Institut GéographiqueNational, Paris. 657 pp.

Marković, S.B., Bokhorst, M.P., Vandenberghe, J., McCoy, W.D., Oches, E.A., Hambach, U.,Gaudenyi, T., Jovanović, M., Zöller, L., Stevens, T., Machalett, B., 2008. LatePleistocene loess-palaeosol sequences in the Vojvodina region, north Serbia.Journal of Quaternary Science 23, 73–84.

Marosi, S., 1970. Belső-Somogy kialakulása és felszínalaktana. Akadémiai Kiadó,Budapest. 158 pp. (in Hungarian).

McCauley, J.F., Grolier, M.J., Breed, C.S., 1977a. Yardangs. In: Doehring, D.O. (Ed.),Geomorphology in arid regions: Proceedings of the Eighth Annual GeomorphologySymposium, Binghamton, New York, pp. 233–269.

McCauley, J., Grolier, M., Breed, C., 1977b. Yardangs of Peru and other desert regions.USGS Interagency Report. : Astrogeology, 81. USGS, Flagstaff, Arizona. 177 pp.

Mersich, I., Práger, T., Ambrózy, P., Hunkár, M., Dunkel, Z., 2003. Climate Atlas ofHungary. Hungarian Meteorological Service, Budapest. 107 pp.

Mészáros, J., 2007. Löszrétegtani és geomorfológiai vizsgálatok Lovasberény környékén.MSc thesis, Eötvös University, Budapest. (in Hungarian).

Miháltz, I., 1953. A Duna-Tisza köze déli részének földtani felvétele. Annual Report ofthe Geological Institute of Hungary from 1950, pp. 113–141 (in Hungarian, withFrench abstract).

Molnár, B., 1961. A Duna-Tisza közi eolikus rétegek felszíni és felszínalatti kiterjedése.Bulletin of the Hungarian Geological Society 91, 300–315 (in Hungarian, withGerman abstract).

Molnár, B., 1965. Adatok a Duna-Tisza köze fiatal harmadidőszaki és negyedkorirétegeinek tagolásához és származásához nehézásvány összetétel alapján. Bulletinof the Hungarian Geological Society 95, 217–225 (in Hungarian, with Germanabstract).

Novothny, A., Horváth, E., Frechen, M., 2002. The loess profile at Albertirsa, Hungary—improvements in loess stratigraphy by luminescence dating. Quaternary Interna-tional 95 (96), 155–163.

Novothny, Á., Frechen, M., Horváth, E., Bradák, B., Oches, E.A., McCoy, W.D., Stevens, T.,2009. Luminescence and amino acid racemization chronology of the loess–paleosolsequence at Süttő, Hungary. Quaternary International 198, 62–76.

Novothny, Á., Frechen, M., Horváth, E., Krbetschek, M., Tsukamoto, S., 2010. Infraredstimulated luminescence and radiofluorescence dating of aeolian sediments fromHungary. Quaternary Geochronology 5, 114–119.

Nyári, D., Kiss, T., Sipos, Gy., 2007. Investigation of Holocene blown-sand movementbased on archaeological findings and OSL dating, Danube–Tisza Interfluve,Hungary. Journal of Maps Student Edition 46–57.

Parsons, A.J., Abrahams, A.D. (Eds.), 2009. Geomorphology of Desert Environments.Springer. 831 pp.

Pécsi, M., 1961. Periglaciális talajfagyjelenségek főbb típusai Magyarországon. FöldrajziKözlemények 9, 1–24 (in Hungarian).

Pécsi, M., 1982. The most typical loess profiles in Hungary. In: Pécsi, M. (Ed.),Quaternary Studies in Hungary. : Elmélet–Módszer–Gyakorlat, 24. GeographicalResearch Institute, Budapest, pp. 145–169.

Poser, H., 1950. Zur Rekonstruktion der spätglazialen Luftdruckverhältnisse in Mittel-und Westeuropa auf Grund der vorzeitlichen Binnendünen. Erdkunde 4, 81–88.

Pye, K., Tsoar, H., 2009. Aeolian Sand and Sand Dunes. Springer, Berlin–Heidelberg.458 pp.

Rea, D.K., 1994. The paleoclimatic record provided by eolian deposition in the deep sea:the geologic history of wind. Reviews of Geophysics 32, 159–195.

Rónai, A., 1985. Az Alföld negyedidőszaki földtana (Quaternary geology of the GreatHungarian Plain). Geologica Hungarica Series Geologica 21 446 pp. (in Hungarianwith English abstract).

Rozycki, S.Z., 1967. Le sens des vents portent la poussiere de loess à la lumière del'analyse des formes d'accumulation du loess en Bulgarie et en Europe Centrale.Revue de Géomorphologie Dynamique 1, 1–9.

Ruszkiczay-Rüdiger, Zs, Fodor, L., Horváth, E., 2007. Neotectonic and landscapeevolution of the Gödöllő Hills, Central Pannonian Basin. Global and PlanetaryChange 58, 181–196.

Ruszkiczay-Rüdiger, Zs, Fodor, L.I., Horváth, E., Telbisz, T., 2009. Discrimination offluvial, eolian and neotectonic features in a low hilly landscape: a DEM-basedmorphotectonic analysis in the Central Pannonian Basin, Hungary. Geomorphology104, 203–217.

Ruszkiczay-Rüdiger, Zs., Braucher, R., Csillag, G., Fodor, L., Dunai, T.J., Bada, G., Bourlés,D., Müller, P., (in press). Dating Pleistocene aeolian landforms in Hungary, CentralEurope, using in situ produced cosmogenic 10Be. Quaternary Geochronology, DOI:10.1016/j.quageo.2011.06.001.

Seppälä, M., 2004. Wind as a Geomorphic Agent in Cold Climates. Cambridge UniversityPress, Cambridge. 358 pp.

Sharp, R.P., 1949. Pleistocene ventifacts east of the Big Horn Mountains, Wyoming.Journal of Geology 57, 175–195.

482 K. Sebe et al. / Geomorphology 134 (2011) 470–482

Siame, L., Bellier, O., Braucher, R., Sébrier, M., Cushing, M., Bourlès, D.L., Hamelin, B.,Baroux, E., de Voogd, B., Raisbeck, G., Yiou, F., 2004. Local erosion rates versus activetectonics: cosmic ray exposure modelling in Provence (South–East France). Earthand Planetary Science Letters 220, 345–364.

Sima, A., Rousseau, D.-D., Kageyama, M., Ramstein, G., Schulz, M., Balkanski, Y., Antoine,P., Dulac, F., Hatté, C., 2009. Imprint of North-Atlantic abrupt climate changes onwestern European loess deposits as viewed in a dust emission model. QuaternaryScience Reviews 28, 2851–2866.

Smith, B.J., Wright, J.S., Whalley, W.B., 1991. Simulated aeolian abrasion of Pannoniansands and its implications for the origins of Hungarian loess. Earth SurfaceProcesses and Landforms 16, 745–752.

Sümegi, P., Krolopp, E., 2002. Quatermalacological analyses for modelling of the UpperWeichselian paleoenvironmental changes in the Carpathian Basin. QuaternaryInternational 91, 53–63.

Svendsen, J.I., Alexanderson, H., Astakhov, V.I., Demidov, I., Dowdeswell, J.A., Funder, S.,Gataullin, V., Henriksen, M., Hjort, C., Houmark-Nielsen, M., Hubberten, H.W.,Ingólfsson, Ó., Jakobsson, M., Kjær, K.H., Larsen, E., Lokrantz, H., Lunkka, J.P., Lyså, A.,Mangerud, J., Matiouchkov, A., Murray, A., Möller, P., Niessen, F., Nikolskaya, O.,Polyak, L., Saarnisto, M., Siegert, C., Siegert, M.J., Spielhagen, R.F., Stein, R., 2004. LateQuaternary ice sheet history of northern Eurasia. Quaternary Science Reviews 23,1229–1271.

Szilárd, J., 1967. Külső-Somogy kialakulása és felszínalaktana. Akadémiai Kiadó,Budapest. 150 pp. (in Hungarian).

Tewes, D.W., Loope, D.B., 1992. Palaeo-yardangs: wind scoured desert landforms at thePermo-Triassic unconformity. Sedimentology 39, 251–261.

Thamó-Bozsó, E., Csillag, G., Fodor, L., Müller, P.M., Nagy, A., 2010a. OSL-dating theQuaternary landscape evolution in the Vértes Hills forelands (Hungary). Quater-nary Geochronology 5, 120–124.

Thamó-Bozsó, E., Magyari, Á., Musitz, B., Nagy, A., 2010b. OSL ages and origin of LateQuaternary sediments in the North Transdanubian Hills (Hungary): timing ofneotectonic activity. Quaternary International 222, 209–220.

Thorson, R.M., Schile, C.A., 1995. Deglacial eolian regimes in New England. GSA Bulletin107, 751–761.

Trachte, K., Nauss, T., Bendix, J., 2010. The impact of different terrain configurations onthe formation and dynamics of katabatic flows: idealised case studies. Boundary-Layer Meteorology 134, 307–325.

Ujházy, K., Gy, Gábris, Frechen, M., 2003. Ages of periods of sand movement in Hungarydetermined through luminescence measurements. Quaternary International 111,91–100.

van Huissteden, K., Pollard, D., 2003. Oxygen isotope stage 3 fluvial and eoliansuccessions in Europe compared with climate model results. Quaternary Research59, 223–233.

van Vliet-Lanoë, B., Magyari, Á., Meilliez, F., 2004. Distinguishing between tectonic andperiglacial deformations of quaternary continental deposits in Europe. Global andPlanetary Change 43, 103–127.

Vincent, P., Kattan, F., 2006. Yardangs on the Cambro-Ordovician Saq Sandstones,North–West Saudi Arabia. Zeitschrift für Geomorphologie 50, 305–320.

Walsh, J.J.,Watterson, J., 1992. Populationof faults and fault displacements and their effects onestimates of fault-related regional extension. Journal of Structural Geology 14, 701–712.

Washington, R., Todd, M.C., Lizcano, G., Tegen, I., 2006. Links between topography,wind, deflation, lakes and dust: the case of the Bodélé Depression, Chad.Geophysical Research Letters 33, L09401.

Willis, K.J., Rudner, E., Sümegi, P., 2000. The full-glacial forests of Central andSoutheastern Europe. Quaternary Research 53, 203–213.

Wintle, A.G., Packman, S.C., 1988. Thermoluminescence ages for three sections inHungary. Quaternary Science Reviews 7, 315–320.