The entrance part of the LCC consists of a main channel and several smaller branches. Morphologically, the channel consists of three parts. The first part, 70 m long, is straight and has a Dinaric orientation (NW–SE). This is followed by two successive bends and a third straight section about 50 m long. The entire channel runs almost parallel to the hill slope at the surface. The present cave entrance is a 1.6 × 1 m door. The entrance was naturally probably much lower but it was artificially slightly modified. The main channel is 6–8 m high, although in some places it is much higher (over 15 m). The width of the channel is generally in the range of 4–7 m. The area cross-sections (Figure 2C) are generally 20–35 m² but may be larger in some places, especially where the ceiling is higher due to some fractures. The depth of the allogenic sediment in the bottom of the cave channel is not known, so the full dimensions and shapes of the cross sections are not completely known. Three sets of fractures were recorded within the channel, which significantly influence its shape and formation. The first part of the channel is dominated by a set of fractures with Dinaric orientation (∼130°–310°), the middle part by a set of fractures with meridional orientation (∼0°–180°), and the last part is the most influential by the set of fractures with orientation of ∼100°–280°. The host-rock is not stratified, so there is no visible influence of bedding planes on the cave morphology. Numerous speleogens were recorded behind the channel walls: scallops, elongated domes, solution pockets, pendants, wall rills, etc. Flowstones appear only about 80 m from the entrance, mostly in the form of wide stalagmites and wall flowstones. Due to low temperatures in winter and occurrence of ice on the cave walls, the cryofracture weathering of flowstones is pronounced in this part. The floor in the almost entire length of the channel is built of sediment and in the upper part it is mostly covered by archaeological layer. At the site of the test trench, the channel is slightly larger. It is 12 m high, 6.6 m wide, and the cross-sectional area is about 45 m². This is most likely caused by the intersection of the two sets of fractures (Figure 2D, Supplementary Material S1). The genesis of the whole cave, including this entrance channel, is related to the denudation effect of the sinking waters of the Gračac karst polje. However, its mechanism is not yet completely clear. Morphological traces (cross-sectional shapes and speleogens indicating saturated conditions) of speleogenesis indicate that the channel was formed mainly under saturated conditions. Erosion traces of water flow in the vadose zone were not recorded because, if present, they are located in the lower parts of the channel covered with sediment.

Within the cave sediment infill found in the passage of the LCC, eight intervals were recognized and described. Based on similar lithological and structural features they were grouped into three lithofacies units and the top archaeological layer. A detailed description of the profile and the individual lithofacies units is given in Table 1 and Figure 3A. According to the results of particle size analysis by the sedimentation method (shown by cumulative granulometric curves, Figure 4A) it can be seen that all described intervals consist of clayey silt to silt with a small amount of very fine sand particles (<9% of sand). According to Trask’s sorting coefficient (S₀), all samples show poor to very poor sorting (1.802–3.067), while the asymmetry coefficient (Sk < 1) shows that grains smaller than the median (average Md value is 0.0154) predominate in the samples. Samples DC-SP 2, 3 and 5 are classified as very poorly sorted clayey silt, while samples DC-SP 6 and 7 are classified as poorly sorted silt. Furthermore, the results of the high-resolution particle size analysis using the laser diffractometer show a decrease in the amount of the clay-sized fraction from bottom to top of the section (Figure 4B), the trend comparable to the results of areometric particle size analysis (Figure 4A).

The results of the modal analysis of the sediment samples are given in Table 2. Within all samples the LMF predominates, represented by grains of monocrystalline quartz (84–92%) which is mostly represented by angular and slightly rounded grains of uniform and undulose extinction. In addition to this dominant group, euhedral quartz grains can be found (Figure 5A). Sporadic occurrences of well-rounded and spherical quartz grains were recorded, as well. The amount of rock fragments ranges from 4 to 8%. Among them, the most common are chert particles. Rare tuffitic particles were found, as well as schist rock fragments. Feldspars are represented mainly in the form of potassium feldspars, and their total amount is up to 9%. Feldspars are most often anhedral (Figure 5B) to subhedral. Mica (muscovite) appears only sporadically (<1%) in the form of transparent plates with a rounded outline (Figure 5C). The LMF is quite uniform throughout the profile (Table 2). A slight decrease in the number of quartz grains toward the top of the profile was observed, and in connection with that, a slight increase in the number of feldspars and lithic particles. The mineral composition of all analyzed samples is uniform. Among the HMF, the amount of opaque minerals is high, about 90%. Completely opaque black grains, often well rounded, are observed. Chromite grains, slightly reddish-brown colored, are present in all analyzed samples and in some samples pyrite (Figure 5D). The THM are very sparse in the samples. Among the THM, the pyroxene predominates. They appear in the form of anhedral to stubby prismatic grains, are green in color and the typical “hacksaw” terminations are often visible (Figures 5F,G). According to the extinction angle, they are classified in the group of clinopyroxene. The second most common translucent heavy mineral is zircon. Zircons, mainly short-prismatic or anhedral (slightly rounded) are present in all samples. Euhedral grains are rare (Figure 5G). Tourmaline is present in roughly the same proportion as zircon. It appears in the form of subhedral grains. It is rounded in some places. Pleochroism in brown to greenish color is visible in places. Other observed varieties belong to a group of hemimorphic grains with multicolored poles (Figures 5E,H). Rutile is rare but still present in all samples, appearing in rounded forms with a slightly prismatic habitus. Their color is usually reddish-brown or dark red. Garnets are rare, occur in the form of weakly rounded grains or angular grains/shards with sharp edges. Colorless garnets predominate. Slightly pink garnets are also present (Figure 5E). Among other THM, grains from the epidote-zoisite group rarely occur. The epidote is greenish, semi-rounded, in the form of irregular grains while mineral grains classified as zoisite/clinozoisite look fresh and show an anomalous blue interference color. Rare occurrences of biotite and greenish anhedral amphiboles are also present.

The results of the XRPD method are shown in Table 2. Sample DC-SP 1 is extremely heterogenous. It consists of bone fragments, and sandy silt size sediments. The analysis was performed on both parts. Sediment sample DC-SP 1 (silt), in addition to clay minerals, contains a significant amount of hydroxylapatite (HA) and quartz (Table 2). Quantities of fractions <2 μm were too small for clay analysis, so the analysis was performed on a fraction <0.063 mm. In that fraction quartz, vermiculite, illite/muscovite, talc-pyrophyllite, kaolinite and a small amount of chlorite are present. Bone fragment sample consists only of HA. The main mineral phases in all other analyzed samples are clay minerals and quartz (Table 2). Samples DC-SP 4 to DC-SP 7 contain a smaller amount of plagioclase. Some of the samples (Table 2) contain a very small amount of gibbsite and hematite, but due to the low content cannot be confirmed with certainty in all samples. Mineral composition of <2 μm fraction of all analyzed samples is similar. In the analyzed samples, among clay minerals, vermiculite, illite/muscovite, kaolinite, and a lesser amount of chlorite and chlorite-vermiculite regularly appear. In some samples, small quantities of low-charge vermiculite or high-charge smectite, talc-pyrophyllite, kaolinite which forms intercalation compounds with DMSO and illite-smectite are also present. Samples DC-SP 3 and DC-SP 5 probably contain secondary chlorite (the 14Å diffraction maximum disappeared after heating to 550°C).

Palynofacies of all studied samples beside the oldest one (DC-SP 1) are dominated by phytoclasts. Sporomorphs occur in a small amount and therefore there is no standard palynological diagram. Instead of that, only organic matter abundance is presented in the diagram (Figure 6). In the oldest analyzed sample, DC-SP 1, the palynofacies is dominated by bacterial amorphous organic matter (AOM) particles (Figure 7A), and non-opaque phytoclasts, mostly amorphous particles, which indicate an increased input of terrigenous material. Only a few Pinaceae pollen (Figure 7C) and Fungi spores occur. Microscopic charcoal remains (around 100 microns in size; Figure 7B) point to the influence of fire (Whitlock and Larsen, 2002). Palynofacies from the sample DC-SP 2 is dominated by opaque phytoclasts, mostly corroded charcoal, while non-opaque phytoclasts decreased. Sporomorphs from conifer Pinaceae (Figure 7D) family as well as herbs of Asteraceae (Figure 7D) and Cichoriaceae family (Figure 7E) dominate in the same ratio (6%). They point to a cold and dry climate, probably to a glacial stage. In sample DC-SP 7 palynofacies is still dominated by the opaque phytoclasts, mostly corroded charcoal. Beside phytoclasts there are a lot of particles resembling cyanobacteria, maybe degraded cyanobacteria (Figures 7F,G) that lived in the cave. Palynofacies from the youngest sample DC-SP 8 is dominated by the phytoclasts, mostly non-opaque phytoclasts—brown wood and amorphous particles. Rare findings of the palynomorph Pseudoschizaea (Figure 7I), probably related to Zygnemataceae, indicate the runoff due to periods of enhanced soil erosion outside the cave (Leroy et al., 2007). The presence of spores from the genus Glomus suggests erosion from a forested upstream slope (van Geel, 2001), possibly due to fires, themselves evidenced by an increase in microcharcoal particles. Aquatic pollen (Typha) and grasses (Graminae) were also present in the similar ratio (ca. 3%). At the same time rare Polygonum persicaria (Figure 7H) points to anthropogenic influence.

The vertebrate remains from the Pleistocene deposits of the test trench (specifically 873 fragments) were recovered at a depth of ∼1.2 m, within a 30 cm thick layer (Figure 3A). Of these, 230 bones and teeth are identified to the genus Ursus (26.4%). The vertebrate remains are therefore documented and presented within Figure 8. Based on morphological and metrical characteristics, all bear remains from LCC are attributed to cave bear (U. spelaeus), making this species the only mammalian taxon identified within the analyzed faunal assemblage from Pleistocene deposits. The vast majority of the remains, however, remained taxonomically undetermined. Based on their relative size and robustness, many of these remains could also come from a cave bear. However, given the mention of rare findings of other large carnivores and herbivores in previous studies (Malez, 1965b; Paunović et al., 1999), this assumption should be considered with caution and the possibility of the existence of other taxa in this assemblage should not be ruled out. Although most of the remains could not be aged precisely, information of the relative age at death shows predominance of adult bears (81.7%), while the rest belongs to juveniles (Figure 8). A single bone is attributed to a fetal or newborn animal.

In order to study the body part representation of bears in the LCC, data for both cave bear and taxonomically indeterminate remains of the similar body size are combined. All major parts of the body are present (Figure 8) suggesting deposition of complete bear carcasses within the cave. However, a closer examination of the differential representation of different body parts revealed the following: the most abundant class are trunk elements (32.2%), closely followed by the elements of the head (teeth included; 29.1%) and feet bones (25.2%). Relative to them, larger bones of the appendicular skeleton are under-represented within the analyzed assemblage. Thus, the upper elements of the fore limbs (scapula, humerus, radius, ulna) and hind limbs (pelvis, femur, tibia, fibula) are represented by only 6.1 and 7.4%, respectively.

With the exception of a few more complete bones, the skeletal material is fragmented. Recent breaks are present but most breakages are dry and attributable to natural processes typical of cave environment (e.g., trampling by other animals, sediment pressure). Looking at the bone surface modifications the material is relatively well preserved. The average bone color varies between pale white to yellowish white. Just a few fragments display small areas of dark brownish coloration, suggesting light staining probably due to exposure to minerals in the sediment. Besides fragmentation, the most common taphonomic modification is very light weathering (fine line fractures and spalling of bone surface), while chemical etching is evidenced on several fragments. In addition, only a few bones were gnawed by large carnivores (e.g., cave lion or hyena) and there is no evidence of modification by hominins.

Using the rejection criteria determined in dose recovery experiments, 43 equivalent dose values were accepted for the IR₅₀ signal, and 33 for the pIRIR₂₂₅ signal, respectively. Dose distributions for both signals are positively skewed and show overdispersion values of 56 ± 7% (IR₅₀) and 42 ± 7% (pIRIR₂₂₅), which in combination can be interpreted as an indication for incomplete bleaching being significant in the sample. Therefore, average equivalent doses for both signals were calculated using a bootstrapped three-parameter minimum age model (Galbraith et al., 1999; Cunningham and Wallinga, 2012), with sigma_b of 0.3 ± 0.2 as a threshold value, based on the overdispersion values from dose recovery experiments and assigned with an uncertainty to account for the lack of a well-bleached natural reference sample. The g-values of 5.6 ± 0.8 for the IR₅₀ and 0.6 ± 0.9 for the pIRIR₂₂₅ signal were obtained after fading corrections. Fading correction was conducted using the approach of Huntley and Lamothe (2001) and calculated using the R-Luminescence package of Kreutzer et al. (2012). Equivalent dose values and resulting ages as well as all luminescence results are summarized in Table 3.

The different characteristics of the two luminescence signals measured with the pIRIR₂₂₅ dose protocol can be used to assess the reliability of the determined ages. The IR₅₀ and pIRIR₂₂₅ signals are known to exhibit different fading rates, as was confirmed by the fading measurements in this study, and different bleaching rates, with the IR₅₀ signal bleaching much faster than the pIRIR₂₂₅ signal (e.g., Murray et al., 2012; Bickel et al., 2015a, 2015b). If incomplete bleaching is significant in a sample, the success of the correction of the effect of incomplete bleaching using the statistical approach of the bootstrapped MAM can be assessed by comparing the fading corrected ages for both signals. If the apparent age of the IR₅₀ is significantly younger than that of the pIRIR₂₂₅ signal, incomplete bleaching was not successfully corrected for. If, however, the fading corrected ages for the two signals are in agreement within error limits, like is the case here, this is a strong argument for a successful correction of the effects of incomplete bleaching. Although fading correction was applied also for the pIRIR₂₂₅ signal for comparative reasons, the fading rate is negligible within error and because of that the pIRIR₂₂₅ based age not corrected for fading of 53.7 ± 6.9 ka can be regarded as the most reliable depositional age for the sample (marked in bold in Table 3). In addition, it is important to note that the luminescence age is in stratigraphic order with the radiocarbon age obtained from a flowstone sample (Chapter 4.6.2), further corroborating the high reliability of the age determination.

The speleothem sample (Z-7352, graphite number A2160 (RBI), ID number UGAMS# 49576 (CAIS)—Figures 9A–C) had F ¹⁴C = 0.1148 ± 0.0006 (17,390 ± 40 BP) and δ ¹³C -5.3 ± 0.1‰. F ¹⁴C and radiocarbon dates without and with DCF of 12.5 and 15%, along with their calibrated dates are presented in Table 4. Compared calibration curves for both 12.5 and 15% DCF and both for using the reservoir function in OxCal and calibrating raw ¹⁴C dates are presented in Figure 9D. Here should be pointed out that there is a large difference between conventional radiocarbon dated (expressed as BP) and calibrated calendar dates (expressed as cal AD and cal BC) in this part of the radiocarbon calibration curve, resulting in difference of about 2,500 years between the conventional and calibrated age. The true age of material with the obtained age of 16 ka BP is therefore ∼19.5 ka old (Table 4).

The investigated sediments represent the clastic filling of cave channels. Earlier research assumed that the clastic filling of cave channels is an accumulation of in situ products of weathering of the host rock (Malez, 1965). According to mineralogical analysis presented within this paper the cave sediment is mainly allochthonous clastic detritus but a part of it is autochthonous chemogenic and collapse material. The overall mixture of cave detrital sediments depends greatly on the weathering products in the source area, transported and deposited by episodic events in different facies types (depending on flow dynamics) inside the cave (Georgiadis et al., 2019). Therefore, our results are compared to the geological units in the river Otuča catchment area (Figures 1B,C).

The results of the LMF and XRD analysis (Table 2) shows that the main components of the analyzed cave sediments are quartz and clay minerals. The sample DC-SP 1 additionally contains bone fragments and significant amounts of hydroxyapatite (HA) (Table 2). HA is the main constituent of mammalian bones and teeth and is often recognized within cave sediments, like e.g., in the Modrič Cave (Miko et al., 2001). Also, the relationship between the habitation of bats and HA formation in caves has been found in limestone caves worldwide (Hill and Forti, 1997). HA is usually formed as a crust that coats speleothems and bedrock substrate surfaces within or near bat habitats, while powdery forms of HA could be found under bat guano deposits (Chang et al., 2010 and references therein). Hence, it is possible that beside the tiny bone fragments in the fine-grained part of the sample DC-SP 1, a part of the HA may also have originated from the in situ bat guano, as indicated by the dark color of the sediment. In samples DC-SP 4 to DC-SP 7 small amount of plagioclase are also present (Table 2). Their preservation in the samples and the absence of gibbsite and hematite indicates that these samples were subjected to less intensive pre-burial weathering compared to samples DC-SP 2 and DC-SP 3. The clay minerals contained in the analyzed cave sediments (Table 2) are similar in composition to the Terra Rossa type paleosols developed on the carbonate rock in the Adriatic region (Durn et al., 2007, 2018). In the wider Mediterranean area, the mineral composition of Terra Rossa type soils and palaeosols may be very variable (Durn, 2003). This variable composition is a consequence of the polygenetic nature of Terra Rossa which can form exclusively from the insoluble residue of limestone and dolomite but much more often encompasses a range of parent materials that arrived on the carbonate terrain by different transport mechanisms (Durn, 2003). Thus, in Istrian Terra Rossa, the dominant mineral phases in the clay fraction are kaolinite, illitic material, Fe-oxides and XRD amorphous inorganic compounds, while vermiculite, low-charge-vermiculite or high-charge smectite, chlorite, mixed-layer clay minerals and quartz are present in smaller quantities (Durn, 2003). Terra Rossa soils from Western Herzegovina have a similar composition where the dominant mineral phases in the clay minerals fraction are kaolinite, Fe-oxides and XRD amorphous inorganic compounds, while vermiculite, smectite, illitic material, chlorite–vermiculite and quartz are present in a subordinate amount (Durn et al., 2014). However, even though cave sediments show similarities to Terra Rossa type soils, they cannot be classified as soils (Zupan Hajna et al., 2020). Cave sediments reveal a good, multi-proxy record of cave and surface environmental conditions in the time of their deposition (Bosák, 2002). Therefore, the red clayey-silty sediments found in the LCC can, to some extent, be considered as redeposited Terra Rossa. Despite the similarities, cave sediments appear to have suffered less advanced stages of weathering compared to Terra Rossa. Iacoviello and Martini (2012) came to the same conclusion comparing the clay mineral composition of cave sediments and Terra Rossa soils in the karst massif of Montagnola Senese in Italy. The source area of the investigated sediment could be in the Gračac karst polje in the close vicinity of the entrance to the LCC. In the catchment area of the Otuča river (Figure 1B), approximately 3.5 km upstream of today’s ponors close to the cave entrance, the river flows through the area covered with Terra Rossa type sediment, as can be seen from the geological map (Figure 1C, unit 10; Ivanović et al., 1973). Beside kaolinite, vermiculite, illite/muscovite, chlorite and mix-layered clay minerals, samples DC-SP 2 to DC-SP 6 contain minerals from the talc-pyrophillite group (Table 2). Talc-pyrophyllite usually occurs as a minor component in soils, which could be inherited from the parent rock, but they can also form as a result of weathering processes (Weaver, 1989). According to Velde and Meunier (2010) talc can be formed directly from pyroxenes. Other secondary minerals that can be formed from pyroxene are vermiculite, smectite, kaolinite or hematite. The analyzed samples contain a small amount of pyroxene group minerals (Table 2) which could have been the parent material for the talc-pyrophyllite group of minerals. Hematite is present only in traces (Table 2), very likely derived from the Terra Rossa. The mineral assemblage determined by the XRD within the analyzed DC-SP samples is comparable to detrital cave sediments within Dinaric karst (Bosák et al., 2012; Zupan Hajna et al., 2021), and e.g., cave sediments from northwestern Oltenia in Romania (Ghenciu, 2017). When comparing cave sediment compositions, it should be noted that the composition of these sediments depends on the clastic source rocks, and factors such as weathering and/or pedogenesis.

To answer the question about the clastic source rocks and source areas, results of LMF and HMF analysis (Table 2; Figure 5) were compared to the main lithological units in today’s catchment area of the Otuča river and Gračac karst polje (Figures 1B,C). Underground passages of the LCC are developed in the Tertiary carbonate breccia host rock (Ivanović et al., 1973) (Figure 1C, unit 8), known as Jelar breccia (Bahun, 1963, 1974) which is locally composed of various lithic fragments, most commonly related to the lithological composition of underlying rocks. Therefore, the breccia is mainly composed of lithic fragments of Jurassic, Cretaceous, and younger Paleogene carbonate rocks (Ivanović et al., 1976). In the investigated sediment, it was not possible to recognize this carbonate source area, except within the clearly in situ formed Bd facies which contains angular boulders of the host rock (Figure 11A). However, in most of the sediment, the siliciclastic detritus predominates (Table 2) which points to the allochthonous origin of the detritus. According to the results of the HMF and LMF analysis (Figure 5; Table 2) it can be concluded that the source of the clastic detritus is connected to the wider river Otuča catchment area. Downstream from the source, river Otuča flows through the upper Carboniferous deposits (Šušnjar et al., 1973) (Figure 1C, unit 1), mostly clayey shales and sandstones, accompanied by conglomerates. In the composition of the clayey shales quartz, muscovite, chlorite, plagioclase, kaolinite, organic substance, and pyrite can be found. In the layered, well-sorted, and fine-grained sandstones the detritus is composed mainly of quartz, muscovite, chlorite, lithic fragments, and a small amount of plagioclase (Sokač et al., 1976). Traces of such mineral assemblages can be recognized within the composition of LMF from LCC where quartz grains prevail (84–92%), together with the occurrence of muscovite and feldspar (Table 2), while the presence of chlorite and kaolinite is confirmed by XRD analysis (Table 2). Potential source area could be found within the clastic-pyroclastic series of middle to upper Triassic (Figure 1C, unit 3), composed of shales, quartz-calcarenites, subgraywacke sandstones, calcilutites, breccias, and tuffitic rocks (Sokač et al., 1976). The subgraywacke shows pronounced domination of quartz, followed by chert, plagioclase, and platy minerals such as biotite, muscovite, and chlorite, all comparable with the composition of the analyzed samples (Table 2; Figure 5). Furthermore, other rocks within the clastic-pyroclastic series contain angular quartz, chert/radiolarite, and in smaller amounts fragments of older sandstones and shales, similar to the investigated samples. The predominance of pyroxenes among the THM in investigated samples points to the upper part of the middle Triassic clastic-pyroclastic series as the possible source rocks (Sokač et al., 1976). Pyroxenes, as chemically unstable detrital constituents, possibly indicate higher erosion rates of the source rocks, rapid transport, and short grain residence time in the river (Sevastjanova et al., 2012; Wacha et al., 2019). Although present, garnet grains are not the dominant phase within the composition of THM in LCC detrital sediments (Chapter 4.3, Figure 5) but the overall mineral assemblage is almost identical except for the absence of corundum and apatite. The apatite absence could be related to the sample preparation process and chemical dissolution. The relatively abundant well-rounded opaque grains within HMF resembles to Fe-Mn nodules, which are common in clastic cave sediments and Terra Rossa type soils and paleosols (Durn et al., 2018; unpublished data) (Figures 5D,E,H). Their occurrence could be related to the earlier mentioned Terra Rossa type soil in the Gračac karst polje (Figure 1B, Ivanović et al., 1973). To conclude, the mineral composition of the clastic detritus in which there is high proportion of quartz, followed by feldspars, muscovite, opaque minerals, biotite, chlorite, chromite, pyroxenes, zircon, rutile, garnets, epidote, zoisite/clinozoisite group minerals and chromite can indicate a diverse source area, but the one that is geographically connected to the nearby river Otuča catchment area.

Due to the unique characteristic of cave environments, it is sometimes difficult to interpret specific depositional conditions within sequences of detrital cave deposits. Cave sediments represent the most complex terrestrial depositional environment where the law of superposition is often violated, facies are usually diachronous and re-deposition along the same cave passage is very common (White, 2007; Zupan Hajna et al., 2020 and references within). However, we were able to describe and interpret three lithofacies types within the sedimentary profile DC-SP in the LCC; the Breakdown facies (Bd), the Diamicton facies (Di) and the Slackwater facies (Sw) (Figure 3A; Table 1).

Sediments of the Bd facies are commonly formed by the gravitational collapse of the host-rock or speleothems from the ceiling of the caves. Such facies type has been described in other caves, and is considered as an autochthonous type of sediment (e.g., Bosch and White, 2004; White, 2007). The transport of material, in this case, is very short, as is confirmed by the angularity of collapsed blocks and poor sorting of the debris (Table 1; Figure 11A). Collapsed sediments are common near the cave entrances (e.g., Bočić et al., 2012; Haddad-Martim et al., 2017), although they can often be seen as big accumulations of unsorted boulders and cobbles on the floor of big chambers within caves (Fornós et al., 2014). Deposition is, besides other processes, often triggered by the cryofracturing process within cave channels during the cold periods (White, 2019). Hence, it can be assumed that sedimentation of the Bd facies in the channel of the LCC could also happen during a relatively cold period. That assumption is confirmed by palynological data which shows that the base of the investigated profile was formed during a period of cold and dry climate (see Chapter 4.4). The silty matrix within the Bd facies is the result of secondary infiltration (e.g., Martini, 2011) due to the existence of occasional and relatively insignificant water flow within the open cave channel. This is evidenced by the preserved/transformed phytoclast ratio according to Sebag et al. (2006). Even though this index was introduced by Sebag et al. (2006) for Holocene alluvial sediments, our data show a distribution of the index (Figure 10) which allows us to distinguish the paleoenvironment, confirming a transition from a fluvial paleoenvironment in the lower (older) part of the section (DC-SP 1—Bd facies) to a non-aquatic terrestrial paleoenvironment toward the youngest part of the section (DC-SP 8—Late Bronze Age). Such interpretation shows a good correlation with the sedimentological data. The appearance of microscopic-sized charcoal remains (Figure 7B) and possible traces of guano and bone remains (Table 2) between collapsed blocks of the Bd facies indicate that during the deposition of the Bd facies the cave channel was already fully developed and well connected to the surface in an open-air environment. Charcoal remains could indicate occurrences of wildfires in the area. Deposits of this type are formed during the vadose, air-filled stages of cave development (Hill, 1999). However, the distribution of palynofacies shown on Figure 10 in the lower part of the section, could also be a reflection of palynomorph transport by hydrological mechanism of cave drip water into the vadose cave channel.

Coevally to the Bd facies, the sediments of the Di facies were deposited (Figures 11A,D), as implied by their lateral contact. The red clayey silt resembles the Terra Rossa type soils and palaeosols whose genesis within the karst has not yet been unambiguously resolved. Polygenetic, detrital, and residual origins are most commonly mentioned (Durn et al., 2007). Similar red clayey-silty sediments in caves are interpreted as a composite of detritus introduced into the cave, and the insoluble residue from the dissolution of the host rock (e.g., Iacoviello and Martini, 2012). Mostly allogenic origin of the Di sediment within LCC is confirmed by mineralogy reflecting the composition of the river Otuča drainage area (Samples DC-SP 2, 3 in Table 2; Chapters 4.3 and 5.1). Although the vast majority of this interval is formed by finer sediment (Samples DC-SP 2, 3; Figure 4A), a significant number of larger clasts and bone fragments were found within (Figures 3A, 11A), could indicate depositional mechanisms by non-selective agents within cave specific environments. The aeolian contribution could also played important role in creating fine-grained sediment sequences within the cave. However, in the close vicinity of the LCC, there are no occurrences of loess on the surface (Figure 1C). Loess and loess-like deposits were widespread in the Adriatic region during oxygen isotope stage 3 (Wacha et al., 2018; Zhang et al., 2018). Nevertheless, aeolian sediments also play a significant role in the forming of polygenetic Terra Rossa type of soils (Durn et al., 2007) which were found in the river Otuča catchment area (Figure 1C, unit 9). According to paleontological data, it is necessary to mention that most probably intertwining of the depositional and biological processes resulted in today’s distribution of clasts and fossil remains within the sediment. Despite the fact that all parts of the U. spelaeus body are represented (Chapter 4.5), there is a lack of articulated sets and high fragmentation with indications of certain movements and destruction of the faunal material within the sampled DC-SP 2 interval of Di facies (Table 1; Figure 3). In density-mediated attrition, low-density elements are quickly destructed and removed from the assemblage (Lyman, 1994), which does not explain why there are very few fragments of major limb bones. The post-depositional disturbance is possible. However, based on data on body parts representation and fragmentation, it is possible that a larger dispersal of the skeletal material occurred even before, when bear carcasses were lying on the fossil surface of the cave. Except for a single questionable finding of cut marks (Trbojević Vukičević and Babić, 2008), no traces of butchering were previously found on cave bear bones from LCC, so it is safe to assume that hominins were not responsible for the accumulation and dispersal of cave bear bones. Since the cave has probably been used for a very long time by several generations of cave bears (Malez, 1965b), lack of articulated skeletons can be also explained as a result of trampling by other cave bears. However, high fragmentation and under-representation of major long bones may be due to scavenging activities of large carnivore predators, whose presence should not be ruled out despite the lack of corresponding taphonomic traces within the analyzed assemblage. Bears used caves as shelter and for hibernation, and were often targeted by other large carnivores, mostly hyenas, who entered deep into caves in search of them (Diedrich, 2012). To conclude, described lack of complete skeletons or articulated sets, as well as high fragmentation undoubtedly indicate certain movements and destruction of the faunal material within the sediment, although taphonomic processes could have a significant influence (Chapter 4.5). Given the chaotic arrangement of clasts and random orientation of long bones with no visible textures in the clayey silt, the facies of the red clayey silt is interpreted as Di facies (e.g., Bosch and White, 2004; White, 2007; Haddad-Martim et al., 2017). This type of sediment could be formed by high-density debris flows within the caves and by the redeposition of older clastic cave deposits. Distinguishing facies types within this type of deposits is not always unambiguous because cave sediments deposited near cave entrances are often transported by nonselective mechanisms such as slumping, creeping, and collapsing. Such processes result in the fact that they are mainly built of similar ratios of silt, clay and sand, which makes it difficult to recognize unconformities (Haddad-Martim et al., 2017). The chaotic character of the Di facies (see description in Table 1) and the possible redeposition of the sediments, is supported by the results of the palynological analysis which shows signs of redeposition, for example high (relative) values of phytoclasts fragments especially corroded charcoal remains (Figure 10). Deposition of the Di facies sediments took place during a period of a colder and drier climate (Chapter 4.4). Such occurrence of poorly sorted fossiliferous sediments coincides with the results of previous research within the LCC (e.g., Malez, 1960b). Namely, the horizon of the Di facies could be related to deposits earlier described in the caves of the nearby Gračac area as sediment composed of reddish-brown phosphate clays with numerous osteological remains of the Late Pleistocene with traces of the Upper Paleolithic culture (Malez, 1960b; Ivanović et al., 1976). Those sediments were correlated to the Last Glacial Period. Although cave bears are one of the most abundant taxon recovered from Pleistocene cave sites in Croatia (Miracle, 1991), with Cerovačke caves being one of the most important cave bear sites (Paunović et al., 1999), our paleontological data are limited by the relatively modest sample size. Therefore, U. spelaeus remains within the investigated sequence cannot be used as a precise stratigraphic marker. The earliest appearance of the U. spelaeus dates back to the end of the Middle-Late Pleistocene transition, and it became extinct in central Europe during the LGM, around 24 ka BP (Kurtén, 1958, 1968; Pacher and Stuart, 2009). Within Dinaric karst there are dated cave bear specific sites (e.g., Križna jama) with ages of cave bear thanatocenoses around 47–45 ka and >94 ka BP (Bosák et al., 2012). To establish a chronological framework of the Di facies of the LCC luminescence dating was performed. Although it is a new approach in the region, the application of luminescence dating techniques has proven to be a suitable chronometer in cave settings (e.g., Montanari et al., 2019). In general, luminescence techniques enable the determination of depositional ages of sediments by determining the point in time when quartz or potassium-rich feldspar grains were last exposed to daylight during transport before final deposition (for the basic principles of luminescence dating see Preusser et al., 2008, Rhodes 2011; Wintle 2008). According to the obtained luminescence age, the sediment directly overlying cave bear remains entered the cave environment 53.7 ± 6.9 ka ago (Table 3), which correlates well with oxygen isotope stage 3 (OIS3). The obtained date does not necessarily indicate the time of deposition at the investigated site but the time when the sediment entered the cave environment. It also indicates that this part of the sequence is not older than ∼54 ka. The end of deposition of the Di facies is marked by a change in color from reddish to gray to yellow and the appearance of a discontinuous flowstone level (Table 1; Figure 11C). Similar it is also described in other caves of the investigated area (reddish-brown phosphate clays phase covered with flowstone—Ivanović et al., 1976).

A thin layer of laminated flowstone composed of columnar sparry calcite was determined with visible crystal growth directions (Supplementary Material S2). Millimeter-sized sparry calcite possibly indicate a relatively high growth rate of the flowstone. We assumed in situ genesis of the flowstone because platy fragments (Figures 9A–C) are distributed at the same horizontal level with visible distinct uneven base/nucleation plane shoving traces of the underlying silty-clayey sediment (Supplementary Material S2). However, due to the thin and discontinuous horizon (Figure 3A), an allochthonous origin is not excluded. If in situ, the flowstone within the sedimentary profile is an indicator of change in the cave environment since their formation implies the absence of the detrital input (Haddad-Martim et al., 2017). Flowstones could be formed from sheets of flowing water derived from fissures or major conduits (Fairchild et al., 2006) or as subaqueous flowstones fed with turbulent underground stream water (Wróblewski et al., 2017). As valuable marker horizons for dating, they were previously used in this region as source material for acquiring ¹⁴C ages within caves (e.g., Bočić et al., 2012). Within this study the obtained ¹⁴C age from the flowstone is supported by the pIRIR₂₂₅ dating of the sediment horizon below the flowstone.

Comparing δ¹³C of the LCC speleothem sample (Table 4) of 5.3‰ to the Modrič cave speleothem (Rudzka et al., 2012) that has a mean value of −7.37‰ (2σ = 1.74‰), it could be concluded that the speleothem in the Modrič cave had a lower amount of DCF. Therefore, the most likely date for the LCC speleothem would be the OxCal Reservoir function date 17,949–17,562 cal BC (median 17,758 cal BC) for DCF = 15%. The formation of a relatively thin flowstone layer within the sedimentary profile in LCC, therefore, could be related to the known period-related phenomena in the border zone between the temperate Mediterranean and the periglacial/glacial parts of Europe (Surić and Juračić, 2010) when the formation of speleothems were rare or slow. Speleothem deposition ceased during the Last Glacial Maximum (LGM) in most of Europe and began again around 15 ka ago (Gascoyne, 1992; Lowe and Walker, 1997; Mihevc, 2001; Surić and Juračić, 2010). According to our data, the pronounced collapse processes and deposition of both Bd and Di facies sediments happened probably during local LGM, sometimes before ∼19.5 ka from today (¹⁴C age 17,949–17,562 cal BC—Table 4, Figure 11C), before postglacial warming commenced. The Dinaric mountains were glaciated during the Middle and Late Pleistocene (Hughes and Woodward, 2009). Therefore, periglacial influence on the LCC could be expected. If the flowstone is allochthonous (e.g., spallation from the ceiling), the obtained age could still be used as the oldest possible date when deposition of the upper section of the DC-SP profile commenced. Sedimentation in a calm aquatic environment therefore set on after ∼19.5 ka ago, as evidenced by the sedimentary filling of the cave channel which covers the flowstone marker horizon. This phase is characterized by gray to yellow laminated silt (Table 1; Figure 3A) which can be interpreted as slackwater deposit (Sw facies) (e.g., Bosch and White, 2004; White, 2007; Iacoviello and Martini, 2012; Ballesteros et al., 2017). These fine-grained sediments (Figure 4) entered the cave after flood events and settled during calmer conditions with very low flow velocities (often submerged cave channels). The allogenic origin of the Sw sediment is confirmed by the mineralogy, reflecting the composition of the river Otuča drainage area (Samples DC-SP 5–7 in Table 2; Chapters 4.3 and 5.1). Although the detritus was transported underground by turbulent flows rich in suspended sediment, the final depositional mechanism of this sediment type is very likely related to deposition from stagnant water suspension, as evidenced by the particle size characteristics and the lamination of the sediment. This is also confirmed by the fining upwards sequence observed in the upper part of the Sw facies interval (Figure 4B; Table 1). Slight reduction of the grain size from the base to the top indicating a decrease in the hydrodynamic energy conditions of the cave, is commonly observed within the Sw facies (Fornós et al., 2014). This is also confirmed by the palynofacies analysis, results of which indicate sedimentation in a calm palustrine environment (Figure 10). However, the laminated structure of the sediment in the lower part of the SW facies followed by a slight increase of the sand content (Figures 3A, 4B), described as vertical alterations of silty laminae and silty-sandy laminae (Table 1), could also resemble sediments deposited in stagnant hydrological conditions in regions prone to glaciation. Nearby areas of southern Velebit Mt. were prone to glaciation, documented with morphological moraine features and with glacial and glacio-fluvial sediments (Nikler, 1973; Krklec et al., 2015; Marjanac and Marjanac, 2016; Velić et al., 2017). The retreat of mountain glaciers in the southern Velebit Mt. is therefore dated roughly around 20.7–22.7 ka ago (Sarıkaya et al., 2020). Even so, there is no data about glacial chronology and processes in the very close vicinity of LCC. It should be mentioned that warming and deglaciation in the inland part of the Dinarides commenced mainly after 12.5 ka BP. It can be seen from increased speleothem growth in the Dinaric karst (Horvatinčić et al., 2003). Hence, we assume that it is highly possible to recognize glacial/periglacial influence on the deposition of Sw facies. According to our data, the Sw sediments were deposited during the Late Pleistocene (starting after ∼19.5 ka). In caves within glaciated regions, similar varve-like deposits are common (Ford and Williams, 2007). Even though the grain-size curves of the Sw facies (Figure 4A) resemble the varve-like sediments (Valen et al., 1997; Ford and Williams, 2007), the glacial-related depositional environment through the whole profile of the Sw facies sediments is ambiguous and needs further proof. Grain-size distribution curves themselves do not point to an unambiguous conclusion, it is necessary to compare them with known sediment ages to resolve the glacial-related origin of the laminae (Tischler et al., 2020).

The termination of allogenic siliciclastic sedimentation in an aquatic environment and consequently the end of the ponor function of this part of the LCC channel cannot be precisely dated within our research. Nevertheless, our data revealed a relatively young age in comparison to the known data within the wider area of the Dinaric karst where cave sediments cover the time span of the last ∼5 Ma years (Zupan Hajna et al., 2020; Zupan Hajna et al., 2021). We can conclude that this pronounced sedimentary environment shift is visible in the profile DC-SP with the onset of a Holocene Late Bronze Age layer (Tresić Pavičić, 2020). As known, the cessation of allogeneic sedimentation in caves is mostly controlled by tectonics and therefore related to changes in the hydrological regimes due to the separation of cave systems from active watercourses (Zupan Hajna et al., 2020). At Postojnska Cave, the sequence follows the series of events: formation of the fault, growth of the initial conduit due to groundwater circulation through the fault, infilling of the conduit with allogenic sediment, abandonment due to regional base level lowering, and continued motion along the fault (Sasowsky et al., 2003). Locally in LCC, it could be related to the permanent neotectonic uplift of this area (Prelogović, 1975). Especially pronounced uplift along the main NW–SE faults, e.g., along the Lika fault (Figure 1C), was recorded during the Pliocene and Quaternary (Prelogović, 1975). Maximum uplift of the local mountains was calculated of up to 1,200 m, with vertical shifts on individual faults averaging from 300 to 500 m (Prelogović, 1975). The sedimentary profile DC-SP, and the obtained data therefore possibly reveal the sedimentation history of the youngest inactive cave level within Mt. Crnopac. The same phases in the development of allogenic cave sediments occurred earlier in the higher cave levels within the mountain. It is evidenced with similar facies types (Sw) of detrital sediments which were found within the hypsometrically highest horizontal level of the todays CCS (Talaja and Kurečić, 2017). Considering that caves and their sediments are often related to the former base level and can conserve this information for long periods (Neuhuber et al., 2021), the average offset rate of local ongoing uplift/base level drop can be roughly estimated. Considering the relative displacement value of ∼80 m between today’s LCC entrance (at 624 m a.s.l.), and the recent active ponor phase on the polje level we have calculated the relative displacement rate. It is based on the obtained ¹⁴C age of 17,949–17,562 cal BC (before approximately 19.5 ka from today) of the flowstone strata identified within the DC-SP profile (Table 4). The displacement rate gives us a relatively high value up to ∼0.004 m per year. The calculated result is higher and not unambiguously comparable with fault slip rates calculated in the area of External Dinarides (Kastelic and Carafa, 2012). Therefore, due to scarce data without numerous variables, we believe that the computed rate could only be used as a rough estimation and basis for further research. Even if we use the combination of absolute and relative methods in order to get an accurate and robust age estimate (Häuselman et al., 2015), the critical problems identified are lack of data regarding the timing of the cessation of the allogenic sediment input. The archeological data suggests that the LCC was used for specific purposes during the Late Bronze Age (Tresić Pavičić, 2020), such as food storage and as a temporary dwelling in specific circumstances which surely indicate already ceased sediment input. However, due to the possible large time gap between the cessation of the Sw deposition and the onset of the Late Bronze Age, we cannot use that data for accurate calculations.