Introduction

An interest in understanding the coevolution of palaeoclimatic conditions and palaeovegetation (e.g. Genise et al. 2022; Salazar-Jaramillo et al. 2022; Develle et al. 2011) as well as past pedogenetic (Hobbs and Fawcett 2022) and weathering (Ekoa Bessa et al. 2021) processes, including karstification (Hartmann and Baker 2017), is growing worldwide. The reconstruction of paleoclimatic conditions and palaeoweathering processes given the tectonic setting form the integrant parts of diagenetic history (Tawfik et al. 2018). Not only is a better understanding of these processes important for the reconstruction of periods of geological history, but may also be of economic significance. For example, unravelling former erosion processes facilitates the exploration of petroleum systems in the Barents Sea (Lasabuda et al. 2018), provides an explanation for the formation of Paleogenic and Neogenic kaolinitic sedimentary rocks in North East Scotland (Hall et al. 2015) as well as accounts for the presence of the karst bauxites in Central Dalmatia, Croatia during the Middle Miocene (Brlek et al. 2021).

Knowledge of paleoclimatic conditions plays a key role in understanding the formation of terrestrial sedimentary raw materials. Fortunately, the reconstruction of Cenozoic palaeoclimates has brought about significant methodological progress over recent decades (Grimm and Potts 2016). In Eurasia, a relatively large number of fossil flora sites have made it possible to build palaeoclimatic models (Bruch et al. 2011; Utescher et al. 2011; Li et al. 2018). However, several global models of the palaeovegetation and palaeoclimate during the Late Paleogene and Neogene have also been created (Henrot et al. 2017; Tardif et al. 2021). The statistical analysis of the palaeovegetation related to the current climatic constraints of the closest living relatives of the fossil plant taxa have enabled the climate to be reconstructed by quantifying its values (e.g., Postigo-Mijarra et al. 2022; Mosbrugger and Utescher 1997). Furthermore, advances in the palaeoclimatology with regard to the oxygen isotope composition has been key to reconstructing changes to the terrestrial palaeoclimate (Huth et al. 2022) and karst development (White 2007). In addition to the fossil plants and palaeoclimatic data provided by foraminifera, minerals indicative of former intense weathering have received special attention because they provide detailed information about the palaeoenvironment (Raphalalani 2017) and are useful as paleoclimate proxies (Dianto et al. 2019; Chanvry et al. 2020; Laita et al. 2020).

No consensus has been reached on the Cenozoic development of haematitic and kaolinitic sediment-forming processes in the Pannonian Basin of Central Europe (Fig. 1a). Questions, for example, about the role of the re-deposition of older materials, the intensity of weathering as well as the effect of former climatic conditions in terms of the formation of red clays and kaolinites remain unanswered because such weathering products like the minerals goethite, haematite and kaolinite can either be degradation products of old bauxites or the results of in situ weathering processes. In fact, some Pannonian haematite-rich formations contain large amounts of re-sedimented bauxitic materials from older, local Cretaceous-Eocene bauxite deposits (Kelemen et al. 2017). Due to these uncertainties, different authors have published various palaeoenvironmental interpretations related to the formation, properties and distribution of Cenozoic sedimentary rocks as well as the related palaeomorphological features in the Pannonian Basin over recent decades (e.g., Kovács et al. 2013, 2015; Ruszkiczay-Rüdiger et al. 2009; Fekete and Stefanovits 2002A).

Fig. 1
figure 1

a, b Sites mentioned in the text (1: Cserszegtomaj, 2: Balatonfüred, 3: Lovas, 4: Meggyespuszta palaeodoline between Veszprém and Szentkirályszabadja, 5: Southern Institution Centre of Veszprém, 6: Veszprém-Hajmáskér, 7: Öskü, 8: Várpalota, 9: Réde, 10: Bodajk, 11: Gánt, 12: Mezőszilas, 13: Csarnóta, 14: Dunaföldvár, 15: Fülöpháza, 16: Répáshuta. A: Austria, BIH: Bosnia and Herzegovina, CZ: Czech Republic, H: Hungary, HR: Croatia, PL: Poland, RO: Romania, SK: Slovakia, SLO: Slovenia, UA: Ukraine); c, d: the Köppen-Geiger climate categories of East and Southeast China and the Southeastern United States based on Beck et al. (2018) (Af: tropical rainforest climate, Am: tropical monsoon climate, Aw/As: tropical savanna climate, BWh: hot desert climate, BWk: cold desert climate, BSh: hot semi-arid climate, BSk: cold semi-arid climate, Cwa: monsoon-influenced humid subtropical climate, Cwb: subtropical highland climate, Cfa: humid subtropical climate, Cfb: temperate oceanic climate, Dwa: monsoon-influenced hot-summer humid continental climate, Dwb: monsoon-influenced warm-summer humid continental climate, Dwc: monsoon-influenced subarctic climate, Dfa: hot-summer humid continental climate, Dfb: warm-summer humid continental climate)

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In connection with several Upper Paleogene and Neogene formations rich in kaolinites and red clays in the Pannonian Basin, an earlier general assumption was that these sediments can be derived from the redeposition or heterotopic facies of the older, Eocene-aged bauxitic formations (Szentes 1957; Csillag 1959; Bárdossy 1961). In accordance with this hypothesis, some paedomorphic formations like the Óbarok and Vöröstó Formations are considered to be partly allogenic and, regarding the formation of their originating materials, allochronic (Kelemen et al. 2017). In contrast, Fekete and Stefanovits (2002B) reported that most red clays are found in old, in situ formations. They pointed out that although the bauxite in the Transdanubian Mountains formed on the surface of Mesozoic limestone and dolomite plateaus in the Eocene epoch, red clays can also be found in the upper layers of eroded bauxite formations, suggesting that their origin is more recent.

The local formation of haematite-rich deposits is also supported by the U–Pb dating and heavy mineral composition of red clays, indicating that haematite and kaolinite in the sedimentary rocks could be the local weathering products of airborne volcanic material (Kelemen et al. 2017). The volcanic-origin starting material of the weathering products was produced by a series of violent eruptions in the Pannonian Basin between 42.4 and14.0 Ma (Benedek et al. 2004; Harangi and Lenkey 2007; Bergomi et al. 2015; Danišík et al. 2015; Lukács et al. 2015), Based on similarities between rare earth elements, the volcanoes situated in the Carpathian Basin and the Periadriatic region could be the source of the volcanic ash (Kelemen et al. 2017). More than 4,000 cubic kilometers of silica material was ejected during eruptions in the Bükkalja Volcanic Field (Southern Bükk Mountains, Fig. 1a), Northern Hungary (Brlek et al. 2023), which was one of many active Neogene volcanic fields in the Carpathian Basin over the short time span of 18.2–14.4 Ma. However, because the chemophysical signals of the minerals found in deposits of red clay and kaolinite might be the result of previous weathering regimes (Kelemen et al. 2017), they cannot be used to reconstruct the former ecological conditions at the time of sedimentation.

Objectives

Based on the aforementioned results from the literature, although many authors have examined the relationship between Cenozoic palaeoclimates and different environmental processes in the Pannonian Basin, the combined, interrelated examination and temporal development of the biotic, climatic, chemical and topographical factors in an area remain research gaps. The main objective of this study is to investigate the environmental development of the post-Eocene Pannonian Basin (Fig. 1a–b), including the nature and effect that shaped the land surface as a result of the Tertiary and Quaternary weathering processes over the last 34 mys.

To achieve this goal, the study aimed to characterize the former biomes, as well as the paedogenic, weathering and karstification processes and patterns, in addition to geomorphological features linked to mean annual temperatures which were compared to East and Southeast China along with the South-eastern United States as analogues (Fig. 1c–d).

Geological setting

The geological record of the Pannonian Basin (Fig. 1a–b) contains several well-investigated Paleogene, Neogene and Quaternary geological strata, providing a basis for the reconstruction of long-term palaeoenvironmental changes in this region. The Pannonian Basin was formed by the coalescence of the Pelso unit of African origin (Szederkényi et al. 2012) and the Tisza unit of European origin when they shifted to their present positions in the Late Paleogene and Early Neogene (Palotai 2013). Prior to the formation of the basin, thick strata of bauxite (Gánt Bauxite Formation) were deposited on the karstic surface of the terrain (Mindszenty 2010; Fig. 1a). However, these bauxitic layers only formed until the Upper Ypresian (Császár 1997) when, following the end of the Early Eocene Climatic Optimum (Lowe et al. 2018), the climatic preconditions for the formation of thick lateritic soils may have no longer been met. However, the climates of the Pelso unit remained warm, even hot, as far as the Middle Miocene, leading to the formation of tropical karst.

Based on the spatial patterns of the red clay sequence that occurred in the Vincepál section of the Lower Miocene Felsőnyárád Formation in Répáshuta (Southeastern Bükk Mountains), Hevesi and Mádai (2007) proposed that the former geomorphological character of the Bükk Mountains (Fig. 1a), consisting mainly of Triassic and Jurassic carbonates, could be very similar to the cone karst regions of South China. Other geological evidence also suggests that subtropical and tropical karst landforms may have characterized the carbonate rock areas of the Pannonian Basin during the Tertiary period. For instance, originally deep but severely eroded, fossilized sinkholes are found in the Keszthely Mountains (Transdanubian Mountains; Fig. 1b) filled with Middle Miocene clay-rich and haematitic sedimentary rocks (Budai et al. 1999). Furthermore, an 800 m-wide, exhumed palaeodoline was found between Veszprém and Szentkirályszabadja in the Balaton Highlands (Fig. 1b) filled with dolomite flour and red clay sediments (Fig. 2a) (Trájer et al. 2015). Exhumed, cone-shaped karst features and relict, wide palaeovalleys of Tertiary origin have also been uncovered in the region (Fig. 2b–c).

Fig. 2
figure 2

Oligocene and Miocene relict geomorphological features and chemical weathering product sediments. Exhumed Tertiary relict karst surfaces: a1 view of the southern half of the Meggyespuszta palaeodoline; a2 in Veszprém-Szentkirályszabadja and its topographical model; b the bedrock is formed by Anisian-Ladinian carbonate formations; c Karst cones in Öskü and an old stream valley in Kádárta, Veszprém-Hajmáskér; d the bedrock in both cases is the Ladinian Budaörs Dolomite Formation. Different post-Eocene bauxite and clay sediments: Rupelian-age Oligocene bauxitic sediment fill (Óbarok Bauxite Formation) in a fossilized karst cavity in Bodajk, Kajmát quarry and e its field occurrence in Várpalota; the bedrock in both cases is the of the late Carnian-Norian Main Dolomite Formation; f Calcareous kaolin deposition (light green layer) between the sandy, gravelly Chattian-age Csatka Formation in Réde and g the middle Miocene-age Cserszegtomaji Kaolinite Formation (white heaps on the quarry slope) next to the fossilized Koponár sinkhole (Cserszegtomaj); h The outcrop of the bauxitic middle-Miocene Vöröstó Formation in Szentkirályszabadja; the bedrock is the early Anisian Iszkahegy Limestone Formation; i Bauxitic karst cavity-filling sediment in the Main Dolomite Formation; the covering sediment is upper Pleistocene loess, the Southern Institution Centre of Veszprém

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As a result of the intensive Late Paleogene and Neogene chemical weathering processes, haematite- or kaolinite-rich formations are frequent in the Transdanubian Mountains. The oldest Oligocenic palaeosols are karstic depression-infilling red clays containing high amount of bauxite debris and bauxitic pebbles (Óbarok Bauxite Formation). These sediments may be the products of the Early Oligocene denudation of the Transdanubian terrain (Fig. 2d–e). The Csatka Formation, originating from Chattian molasse, contains kaolin-rich deposits (Fig. 2f) (Budai et al. 2015). Similar kaolinite-rich soils currently form under warm and humid subtropical climatic conditions (Bortoluzzi et al. 2008). The Cserszegtomaj Kaolinite Formation (Fig. 2g) which is rich in halloysite and kaolinite, was formed during the Middle Badenian-Sarmatian (Lower–Upper Burdigalian to the Middle Serravallian) and accumulated in former deep sinkholes in the Keszthely Mountains (Budai et al. 1999). The Vöröstó Clay Formation of similar age is a red clay matrix-supported conglomerate (Budai and Csillag 1998; Budai et al. 1999) which contains embedded bauxite pebbles, angular bauxite, dolomite debris as well as grains of sand and gravel. Moreover, it is usually rich in kaolinite. The Vöröstó Clay Formation is located in an area covered by Neogene layers in the depressions of the Triassic terrain and at some sites, it appears in the form of crack-filling sediment (Fig. 2h). In the Badenian stage (Lower–Upper Burdigalian to the Middle Serravallian), similar blankets of red clay were formed on Miocene basalts in Central Germany (Schwarz 1997). Other red clay strata consisting of karst cavities and infillings can be found under the sediments of the last glacial-interglacial cycle at several localities in the Transdanubian Mountains (Fig. 2i).

Haematite-rich chemical weathering products are also known from geological periods other than the Paleogene or Early and Middle Neogene. Fekete and Stefanovits (2002A-B) identified a young group of red clays, which can be found, for example, in Atkár, Northern Hungary (Kovács et al. 2013) as well as in the karst cavities of the Mecsek and Villány Hills (Southwest Hungary; Fig. 1a), for instance, in Csarnóta (Fig. 3a). These sediments can be the remains of fossilized terra rossa-like soils formed in a Mediterranean climate during the Plio-Pleistocene era. Red clays and kaolinitic deposits are not the only signs of chemical weathering processes from the Neogene and Early Quaternary periods. Before the Upper Quaternary, the karst plateaus of the Transdanubian Mountains were covered by a thick layer of dolomite debris and fine-grained dolomite flour (Trájer et al. 2015), which covered and preserved Tertiary karst landforms. Although currently, the intensive pulverization of the dolomite as a result of chemical weathering is not characteristic due to the relatively cool continental climate in the Pannonian Basin, under the humid and warmer climate of the Tertiary, the chemical weathering was much more intensive than observed today. A similar, intensive chemical weathering of dolomite currently can be observed, e.g., on the Yuntai Mountains in Guizhou, China (Liu et al. 2015).

Fig. 3
figure 3

Pliocene and Quaternary geomorphological features and sediments; a middle Pliocene-age red clay infilling (Tengelic Red Clay Formation) in a karst cavity of the middle Triassic Lapis Limestone Formation, Csarnóta. b The lateral view of a Pleistocene mega-yardang from the west in Mezőszilas. c Loess wall with two mid-Pleistocene palaeosol horizons (reddish-brown layers) in Dunaföldvár. d The palaeosol-loess border between the Last Interglacial (130 ka; lower reddish-brown layer) and the Last Glacial periods’ (21 ka; upper light brown layer) sequences in Lovas, Királykút valley (Paks Loess Formation). e firm bedrock and periglacial rock flow in Balatonfüred, Öreg peak, produced by the mechanical weathering of the middle Triassic Füred Limestone Formation. f mobile dune in Fülöpháza, built up upper Pleistocene-Holocene aeolian sand

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Aeolian processes and other forms of mechanical erosion became increasingly important with regard to the formation of the former landscape of the Pannonian Basin in the Pliocene and the Quaternary. In the western part of the basin, ventifacts and large-scale features resulting from deflation, including mega-yardangs (Fig. 3b) were formed (Sebe et al. 2011). The strong northwesterly winds eroded the thick layer of dolomite flour, which covered the Tertiary relict karst relief forms, exhuming the fossilized karst forms. This aeolian denudation resulted in a deflation rate of 40–80 m mys-1 (Ruszkiczay-Rüdiger et al. 2011). The former intensive aeolian processes may also explain why only sediments from the youngest glacial-interglacial cycle cover the old pre-Quaternary bedrocks in the Transdanubian Mountains (see Fig. 3j).

However, the formation of haematite-rich clayey sediments did not come to an end in the Early Pleistocene. Red clay and reddish-brown paleosol horizons were found in the early loess strata (Fig. 3c–d), indicating a significantly warmer and more humid climate during the interglacial periods than today. The gradual shift from the Early Pleistocene terra rossa-like palaeosols to the forest-steppe, steppe soils of the Late Pleistocene interglacials suggests a continuous decrease in the extent of chemical erosion processes over the last 2.58 mys (Varga 2016). In colder glacial periods, rock flows (Fig. 3e) were formed on the steeper foothills of the middle mountainous regions, indicating expressed, frost-induced mechanical weathering processes, including cryoturbation (Schweitzer 2015). The aeolian denudation of the Transdanubian Middle Mountain Range might last until the end of the Pleistocene or even into the Early Holocene. This observation supports that the sediments at the foothills of the Transdanubian Mountains, the karst plateaus of the Transdanubian Hills as well as the dunes of the Danube-Tisza Interfluve belong to the same geochemical region (Fügedi et al. 2014). It should be noted that the dunes of the Danube-Tisza Interfluve contain Upper Pleistocene-Holocene aeolian sand materials (Gębris et al. 2002) (Fig. 3f) and the majority of the volume of sand located in the sand dunes is of fluvial origin (Nyári and Kiss 2009). The aforementioned sites are shown in Fig. 1a, b.

Materials and methods

Climatic data and models

The analysis of the former environmental conditions was based on two types of paleoclimatic data sources. The first group consists of site-related climatic data. In this case, the Coexistence Approach (Mosbrugger and Utescher 1997) provided the theoretical background behind the palaeoclimatic reconstructions. This method utilizes the fossil record of plants to quantify the climate by taking into consideration the typical range of the climatic limits of the closest living relatives of the former species. The palaeoflora-based data represent the Rupelian–Zanclean time interval and consist of the reconstructed palaeoclimatic data (Suppl. Table 1) of 22 Hungarian palaeoflora-bearing fossil sites (pink points in Suppl. Fig. 1). The Tortonian palaeoflora-related palaeoclimatic data were based on the study of Bruch et al. (2006). The Oligocene, most of the Miocene and the Lower Pliocene climatic data were based on a paper by Erdei et al. (2007).

The second kind of palaeoclimatic dataset which covers the Piacenzian-Anthropocene time interval was based on existing climatic models. Climatic models related to two Pliocene (mid-Pliocene M2 cold and warm periods: 3.300 and 3.205 Ma; Dolan et al. 2015; Hill 2015), six Pleistocene (MIS-19 interglacial, Last Interglacial, Last Glacial Maximum, Heinrich Stadial 1, Bølling–Allerød Interstadial, Younger Dryas Stadial: 787, 130, 21, 17.0–14.7, 14.7–12.9 and 12.9–11.7 Ka; Brown et al. 2018; Otto-Bliesner et al. 2006; Fordham et al. 2017; Karger et al. 2017, 2021) and four Holocene periods (Greenlandian, Northgrippian, Meghalayan, Anthropocene: 11.7–8.326, 8.326–4.2, 4.2–0.3 Ka 1979–2003; Fordham et al. 2017; Karger et al. 2017) were used for modelling purposes (Suppl. Tab. 2). The applied palaeoclimatic models originated from the PaleoClim.org website (Brown et al. 2018). Three cities, namely Keszthely, Budapest and Miskolc, were selected to compare the data of the point-like palaeoflora-based reconstructions, moreover, the georeferenced climatic models represent the climate of the Transdanubian Middle Mountain Range (denoted by white rectangles in Suppl. Fig. 1).

Environmental classifications

The nature of weathering was characterized using the Peltier weathering diagram (Peltier 1950). On the diagram, the x-axis denotes the average temperature, while the y-axis represents the annual precipitation sum. The intensity and nature of weathering are strongly related to pedogenesis because the degree of chemical weathering in soils increases as both the mean annual temperature and annual precipitation sum rise (Sheldon et al. 2002). Because Peltier (1950) did not incorporate the effects of biological processes into his weathering model, the terrestrial biome model by Whittaker (Deshmukh and Singh 2019) was added to the study. For this purpose, the revised version produced by Stefan and Levin (2020) was used. Since this plot is also based on the mean annual temperature and annual precipitation sum, it can be accurately compared with Peltier’s diagram (Peltier 1950).

The possible karstification type of an area based on thermal criteria was also added to the study to supplement the information on weathering processes and their landscape-shaping effect. Based on the approach of Baker et al. (2019), the global karst regions were classified into three types depending on the mean annual temperature:

  1. 1.

    Temperate (continental) and subpolar karst regions are located in areas where:

    $$MAT<10^\circ C$$
    (1)
  2. 2.

    Subtropical karstification processes are characteristic of a region if:

    $$10^\circ C\le MAT<16^\circ C$$
    (2)
  3. 3.

    Tropical karst regions can be found in those regions where:

    $$16^\circ C\le MAT$$
    (3)

    where MAT denotes the mean annual temperature (°C).

The modelled karst regions were superimposed on the karst aquifer map of Europe (Chen et al. 2017). In the case of the Tortonian model, the karst aquifers of Hungary were georeferenced based on the geological map of the Mining and Geological Survey of Hungary (MBFSZ 2023).

Two methods were used to classify the former climates and potential types of soils. In the first case, the earth's surface was sampled at 20,000 random points. The sampled sites were characterized by the Köppen–Geiger Climate Classification System (Köppen 1884; Geiger 1954) and the United States Department of Agriculture (USDA) Soil Classification System. The Köppen–Geiger climate classification maps used were published by Beck et al. (2018) and the global soil regions retrieved from the home page of USDA (USDA-NRCS 1997). The applied actualistic approach (Ribeiro et al. 2021) assumes that in similar climates, similar vegetation as well as weathering and paedogenic processes occurred in the past to what are observed to this day. Since large humid subtropical regions are not found in the western areas of contemporary Eurasia, the Mediterranean regions were not included in the study. This decision is also confirmed by the fact that the Late Paleogene-Neogene distribution in addition to the extension of the seas and drylands in Europe were marginally different compared to what is currently experienced on the continent. From the Eocene to the Late Miocene epoch, the Paratethys Sea formed a large waterbody between the central and eastern parts of continental and Mediterranean Europe in a north–south direction. From west-south, this subtropical sea extended from Central Europe to Central Asia (Popov et al. 2004). Therefore, the Paratethys Sea could influence the climate of Western Eurasia, noticeably reducing the continentality of the neighbouring regions as was shown, e.g., in the case of the coastal regions of the Central Paratethys (Kováč et al. 2017). Since the amount of precipitation in the Paleogene and Neogene generally significantly exceeds the present-day values at the same localities in Europe (Erdei et al. 2007), it cannot be excluded that Europe experienced a monsoon back then.

In the second case, two model regions were selected where large subtropical areas are found to this day in the Northern Hemisphere to identify climatically analogous areas, namely the Southeastern United States and the eastern regions of China. The Southeastern United States represents a subtropical region that is not considered to be influenced by monsoons. In contrast, the eastern regions of China are affected by Asian monsoons (Wang et al. 2018). The present-day, climatically analogous regions were identified using the logic of Boolean algebra. Deterministic unit step functions can be written in the following form using the related upper and lower extrema:

$$1\left(cf\right)=\left\{\begin{array}{c}0\,\,\, if \, {cf}_{upper extrema}<cf \,and \,{cf<cf}_{lower extrema} \\ 1\,\,\,\, if\, cf\le {bio}_{upper extrema}\, and\, {cf}_{lower extrema}\le cf\end{array}\right.$$
(4)

The used climatic factors (cf) related to the temperature were as follows: the mean annual temperature in January (MAT; °C), the January mean temperature (Tm01; °C) and the July mean temperature (Tm07; °C). The annual precipitation sum (MAP; mm) represents climatic factors related to the amount of precipitation. Since the lower and upper extrema related to the factors were used, 2 × 4 climatic limits were applied in each case.

The areas excluded by the bioclimatic limiting factors can be totalled and the intersection of the potential area designated by the factors yields the aggregated distribution area as follows:

$$A\left({fT}_{upper;lower};\,{fP}_{upper;lower}\right)=fT-\sum_{i=1}^{3}0({bioT}_{upper;lower})\bigcap fP-{fP}_{upper;lower}$$
(5)

where A(fTupper;lower;fPupper;lower) shows the climatically analogous areas, which contain the remaining areas after taking into consideration limitations concerning the temperature (fT) or precipitation-related (fP) climatic factors. Due to the uncertainties of the palaeoclimate reconstructions, such current areas were the present-day climatic analogy of the past locality if 7 of the 8 limits are within the required climatic constraints. Since current analogous regions were not found in some instances based on the reconstructed palaeoclimatic values, these sites were omitted from the analysis. For statistical purposes, the statistical software PAST v.4.03 was used (Hammer et al. 2001).

Biome modelling

The biome modelling was based on the approach published by Trájer (2022a, b). To reconstruct the former biome patterns, the author created the area-delimiting curves of the revised Whittaker biome plot (Stefan and Levin 2020; Deshmukh and Singh 2019). For modelling purposes, the equations of the upper or lower delimiting curves related to the biome areas were applied. The selected biomes were as follows: 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 tundra) biomes. The equation of the upper delimiting curve of the temperate grassland/cold desert biome is as follows:

$$-0.3498\times {MAT}^{2}+24.436\times MAT+159.92=MAP$$
(6)

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

$$-0.0086\times {MAT}^{4}+0.2955\times {MAT}^{3}-2.1072\times {MAT}^{2}+41.12\times MAT+242.93=MAP$$
(7)

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

$$2235.23+(47.91482-2235.23)/(1+{\left(\frac{MAT}{7.782309}\right)}^{6.265971}=MAP\, and\, if\, MAT>-2.5^\circ C \,and\, MAT<22.5^\circ C$$
(8)

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

$$4.8254\times {MAT}^{3}-77.874\times {MAT}^{2}+519.17\times MAT+318.67=MAP$$
(9)

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

$$4.8254\times {MAT}^{3}-77.874\times {MAT}^{2}+519.17\times MAT+318.67=MAP$$
(10)

where MAT denotes the mean annual temperature (°C) and MAP the annual precipitation sum (mm).

The use of δ18O data in the modelling of palaeotemperatures

To characterize the karstification regimes over the last 34 Mys, the palaeothermal conditions had to be reconstructed. For this purpose, a correlation between δ18O and the reconstructed mean annual temperature values was identified. Given that the δ18O values of the Atlantic benthic foraminifera are indicative of both the water and global temperatures (Mix et al. 1985), in terms of applying past regional temperature trends, the data series are suitable for compiling local reconstructions (Trájer 2022a, b). The δ18O data were compiled by Lisiecki and Raymo (2005), Tiedemann et al. (1994) and Zimmerman et al. (1984). The primary data related to the correlation between the δ18O and reconstructed mean annual temperature values are presented in Suppl. Table 3.

Results

Former biome patterns

In the Lower Oligocene epoch (regional Kiscellian stage), temperate seasonal forest was the characteristic biome of the Pannonian Basin. During the Upper Oligocene (regional Egerian stage), temperate seasonal forest or a transitional biome between this one and the woodland/shrubland (forest-steppe) biomes could have been the characteristic biome, depending on its locality. Parallel to the cooling and aridification trends, forest density continued to decrease from the Upper Oligocene to the Late Miocene. From the Lower Miocene (regional Eggenburgian and Karpatian stages) to the Middle Miocene (Sarmatian regional stage), the biomes in the Pannonian Basin could be transitional between the temperate seasonal forest and woodland/shrubland biomes. In the Tortonian, when the Pannonian Lake was most extensive, the temperate seasonal forest biome once more became characteristic of the Pannonian Basin. During the Pliocene, the climate became drier and colder again. The distinct biomes varied between temperate seasonal forest and woodland/shrubland biomes in the warmer stages (e.g., in the Early Pliocene and the warmest periods of the mid-Pliocene). However, in the colder periods, namely during the M2 mid-Pliocene glaciation, the biome approached the temperate grassland/cold desert biome, indicating that in the coldest and driest periods of the Pliocene, steppe vegetation could have covered the Pannonian Basin. During the interglacial periods of the Quaternary Period, temperate (deciduous) seasonal forests (e.g., in the Early Holocene), the transitional biome of the temperate seasonal forest and the woodland/shrubland biome (e.g., in the Last Interglacial or MIS 19 interglacial) covered the Pannonian Basin. Boreal forests covered the Pannonian Basin during the transitional periods between the interglacial and glacial maxima. In such glacial maxima, like the Last Glacial Maximum, even the tundra biome may have existed in the northeastern parts of the mountain ranges. It can be concluded that the reduction in temperature was usually accompanied by a decrease in the amount of precipitation during the Paleogene and Neogene. This caused an environmental shift from forested to forest-steppes or even to types of steppe vegetation. In contrast, during the Quaternary, as a result of climate change cycles that occurred roughly every 100,000 years, the dominant biome fluctuated between forest/forest-steppe and boreal forest/tundra biome in the lower elevation regions (Fig. 4).

Fig. 4
figure 4

The possible biomes between the Rupelian and the Anthropocene (model period: 1979–2003) stages in the Pannonian Basin. O1: Kiscellian (Rupelian-early Chattian) stage, O2: Egerian (mid-and late Chattian) stage, M1: Eggenburgian (late Aquitanian-lower and mid-early Burdigalian) stage, M2: Karpatian (early-upper Burdigalian) stage, M3: Badenian (lower–upper Burdigalian to the mid- Serravallian) stage, M4: Sarmatian (upper Serravallian) stage, M5: Pannonian (Tortonian) stage, P1: mid-Zanclean stage, P2: M2 Pliocene cold period, lower Piacenzian, P3: mid-Pliocene warm period, lower Piacenzian, Q1: MIS19 interglacial, Middle Pleistocene, Q2: Last Interglacial period, Middle Pleistocene, Q3: Last Glacial Maximum (21 ka), late Pleistocene, Q4: Heinrich Stadial 1, late Pleistocene, Q5: Bølling-Allerød Interstadial, late Pleistocene, Q6: Younger Dryas Stadial, late Pleistocene, Q7: Greenlandian, early Holocene, Q8: Northgrippian, mid-Holocene, Q9: Meghalayan, late Holocene and Q10: Anthropocene (model period: 1979–2013). Tu: tundra, BoFo: Boreal Forest, TeSeFo: temperate seasonal forest, Wo/Sh: woodland or shrubland, TeGr/CoDe: temperate grassland or cold desert, TeRaFo: temperate rainforest, TrRaFo: tropical rainforest, TrSeFo/Sa: tropical seasonal forest or savanna, SuDe: subtropical desert

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Palaeoclimates

The reconstructed mean annual temperature and precipitation sum values of the Oligocene and Miocene, as well as the warmer Pliocene-aged sites were similar to what can be measured in the humid subtropical without dry season (Cfa), marine west coast temperate oceanic (Cfb), dry winter humid subtropical (Cwa) and subtropical highland (Cwb) climate regions of the world today. During the mid-Pliocene cold period, the Pannonian Basin had a cold, semi-arid steppe (BSk) climate. The Quaternary interglacial temperature-precipitation data pairs of points fit to the point clouds of temperate oceanic (Cfb), subtropical highland (Cwb), hot-summer humid continental (Dfa) and the dry-winter, hot-summer humid continental (Dwa) climates. The climate of the glacial and transitional periods (between the glacial-interglacial cycles) could be like the current dry-winter, warm-summer humid continental (Dwb), dry-winter regular boreal (subarctic) (Dwc), humid regular boreal (subarctic) (Dfc) and tundra (ET) climates (Fig. 5).

Fig. 5
figure 5

The fitting of past climate values to the temperature-precipitation points to clouds of the present-day Köppen-Geiger climate classes related to similar climatic values (Af: tropical rainforest climate, Am: tropical monsoon climate, Aw/As: tropical savanna climate, BSh: hot semi-arid climate, BSk: cold semi-arid climate, BWk: cold desert climate, Cfa: humid subtropical climate, Cfb: temperate oceanic climate, Cwa: monsoon-influenced humid subtropical climate, Cwb: subtropical highland climate, Dfa: hot-summer humid continental climate, Dwa: monsoon-influenced hot-summer humid continental climate, Dwb: monsoon-influenced warm-summer humid continental climate, Dwc: monsoon-influenced subarctic climate, Dfc: subarctic climate, ET: tundra climate)

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Monsoon-influenced subtropical regional analogies of the Late Paleogene to mid-Pliocene.

Considering the reconstructed mean annual temperature and annual precipitation values of the Rupelian–Zanclean-aged sites, it can be concluded that during the Late Paleogene and Neogene, the climate of the Pannonian Basin could have been like the present-day humid subtropical, monsoon-influenced humid subtropical and subtropical highland climate regions of Southeast China. Although monsoon-influenced hot-summer humid continental regions also appear on the maps prior to the Zanclean, a clear cooling tendency can only be seen from the Late Pliocene. From this time, the extension of the subtropical highland and monsoon-influenced hot-summer humid continental regions exceeds the area of the humid subtropical territories (Fig. 6).

Fig. 6
figure 6

Köppen-Geiger climatic categories in the areas of present-day East and Southeast China that are climatically analogous to the Oligocene-Pliocene Pannonian Basin according to the reconstructed mean annual temperature and annual precipitation sum values. Kiscellian (Rupelian-early Chattian) stage (a: Budapest; b: Eger-Kiseged; c: Nagybátony-Újlak), Egerian (mid-and late Chattian) stage (d: Eger, Andornaktálya; e: Eger-Wind; f: Kesztölc; g: Pomáz; h: Vértesszőlős), Eggenburgian (late Aquitanian-lower and mid-early Burdigalian) stage (i: Ipolytarnóc), Karpatian (early-upper Burdigalian) stage (j: Magyaregregy), Badenian (lower–upper Burdigalian to the mid- Serravallian) stage (k: Nógrádszakál), Sarmatian (upper Serravallian) stage (l: Erdőbénye), Pannonian (Tortonian) stage (m: Bükkábrány; n: Hidas; o: Visonta), Pliocene1: mid-Zanclean stage, q: overview map)

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Non-monsoonal subtropical regional analogies from the late Paleogene to mid-Pliocene

It can be said that during the Rupelian–Zanclean period, the climate of the Pannonian Basin could have resembled the present-day humid subtropical (Cfa) climatic regions of the Southeastern United States, apart from some minor, climatically similar hot semi-arid (BSh) and hot-summer humid continental (Dfa) climatic regions. The only clear trend-like change is the northern shift of the climatically analogous regions during the Rupelian–Zanclean time interval (Fig. 7).

Fig. 7
figure 7

Köppen-Geiger climatic categories in the present-day Southeastern United States of America are climatically analogous to the Oligocene-Pliocene Pannonian Basin according to the reconstructed mean annual temperature and annual precipitation sum values Kiscellian (Rupelian-early Chattian) stage (a: Budapest; b: Eger-Kiseged; c: Nagybátony-Újlak), Egerian (mid-and late Chattian) stage (d: Eger, Andornaktálya; e: Eger-Wind; f: Kesztölc; g: Pomáz; h: Vértesszőlős), Eggenburgian (late Aquitanian-lower and mid-early Burdigalian) stage (i: Ipolytarnóc), Karpatian (early-upper Burdigalian) stage (j: Magyaregregy), Badenian (lower–upper Burdigalian to the mid- Serravallian) stage (k: Nógrádszakál), Sarmatian (upper Serravallian) stage (l: Erdőbénye), Pannonian (Tortonian) stage (m: Bükkábrány; n: Hidas; o: Visonta), Pliocene1: mid-Zanclean stage, q: overview map)

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Mapped biome and karstification process models

In the Tortonian, the coast of the Pannonian Sea may have been covered by woodland/shrubland or temperate seasonal forest vegetation. The first biome could have been dominant in the regions of low elevation, while the second in those of middle and high elevations (Fig. 8A). Generally, although subtropical karsts were characteristic of the Pannonian archipelago, in certain parts of the present-day Balaton Highlands and Keszthely Mountains, tropical karstification processes could also be characteristic. The temperate karst zones could have only occurred in the former regions of higher elevation in the Carpathian and Apuseni Mountains (Fig. 8B).

Fig. 8
figure 8

a The Tortonian (11.608–7.246 Ma) potential biome and b karstification type patterns in the Carpathian Basin. TeGr/CoDE: temperate grassland or cold desert, Wo/Sh: woodland or shrubland, TeSeFo: temperate seasonal forest, BoFo: boreal forest, Tu + Ic: tundra and ice. Temp + SubPol: temperate (continental) and subpolar karst regions, SubTrop: subtropical karst regions, Trop: tropical karst regions

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During the mid-Pliocene glaciation event, the climate of the lowland and mid-elevation regions of Europe could have been relatively dry. The model suggests a general retreat of the forest biome at that time. In addition, the forest-steppe was replaced by grassland vegetation in the eastern part of the Pannonian Basin. Thermophilous forest steppe was the characteristic biome in the higher regions and the western part of the Pannonian Basin. In the warm period of the mid-Pliocene, the climate was somewhat more humid, and the forest-steppe biome was limited to the northern part of the Balkan Peninsula in the vicinity of the Pannonian Basin. At that time, temperate forests covered the western part of the Pannonian Basin and the regions of higher elevation in the mountainous regions of North Hungary. In the MIS 19 interglacial, the climate was drier and colder than during the previous period, resulting in the appearance of the boreal biome in the Eastern Alps as well as the Carpathian and Apuseni Mountains surrounding the Pannonian Basin. Although the make-up of the vegetation was as it is nowadays, the forest cover of the Carpathian Basin could have been lower. Temperate deciduous forests covered the regions of middle elevation in the southwestern part of the Pannonian Basin and forest-steppe biomes over a significant proportion of the lower elevation regions. The climate of the Last Glacial Period was colder and more humid than that during the MIS 19 interglacial; moreover, the potential vegetation patterns strongly resembled that of the present day. Large regions of Western and Central Europe were covered by deciduous forests, including the western part and mid-elevation regions in the eastern Pannonian Basin.

In contrast, the biome patterns of the Last Glacial Period exhibited merely different features. The vegetation of most of Europe's western, central, and eastern regions was arctic or subarctic. Boreal forests covered the Transdanubian Mountains, and tundra vegetation covered the higher regions in the North Hungarian and Carpathian Mountains (Fig. 9).

Fig. 9
figure 9

The potential biome patterns between the mid-Pliocene to the present era in Europe. a the M2 Pliocene cold period, lower Piacenzian; b the mid-Pliocene warm period, lower Piacenzian; c the MIS19 interglacial, Middle Pleistocene d the Last Interglacial period, Middle Pleistocene; e the Last Glacial Maximum, late Pleistocene; f the Anthropocene (model period: 1979–2013). TeGr/CoDE: temperate grassland or cold desert, Wo/Sh: woodland or shrubland, TeSeFo: temperate seasonal forest, BoFo: boreal forest, Tu + Ic: tundra and ice

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In the mid-Pliocene, a significant part of Western and Central Europe belonged to the subtropical karst region based on the former thermal conditions. The formation of tropical-like karsts could have persisted in Southern Europe. Large parts of the Pannonian Basin belonged to the subtropical karst zone; however, in the Southern Pannonian Basin and Northern Balkans, the thermal conditions of tropical karstification persisted. Like today, the subtropical karst zone covered a significant part of the Pannonian Basin and temperate karstification was dominant in the territory of present-day North Hungary during the MIS 19 interglacial. Like nowadays, in the Last Interglacial, most of the Pannonian Basin belonged to the temperate karst zone, and subtropical karstification was only characteristic of low-elevation regions in the Southern Pannonian Basin. Under the cool climate of the Last Glacial Maximum, most territories in Europe and the entire Pannonian Basin fell into the temperate-subpolar karst zone (Fig. 10).

Fig. 10
figure 10

The potential karstification type patterns between the mid-Pliocene to the present era in Europe. a the M2 Pliocene cold period, lower Piacenzian; b the mid-Pliocene warm period, lower Piacenzian; c the MIS19 interglacial, Middle Pleistocene d the Last Interglacial period, Middle Pleistocene; e the Last Glacial Maximum, late Pleistocene; f the Anthropocene (model period: 1979–2013). Temp + SubPol: temperate (continental) and subpolar karst regions, SubTrop: subtropical karst regions, Trop: tropical karst regions

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A strong (r2 = 0.8484) negative polynomial correlation was found between δ18O and the mean annual temperature values related to the modelled and reconstructed palaeothermal values of the Pannonian Basin (Fig. 11A). Using the obtained equation, the average palaeotemperature values over the last 34 mys were modelled for the Pannonian Basin. Based on the classification approach of the climazonal karst zone proposed by Baker et al. (2019), between the Rupelian and Late Sarmatian (Late-Upper Serravallian) ages, the Pannonian Basin generally belonged to the tropical karst zone and in the short, cooler periods to the subtropical karst one. Subsequently, until the Late Pliocene epoch, the region belonged to the global subtropical karst area. From the Late Pliocene until the present day, the temperate (continental) boreal karstification was characteristic of the Pannonian Basin. Moreover, the climate could only support subtropical karstification processes during the warmest interglacial periods in the Quaternary Period (Fig. 11B).

Fig. 11
figure 11

a Correlation between the Atlantic foraminifera δ18O and the reconstructed annual mean temperature values of the Hungarian Middle Mountain localities; b the reconstructed 20-period moving average of the temperature values in the last 34 million years in the Pannonian Basin and their relation to karstification types. Stages and abbreviations: Rupelian, Chattian, Aquitanian, Burdigalian, La.: Langhian, Serra.: Serravallian, Tor.: Tortonian, Messin.: Messinian, Zanclean, Piacen.: Piacenzian, Gelas.: Gelasian, Calab.: Calabrian, Chi.: Chibanian (0.774–0.129 Ma), L.: late Quaternary period (0.129–0 Ma)

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Changes in weathering processes

In the Lower Oligocene epoch, the nature of the weathering processes was in transition from intense to moderate chemical weathering. Throughout the Neogene, moderate chemical weathering was characteristic in the Pannonian Basin. The rate of chemical erosion tended to decrease throughout the entire period, except for during the Tortonian stage of the Late Miocene. Even though moderate chemical weathering may also have been typical during the warmer interglacial stages of the Quaternary, it was weaker than during the Neogene. In the transitional periods between the interglacial and glacial maxima, moderate chemical weathering with frost action may have been the dominant weathering process. In contrast, during the coldest glacial periods, slight mechanical weathering with significant frost action may have resulted in the freeze–thaw weathering of rocks and the cryogenic disturbance of soils in the Pannonian Basin (Fig. 12).

Fig. 12
figure 12

Weathering regimes between the Rupelian and the Anthropocene (model period: 1979–2003) stages in the Pannonian Basin. The possible biomes between the Rupelian and the Anthropocene stages (model period: 1979–2003) in the Pannonian Basin. O1: Kiscellian (Rupelian-early Chattian) stage, O2: Egerian (mid-and late Chattian) stage, M1: Eggenburgian (late Aquitanian-lower and mid-early Burdigalian) stage, M2: Karpatian (early-upper Burdigalian;) stage, M3: Badenian (lower–upper Burdigalian to the mid- Serravallian) stage, M4: Sarmatian (upper Serravallian) stage, M5: Pannonian (Tortonian) stage, P1: mid-Zanclean stage, P2: M2 Pliocene cold period, lower Piacenzian, P3: mid-Pliocene warm period, lower Piacenzian, Q1: MIS19 interglacial, Middle Pleistocene, Q2: Last Interglacial period, Middle Pleistocene, Q3: Last Glacial Maximum, late Pleistocene, Q4: Heinrich Stadial 1, late Pleistocene, Q5: Bølling-Allerød Interstadial, late Pleistocene, Q6: Younger Dryas Stadial, late Pleistocene, Q7: Greenlandian, early Holocene, Q8: Northgrippian, mid-Holocene, Q9: Meghalayan, late Holocene and Q10: Anthropocene (model period: 1979–2013). StChW: strong chemical weathering, MoChW: moderate chemical weathering, VSlW: very slow weathering, MoChW + FA: moderate chemical weathering with frost action, MoMeW: moderate mechanical weathering, SlMeW: slight mechanical weathering, StMeW: strong mechanical weathering

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The potential palaeosols

The reconstructed Oligocene-mid Pliocene mean annual temperature-precipitation sum data pairs fit the alfisol (excluding the Rupelian data pairs), andisol, entisol, inceptisol, mollisol and ultisol point clouds. In the Rupelian, the climatic conditions were at the limit of the formation of oxisols. Certain temperature-precipitation data pairs from the Oligocene and Lower Miocene also match the vertisol point cloud. The points related to the mid-Pliocene epoch and the Last Glacial Maximum partly overlap with the aridisol point cloud. The modelled temperature-precipitation data-pair points from the Quaternary Period generally fit the andisol, entisol, histosol, inceptisol, mollisol and spodosol soil types of point clouds. Furthermore, data-pair points from the glacial and transitional (between the glacial-interglacial cycles) climatic periods overlap with the gelisols point cloud (Fig. 13).

Fig. 13
figure 13

The fitting of past climate values to the temperature-precipitation points to present-day sol-type clouds is related to similar climatic values (a: ultisols, b: mollisols, c: aridisols, d: gelisols, e: entisols, f: inceptisols, g: alfisols, h: vertisols, i: oxisols, j: andisols, k: spodosols, l: histosols)

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Monsoon-influenced subtropical regional analogies of the Late Paleogene to mid-Pliocene

In East and Southeast China, regions with similar average annual temperature-precipitation sum values to palaeoclimates from the Rupelian-Zanclean in the Pannonian Basin are dominantly covered by entisols, inceptisols and ultisols. However, alfisols, mollisols and vertisols also occur in this area. Over time, the area where entisols are found in the climatically analogous regions has been expanding at the expense of that of ultisols (Fig. 14).

Fig. 14
figure 14

Soil orders in the present-day East and Southeast China are climatically analogous to the Oligocene-Pliocene Pannonian Basin. The classification was based on the Köppen-Geiger climatic categories according to the reconstructed mean annual temperature and annual precipitation values (a: Budapest; b: Eger-Kiseged; c: Nagybátony-Újlak), Egerian (mid-and late Chattian) stage (d: Eger, Andornaktálya; e: Eger-Wind; f: Kesztölc; g: Pomáz; h: Vértesszőlős), Eggenburgian (late Aquitanian-lower and mid-early Burdigalian) stage (i: Ipolytarnóc), Karpatian (early-upper Burdigalian) stage (j: Magyaregregy), Badenian (lower–upper Burdigalian to the mid- Serravallian) stage (k: Nógrádszakál), Sarmatian (upper Serravallian) stage (l: Erdőbénye), Pannonian (Tortonian) stage (m: Bükkábrány; n: Hidas; o: Visonta), Pliocene1: mid-Zanclean stage, q: overview map)

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Non-monsoonal subtropical regional analogies of the Late Paleogene to mid-Pliocene

Considering the Southeastern United States as a climatic model environment, it can be hypothesized that ultisols were the most dominant soils from the Rupelian to Zanclean in the Pannonian Basin. However, alfisols, inceptisols, mollisols, spodosols and vertisols could also have occurred in the regions of low to mid-elevations in the Carpathian Basin. It should be noted that no trend is observed with regard to the change in the relative expansion of the different soil types over time when comparing the climatically analogous regions of the Southeastern United States with the estimated annual temperature and precipitation sum values from the past at the plant-fossil sites in the Pannonian Basin (Fig. 15).

Fig. 15
figure 15

Soil orders in the present-day southeastern part of the United States are climatically analogous to the Oligocene-Pliocene Pannonian Basin. The classification was based on the Köppen-Geiger climatic categories according to the reconstructed mean annual temperature and annual precipitation values (a Budapest; b Eger-Kiseged; c Nagybátony-Újlak), Egerian (mid-and late Chattian) stage (d Eger, Andornaktálya; e Eger-Wind; f Kesztölc; g Pomáz; h Vértesszőlős), Eggenburgian (late Aquitanian-lower and mid-early Burdigalian) stage (i Ipolytarnóc), Karpatian (early-upper Burdigalian) stage (j Magyaregregy), Badenian (lower–upper Burdigalian to the mid- Serravallian) stage (k Nógrádszakál), Sarmatian (upper Serravallian) stage (l Erdőbénye), Pannonian (Tortonian) stage (m Bükkábrány; n Hidas; o Visonta), Pliocene1: mid-Zanclean stage, q overview map)

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Discussion

Early Oligocene

The results suggest that the palaeoenvironment in the Pannonian Basin during the Early Oligocene resembled the present-day monsoon-influenced humid subtropical climatic (Cwa) regions in Southeast China or the humid subtropical (Cfa) ones in the Southeastern United States. Based on the results, temperate seasonal forests covered the region during this epoch. It cannot be ruled out that the climate in the Pannonian Basin during the Rupelian and even the Chattian was affected by mild monsoonal influence. Alternatively, the presence of former intense Mediterranean tropical-like cyclone activity (Miglietta and Rotunno 2019) can also explain the estimated relatively high annual precipitation sum values (ca. 1321–1420 mm according to Erdei et al. 2007). Due to the warm and humid climate, the weathering processes could have been very intensive in the Rupelian, reaching the transition zone between strong and moderately strong chemical weathering (Fig. 12). Today, similar climatic conditions support the formation of reddish haematite and kaolinite-rich ultisols or alfisols. Concerning the relatively smaller amount of ferric oxide of certain red-clay sediments from the Early Oligocene, Fekete and Stefanovits (2002A, B) concluded that originally the related paleosols were closer to the present-day tropical and subtropical oxisols than ultisols. Furthermore, the oldest Late Paleogene palaeosols contain a large amount of boehmite and gibbsite. Although these minerals occur mainly in oxisols (ferralsols) and ultisols (Allen and Fanning 1983), they are often the products of ferralitic weathering processes which involve a strong depletion of basic cations (van Breemen and Buurman 2002). As was shown, even during the relatively warm Rupelian stage, the climatic conditions were only barely suitable for the formation of oxisols in the Pannonian Basin. Nevertheless, it cannot be ruled out that oxisols were also formed during the warmest periods of the Early Oligocene in the Pannonian Basin.

In the field, the coexistence of oxisols and ultisols is an existing phenomenon and cannot only depend on the climate but also on the relief and hydrological conditions (Moniz and Buol 1982). Nevertheless, considering the present-day climatic conditions resulting in the occurrence of different soils (Fig. 13), it can be said that the climate of the Carpathian Basin during the Oligocene may have rather supported the formation of ultisols than oxisols. The dense forest vegetation might also support the formation of ultisols because these reddish, clay-rich, acidic soils are formed on mixed forest vegetation in humid temperate regions (Adams et al 2019). In this regard, it should be added that, e.g., in China, the distribution limit of mixed, evergreen-deciduous broad-leaved forests corresponds to the zone of 11–18.1 °C mean annual temperature and 1000–1630 mm annual precipitation (Ge and Xie 2017). It may indicate that most of the Late Paleogene and Early Neogene forests of the Pannonian Basin belonged to this kind of temperate seasonal forest type. The extensively leached soil horizons of soils formed under subtropical mixed forests are well-developed, contain a high amount of clay minerals from the kaolin group and are rich in iron oxides, which explains the reddish colour of these soils (West et al. 1997). Because ultisols could be expected in the adequate terrestrial palaeoenvironments of the Pannonian Basin between the Early Oligocene and the end of the Neogene, it can be hypothesized that one of the most frequent precursors of red clays formed in this era could be these kinds of soils.

Alfisols and mollisols could also have been present on the Rupelian landmasses of the Pannonian Basin. Alfisols are generally formed under hardwood forest cover (Rust 1983). However, the existence of the second soil type in the Oligocene Pannonian Basin is dubious because mollisols were almost always formed under native grass vegetation (Bhattacharyya et al. 2006) and the mixed grassland-forest vegetation only appeared around 8 Ma parallel to the global cooling and expansion of C4 grasses (Beerling and Osborne 2006).

The results raise the possibility of the existence of tropical-subtropical karstification in the Pannonian basin in the late Paleogene and early Neogene periods. In the early Oligocene, the karst morphology may have been fengcong/fenglin-like because on the one hand, the high palaeotemperatures supported intensive chemical weathering and, on the other hand, the CO2 concentration was ca. 1.9–2.5 times higher in the Rupelian than today (Zhang et al. 2013). Due to these conditions, the dissolution of carbonates could also have been very intensive since positive correlations were observed between the temperature and atmospheric CO2 concentration as well as carbonate weathering fluxes (Romero-Mujalli et al. 2019). Furthermore, ultisols are generally acidic (Allen and Fanning 1983). Their high H + activity could also lead to the intensive karstification of the carbonate bedrock in the Carpathian Basin during the Early Oligocene, creating tropical-like karst landscapes with expressed karst features as hypothesized by Hevesi and Mádai (2007).

Late Oligocene to middle Miocene

The second younger “bauxitic” red-clay group found in the Pannonian Basin has allitic characteristics. These red-clays contain a high quantity of quartz as well as a substantial amount of kaolinite and chlorite, but in relatively small amounts of boehmite, gibbsite, and hematite. This chemical composition indicates that this relatively kaolin-rich but not bauxitic-like clay group was formed under moderate weathering conditions (Fekete and Stefanovits 2002A, B). The alteration in the chemical composition of red clays in the Neogene/Paleogene boundary could be related to the decreasing annual precipitation in the Pannonian Basin which resulted in the increase of the territory of woodland/shrubland regions at the expense of forested areas. This environment change created a more open vegetation picture and less shading on the ground surface. The greater amount of solar radiation led to more intensive physical erosion due to greater fluctuations in the temperature and humidity of soils and created more oxidative soil conditions. This supported the formation of reddish palaeosols because the red colour of palaeosols often indicates the presence of coarsely crystalline hematite or goethite, generally resulting from progressive dehydration, oxidation as well as the transformation of hydrated minerals under well-aerated and warm climates (Tabor et al. 2017). It should be added that the red clays of the Middle and Late Miocene (e.g., Fig. 2h) do not necessarily represent purely reaccumulated, older materials but can also be the relict of a regolith member of a former ultisol (C horizon).

The next change in the composition of red-clays during the Middle Miocene may have resulted from the Monterey Carbon Isotope Excursion (MCIE). It is known that before the MCIE, which was related to the development of the East Antarctic Ice Sheet, the climate of the Carpathian Basin was warm and humid (Jiménez-Moreno et al. 2005, 2006). Although the climate became colder after the MCIE, in the Tortonian, it became wetter again, presumably due to the regional mesoclimatic influence of the extensive Pannonian Lake. In this regard, it should be noted that most reconstructions postulate archipelago-like palaeogeographical conditions in the Pannonian Basin between the Late Paleogene and Early Neogene epochs (Kováč et al. 2017; Popov et al. 2004). The archipelago effect might have influenced the environmental conditions on the former islands of the Pannonian Basin, which might result in a somewhat drier climate in the territory of the former islands and peninsulas. Paleovegetation observations may well support this hypothesis. For instance, the Tortonian Kálla Formation (Transdanubian Mountains) contains the fossils of Pinus halepensis Mill. and Rhamnus alaternus L. (Csillag 2003), which are currently characteristic of the dry- and hot-summer Mediterranean regions of Europe, the Levant and North Africa, including the thermomediterranean zone of the Mediterranean mountain ranges (Varone and Gratan 2009; Maestre and Cortina 2004). In contrast, it can be assumed that mesic fossil plants of the Kálla Formation like Myrica and Liriodendron sps. (Csillag 2003) lived in the more humid mesomediterranean zone of the former islands of the Pannonian Basin-filling Central Paratethys.

Pliocene and quaternary

The chemical composition of the young red clays from the Late Neogene–Early Pleistocene, formed on Tortonian palaeosurfaces, is different to the older ones from the Miocene. These red clays contain substantial quantities of illite, chlorite, montmorillonite and kaolinite. The red colour of these sediments is generally a consequence of their amorphous iron compounds (Fekete and Stefanovits 2002A, B). These red clays can be the relicts of typical Mediterranean terra rossa-like palaeosols, formed during the Pliocene's episodic, dry-summer subtropical conditions and the first part of the Pleistocene (Fekete 2002). However, it is not clear under what climatic conditions these sediments were formed, as the Pliocene palaeoenvironments seem to be heterogeneous. For instance, in the Zanclean, temperate deciduous forests covered at least the western regions of the Pannonian Basin under the former mild, humid continental climate. Fossil evidence suggests that the contemporal palaeoflora was dominated by diverse deciduous trees including, e.g., Ginkgo and Zelkova sps. (Kovács et al. 2015). The mild and humid subtropical (Cfa) climatic conditions abruptly changed during the mid-Pliocene glaciation when the climate became cold and semi-arid (Bsk). It is plausible that this kind of steppe environment also returned during the warm but dry periods of the Lower Pleistocene because such elements of steppe-dwelling fauna like Camelus and Pachystruthio were found in the Lower Pleistocene sequence in Kisláng, Central Transdanubia (Kordos and Krolopp 1991). Relatively mild climate, moderate chemical weathering along with subtropical karstification can explain the formation of the Tengelic Red Clay Formation from the Late Pliocene to mid-Pleistocene (Koloszár 2004) as well as the development of the sediment-receiving karst depressions. It can be hypothesized that a semi-arid climate could not only have been characteristic during the mid-Pliocene cold period in lower elevation regions of the Pannonian Basin but also later in the Pleistocene because sand and sandy loess layers were also found within the loess strata (Újvári 2004).

The palaeoenvironmental conditions of the Quaternary varied much more than in the Neogene and Late Paleogene, following the cyclic climatic changes. During the coldest glacial periods, the generally boreal forest was the dominant biome in the subarctic climate. Tundra-like palaeoenvironments were present in the higher and mid-elevation regions of the northeastern part of the Carpathian Basin. However, the dry and cold so-called “mammoth steppe” with predominantly grass-dominated vegetation could not have been the only dominant environment because the temperature-precipitation pairs fall outside of the range of the temperate grassland/cold desert biome. This finding is in good agreement with the gastropod fauna-indicated paleothermal conditions of the Pannonian Basin during the Pleistocene (Molnár et al. 2019). In the coldest Quaternary periods, slight mechanical weathering could have been the dominant erosion action, resulting in periglacial erosion at the peaks and foothills of the mountainous regions of the Pannonian Basin (Fehér et al. 2006) and created cryoturbation-induced palaeosol structures (Horváth et al. 2005). Moderate chemical weathering with frost action could have been characteristic during the transitional periods between the glacials and interglacials. In these periods, spodosols were formed under the former boreal forest vegetation of the Pannonian Basin. As has been shown, moderate chemical weathering conditions lasted throughout the mid-Pliocene glaciation. However, suitable, moderately warm and humid climates also occurred in the warmest interglacials of the Quaternary, which explains the presence of terra rossa-like palaeosols in the Lower and reddish-brown palaeosols in the Upper Pleistocene loess sequences (Kovács et al. 2013). Others suggest that, in addition to the haematite-rich clay products of the former chemical weathering processes, Saharan dust deposition also significantly affected the formation of interglacial reddish soils in the Pannonian Basin (Varga et al. 2016). The intensity of the karstification processes could have been low during the Quaternary period.

Conclusions

The model results suggest that, except for during the second half of the Pliocene, the Late Paleogene–Neogene palaeoenvironment of the Carpathian Basin resembled what could be found in the subtropical regions of the Southeastern United States as well as in East and Southeast China during the Anthropocene. Until the end of the Middle Miocene, relatively intense chemical weathering processes produced well-developed karstic geomorphological features, iron oxide and kaolinite-rich palaeosols, which were redeposited and sometimes preserved at the bottom of the former karst depressions. The dense subtropical vegetation might facilitate chemical weathering and a reduction in the rate of physical weathering caused by the surface-rinsing effect of rainfall. During the Late Miocene and Early Pliocene, due to the decreasing temperature and amount of precipitation, the rate of chemical weathering, including karstification, reduced. However, terra rossa-like deposits and larger-scale karst structures still developed until the mid-Pliocene. The nature of the environmental conditions changed during the second half of the Pliocene, from which time onwards, physical weathering and temperate karstification processes became characteristic in the Carpathian Basin.