3.1 Loess records (Fig. 2)

Numerous loess sequences within the Carpathian and lower Danube basins show the identified penultimate loess unit assigned to MIS 6 or early MIS 5e, either based on luminescence ages or on magnetic susceptibility variations, in Bulgaria (Jordanova and Petersen, 1999; Jordanova et al., 2008), in Serbia (Fuchs et al., 2008; Markovic et al., 2009; Murray et al., 2014), in Ukraine (Buggle et al., 2009), in Romania (Balescu et al., 2010; Timar-Gabor et al., 2011; Vasiliniuc et al., 2012), and in Hungary (Novothny et al., 2010, 2011; Ujvari et al., 2014). Some of them show the occurrence of a tephra layer, e.g. in Batajnica, Ruma, and Stalać in Serbia (Buggle et al., 2009; Markovic et al., 2009; Novothny et al., 2010, 2011; Obreht et al., 2016; Ujvari et al., 2014; Vandenberghe et al., 2014) and at Mostiştea in Romania (Balescu et al., 2010; Panaiotu et al., 2001). Markovic et al. (2015) did not assign this tephra to any particular eruption, but considering the age model proposed, it rather corresponds to either the Vico B ignimbrite or the Pitigliano tuff (Giraudi and Giaccio, 2017) at about −170 ka. Furthermore, in the review of Danube loess sequences, Markovic et al. (2015) also refer to an embryonic pedogenic layer older than the identified −170 ka tephra but younger than the first loess subunit identified at the start of the L2 loess unit correlated with MIS 6, above the interglacial paleosol S2SS1. Because of its position in the stratigraphy scheme, this embryonic paleosol is correlated with the first paleosol identified in Harletz between about −15 and −14.5 m of depth (Antoine et al., 2019; Fig. S1 in the Supplement). The other nine paleosols identified in Harletz have not been mentioned or described westwards in the Serbian and eastwards in the lower Danube loess sequences, making this highly developed loess record of MIS 6 a remarkable reference sequence for this region (Antoine et al., 2019).

Figure 2Correlation between the last two climate cycles from the southern Europe loess sequences along the Danube, modified after Antoine et al. (2019). Map of the locations (see Fig. 1b); the colored rectangles on the map (left) refer to the pedostratigraphic sequences correlated on the correlation sketch (right). Interstadials ISp and ISm (this study; see Fig. 1c) are reported on the Harletz record.

3.2 Pollen records (Fig. 3, Table 1)

Located at 40^∘54^′–41^∘10^′ N and 20^∘38^′–20^∘48^′ E in the Balkan Peninsula within the Dinaride–Albanide–Hellenide mountain belt, Lake Ohrid is a transboundary (Macedonia and Albania) lake. This site is the closest pollen record to the Harletz sequence, where two different open vegetation zones represent MIS 6, namely OD-5 (−190 to −160 ka) and OD-4 (−160 to −129 ka) (Sadori et al., 2016). The former corresponds to a grassland-dominated environment and the latter to a rather steppe-dominated environment. A tephra layer also marks the boundary between the two vegetal formations, indicating that OD-5 and OD-4 could be correlated with the lower and the upper parts of the Harletz MIS 6 record, respectively (Antoine et al., 2019). Considering the variations in both the percentage of total arboreal (AP) or total arboreal minus pine + juniper + birch pollen grains in OD-5, only four tree expansion phases can be noticed, which could correspond to four out of the five interstadials identified in the lower part of the Harletz MIS 6 record. Their onsets are dated, according to the age model used (Sadori et al., 2016), at about −185, −178, −175, and −169 ka below the tephra layer. In the following OD-4, four other interstadial onsets are dated at about −159, −151, −145, and −139 ka.

Figure 3Western (ODP 977 SST – alkenone; Martrat et al., 2004a, b) to eastern (Soreq speleothem δ¹⁸O; Bar-Matthews et al., 2003a, b) Mediterranean records of MIS 6, including a pollen record from Ioannina I-284 (AP, AP − pine + birch + juniper; Roucoux et al., 2011), compared with summer insolation with an indication of the time occurrence of Sapropel 6 (Bard et al., 2002). The pedostratigraphy of the Harletz MIS 6 record is added with the newly described paleosols and pedogenic horizons for comparison. Labels of the interstadials are as identified in the Alboran Sea. Complementary labels assign warm events to the ODP 977 (AI-5'a, b, AI-9'a, b), Ioannina (I1–12), and Soreq (S1–9) Mediterranean MIS 6 records used in Table 1. MIS 6 ice sheet extent by Florence Colleoni in Dendy et al. (2017).

Table 1Synchronism of the various interstadials discussed in this study. Indication of the predicted DO event occurrence by Barker et al. (2011) for the various assigned interstadials from Sanbao11 and Soreq speleothem δ¹⁸O, ODP 984, ODP 983, U1308 remaining percentage in N. pachyderma s., ODP 977 and MD01-2443 sea surface temperature (SST), and Ioannina I-284 arboreal pollen variations. Marine Isotope Stage (MIS) 6 is decomposed into an early part corresponding to the occurrence of the Harletz paleosols (192–170 ka), and a later part (170–130 ka) with incipient weathered horizons.

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At lower latitude, the Tenaghi Philippon site is located in Greece at 41^∘10^′ N, 24^∘20^′ E; 40 m a.s.l. This is a pollen record famous for its long continental climate record covering about eight climate cycles (Tzedakis et al., 2006). This record indicates woody taxa expansions, which can be considered interstadials, during the interval equivalent to MIS 6. Following the age model used (Tzedakis et al., 2006), they occurred at about −185, −180, −171, −163, −153, −151, and −138 ka. Therefore, some of them seem potentially synchronous with the tree expansions described at higher resolution in Lake Ohrid and by extension to Harletz interstadials. Another long pollen core retrieved in Greece, the Ioannina I-284 core (39^∘45^′ N, 20^∘51^′ E; 472.69 m a.s.l.), which covers several climate cycles as well, reveals the record of MIS 6 in the section −138 to −99 m. Roucoux et al. (2011) have described the high-resolution variations of the vegetation during this time interval. Phases of temperate tree expansion, interpreted as interstadials, and phases of tree contraction, interpreted as stadials, have been identified. Based on the employed age model (Roucoux et al., 2011), these temperate tree expansions occurred at about −186, −181, −177, −173, −170, −165, −159, −156, −145, −138, and −135 ka that we labeled I11 to I1 (Fig. 3). They appear more numerous than in Tenaghi Philippon and in Lake Ohrid, probably due to the higher resolution of the pollen analysis of the Ioannina core, and similarly some of these tree expansions could be correlated with the Harletz MIS 6 interstadials, five expansions being recorded older than −165 ka (Table 1).

3.3 Marine records (Fig. 3, Table 1)

On both sides of the Strait of Gibraltar, Martrat et al. (2004a, 2007a) have described the climate variations expressed during at least the past two climate cycles from cores retrieved from the Alboran (ODP 977A; 36^∘1.9^′ N, 1^∘57.3^′ W; 1984 m below sea level) and the Iberian seas (MD01-2443; 37^∘52.85^′ N, 10^∘10.57^′ W; 2925 m below sea level). The analysis of δ¹⁸O from the planktonic foraminifera Globigerina bulloides shows variations with abrupt changes mimicking – for the last climate cycle (MIS 5 to 2) – those described in the Greenland ice core records. During the penultimate cycle (MIS 7–6, −245 to −130 ka) the foraminifera δ¹⁸O record also shows abrupt warmings similar to the DO interstadials, which have also been observed in the variations of the ${\mathrm{U}}_{\mathrm{37}}^{{\mathrm{k}}^{\prime }}$ alkenone index, a proxy for the sea surface temperature. Nine interstadials are therefore identified, named Alboran or Iberian margin interstadials, although the older one, AI-9', seems to be a composite in which three different events can be discriminated. The onsets of these penultimate cycle interstadials are dated at about −186 (−186, −182, −179), −176, −170, −165, −159, −152, −150, −141, and −133 ka. They were labeled AI-9', AI-8', AI-7', AI-6', AI-5', AI-4', AI-3', AI-2', and AI-1', respectively. Similarities in terms of the number of warming events appear with the long Mediterranean pollen records mentioned previously. Although no tephra is identified, five events can once more be identified prior to −165 ka, as also observed in Harletz (Fig. 3, Table 1). Martrat et al. (2007a) stated that the observed variations in sea surface temperature, expressed by ${\mathrm{U}}_{\mathrm{37}}^{{\mathrm{k}}^{\prime }}$ for the last four climate cycles in the western Mediterranean, show a nonlinear response to external triggers of climate: obliquity and precession. However, some interstadials are synchronous with Mediterranean sapropels, which are a direct response to orbital forcings. Rohling et al. (2015) presented an updated review of the present and past Mediterranean climate and oceanography with a clear differentiation between the western and eastern basins, which may have impacted the adjacent terrestrial regions.

3.4 Speleothem records (Fig. 3, Table 1)

Two Mediterranean speleothems provide records for MIS 6. The Argentarola Cave, located (42^∘23^′8.79^′′ N, 11^∘3^′59.17^′′ E) by the Tyrrhenian Sea, shows an interval of low δ¹⁸O values between −180 and −165 ka with minima at −178, −172, and −168 ka, the latter being in phase with high-latitude insolation. The −180 to −165 ka interval is interpreted as corresponding to the Mediterranean Sapropel S6 (Bard et al., 2002) (Fig. 3). The recorded variations are correlated with total organic content variations recorded in the two Mediterranean cores MD84641 and KC19C, indicating rather warm and pluvial conditions during this interval over the western Mediterranean. Eastward, in Soreq Cave in Israel (31.458^∘ N, 35.038^∘ E; 400 m a.s.l.), a speleothem record covering the last 180 kyr was obtained. The measured δ¹⁸O record of MIS 6 shows variations at about −178, −170, −165, −160, −152, −149, −144, −138, and −134 ka (Fig. 2). The peaks at −178 and −152 ka show very low values, interpreted as corresponding to intense pluvial intervals over southern Israel, with the former correlated with the start of Sapropel S6, while no particular sapropel is associated with the latter (Ayalon et al., 2002; Bar-Matthews et al., 2003a), although it is referred to as a monsoon index maximum (Melières et al., 1997). These records indicate high precipitation rates at least during the −180 to −165 ka interval, which also corresponds to the most clearly expressed interstadials and paleosols in Harletz. Describing Soreq and another record in northern Israel, Bar-Matthews et al. (2003a) concluded that these speleothems recorded the proxy signal of global and regional eastern Mediterranean climate over the last 250 kyr. As Fig. 3 shows, the age model of Soreq is based on U−Th dating and agrees with the Bard et al. (2002) results, while the marine and continental cores are partly orbitally tuned.

Although the interstadials identified in the Harletz MIS 6 sequence have not been described in the closest loess sections from the Carpathian Basin and lower Danube region, the overview of the various and diversified MIS 6 records presented above clearly indicates that these paleosols and incipient weathered horizons correspond to climatic events that are recognized all along the northern Mediterranean and therefore have more than a local to regional significance. This shows that the climate mechanism proposed to explain the paleosol–loess alternations of the last climate cycle (Antoine et al., 2009; Boers et al., 2017, 2018; Rousseau et al., 2002, 2007, 2017a, b), in northern European sequences at about 50^∘ N, seems to have already prevailed during the penultimate climate cycle. This suggests that these alternations are part of more global dynamics. Therefore, in the next step of our study we will look for similar events in other records outside the Mediterranean region at a more global scale.

4.1 Sanbao11 (Figs. 4, S1)

The Sanbao cave is located at 110^∘26^′ E–31^∘40^′ N, 1900 m a.s.l., in central China, south of the Chinese Loess Plateau. It is strongly influenced by the East Asian monsoon variations and correlates with the summer insolation gradient between 65 and 15^∘ N. Moreover, similarly to the last climate cycle record measured from the Hulu Cave, the δ¹⁸O record of Sanbao speleothem B11 indicates abrupt changes at a millennial scale during MIS 7 and 6, which have been interpreted as interstadials by Wang et al. (2008). Considering the age boundary between MIS 7 and 6 (Lisiecki and Raymo, 2005) at 191 ka, the Sanbao speleothem preserved five interstadials, named B24 to B20, in MIS 7 (although missing the lower 19 ka; limit at −243 ka) and 13, named B19 to B1, in MIS 6. The δ¹⁸O minima in MIS 6 are dated at −189 ka for B19, −181 ka for B18, −175.5 ka for B17, −172 ka for B16, −168 ka for B15, −166 ka for B14, −162 ka for B13, −160 ka for B12, −157 ka for B11, −151 ka for B10, −149 ka for B9, −147.5 ka for B8, −145 ka for B7, −136 ka for B2, and −134 ka for B1 (Fig. 3); note that B3 to B6 are better recorded in the Hulu Cave MSP speleothem (Cheng et al., 2006). According to the prevailing interpretation, the lower the δ¹⁸O values from the speleothem record, the more intense the East Asian summer monsoon intensity. Interestingly, the minimum δ¹⁸O values for B17, B16, and B15 in MIS 6 are much lower than the other minimum values identified in the past 224 kyr record, including those characterizing the past two interglacials (MIS 7 and 5e) and the Holocene one. Hence, the time interval between −178 and −165 ka must have corresponded to extremely high precipitation rates over the East Asian region. This observation is in accordance with the high precipitation interval derived from the speleothems from Soreq and Argentarola caves in the Mediterranean region during Sapropel 6. Interestingly, the Sanbao cave record yields five interstadials between −190 and −167 ka and hence the same number of interstadials as the number of paleosols observed in Harletz below the tephra layer. Moreover, the B21, B16, and B8 δ¹⁸O minima correspond to peaks of decreasing magnitude in summer insolation at 15^∘ N. However, although summer insolation at both 15^∘ N (tropics) and 65^∘ N (arctic circle) show maximum values corresponding to the Sanbao δ¹⁸O main minima, the insolation gradient (15–65^∘), which induces a potentially strong meridional heat transport in the atmosphere and a possible excess in precipitation over the continent, indicates the weakest maximum for the time interval B17–B15 (about −177 to −166 ka) (Fig. 4).

4.2 Greenland GL_t_syn (Figs. 1, 3, Table 1)

Using the argument of the bipolar seesaw interpretation of the antiphased climate variation prevailing between Antarctica and Greenland, Barker et al. (2011) have developed a synthetic δ¹⁸O record of Greenland for the last 800 kyr based on the EPICA results and timescale, named GL_t_syn. This synthetic record nicely replicates the millennial-scale variability observed during the last climate cycle in the NGRIP ice core and therefore makes the authors confident that it reliably reproduces the last 800 kyr of Greenland δ¹⁸O abrupt variability. Two timescales have been proposed: one based on EPICA DC3 and another based on the Sanbao speleothem, with some discrepancies. Because of the more precise dates obtained from the Chinese speleothems, we will use the GL_t_syn with the Sanbao timescale for our comparisons. Plotted versus the Sanbao δ¹⁸O record, the GL_t_syn record indicates intervals showing peak maximum values aligned with most of the Chinese speleothem interstadials. Among these peak values, Barker et al. (2011) predicted the occurrence of 11 DO-like events, which should correspond to as many interstadials (Fig. 4, Table 1). Five of them can be clearly identified between −190 and −166 ka, which align with the Sanbao ones, therefore supporting the global value of Harletz paleosols interpreted as interstadials.

4.3 North Atlantic (Figs. 1, 4, S1, Table 1)

South of Greenland, the North Atlantic Ocean is the classical region in which to observe the Greenland warmings and interstadials, as this was demonstrated for the last climate cycle (Bond et al., 1992; McManus et al., 1999; Henry et al., 2016; Hodell and Channell, 2016).

For the penultimate cycle, several records also indicate warming events interpreted by the respective authors as interstadials. The warming events observed in the sea surface temperature (SST), deduced from alkenone studies from the Iberian margin, are of similar magnitude as those observed in the Alboran Sea (ODP 977/MD12443) (Fig. 3). Furthermore, the Iberian margin core MD01-2444 shows temperate tree pollen percentages, which vary in line with the planktonic δ¹⁸O measured in the same core (Margari et al., 2014). Such high percentages of arboreal pollen during the base of MIS 6 (−185 to −160 ka) are interpreted in terms of reduced summer aridity and increased pluvial conditions. Back to the sea surface temperature, which has been interpreted as mimicking the Greenland δ¹⁸O variation (Shackleton et al., 2000), plotting ODP 977 alkenone SST reconstructions of MIS 7 and 6 against GL_t_syn, one can notice that numerous discrepancies appear in the number of interstadials and for those identical in their occurrences, which seems to be due to the age model used for ODP 977. Further north, in core U1308 (49^∘53^′ N–24^∘14^′ W, 3871 m n.s.l – near sea level) Obrochta et al. (2014a) report important variations in Neogloboquadrina pachyderma sinistral, an indicator of cold surface water, during MIS 6. The minimum counts are much stronger at the base of the MIS 6 record, between −190 and −160 ka, than in its upper part. Although dating these warming events can be proposed from the published material, at about −190, −175.5, −167.5, −160, −166.5, −143, and −137 ka, these dates remain preliminary as the time resolution is not high enough compared to GL_t_syn or ODP 977/MD01-2443 marine cores (Fig. 4). The M23414-9 gravity core (53.537^∘ N, 20.288^∘ W; water depth 2199 m), selected to study variations in the North Atlantic drift, indicates the highest values in N. pachyderma s. percentages between −165 and −130 ka. However, minima correspond to warmings in both summer and winter sea surface temperature estimates (Kandiano et al., 2004). A total of 11 of such events can be dated at about −192, −187.5, −180.5, −175.5, −172, −163.5, −158.5, −141, −140, −137, and −133.5 ka BP, but they show very little overlap with the other marine cores previously mentioned, probably because of a lower resolution. Still, the five lowest occurred before −165 ka. A similar pattern is noticed further north in ODP 983 (60.4^∘ N–23.6^∘ W, 1984 m n.s.l), with stronger minimum counts of N. pachyderma s. between 190 and 160 ka than between −160 and −130 ka. Using the GICC05/NALPS/China timescale (Barker et al., 2015), warming events are observed at about −189.9, −178.7, −173.2, −169.5, −168.2, −163, −160.7, −151, −149.1, −147.4, −140.4, and −134.5 ka (Fig. 4). These dates vary slightly when using the other proposed timescale based on the EDC3 age model. Closer to Greenland and south of Iceland, ODP 984 (drilled at 61.25^∘ N and 24.04^∘ W, 1648 m) yielded a record of the past 220 kyr (Mokeddem and McManus, 2016). It includes the complete MIS 6, showing millennial-scale variability that is best expressed during the interval −170 to −130 ka. The foraminifer composition indicates alternation between episodes of northward polar-front retreats (low values of N. pachyderma s.), synchronous with warm and salty water inflows (high values of Turborotalia quinqueloba), and episodes of southward polar-front advance (high values of N. pachyderma s.), with fresh and cold water inflows (low T. quinqueloba) and high ice-rafted debris (IRD) content. The cold inflow episodes are interpreted as stadials, while the warm ones are considered interstadials equivalent to those observed during the last climate cycle. They occurred at about −165 to −163, −156 to −152, −148 to −146, −139 to −137, and −135 to −132 ka, and they are named IS6-5, IS6-4, IS6-3, IS6-2, and IS6-1, respectively (Fig. 4). Interestingly, these interstadials are the nearest geographically described to Greenland during MIS 6, but considering the resolution, more events could be considered when looking at the percentage of N. pachyderma s. (see Fig. 4 and Table 1).

When comparing the MIS 6 records described previously, e.g., the reconstructed Greenland δ¹⁸O, the Sanbao speleothem δ¹⁸O, and the foraminifer counts and δ¹⁸O measurements from Sites 984 and 983, one can notice that the interstadials observed in the North Atlantic records can have an equivalent in both the reconstructed Greenland synthetic and Sanbao δ¹⁸O interstadials (Fig. 4). Plotting every record used in the present study on its individual timescale, a global synchronism is far from being evident, except for the early MIS 6. When reconstructing the Greenland δ¹⁸O variations, Barker et al. (2011) predicted DO event occurrences for the last climate cycle, which fit with the DO events described from the Greenland ice cores. They expanded the occurrence prediction to the previous climate cycles covering the past 800 kyr. Plotted versus the Sanbao interstadials, although these predicted DO-like events fit for the early part of MIS 6, between −192 and −166 ka, fewer abrupt changes are assigned DO-like labels than in the Sanbao record of interstadials for the −166 to −130 ka interval, especially in the upper (more recent) MIS 6. This is demonstrated when comparing the reconstructed Greenland δ¹⁸O, the variations in N. pachyderma s. percentages in the North Atlantic ODP 984, the closest to Greenland, and the U1308 records. Reporting all the interstadials in Table 1 supports the previous interpretation that these interstadials are hemispheric and can be interpreted as DO-like events, without any doubt at least for the early MIS 6. However, the question remains as to why these marine records did not show all of the interstadials described in the other domains, contrary to what has been observed for the last climatic cycle. Is this related to the time resolution of the studied cores, which would not be high enough? This is an ongoing problem that this paper cannot address, as it requires further investigation and additional data with higher resolution.

4.4 Chinese loess records (Fig. S1)

One potential record of all these events has been proposed from sequences from the Chinese Loess Plateau, located north of the high-resolution speleothems described previously.

For the last glacial cycle, the study of the grain size variations from temporally highly resolved loess sequences indicates changes in the size of the deposited particles ,which are aligned, using a timescale based on luminescence dates, with the DO events (Sun et al., 2012). As the deposited material originates from northern Chinese deserts, during interstadials, when the wind velocity is reduced due to stronger East Asian monsoon blowing from the south, finer material is deposited. In contrast, during stadials, the wind velocity is stronger and the East Asian summer monsoon weakens, allowing coarser material to deposit. Following this idea, Yang and Ding (2014) have proposed a stack of several loess sequences covering the last two climate cycles (Fig. S1). They show that for the last 130 kyr, the median grain size variations match the former results by Sun et al. (2012) and correlate with the Hulu Cave speleothem δ¹⁸O record (Cheng et al., 2006; Wang et al., 2008). Expanding to the penultimate cycle, Yang and Ding (2014) indicate that the median grain size variations during MIS 7 and 6 also mimic the Sanbao cave δ¹⁸O variations within the uncertainties related to the dating of these sequences. Interestingly, the authors refer to interstadial paleosols that do not correspond to the identified individual interstadials but rather to a group of them (L2-4 groups B17, B16, and B15, while L2-2 gathers B11 to B5). This observation is similar to what was described for the last climate cycle record by Sun et al. (2012): not every interstadial is individually resolved in the stratigraphy, contrary to the various paleosols in the European loess sequences at about 50^∘ N, which correspond one to one with the Greenland interstadials.