Introduction

The term ‘yardang’ is the transliteration of Uyghur ‘yardin’, derived from the ablative form of ‘yar’, meaning ridge or steep bank. It was coined by Hedin (1903) in his book on Central Asia and Tibet, where he first named and identified this geomorphological structure. In a geographical sense, a yardang is a type of landform primarily formed by wind abrasion (Goudie 2007) and hydro-aeolian processes (Maghsoudi and Maghsoudi 2021a, b). These formations are predominantly found in desert environments (Dong et al. 2012) and are characterized by elongated ridges aligned longitudinally with the prevailing wind direction (Niu 2011). Consequently, yardangs often manifest as parallel systems (Mandt and Leone 2015), with the wing regimes influencing not only their orientation but also their length-to-width ratio (Hu et al. 2017). The formation of yardangs is contingent upon the presence of relatively dry, porous, and wind-erodible sediments, leading to their predominant distribution in arid and semiarid regions (Liang et al. 2019). Yardangs typically develop on soft sedimentary substrates such as alluvial and lacustrine deposits (Yang et al. 2021) and occasionally on volcanic tuffs (Wierzchos et al. 2013). With the exception of Antarctica, yardangs are found on all continents, including Eurasia (e.g., the Tarim Basin in China and the Lut Desert in Iran), the Americas (e.g., the Atacama Desert in Chile and Utah in the United States), Africa (e.g., south of Mogadishu in Somalia and along the coastline of Namibia), and Australia (e.g., along the coast of Lake Mungo) (Brook 2002; Brookfield 2011; Dobrea et al. 2020; McCauley et al. 2020). Beyond Earth, yardangs have been observed on celestial bodies within the solar system possessing a significant atmosphere, including Venus (Ganey et al. 2023), Mars (Ding et al. 2020), and Saturn's moon Titan (Paillou et al. 2015).

Yardangs serve as unique paleoenvironmental indicators (Heermance et al. 2013; Ding et al. 2020), offering valuable insights into past climatic and geological processes that have shaped them over time. This significantly enhances our comprehension of landscape evolution on both regional and global scales (Pötter et al. 2023). Beyond merely reflecting paleoclimatic conditions, vertical tectonic movements may also play a crucial role in yardang formation, as these geomorphological structures result from the intricate interplay between geological processes and climatic factors (Sebe et al. 2011). This complex interaction among various factors also influences the nature and structure of basin-filling deposits from which yardangs emerge (Sepehr and Almodaresi 2013). For instance, yardangs frequently manifest in endorheic basins subjected to fluctuating precipitation conditions over time, representing eroded remnants of former paleolakes (Shoaee et al. 2023). These observations hold significant implications, as fossilized yardang structures can preserve evidence of transient, sometimes only millennia-long paleoclimatic conditions (Sebe et al. 2011), information that might otherwise be scarce or nonexistent. Consequently, yardangs can serve as foundational elements in paleohydrological models, while the data provided by yardangs can also facilitate a better understanding of ancient biotic processes (Shoaee et al. 2023) and their ecological implications.

The Pannonian mega-yardang system, recognized as the northernmost mega-yardang field in Europe (Csillag et al. 2010; Sebe et al. 2011; Ding et al. 2020; Fig. 1a, b), is believed to have formed in response to a paleoclimatic event within the Carpathian Basin, which is currently not sufficiently explored. However, it’s worth noting that within the broader context of Eurasia, the mega-yardang system of the Ob loess plateau in the Russian Altai occupies a more northerly geographical position (Pötter et al. 2023). In Europe, only a few other yardang fields are known, such as those near Avignon, France (Bosq et al. 2018), and in the Ebro Depression, Spain (Gutiérrez-Elorza et al. 2002). However, these yardangs typically reach dimensions of only around 10 m in height and width. In contrast, the mega-yardangs of Western Hungary span several tens of meters in high and stretch across several tens of kilometers. Sebe et al. (2011) proposed that these yardangs formed during the Pleistocene glacial periods, possibly during the last glacial period, under cold and arid periglacial climatic conditions. Following the logic of this assumption, late Pleistocene (20.48, 17.42, and 15.62 kys) age fossilized sand dunes found in South Transdanubia (Györgyövics 2019; Transdanubia is the western part of Hungary) also support the similar age of the mega-yardang system. It is conjectured suggested that during the cold and arid glacial phases of the Pleistocene epoch, the sparse vegetation in the Carpathian Basin facilitated intense aeolian processes (Ruszkiczay-Rüdiger et al. 2011; Sebe et al. 2011, 2015). Climatic reconstructions lend support to this hypothesis; for instance, Kirschner et al. (2020) proposed that during the peak of the Last Glacial Maximum, the tundra-steppe biome extended across the plains of the Pannonian Basin.

Fig. 1
figure 1

(Source: Google Earth imagery). The region depicted in the upper part of this figure corresponds to a specific area (area No. 2) shown in Suppl. Figure 4

a Elevation and b shaded relief map of the Pannonian mega-yardang system and the deflation basins (deflation basins: 1: Basin of Lake Neusiedl (Fertő-tó), 2: Tapolca Basin, 3: Balaton Basin, 4: Basin of Lake Velence; mountains: 5: Keszthely Mts, 6: Balaton Highland, 7: Bakony Mts, 8: Vértes Mts, 9: Gerecse Mts, 10: Buda Mts, 11: Pilis and Visegrád Mts, 12: Börzsöny Mts, 13: Mecsek Mts, 14: Villány Mts; Hill lands: 15: Sokoró Hills, 16: Zala Hills, 17: Somogy Hills, 18: Tolna Hills, 19: Gödöllő Hills; Plains: Little Hungarian Plain, 21: Great Hungarian Plain (Duna-Tisza Interfluve section), 22: Dráva Plain). c Examples for large-scale geomorphological features created by dominantly aeolian processes: Csobánc butte in Gyulakeszi; (Tapolca Basin); d large deflation basin: the Balaton Basin from Halom Hill, Mencshely; e yardang: mega-yardang in Pannonhalma (Sokoró Hills) Cross sections of mega-yardangs: f Zala Hills (‘x’ line) and g Sokoró Hills (‘y’ line)

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However, alternative climatic reconstructions contradict the picture of the dry, steppe-like Carpathian Basin in the Last Glacial Maximum. For instance, the CCSM4 paleoclimate simulations suggest relatively high annual rainfall ranging from 500 to 600 mm in the region (Karger et al. 2021). This moderate annual precipitation, coupled with low annual mean temperatures, likely resulted in comparatively humid climatic conditions. Consequently, the impact of deflation processes may not have been as significant due to the presence of vegetation and permanently moist topsoil conditions. This perspective finds support in the Late Pleistocene gastropod fauna of the Pannonian Basin, which includes species, such as Pupilla alpicola Charpentier, 1837, throughout the region (Molnár et al. 2019). Presently, P. alpicola thrives in cool and moist macroclimates across Eurasia (Horsák et al. 2015) but avoids tundra-like environments. For example, relict populations of this species in the Western Carpathians, Slovakia, inhabit areas under a consistently humid boreal (Dfc) climate (Horsák et al. 2011). Notably, P. alpicola specimens were discovered in loess sections rather than paleosol sequences from the Last Interglacial Period. Additional evidence supporting the presence of ligneous vegetation in the Carpathian Basin during the Last Glacial period comes from the microwear analysis of fossil cervids conducted by Szabó et al. (2021). Their findings suggest that even during the glacial periods, tree cover persisted in the Carpathian Basin, likely extending beyond riverine habitats. This assertion may be particularly applicable to Southwest Transdanubia, recognized as the mildest and most humid territory within the broader Carpathian Basin.

Regional setting

Geological context of the Pannonian mega-yardang system

The Pannonian mega-yardang system is situated in the western part of Hungary (Fig. 1a), comprising an extensive hilly terrain with an average elevation ranging from 120 to 300 m above the sea level. This system consists of streamlined features and less aerodynamic crests, delineated by distinct corridors (Sebe et al. 2011). Mega-yardangs manifest as relatively straight or slightly curved formations scattered across the Transdanubian landscape. The largest of these formations extend beyond 40–60 km in length and 8–12 km in width, with mean heights averaging around 150–200 m. However, the backs of the larger mega-yardangs can occasionally rise slightly above the 300-m elevation mark. Alongside ventifacts, deflation basins, and wind-eroded sandy regions, the yardang expanse constitutes an integrated aeolian feature complex influenced by prevailing north-to-northwest winds. Many mega-yardangs exhibit distinct signs of post-formation hydrological erosion, evidence by side valleys often cutting into their flanks. Except for the alluvial plains and the mountainous regions, the yardang system extends over nearly the entire Transdanubia and even reaches into the northwestern part of the Great Plain. The structure of the mega-yardang system exhibits radial topography (Fig. 1b). The azimuth angle of mega-yardangs measures approximately 175–180° in the southwest and about 142–150° in the northwest section of the mega-yardang field. The focal point of the wind vectors lies northwest of Hungary (Suppl. Figure 1a), oriented toward the Devín Gate on the border of Slovakia and Austria. This gate not only marks the entrance of the Danube into the Carpathian Basin, but also serves as a passage for the westerly winds. The alignment of the yardangs corresponds to the prevailing north-northwesterly winds of the region (Mezősi et al. 2015).

However, due to the influence of the local relief patterns, deviations in wind direction occur. Consequently, the axis of the yardangs generally trends north-to-south in Southwest Transdanubia and a northwest-to-southeast in Northeast Transdanubia (Sebe et al. 2011). The hydrological network largely conforms to the topography of the yardangs. Furthermore, the region encompassing the mega-yardangs is often categorized as having high-to-very high erodibility due to wind in Hungary (Négyesi et al. 2019). However, current active wind erosion primarily occurs on agriculturally modified surfaces where permanent vegetation cover is missing. Natural wind erosion is presently confined to semi-arid sandy grasslands in specific locations within the Tisza-Danube Interfluve in Hungary (Veres et al. 2021). This limited occurrence of natural wind erosion can be attributed to the prevailing climate of the Pannonian mega-yardang system, characterized by a warm-summer humid continental climate (Köppen Dfb) (Beck et al. 2018). This region falls within the broadleaf biome (Karnieli et al. 2019) and is situated within the Pannonian Mixed Forest Ecoregion (Dinerstein et al. 2017). The annual mean temperature ranges between 10.1 and 11.4 °C, while annual precipitation varies between 520 and 820 mm in the area (Fick and Hijmans 2017).

Aeolian-origin geomorphologic features are widely recognized within the Pannonian Basin. For instance, several buttes are documented in the Tapolca Basin of Western Transdanubia (Fig. 1c) attributed the eolian deflation process (Gadányi 2015; Sebe et al. 2020). Although Lake Balaton, the largest lake in Central Europe, exhibits a mixed tectonic-deflation origin (Cserny and Nagy-Bodor 2000), deflation likely played a significant role in the formation of the present-day Balaton Basin (Fig. 1d) during dry periods characterized by sparse vegetation and the intense aeolian erosion processes (Cserny 2002). The presence of ventifacts and large-scale deflation forms such as mega-yardangs (Fig. 1e) in the Western Pannonian Basin suggests robust wind activity in the region's past (Ruszkiczay-Rüdiger et al. 2011; Sebe et al. 2015). Ruszkiczay-Rüdiger et al. (2011), through measurements of 10Be isotope concentrations along depth profiles of exposed, wind-faceted rocks, determined a denudation rate via deflation of 40–80 m/Ma in the Balaton Highland over the last 1.5 million years. The yardang system comprises several tens of kilometers, alongside deep and wide troughs (Suppl. Figure 1f, g).

Evidence for wind erosion in the studied region

Evidence of wind erosion in the studied region is abundant. Ventifacts, such as mushroom rocks (Suppl. Figure 2a), sharp keels (Suppl. Figure 2b), wind striations on horizontal surfaces (Suppl. Figure 2c), sphinx rocks (Suppl. Figure 2d), draikanters (Suppl. Figure 2e), wind-abraded faces in volcanic rocks (Suppl. Figure 2f), and ventifacted pebbles, have been extensively documented in Hungary (Jámbor 2002; Ruszkiczay-Rüdiger et al. 2009; Csillag 2010; Sebe et al. 2011). Ventifacts can be found in both old rocks, such as Paleozoic sandstones and Mesozoic carbonates, as well as in young, Neogene rocks, which include silica-cemented late Miocene sandstones, lacustrine breccias, and volcanic rocks. Wind striations on horizontal surfaces consistently align parallel to the prevailing northwest wind direction. Supplementary Fig. 1 provides visual documentation of the mentioned sites.

Sedimentary conditions in the Pannonian mega-yardang field

In addition to strong winds, the presence of erodible and relatively soft sediments is crucial in the formation of yardangs (Pelletier et al. 2018). Therefore, understanding the composition of surface sediments in the region where yardangs are found is paramount. In non-volcanic but arid geological environments, lacustrine and alluvial sediments may originate from previous humid climates and/or flood events (Ding et al. 2021; Yang et al. 2021). In the Western Pannonian Basin, both lacustrine and alluvial sediments could serve as the substrate for mega-yardang fields. During the Tortonian period, the basin was inundated by a large endorheic lake known as Lake Pannon, which deposited thick layers of lacustrine sediments, reaching hundreds or even thousands of meters in certain areas. For instance, extensive deltaic-lacustrine sequences are prevalent across many parts of the basin (Philips et al. 1994).

These deposits primarily consist of soft sedimentary rocks, including clay, silt, sand, and/or aleuritic sequences, such as the Tihany Formation (Sztanó et al. 2013; Suppl. Figure 3a, b) or the Somló Formation (Sztanó and Magyar 2010; Suppl. Figure 3c). Apart from cemented aleurites, these formations are generally highly erodible. Abrasion gravel layers, known as the Diás Gravel Formation (Suppl. Figure 3d), were deposited along the former shores of Lake Pannon (Csillag et al. 1996; see also Suppl. Figure 3a). In specific lagoons of Lake Pannon, lime mud was deposited (Jámbor 1990), which underwent partial diagenesis. Presently, the hard limestone layers serve to protect the soft mud from erosion (Suppl. Figure 3e). Following the desiccation of the expansive lake, bedded siltstones, sandstones, and claystones attributed to the Zagyva Formation (Györgyi 1992) were deposited within the alluvial plain of the Pannonian Basin. These sediments were predominantly formed in fluvial settings and partly in lacustrine environments (Budai et al. 2015) (Suppl. Figure 3f). Supplementary Fig. 1 shows the mentioned sites.

Considerations regarding the formation time of the Pannonian mega-yardang system

There is currently no consensus regarding the timing of the formation of the Pannonian mega-yardang system. While it is widely accepted that its development occurred post-Tortonian period, within the sediments of the Pannonian Lake, there are diverging hypotheses regarding the exact timeframe. Research indicates that during the early Pliocene, the climate of the Pannonian Basin was characterized as temperate and humid (Kovács et al. 2015). However, subsequent analysis of oxygen isotope data suggests a transition towards a more arid climate by the mid-Pliocene (Kovács et al. 2015). This transition is marked by the persistence of characteristic red clay formations until the end of the early Pleistocene in the Pannonian Basin (Kovács et al. 2011) indicating relatively dry and warm climatic conditions during this period. Furthermore, ecological studies focusing on herpetological assemblages suggest the occurrence of brief, yet relatively warm, and dry episodes extending into the Middle Pleistocene in Central Europe (Ivanov 2007).

Conversely, some researchers argue that the formation of the Pannonian mega-yardangs occurred under cold, periglacial climatic conditions (Sebe et al. 2011). Analysis of late Pleistocene and Holocene aeolian landforms, alongside investigations into phases of sand movement in the southern part of the mega-yardang field, provides evidence of active aeolian processes in recent geological eras as well (Györgyövics 2019). The range of evidence emphasizes the need to incorporate various periods between the Pliocene and the Holocene into climate models for the Pannonian Basin to adequately address potential aeolian processes. Thus, the objective of this study was to determine the probable age of the formation of the Pannonian mega-yardang system and elucidate the climatic conditions conducive to the development of such significant aeolian land features, such as mega-yardangs.

Materials and methods

Based on the aforementioned considerations, the logical frame of the study unfolded as follows: initially, two regional paleoclimatic models were constructed for two Zanclean periods, incorporating mean annual temperature and precipitation sum values. Subsequently, to delineate the past climatic conditions of the Pannonian mega-yardang system area, climatic models for two Pliocene periods (both Piacenzian) and 10 Quaternary periods were utilized. The climatically analogous regions corresponding to the past conditions of the Pannonian mega-yardangs were identified using Köppen–Geiger climatic classifications. Subsequently, the humidity zones of Western Eurasia, encompassing the studied region, were delineated based on the reconstructed climates of the two Pliocene (both Piacenzian) and four Quaternary stages. Following this, the former biome distributions within the Carpathian Basin were simulated using the reconstructed climatic parameters for four Pliocene periods (2 Zanclean and 2 Piecenzian) and ten Quaternary stages. Subsequent to these analyses, the biome distributions of Western Eurasia, Central Asia, and Asia Minor during the mid-Pliocene era (explained further below) were compared to the Pliocene to earliest Pleistocene occurrences of fossil camelid and struthionid taxa. Finally, the changing Köppen Aridity Indices of yardang sites across continents over the last 3.3 mys were compared with three sites within the Pannonian mega-yardang system.

The used climatic data

The study utilized three types of climatic data:

  1. 1)

    Fossil flora-based paleoclimatic data.

  2. 2)

    Georeferenced paleoclimatic reconstructions.

  3. 3)

    Mapped climatic data from 1979 to 2013, serving both as modelling tool and a reference period.

For modelling purposes, QGIS 3.16.13 software with Grass Gis 7.8.5 was employed.

Supplementary Fig. 4 illustrates the grids utilized in presenting the model results.

As a georeferenced Zanclean paleoclimatic model was unavailable for the Carpathian Basin, a new georeferenced model was developed. The two Zanclean fossil sites, Gérce and Pula (Suppl. Figure 2a), located in Transdanubia, Western Hungary, were utilized. The K/Ar radiometric ages were determined to be 4.55 for Gérce and 4.25 for Pula (Pápay 2001). Paleoclimatic values were based on the paleoclimatic reconstructions of Erdei et al. (2007), who employed the Coexistence Approach (Mosbrugger and Utescher 1997) to reconstruct past climatic conditions. This approach involves quantifying climate based on the plant fossil record, considering the typical range and the climatic limits of the closest extant relatives of the extinct plant species. The climatic values used included the annual mean temperature (°C) and annual precipitation sums (mm).

To create the paleoclimatic models, differences between the former and present-day climatic values were calculated. Subsequently, these difference values were added to the corresponding values of the region (see Table 1):

$$fz = f_{1970 - 2000} + (f_{{\text{Gerce}}/{\text{Pula}}} - f_{1970 - 2000} )$$
(1)

where fz is the reconstructed Zanclean annual mean temperature (°C) or annual precipitation sum values (mm), fz is the value of the identical climatic factor in the reference period (1970–2000), and fGérce/Pula are the reconstructed annual mean temperature (°C) or annual precipitation sum values (mm) values in Gérce or Pula.

Table 1 Reconstructed annual mean temperature (TM) and annual precipitation sum values (Ps) in Gérce in 455 mya and in Pula in 425 mya according to Erdei et al. (2007) and the differences of the Zanclean values from the moderns
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Two of the georeferenced paleoclimatic models depict the climatic conditions of the mid-Pliocene cold and warm periods. In addition, one model represents the climate of the mid-Pleistocene MIS19 interglacial, while another illustrates the climatic conditions of the Late Pleistocene (MIS 2) Last Interglacial Maximum. The late Pleistocene paleoclimatic models encompass the Last Glacial Maximum, Heinrich Stadial 1, the Bølling–Allerød, and the Younger Dryas Stadial. Three Holocene paleoclimatic models were also incorporated: the Greenlandian, the Northgrippian, and the Meghalayan climate models, along with the climate model spanning from 1970 to 2000 (Table 2). In these georeferenced climatic models, the climatic factors considered were the annual mean temperature (Tm, °C), the January mean temperature (Tm01, °C), and the annual precipitation sum (Ps, mm).

Table 2 Used site-level reconstructed paleoclimate data and the utilized georeferenced paleoclimatic models
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Pliocene fossil data

Extant camelid and struthionid species are predominantly found in open, arid, and semi-arid habitats. For instance, wild Bactrian camels (Camelus bactrianus L.) inhabit arid and semi-arid regions of Central Asia (Kaczensky et al. 2014). Present-day Struthio camelus L. preferentially occupies open, short-grass plains, semi-deserts, and true deserts in Africa (Cooper et al. 2009). Plausibly, these species inhabited very similar habitats in Eurasia before the anthropogenic extinction on this continent (Kuhn 2005). Considering that it can be hypothesized that the Pliocene relatives of extant camelid and struthionid taxa also lived in relatively arid habitats, their fossil relatives can serve as indicator species of semi-arid and arid climatic conditions. The Pliocene–early Pleistocene occurrence data of Camelidae and Struthionidae in Central and Western Eurasia are presented in Suppl. Table 1.

Climatic and environmental classifications

In the subsequent step, the area corresponding to the present-day climatic analogues of the reconstructed paleoclimates was determined. Western Eurasia was chosen as the region to identify climatically analogous areas. The determination of present-day climatically analogous regions was carried out using the logic of Boolean algebra. Deterministic unit step functions were formulated using the respective upper and lower extrema as follows:

$$1(cf) = \left\{ \begin{gathered} 0\;\;if\;\;cf_{{\text{upper}}\;{\text{extrema}}} < cf\;{\text{and}}\;cf\; < \;c_{{\text{flower}}\;{\text{extrema}}} \hfill \\ 1\;\;if\;\;cf\; \le \;bio_{{\text{upper}}\;{\text{extrema}}} \;{\text{and}}\;c_{{\text{flower}}\;{\text{extrema}}} \le cf \hfill \\ \end{gathered} \right.$$
(2)

The temperature-like climatic factors (cf) employed in the analysis included the annual mean temperature (°C) and the mean temperature of the coldest quarter (°C). The annual precipitation sum (mm) represented precipitation-like climatic factors. By utilizing both the lower and upper extrema associated with these factors, a total of 2 × 3 climatic limits were considered in each case.

The areas excluded by the bioclimatic limiting factors can be summed, and the intersection of the potential areas designated by the factors provides the aggregated distribution area, formally expressed as

$$A(fT_{{\text{upper}};{\text{lower}}} ;fP_{{\text{upper}};{\text{lower}}} ) = fT\; - \;\sum_{i = 1}^3 {0({\text{bio}}T_{{\text{upper}};{\text{lower}}} )} \cap fP - fP_{{\text{upper}};{\text{lower}}}$$
(3)

where A(fTupper;lower;fPlupper;ower) represents the climatically analogous areas, which contain the remaining areas after considering the temperature (fT) or precipitation nature (fP).

As a climatic classification system, the Köppen–Geiger climate classification (Rubel and Kottek 2011) was employed, comprising five main climate groups (A: tropical, B: arid, C: temperate, D: continental, and E: polar). The climatic classes are determined based on seasonal precipitation and temperature patterns. The present Köppen–Geiger climate classification at 1-km resolution was derived from the map provided by Beck et al. (2018).

Biome modelling

The biome modelling was conducted based on the methodology outlined by Trájer (2022). To reconstruct the biome patterns, the author employed the area-delimiting curves of the revised Whittaker biome plot as described by Stefan and Levin (2020) and Deshmukh and Singh (2019). For modelling purposes, equations representing the upper or lower delimiting curves associated with biome areas were utilized. The selected biomes included temperate grassland or cold desert (TeGr/CoDE), woodland or shrubland (Wo/Sh), temperate seasonal forest (TeSeFo), boreal forest (BoFo), tundra, and ice (Tu + Ic; including the Alpine) biomes. The equation for the upper delimiting curve of the temperate grassland/cold desert biome is as follows:

$$- 0.3498 \times Tm^2 + 24.436 \times Tm + 159.92 = Ps$$
(4)

The equation for the upper delimiting curve of the woodland/shrubland biome is as follows:

$$- 0.0086 \times {\text{Tm}}^4 + \;0.2955\; \times \;{\text{Tm}}^3 - 2.1072\; \times \;{\text{Tm}}^2 + 41.12\; \times \;{\text{Tm}}\;\; + \;242.93\; = \;{\text{Ps}}$$
(5)

The equation for the upper delimiting curve of the temperate seasonal forest biome is as follows:

$$\begin{gathered} 2235.23\; + \;(47.91482 - 2235.23)/\left( {1\; + \;\left( {\frac{{{\text{Tm}}}}{7.782309}} \right)^{6.265971}} \right)= {\text{Ps}}\;{\text{and}}\;{\text{if}}\;{\text{Tm}} > \hfill \\ - 2.5^\circ C\;{\text{and}}\;{\text{Tm}}\; < \;22.5^\circ C \hfill \\ \end{gathered}$$
(6)

The equation for the upper delimiting curve of the boreal forest biome is as follows:

$$4.8254\; \times \;{\text{Tm}}^3 - 77.874\; \times \;{\text{Tm}}^2 + 519.17 \times {\text{Tm}} + 318.67 = Ps$$
(7)

The equation for the lower delimiting curve of the tundra and ice biome is as follows:

$$4.8254\; \times \;{\text{Tm}}^3 - 77.874\; \times \;{\text{Tm}}^2 + 519.17 \times {\text{Tm}} + 318.67 = Ps$$
(8)

where Tm is the annual mean temperature (°C), and Ps is the annual precipitation sum (mm).

Modelling of aridity indices

Furthermore, to evaluate paleohumidity conditions, the Köppen Aridity Index (KAI) (Köppen 1923, 2011) was employed. This aridity-related metric is calculated by dividing the mean annual precipitation (Ps) by the sum of the mean annual temperature (Tm) and a constant of 33 (Quan et al. 2013, 2014):

$${\text{KAI}}\; = \;\frac{Ps}{{{\text{Tm}} + 33}}$$
(9)

where KAI (mm°C−1) denotes the Köppen Aridity Index, Ps (mm) represents the total mean annual precipitation, Tm (°C) stands for the mean annual temperature, and 33 is a constant.

The range of KAI values characterizes climate into five distinct categories (Quan et al. 2013, 2014); those are: hyper-arid (< 0.9), arid (0.9–5.7), semi-arid and dry semi-humid (5.7–13.6), semi-humid (13.6–15.6), and humid (> 15.6).

Yardang sites

Apart from the sites within the Pannonian mega-yardang system, 14 yardang sites were selected from five continents of the Earth, distributed as follows: 3 from the Americas, 3 from Africa, 7 from Eurasia, and 1 from Australia. The chosen yardang sites are as follows: Rogers Lake (USA; Ward et al. 1984), Ocotillo Wells State Vehicular Recreation Area (USA; Pelletier et al. 2018), Campo Piedra Pomez (Argentina; de Silva et al. 2008), Bujaraloz village (Spain; Gutiérrez-Elorza et al. 2002), Karas Region (Namibia; Goudie and Viles 2014), Zalaegerszeg, Kisláng, and Ercsi (these three sites characterize the Pannonian mega-yardang field; see Suppl. Figure 1). Aorounga (Chad; DiMartino et al. 2019), El Qasr, Dakhla Oasis (Egypt; Brookes 1993), Um Al-Rimam depressions (Kuwait; Al-Dousari et al. 2009), Shahdad Desert (Iran; Maghsoudi and Maghsoudi 2021a, b), Dunhuang (China; Liang et al. 2019), between Mahai and Xiaoqaidam (China; Wang et al. 2021), north of Dabuxun Lake (China; Wang et al. 2018), southeast of Lenghu (China; Wang et al. 2018), Lake Mungo (Australia; Goudie and Wells 1995) (Fig. 2).

Fig. 2
figure 2

Yardang sites: 1: Rogers Lake (USA), 2: Ocotillo Wells State Vehicular Recreation Area (USA), 3: Campo Piedra Pomez (Argentina), 4: Bujaraloz village (Spain), 5: Karas Region (Namibia), 6: Zalaegerszeg, 7: Kisláng, 8: Ercsi, 9: Aorounga (Chad), 10: El Qasr, Dakhla Oasis (Egypt), 11: Um Al-Rimam depressions (Kuwait), 12: Shahdad Desert (Iran), 13: Dunhuang (China), 14: between Mahai and Xiaoqaidam (China), 15: north of Dabuxun Lake (China), 16: southeast of Lenghu (China), and 17: Lake Mungo (Australia) projected to the global pattern of humidity regimes based on the Köppen Aridity Index values

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Results

Geological evidence related to the age of the mega-yardangs

To the west of Lake Balaton (Fig. 3a; also referenced in Fig. 1f), where the prevailing wind direction is northerly, the mega-yardangs are predominantly enveloped by Pleistocene loess. However, in the wind-exposed northern regions and at the elevated ridges of the mega-yardangs, the loess has undergone erosion, revealing the underlying bedrock composed of Tortonian lacustrine formations and post-Tortonian, pre-Quaternary sediments (Zagyva Formation) at the surface. Moving northward towards the Transdanubuan Mts., specifically in the Sokoró Hills (Fig. 3b, as depicted in Fig. 1g), where the prevailing wind direction veers northwest, the mega-yardangs have developed within Tortonian lacustrine sediments. Pleistocene loess and sand have either been conserved or accumulated solely on the wind-sheltered, eastern, and southeastern slopes of the mega-yardangs. Notably, the troughs within both regions are overlain by fluvial–deluvial sediments dating back to the late and mid-Pleistocene epochs.

Fig. 3
figure 3

Surface geological formations of mega-yardangs in Zala county, west to Lake Balaton (a) and the Sokoró Hills, north to the Transdanubian Mts (b) (see Fig. 1f, g), based on the geological map server of the Mining and Geological Survey of Hungary Blue arrow indicates the prevailing wind direction Tu: Upper Triassic carbonates Mt: Upper Miocene, Tortonian sediments; MPv: Miocene–Pliocene volcanites; MPf: Miocene–Pliocene fluvial sediments; Ql: Quaternary loess, Qs: Quaternary sand, Qa–d: Quaternary: alluvial–deluvial sediments The blue arrow shows the prevailing wind direction, Cl: contour liney by 50 m, pwd: prevailing wind direction. The regions depicted in this figure correspond to specific areas shown in Fig. 1b

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Paleoclimates

Climatically analogue regions

Based on the reconstructed paleoclimatic parameters such as annual mean temperature, mean temperature of the coldest quarter, and annual precipitation sum during the mid-Pliocene cold period (3.3 Ma), it can be concluded that the climate characterizing the Pannonian mega-yardang field was dominantly akin to a hot-summer Mediterranean (Csa) and semi-arid (Bsh) climate. Regions exhibiting humid subtropical (Cfa) and continental climates were primarily confined to the higher elevations of the central mountains. Present-day climatically analogous regions can primarily be identified in Southeast Europe, encompassing the Balkans and Asia Minor (Fig. 4a).

Fig. 4
figure 4

Köppen–Geiger climatic categories in present-day Europe are climatically analogous to the mid-Pliocene and Pleistocene periods according to the reconstructed mean annual temperature and annual precipitation sum values; a mid-Pliocene cold period (33 Ma), b mid-Pliocene warm period (3205 Ma), c MIS19 interstadial (787 ka), d Last Interglacial (130 ka); Köppen–Geiger classes: BSk: cold semi-arid climate, Csa: hot-summer Mediterranean climate, Csb: warm-summer Mediterranean climate, Cfa: fully humid subtropical climate, Cfb: temperate oceanic climate, Dfa: hot summer humid continental climate, Dfb: warm-summer humid continental climate, Dsa: Mediterranean-influenced hot-summer humid continental climate, Dsb: Mediterranean-influenced warm-summer humid continental climate, Dfc: subarctic climate, ET: tundra. The region depicted in this figure corresponds to a specific area (area No. 5) shown in Suppl. Figure 4

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In contrast, during the mid-Pliocene warm period (3.205 Ma) in Western Eurasia, climatic analogs correspond to the temperate categories, including temperate oceanic (Cfb), permanently humid subtropical (Cfa), and hot or warm-summer Mediterranean (Csa, Csb) climates. The climatically analogous regions during this period are predominantly observable in present-day Southwest Europe (Csb; Fig. 4b).

During the MIS19 interglacial period (787 ka), the climate prevalent in the Pannonian Basin was characterized by hot- and warm-summer humid continental conditions (Dfa and Dfb). In addition, there exists a degree of resemblance to present-day cold semi-arid (BSk) and warm-summer humid continental (Dsb) climate regions (Fig. 4c).

In the Last Interglacial Period (130 ka), the area encompassing the Pannonian mega-yardangs experienced a climate characterized by hot- and/or warm-summer humid continental conditions (Dfa and Dfb; Fig. 4d). Climatic analogs for both the MIS19 and the Last Interglacial periods can be identified in Central Europe, as well as in the western part of Eastern Europe (Fig. 4c, d).

Based on paleoclimatic data, the late Pleistocene epoch witnessed warm-summer humid continental (Dfb) and subarctic (Dfc) climates prevailing in the vicinity of the Pannonian mega-yardangs. However, the distribution of these climates varied over time. During the Last Glacial Maximum (21 ka; Fig. 5a) and Heinrich Stadial 1 (17.0–14.7 ka; Fig. 5b), the subarctic climate (Dfc) was predominant. Conversely, during the Bølling–Allerød period (14.7–12.9 ka), Transdanubia (Fig. 5c) experienced a warm-summer humid continental (Dfb) climate. In the Younger Dryas Stadial (12.9–11.7 ka), the dominance of these two climates was roughly equal (Fig. 5d). Present-day climatic analogs for these late Pleistocene stages can be found in Northern Europe, the Eastern and Northeastern parts of Eastern Europe, as well as in Western Siberia (Fig. 5a–d).

Fig. 5
figure 5

Köppen–Geiger climatic categories in present-day Europe are climatically analogous to the Late Pleistocene according to the reconstructed mean annual temperature and annual precipitation sum values; a Last Glacial Maximum (21 ka), b Heirich Stadial 1 (170–147 ka), c Bølling–Allerød (147–129 ka), d Younger Dryas Stadial (129–117 ka); Köppen–Geiger classes: BSk: cold semi-arid climate, Csa: hot-summer Mediterranean climate, Csb: warm-summer Mediterranean climate, Cfa: fully humid humid subtropical climate, Cfb: temperate oceanic climate, Dfa: hot summer humid continental climate, Dfb: warm-summer humid continental climate, Dsa: Mediterranean-influenced hot-summer humid continental climate, Dsb: Mediterranean-influenced warm-summer humid continental climate, Dfc: subarctic climate, ET: tundra. The region depicted in this figure corresponds to a specific area (area No. 5) shown in Suppl. Figure 4

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Based on reconstructed paleoclimatic data, the post-glacial warming period spanning from the Greenlandian (11.7–8.326 ka) to the reference period (1970–2000) resulted in a southwestward shift of climatic analogs for the Pannonian mega-yardangs (Fig. 6a–d). Despite this shift, the prevailing climate remained warm-summer humid continental (Dfb), with influences of hot-summer humid continental (Dfa) and warm, dry-summer continental climates (Dsb) persisting throughout the Holocene. In addition, in the contemporary era, some influences of temperate oceanic (Cfb) and cold semi-arid (BSk) climates can be observed based on climatically analogous regions (Fig. 6d). During the Greenlandian and Northgrippian Stages of the Holocene, climatically analogous regions are identifiable primarily in the western regions of Eastern Europe (Fig. 6a, b). Conversely, for the Meghalayan and the reference period (1970–2000), climatically analogous regions are situated in Eastern and East-Central Europe, as well as in the Northern Balkans and Asia Minor (Fig. 6c, d).

Fig. 6
figure 6

Köppen–Geiger climatic categories in present-day Europe are climatically analogous to the Holocene stages according to the reconstructed mean annual temperature and annual precipitation sum values; a Greenlandian (117–8326 ka), b Northgrippian (8326–42 ka), c Meghalayan (42–03 ka), d 1970–2000; Köppen–Geiger classes: BSk: cold semi-arid climate, Csa: hot-summer Mediterranean climate, Csb: warm-summer Mediterranean climate, Cfa: fully humid subtropical climate, Cfb: temperate oceanic climate, Dfa: hot summer humid continental climate, Dfb: warm-summer humid continental climate, Dsa: Mediterranean-influenced hot-summer humid continental climate, Dsb: Mediterranean-influenced warm-summer humid continental climate, Dfc: subarctic climate, ET: tundra. The region depicted in this figure corresponds to a specific area (area No. 5) shown in Suppl. Figure 4

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Humidity patterns in Western Eurasia

During the mid-Pliocene warm period, the broader region encompassing the Pannonian mega-yardangs exhibited a humid climate. Semi-arid and dry sub-humid climatic regions were confined mainly to the eastern part of the Pannonian basin. Across Europe during this period, most areas experienced a humid or subhumid climate. In the MIS19 period, semi-arid regions were observed in the eastern and southeastern parts of the Pannonian Basin, as well as in the East Balkans, North Pontic, and North Caspian regions in Western and Central Eurasia (Fig. 7b). Similar patterns were evident in Western Eurasia during the Last Interglacial and the reference period, although during these periods, sub-humid regions extended over larger areas of the Pannonian basin, excluding the primary territory of the Pannonian mega-yardang system (Fig. 7d, f). During the Last Glacial Maximum, neither semi-arid nor dry sub-humid regions were discernible in the Carpathian Basin (Fig. 7e).

Fig. 7
figure 7

Humidity regimes in western Eurasia and northern Africa based on the Köppen Aridity Index values; a mid-Pliocene cold period, b mid-Pliocene warm period, c MIS19 interstadial, d Last Interglacial, e Last Glacial Maximum, f 1970–2000. The region depicted in this figure corresponds to a specific area (area No. 6) shown in Suppl. Figure 4

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Pliocene and Quaternary biome patterns

From the Rupelian to the Last Interglacial

During the two Zanclean periods (4.55 and 4.25 Ma), a significant portion of the Pannonian Basin, including its northern and eastern territories within the realm of the Pannonian mega-yardang system, was enveloped by a woodland–shrubland biome. Temperate seasonal forests dominated the vegetation in mid-elevation areas and southwestern Transdanubia (Fig. 8a, b). In the mid-Pliocene cold period (3.3 Ma), the central region of the Pannonian Basin, including the eastern part of Transdanubia, transitioned to a grassland–cold desert biome. However, woodland–shrubland vegetation persisted across the rest of the Carpathian basin, including mid-elevation regions (Fig. 8c). During the mid-Pliocene warm period (3.205 Ma), the temperate seasonal forest biome re-established itself in Transdanubia, while the grassland–cold desert biome remained confined primarily to the Northwest Balkans, with mid-elevation regions of the Pannonian Basin transitioning back to woodland–shrubland (Fig. 8d). Similar biome patterns emerge in the MIS19 interglacial (787 ka) and the Last Interglacial (130 ka) periods, albeit with an increased presence of the boreal biome in high-elevation regions of the Carpathian Basin (Fig. 8e, f).

Fig. 8
figure 8

Modelled biomes in the Pannonian Basin between the Zanclean and the Last Interglacial TeGrCoDE: temperate grassland or cold desert, WoSh: woodland or shrubland, TeSeFo: temperate seasonal forest, BoFo: boreal forest, Tu + Ic: tundra and ice. a Zanclean, 425 Ma; b Zanclean, 455 Ma, c mid-Pliocene cold period, d mid-Pliocene warm period, e MIS19 interstadial, f Last Interglacial. The region depicted in this figure corresponds to a specific area (area No. 3) shown in Suppl. Figure 4

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Late Pleistocene biome patterns

During the Last Glacial Maximum (21 ka) and the Heinrich Stadial (17.0–14.7 ka), the Carpathian Basin was characterized by boreal, tundra, and ice biomes (Fig. 9a, b). In the lower elevation regions of Transdanubia, the boreal biome prevailed, while in mid-elevation areas, primarily in the northern half of the basin, the tundra and ice biome also emerged. Transdanubia exhibited a similar biome pattern during the Bølling–Allerød (14.7–12.9 ka) and the Younger Dryas (12.9–11.7 ka) periods, with the boreal forest biome dominating (Fig. 9c, d). During the Last Glacial Maximum and the Heinrich Stadial, the temperate seasonal forest and woodland–shrubland biomes were confined mainly to the Northern Balkans, extending into the southern region of the Pannonian Basin during the Bølling–Allerød period (14.7–12.9 ka; Fig. 9c). However, these biomes did not experience significant areal expansion during the Younger Dryas Stadial (12.9–11.7 ka; Fig. 9d).

Fig. 9
figure 9

Modelled biomes in the Pannonian Basin in the Late Pleistocene TeGrCoDE: temperate grassland or cold desert, WoSh: woodland or shrubland, TeSeFo: temperate seasonal forest, BoFo: boreal forest, Tu + Ic: tundra and ice. a Last Glacial Maximum, b Heinrich Stadial, c Bølling–Allerød interstadial, d Young Dryas Stadial. The region depicted in this figure corresponds to a specific area (area No. 3) shown in Suppl. Figure 4

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Holocene biomes in the Carpathian basin

In parallel with post-glacial warming, there has been a gradual retreat of the tundra, ice, and boreal biome during the Holocene in the Carpathian Basin. During the Greenlandian period (11.7–8.326 ka; Fig. 10a), deciduous forests predominantly covered most of Transdanubia, while boreal forests still persisted in the mid-elevation regions. In addition, the appearance of the woodland–shrubland biome was noted in eastern Transdanubia during this period. Subsequently, during the Northgrippian (8.326–4.2 ka; Fig. 10b) and Meghalayan (4.2–0.3 ka; Fig. 10c) periods, there was a gradual westward expansion of the woodland–shrubland biome. In the contemporary era, deciduous forest constitutes the natural vegetation of the area surrounding the Pannonian mega-yardangs in the southwest, while the eastern and low-elevation regions are characterized by the woodland–shrubland biome (forest steppe; Fig. 10d).

Fig. 10
figure 10

Modelled biomes in the Pannonian Basin in the Holocene TeGrCoDE: temperate grassland or cold desert, WoSh: woodland or shrubland, TeSeFo: temperate seasonal forest, BoFo: boreal forest, Tu + Ic: tundra and ice a Greenlandian, b Northgrippian, c Meghalayan, d 1970–2000. The region depicted in this figure corresponds to a specific area (area No. 3) shown in Suppl. Figure 4

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Comparison of the mid-Pliocene biome patterns and the occurrence of fossil camelids and struthionids

Comparing the Pliocene occurrence of fossil camelid (Fig. 11a) and struthionid (Fig. 11b) species with the biome patterns of the mid-Pliocene cold period reveals an interesting correlation. Ancient relatives of camelids and ostriches were discovered in sites located approximately on the border of woodland–shrubland and grassland–cold desert biomes. These fossil sites spanned from the western part of the Pannonian Basin (Transdanubia) to Central Asia, Asia Minor, and the Southeast Balkans, including the North Pontic region. This distribution suggests that during the mid-Pliocene cold period, a similar relatively dry, steppe-like environment may have characterized the Pannonian Basin, akin to present-day western regions of Central Asia. The presence of camelid and struthionid fossils in areas bordering woodland–shrubland and grassland–cold desert biomes implies that these species likely thrived in transitional habitats with a mix of vegetation types, supporting the idea of a diverse mosaic of ecosystems during this period.

Fig. 11
figure 11

Fossil occurrence of camelids and struthionids in the mid-Pliocene cold period (33 Ma) in Central Asia, Asia Minor, and Eastern Europe TeGrCoDE: temperate grassland or cold desert, WoSh: woodland or shrubland, TeSeFo: temperate seasonal forest, BoFo: boreal forest, Tu + Ic: tundra and ice a the known occurrences of genera Camelidae (inc genera Camelus and Paracamelus) in the Pliocene–early Pleistocene period: 1: Kisláng and 2: Ercsi, Hungary; 3: Catacombs of Odessa and 4: Odessa sl, Ukraine; 5: Yevpatoriya, Ukraine; 6: Novo-Nikolayevka, Ukraine; 7: Lyapino, Ukraine; 8: Vinodel'noye and 9: Rostov-na-donu, Russian Federation, 10: Selim-Dzhevar, Kazakhstan. b Known occurrences of Struthionidae (incl genera Struthio and Pachystruthio) in the Pliocene–early Pleistocene period: 1: Kisláng, Hungary; 2: Priozernoe, Moldova; 3: Catacombs of Odessa, Ukraine; 4: Liventsovka, Russian Federation; 5: Sesklon, Greece; 6: Çalta, Turkey; 7: Pawlodar City, Kazakhstan; 8: Akterek, Kyrgyzstan. The region depicted in this figure corresponds to a specific area (area No. 4) shown in Suppl. Figure 4

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Comparison of Köppen Aridity Indices in yardang sites in the last 3.3. Ma

Over the past 3.3 million years, the Köppen Aridity Index (KAI) values at the examined yardang sites, except for the Pannonian mega-yardang system, generally ranged between 0 and 10 mm°C−1. However, in many of the yardang sites situated in regions where current KAI values are less than 0.5 mm°C−1, there were periods within the last 3.3 million years when KAI values exceeded 1.5 mm°C−1. Presently, apart from the Transdanubian sites, KAI values at all yardang locations are below 7 mm°C−1. In the three Transdanubian sites, current KAI values range from approximately 15 to 12 mm°C−1, and except for the mid-Pliocene warm period (3.3 Ma), KAI values exceeded 11.5 mm°C−1 for all studied periods. During the mid-Pliocene cold period, the KAI values of the three Pannonian sites ranged between 7 and 11.5 mm°C−1 (Fig. 12).

Fig. 12
figure 12

Köppen Aridity Index (KAI) values between the Zanclean and 1970–2000 MH: mid-Pliocene period, NG: Northgrippian period, GL: Greenlandian period, YDS: Younger Dryas Glacial, BA: Bølling–Allerød, HS: Heinrich Stadial, LGM: Last Glacial Maximum, LIP: Last Interglacial, MIS19: MIS19 interglacial, MPWP: mid-Pliocene warm period, MPCP: mid-Pliocene cold period

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Discussion

Quaternary glacial periods

The results prompt an inquiry into the extent to which aeolian processes contributed to the formation of the relief during glacial periods in Transdanubia, assuming the glacial period origin of the Pannonian mega-yardang field (Ruszkiczay-Rüdiger et al. 2011; Sebe et al. 2011). Aeolian landforms are known to exist in arctic and subarctic regions. For instance, strong katabatic winds in Greenland create deflation patches near glaciers (Heindel et al. 2017); coastal sand dune systems are present in Quebec, Canada, under subarctic climatic conditions (Bélanger and Filion 1991); and active niveo-aeolian transport processes have been observed in Ellesmere Island, Canadian Arctic (Lewkowicz 1998). Yardangs also occur in cold desert climate territories such as the Tavan Har in Mongolia (Carson et al. 2004) or the Qaidam Basin in China (Ding et al. 2020). Although low annual mean temperatures do not preclude the formation of yardangs, as evidenced by their occurrence in the cold Martian environment (Ding et al. 2020), the presence of yardangs is unusual in regions with boreal, humid continental climatic conditions. This uniqueness, combined with their size, distinguishes the Pannonian mega-yardangs.

Based on climatically analogous regions and modelled biomes of the Pannonian Basin, it was determined that during the Last Glacial Maximum and Heinrich Stadial 1, the area of the Pannonian mega-yardangs experienced boreal and/or cold, humid continental climates, excluding mountainous territories. This contrasts with findings from some authors who suggested that the dominant vegetation in the Pannonian Region during the glacial period consisted of cold and dry steppe or steppe tundra, interspersed with varying amounts of arboreal vegetation patches (Willis et al. 2000; Sümegi et al. 2002; Járainé Komlódi 2003; Gábris and Nádor 2007). Contrary to the previously proposed vegetation picture, the present study indicates that boreal forests characterized the territory of the Pannonian mega-yardangs during glacial maximums, with a fully humid climate prevailing in the Carpathian Basin during this period. For instance, during Heinrich Stadial 1, even the coldest parts of the Pannonian mega-yardangs system exhibited a climate like that of the southwestern Zapolyarny District of Northwest Russia, where spruce and fir species-formed forests and extensive marshes are prevalent (Schepaschenko et al. 2011). In this region, typical tundra vegetation is only found on the Kanin Peninsula, north of the climatically analogous regions of the Last Glacial Maximum. This suggests that while tundra-like paleoenvironments could have been present in the higher and mid-elevation regions of the northeastern part of the Carpathian Basin during the glacials, the hypothesized dry and cold, so-called "mammoth steppe" with predominantly grass-dominated vegetation could not have been the dominant biome in the Pannonian Basin.

Quaternary post-glacial and interglacial periods

While evidence exists regarding sand dune movement in the Holocene in the Duna-Tisza Interfluve, Hungary (Kiss et al. 2006), both climatically analogous territories and results from biome modelling, along with palynological evidence (Medzihradszky 2005), suggest that throughout the entire Holocene epoch, the climate of Transdanubia maintained a humid continental (Dfa or Dfb) character. During these stages, the Transdanubian part of the Pannonian Basin was enveloped by temperate seasonal forests, indicating that the environmental conditions necessary for the formation of aeolian landforms were likely absent due to dense vegetation cover. Model results from the latest Pleistocene, along with climatic analogs, propose that Transdanubia experienced a cool but relatively humid transitional boreal/continental climate during this period. Boreal coniferous forests and, in warmer areas, mixed coniferous–deciduous forests may have covered the territory of the Pannonian mega-yardangs during this time. This finding aligns with climatic proxies inferred from palynological data by Nagy-Bodor et al. (2020). In addition, it is worth noting that in the Szigliget Bay of the Balaton Basin, Transdanubia, peat formation commenced around 16.79–16.39 kya BP (Sümegi et al. 2008; Jakab et al. 2005), indicating humid climatic conditions shortly after the Last Glacial Maximum in the region.

In their investigation of the stable isotope compositions of speleothems from the last interglacial period, Demény et al. (2017) discovered evidence suggesting the cessation of stalagmite formation in the Baradla Cave, Northeast Hungary, around 110.1 thousand years ago (kya). Siklósy et al. (2011) ruled out the possibility of stalagmite formation being halted due to freezing and instead proposed a hypothesis linking the cessation to decreasing soil activity concurrent with cooling and increasing aridity after the Last Interglacial Period. It is pertinent to note that the dissolution of carbonates shows positive correlations with temperature (Romero-Mujalli et al. 2019). However, during this period, the annual mean temperature in the Baradla Cave was only approximately 3 °C (Demény et al. 2017). Such a low temperature indicates very low-intensity subpolar karst formation processes (Baker et al. 2019). This factor alone could plausibly explain the cessation of stalagmite formation, even without a significant decrease in annual rainfall.

Semi-arid and arid conditions were not characteristic of earlier interglacial periods, including the MIS19 interglacial and the Last Interglacial, nor during the Zanclean and mid-Pliocene warm periods. Evidence from fossil flora at Zanclean-age sites suggests mesophytic conditions (Erdei et al. 2007). However, the climatic model for the mid-Pliocene cold period indicates the emergence of semi-arid climatic conditions in the Carpathian Basin. Comparing the Köppen Aridity Indices of yardang sites on five continents with values from three Transdanubian localities over the last 3.3 million years also suggests that the climate of Transdanubia was conducive to the formation of yardangs only around the Neogene/Quaternary boundary. During this era, semi-arid and sub-humid regions extended far to the west, reaching the shores of the Atlantic Ocean and Western Europe.

Late Neogene periods and the impact of the retreat of Lake Pannon

Ding et al. (2020) noted that although the mega-yardangs of the Pannonian Basin do not exhibit an obvious relationship with lacustrine sediments, the presence of deflation hollows in the western Pannonian Basin suggests potential geomorphological evolution during dry–wet climatic fluctuations. However, contrary to this assertion, it can be argued that the Pannonian mega-yardang system actually formed within the desiccated basin of a large Tortonian lake, known as Lake Pannon. This paleolake occupied approximately 250 km2 of area in the Carpathian Basin (Balázs et al. 2018) and, in certain localities, reached depths of 1–2 hundred meters (Sztanó et al. 2020), akin to the present-day Black Sea. The prograding delta of the paleo-Danube River system reached the northwestern foreland of the Transdanubian Mountains (it corresponds to a specific area, No. 1, shown in Suppl. Figure 4) around 8.9 million years ago in the latter half of the Tortonian period and retreated from the area of the Pannonian mega-yardang system around 7.5 million years ago in the latest Tortonian stage (Magyar 2021). As depicted in Suppl. Figure 3, the paleolake left behind various sediment types, with most consisting of sandy materials, clays, and silts, possessing a well-erodable, soft nature. Although limited data is available regarding the climate evolution from the late Tortonian to the Messinian stage in the region, existing paleoflora evidence suggests a somewhat drier climate (Bruch et al. 2011) compared to the warm and humid conditions of the preceding Tortonian stage (Erdei et al. 2007). However, Csató et al. (2007) found that the relative lake-level decline during the Messinian period, the successor of Lake Pannon, was more likely attributable to tectonic processes than a significant decrease in annual precipitation.

Both the Late Miocene–Pliocene sedimentary environments in the Slovakian part of the Danube Basin (Kováč et al. 2006) and contemporaneous fluvial sediments in the southeastern foreland of the Vértes Mountains (Fodor et al. 2005) suggest the past existence of braided paleochannels. Braided channels typically form in non-cohesive sediments such as sand and gravel and can be associated with alluvial fans in glaciated areas (Malard et al. 2006), mountainous regions (Lu et al. 2022), coastal plains (Leigh et al. 2004), and arid regions (Laronne and Shlomi 2007). While the presence of braided paleochannels in non-glaciated environments supports the hypothesis of a relatively dry post-Tortonian climate in the Pannonian Basin, it's important to consider that the entire late Miocene–Zanclean Transdanubia could be regarded as a coastal plain of Lake Pannonian. This complicates determining under which climatic conditions this braided river system formed.

However, certain geological and fossil findings provide evidence supporting the former existence of late Neogene to early Quaternary dry paleoenvironmental conditions in the Pannonian Basin. The presence of camelid and struthionid fauna elements in the late Pliocene fossil record of the Pannonian Basin has been documented. In addition, the recurring appearance of 'red beds' in the latest Miocene–Pliocene geological strata of the Carpathian Basin suggests fluctuating humidity conditions in the post-Pannonian era (Kováč et al. 2006). Further evidence comes from the discovery of 4.2–3.6 million-year-old reddish karst fissure-filling sediments in Ivanovce, Slovakia, which are associated with the global climate transition interval (Zhang et al. 2020). The red beds of the Lukáčovce Member, Slovakia, deposited at the Pliocene/Quaternary boundary, could represent remnants of fossilized, terra-rossa-like paleosoils (Šarinová 2005). Red clays and kaolinitic red clays are also known from the Villány Mountains in southwestern Hungary (Varga 2011). The chemical composition of the mid-Pliocene to early Pleistocene Tengelic Formation (Villány Mountains) indicates generally warm and dry environmental conditions with former savannah, steppe, or forest steppe biomes (Viczián 2007). In addition, the red clays of the Middle Pleistocene Tengelic member of the Paks Loess Formation (Transdanubia, Hungary) suggest semi-arid, steppe-like biomic conditions during interglacial periods (Viczián 2007). These observations suggest that the sub-humid, semi-arid climate characteristic of the mid-Pliocene cold period (Dolan et al. 2015) may have persisted between the mid-Pliocene and the early Pleistocene epochs, alternating with wetter periods. This variation between relatively dry and humid periods aligns with the principle proposed by Ding et al. (2020), who stated that the formation of yardangs requires fluctuating dry and wet climatic conditions.

Other evidence related to the age of the mega-yardangs

The geological composition of the Pannonian mega-yardangs indicates that they were formed within pre-Quaternary formations, such as the Tihany Formation, Somló Formation, and Zagyva Formation, which were deposited during the Tortonian–Pliocene interval. The valleys between the yardangs are partly filled with young, late, and partly mid-Pleistocene fluvial–deluvial sediments, suggesting partial sedimentation in the streams and, in some cases, further deepening of the troughs after the formation of the mega-yardang fields. Although Pleistocene loess covers much of the mega-yardangs except for the wind-exposed sides, these challenge hypotheses proposing a late Pleistocene origin for the mega-yardangs. Accepting a late Pleistocene origin would imply simultaneous aeolian erosion and deposition processes on the ridges at the same time and place.

Paleoclimatic assessments indicate that during the M2 Pliocene cold period (lower Piacenzian), the climate of the Pannonian Basin fluctuated between woodland/shrubland and cold desert/steppe biomes (Trájer 2023). These intermittently dry climatic conditions persisted at least until the early Pleistocene, as evidenced by fossil assemblages such as that of Kisláng, Hungary. The Kisláng assemblage, which includes a mix of Pliocene–early Pleistocene taxa, contains indicators of arid and semi-arid environments, such as ancient camels and ostriches (Jánossy 2011; Zelenkov et al. 2019). Cosmogenic dating of ventifact surfaces by Sebe et al. (2011) suggests a significant wind erosion event dating back to the Early Pleistocene, approximately 1.5 million years ago. However, this age might instead reflect a period of bedrock exhumation previously concealed by late Miocene formations. Reconstruction of denudation rates in the Tapolca Basin, Balaton Highland, suggests that around 3–4 million years ago, the mean surface elevation could have ranged between 260 and 300 m above sea level (a.s.l.) (Ruszkiczay-Rüdiger et al. 2011). This finding is consistent with the approximately 300-m-high backs of better-preserved mega-yardangs (Sebe et al. 2011), implying these surfaces may be remnants of the post-Pannonian fluvial plain. The presence of early Pleistocene–age fluvial gravels found at relatively high elevations above sea level, overlaying Tortonian strata with sedimentary hiatus, is well-documented beneath late Pleistocene loess layers in the broader area of the Balaton Basin (Bada et al. 2010).

The pre-middle Pleistocene age of the mega-yardangs is further supported by the age of river terraces formed between them. Paleohydrological reconstructions suggest that ancient rivers to the east and southeast of the present-day Balaton Basin followed the direction of the mega-yardangs, possibly as early as the late Pliocene–lowermost Pleistocene, but certainly from the late early Pleistocene onwards (Gábris and Mari 2007). This hydrological system pattern persisted until about 140 kya (Ruszkiczay-Rüdiger et al. 2020), when the rising uplift of the Transdanubian Range altered the landscape, leading to events such as the beheading of the Zala River (Gábris and Mari 2007). In addition, optically stimulated luminescence (OSL) dating, as indicated by Thamó-Bozsó et al. (2010), reveals evidence of relatively recent neotectonic activity occurring between 49 and 9 kya in and around the Balaton Basin. This activity modified the geomorphology of pre-existing mega-yardangs within the North Transdanubian Hills region and played a role in the formation of Lake Balaton (Visnovitz et al. 2021). As the formation of yardangs had to precede the creation of river valleys and Lake Balaton, it is inferred that the age of the yardangs predates the late Pleistocene. Overall, it can be concluded that the Pannonian mega-yardang system is older than previously assumed, and its formation age can be placed at the earliest, in the middle Pliocene, and, at the latest, by the end of the early Pleistocene.

The yardangs-determined relief-based development of the Transdanubian hydrological system also indicates the relatively pre-late Pleistocene origin of the mega-yardangs. For example, the Zala River, which flows into Lake Balaton today, has five terraces between the wide ridges of the surrounding mega-yardangs. The age of the oldest, 5th terrace dates to the MIS16–12 period, and the younger, 4th terrace has MIS12 age (Gábris and Mari 2007). Moreover, a similar pre-last glacial age of the yardangs and large deflation basins can be assumed based on this. The syngenesis of the mega-yardangs and the large deflation basins is supported by the observation that post-Pannonian and pre-MIS2 sediments are missing in the floor of Lake Balaton, and under the Holocene and late Pleistocene young sediments, there is no sign of any yardang-like structures (Zlinszky et al. 2010).

Reconstruction of the development of the Pannonian mega-yardang system

Based on the above-mentioned facts, by integrating the age, geology, and climate background of the mega-yardang system, the following development history can be reconstructed:

  1. 1)

    During the Tortonian epoch, a large, endorheic basin housed the Pannonian Lake. Sediments with varying degrees of cementation and hardness were deposited in this lake (Fig. 13a).

  2. 2)

    In the Messinian period, the Pannonian Lake receded from the area of the later mega-yardangs. Initially, deltaic sedimentation occurred, followed by fluvial–lacustrine sedimentation on the relatively flat plain. These processes occurred under temperate but variably humid, subhumid climatic conditions (Fig. 13b).

  3. 3)

    Between the mid-Pliocene and the mid-Pleistocene epochs, transiently arid or semi-arid conditions prevailed in the western part of the Pannonian Basin. During this time, wind erosion was notably strong, leading to the formation of yardangs on the desiccated plain of the former Pannonian Lake (Fig. 13c).

  4. 4)

    From about the Middle Pleistocene onwards, the climate gradually cooled and became more humid. Concurrently, tectonic processes, triggered by the northward shift and clockwise rotation of the Adriatic Tectonic Plate, initiated the slow uplift of the area. Rivers, following pre-existing depressions, carved valleys into the yardangs, forming an almost parallel water network. During interglacial periods, the climate transitioned to oceanic or humid continental conditions (Fig. 13d).

  5. 5)

    In the cold phases of the latter half of the Pleistocene, under subpolar climate conditions, loess accumulated over the area, covering surfaces primarily composed of eroded late Miocene sediments (Fig. 13e).

  6. 6)

    In the Holocene epoch, humid continental conditions returned, and following the Neolithic period, anthropogenic modifications began to impact the landscape of the Pannonian mega-yardangs increasingly. Ultimately, natural deciduous forest vegetation, destroyed by agriculture, may have given way to intensified wind erosion (Fig. 13f).

Fig. 13
figure 13

Model depicting landscape development in the region of the Pannonian mega-yardang system; a formation of the Pannonian Lake during the Tortonian, b establishment of a fluvial plain during the Messinian and early Pliocene under a humid or sub-humid temperate climate, c formation of the yardangs likely during a dry phase in the mid-to-late Pliocene under cold steppe/desert climatic conditions, d fluvial erosion occurring during humid periods in the Pleistocene under either oceanic or humid continental climates, e deposition of loess during glacial phases in the Pleistocene under a subarctic climate, f contemporary anthropogenic landscape modifications under humid continental climatic conditions Dark bedrock displays the Neogene sediments, and light-yellow marks the Pleistocene loess Other Quaternary sediments were not indicated on the sub-figures

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It is noteworthy that in various regions around the world, the process of yardang formation exhibits similar developmental characteristics to those observed in the Pannonian mega-yardangs, albeit with variations in the timeline. For instance, in the Lop Nur Region of Northwestern China, palynological evidence indicates that during the early Pleistocene, the dominant vegetation comprised forests and steppe, transitioning to a desert steppe and ultimately a desert environment by the mid-Pleistocene. The formation of yardangs occurred within sediments deposited in a lake, which existed under relatively humid conditions during the early Pleistocene (Yan et al. 1997). The findings of the present study suggest a comparable geomorphological evolutionary scheme; however, it is proposed that the arid or semi-arid conditions were intermittent and likely recurred between the mid-Pliocene and early Pleistocene in the Pannonian Basin. Fluvial processes may have also played a significant role in the formation of yardangs. For instance, fluvial erosion was particularly influential in shaping the yardangs of the Kumtagh Desert in Northwest China (Dong et al. 2012), and a strong correlation can be inferred between the formation of yardangs and the presence of aeolian deposits in this region (Warren 2019). Similar conditions may be hypothesized in Hungary, with yardang formation occurring concurrently with the deposition of aeolian sediments. It is established that the Pliocene–early Pleistocene aeolian red clays in Hungary constitute wind-blown dust deposits (Varga et al. 2011), mirroring the findings in China, where analytical data have also indicated a wind-blown origin for red clay deposits of similar age and type (Lu et al. 2001). In addition, it should be noted that these red clays are widely distributed in the Pannonian Basin and have been preserved in the wind-protected geomorphological features of the karst areas of the Transdanubian Mountains (Trájer et al. 2015, 2020).

Conclusion

The results indicate that instead of forming during the Late Pleistocene, the Pannonian mega-yardang system originated in the mid-Pliocene to the early (or mid-) Pleistocene interval, situated on the sediments of Lake Pannon and the fluvial sediments of post-Tortonian braided rivers. In its current state, the mega-yardang system exhibits a mixed, fluvio-aeolian origin. The development of the hydrological system adapted to the large-scale aeolian geomorphological structures, with rivers further deepening the troughs during humid periods. In addition, large deflation basins, such as the Balaton Basin, are integral parts of the aeolian complex, forming syngenetically with the yardangs and the buttes. Importantly, these results do not contradict the formation of fluvial-origin sand dune fields during the Late Quaternary epoch.