Due to the promising glacier archive conditions obtained from the shallow ice coring in 2004 (see Sect. 2.2), a full-depth ice-core drilling was completed on the western plateau from 27 August to 6 September 2009 (Mikhalenko, 2010). Bedrock was reached at the depth of 181.80 m. Drilling was done in a dry borehole using the lightweight electromechanical drilling system developed by Geotech Co. Ltd., Nagoya, Japan. Technical details of the drill are described in Takeuchi et al. (2004). The recovered ice cores were subjected to stratigraphic observations, packed into plastic sleeves, and stored in a snow pit with temperatures of -10 ∘C. Ice-core drilling was accompanied by borehole temperature measurements (using thermistor chains which were left for 3 days in the borehole and calibrated before and after the study with an error of ±0.1 ∘C), and snow pit sampling conducted 30 m to the south of the drilling site. The ice core was shipped in a frozen condition to the cold laboratory of the Lomonosov Moscow State University where detailed stratigraphic descriptions, including photographing each piece of the core, and bulk density measurements, were conducted.

In addition to the 2009 deep ice core, a subsequent 12 m ice core was extracted in June 2012 at the same site to expand the existing ice-core sample set from 2009 to 2012. The 2012 ice core was also used for the dust study of Kutuzov et al. (2013). Finally, in 27–30 June 2013, a 20.36 m ice core was recovered at the same location.

Stratigraphic descriptions of the ice core were carried out using transmitted-light illumination, resulting in 1 mm resolution details of the depths and thickness of individual horizons. The density of firn and ice were measured on 457 individual samples. Figure 5 shows the bulk density distribution with depth. The sharp random outliers from the general profile, especially with the lower values, could result from uncertainties in estimations of samples lengths. The uncertainty increases for the denser and smaller samples.

Ionic species such as ammonium (NH4+) succinate (HOOCCH2COO-, also denoted succinic acid) were investigated along the uppermost 157 m of the Elbrus core (122 mw.e.) with the aim of aiding in the sampling of ice core for alpine firn and ice (Legrand et al., 2007a). Pieces of firn and ice were decontaminated in a clean air bench located in a cold room using a pre-cleaned electric plane tool. A total of 3350 subsamples were obtained along the 157 m borehole, with a sample resolution decreasing from 10 cm at the top to 2 cm at 157 m depth.

For cations (Na+, K+, Mg2+, Ca2+, and NH4+), a Dionex ICS-1000 chromatograph equipped with a CS12 separator column was employed. For anions, a Dionex 600 equipped with an AS11 separator column was used with an eluent mixture of H2O, NaOH at 2.5, 100 mM, and CH3OH. A gradient pump system allows determining inorganic species (F-, Cl-, NO3-, and SO42-) as well as short-chain monocarboxylates (denoted MonoAc-) and dicarboxylates (denoted DiAc2-). For all investigated species, ion chromatography and ice-core decontamination blanks were found to be insignificant with respect to respective levels found in the ice-core samples.

As will be discussed in greater detail in Sect. 3.3.5, the search for volcanic horizons in the Elbrus ice cores requires examining the acidity (or alkalinity) of samples by evaluating the ionic balance between anions and cations where concentrations are expressed in micro-equivalents per liter, µEqL-1): [H+]=[F-]+[Cl-]+[NO3-]+[SO42-]+[MonoAc-]+[DiAc2-]-[Na+]+[K+]+[Mg2+]+[Ca2+]+[NH4+].

Within the present study, we focus (Sect. 3.3.4) on the NH4+ and succinic acid profiles, in order to (1) define a criterion which allows the separation of winter and summer snow deposition and (2) to apply this criterion to the first 157 m of the Elbrus ice core, in order to establish a depth–age relationship based on both annual layer counting and the NH4+ and succinic acid depth profiles.

The shallow 2012 and 2013 ice cores and the deep (down to 106.7 m) 2009 ice core were analyzed for deuterium–hydrogen (D / H) and oxygen (18O/16O) isotope ratios using a Picarro L1102-i instrument in the Climate and Environmental Research Laboratory (CERL), Arctic and Antarctic Research Institute, St. Petersburg, Russia. The instrument was calibrated on a regular basis with the isotopic standards V-SMOW, GISP, and SLAP provided by the International Atomic Energy Agency (IAEA) to estimate the precision of the measurements and to minimize the memory effect associated with continuous measurements. The reproducibility of the measurements was ±0.07 ‰ for oxygen isotopes (δ18O) and ±0.3 ‰ for deuterium (δD). The CERL laboratory work standard SPB was measured after every five samples. The δ18O and δD values are expressed in ‰ units relative to the V-SMOW value.

Figure 6 shows the vertical temperature profile measured along the 181 m long borehole drilled in 2009. Temperatures ranged from -17 ∘C at 10 m depth to -2.4 ∘C at 181.8 m. The temperature profile can be divided into three parts based on different temperature gradients: from the surface down to 10 m, from 10 to 100 m, and from 100 m to the glacier bottom. The upper section of the temperature profile reflects seasonal changes at the surface. The borehole temperature ranges from -17 to -12 ∘C between 10 and 100 m, and most accurately reflects past temperature fluctuations. Temperature changes are almost rectilinear from 100 m depth to the glacier bottom and provide evidence of a steady heat transfer regime. The heat flux of 0.34 Wm-2 at the glacier bottom was calculated from the measured temperature gradient and the coefficient of the heat conductivity of ice (2.25 Wm-2). This value is 4–5 times higher than the average heat flux density for the Earth's surface and higher than the mean value for central Caucasus, and may be associated with a heat magma chamber of the Elbrus volcano. Figure 6 also shows the temperature profile measured in the 19 m borehole in 2013 and temperature records obtained in 2004 after the 22 m depth shallow ice-core drilling on the western plateau (Mikhalenko et al., 2005). The good match between records is indicative of the stable temperature regime on the western Elbrus plateau for the last decade.

Using the altitudinal temperature gradient estimated in Sect. 3.1 based on western plateau AWS temperature data close to the 2009 drill site and the low-elevation station Mineralnye Vody, we estimate that the annual mean air temperature at the drill site is approximately -19 ∘C. This value is close to the mean annual air temperature of -19.4 ∘C, calculated using the general relationship with the ice temperature at the bottom of the active layer (Zagorodnov et al., 2006) and only slightly enhanced compared to the 10 m firn temperatures.

The measured temperature profile shows that basal melting can occur due to ice pressure at the deepest part of the glacier. Potential bottom melting has been estimated using a mathematical model of temperature regimes (Salamatin et al., 2001). Our modeling results demonstrate that basal melting occurs under ice thicknesses of more than 220 m, but that its value does not exceed 10 mmw.e.a-1.

The bulk density profile suggests a change in densification around the critical densities (Maeno and Ebinuma, 1983) of 550 and 840 kgm-3, and no visible change at 730 kgm-3, which is consistent with other analyses of density profiles (Hörhold et al., 2011; Ligtenberg et al., 2011). However, the slight decreasing trend in density at depths below the maximum values at ∼ 80 m (Fig. 5), close to the critical density across the whole depth interval, is unlikely to be a systematic error in measurements and needs further investigation. Future research should account for the ice flow characterization and the possible effects related to the “intervening depth interval” (the alternation of layers which have already reached the close-off density, with layers that are still permeable) due to seasonal (Bender et al., 1997) or wind-induced (due to seasonal differences in wind speeds) snow density variability at high accumulation sites. Unlike polar ice cores where the “intervening depth interval” is just a fraction of the whole length of the ice core (Bender et al., 1997), the measured bulk density in the Elbrus ice core spans a wide interval between 800 and 915 kgm-3 towards the bottom of the glacier (Fig. 5). Comparing Elbrus's density profile with the results from the Salamatin et al. (2009) densification model demonstrates that there is an increase in the accumulation rate over the past several years. The minimum deviation between simulated and measured ice-core density profiles occurred when the accumulation history was assumed to be similar to the long-term precipitation changes observed at meteorological stations (Nosenko et al., 2013).

According to the morphogenetic classification of stratigraphic features (Arkhipov et al., 2001), two distinct types of layering were observed in the core: firn layers which have not been affected by melting, and ice layers formed by the refreezing of meltwater in the surface snow. The thickness of the infiltration ice layers, which do not form every year, does not exceed 10 mm. Ice formation occurs in cold, dry conditions, as already concluded on the basis of borehole and air temperatures at the drill site (see Sects. 3.1, 3.3.2). The pore close-off depth occurs at around 55 m, where the air bubbles separate from the surrounding ice matrix. This depth coincides with a measured bulk density of around 840 kgm-3. This density is consistent with the presence of ice layers, as these layers increase the close-off density value above what is expected in ice in which no melting occurs (i.e., 830 kgm-3).

The seasonal cycle of the isotopic composition is detectable over the entire measured part of the core (Fig. 7). Mean seasonal values of δD are -200 ‰ for the winter and -25 ‰ for the summer. Values of δ18O are about -5 to -10 ‰ in summer and -30 ‰ in winter. Using isotope values to determine annual cycles over 106.7 m of the ice core suggests that this 106.7 m covers 86 years (AD 1924–2009). The mean accumulation rate for this period, based on this dating and taking the firn density and layer thinning into account, was 1455 mmw.e. Figure 7 shows results of isotopic measurements of four different ice cores obtained from the western Elbrus plateau. While 2009, 2012, and 2013 cores were obtained from almost the same location; the 2004 core was drilled 120 m to the southwest. Good agreement in isotopic variations of all cores suggests a relatively homogeneous snow deposition on the plateau.

We used the isotope diffusion model of Johnsen et al. (2000) to estimate the preservation of the isotopic signal with diffusive smoothing. Although the drilling site is located in a relatively warm place (-17 ∘C), the high snow accumulation rate does not favor strong diffusion, since any firn layer rapidly reaches the pore close-off depth. The maximum “diffusion length” at the pore close-off depth is estimated as 5 cm in ice equivalent (IE). The effective diffusion length may be even smaller if we take into account the ice lenses in the firn that prevent the vertical exchange of the water molecules.

Such a diffusion length means that all oscillations shorter than 13 cm IE, will be completely erased due to the diffusion, while oscillations between 13 and 70 cm IE will survive but will be damped to some extent, and the cycles longer than 70 cm (e.g., the annual cycle) IE will not be affected by the diffusion. Thus, if during a single snowfall results in a 35 cm snow layer (that corresponds to 13 cm IE), the isotopic signal of this layer will survive during the diffusion processes and will be seen in the ice core.

Diffusion occurs below the pore close-off depth, but the in ice is much slower than in firn. The final diffusion length solely depends on the time and temperature of the firn-ice thickness. Even if we take a maximum possible temperature (-2.4 ∘C) and an age estimate of a few hundred years, the additional diffusion in ice will still be very small. This combination leads us to an important conclusion; we may expect to obtain seasonal cycles in the isotope profile down to the very bottom of the core, and our ability to detect the annual cycle in the core depend on the sampling resolution, as well as on such basal processes such as layer folding and mixing.

We attempted to date the core by counting annual layers based on chemical ice-core stratigraphical records, as we previously successfully applied such layer counting to mid-latitude Alpine ice cores using the NH4+ signal (Preunkert et al., 2000). Since NH4+ experiences strong maximum emissions in phase with strengthened summer upward transport of air masses, a particularly well-pronounced seasonal cycle is expected, such as observed at the Col du Dôme Alpine site (Preunkert et al., 2000; Fagerli et al., 2007). However, it appears that the NH4+ seasonal cycle at Elbrus is less pronounced than in the Alps. Whereas recent summer NH4+ levels are comparable at both sites, recent winter concentrations at Elbrus are significantly higher than at Col du Dôme.

The first study regarding the seasonality of Elbrus snow accumulation was conducted by Kutuzov et al. (2013) along a short firn core spanning the years 2012–2009. Based on the dust layer stratigraphy of absolute-dated dust events and the stable isotope record the authors showed that the annual deposition at Elbrus has a mean δ18O signature of -15 ‰. The isotopic signature consists of nearly equal deposition amounts from the warm season (45 % of total accumulation), where δ18O values vary between -5.5 and -10 ‰, and from the cold season (55 % of total accumulation), for which values vary between -17 and -27 ‰.

The concentration distribution of NH4+ values was inspected in recent firn layers (0–12 mw.e.), and the 50 % concentration limit of 100 ppb was a first approach to separate snow deposition arriving from summer and winter precipitation at Elbrus. However this criterion is not conserved in time as the NH4+ sources are mainly anthropogenic in origin, with an expected NH4+ trend in summer as well as in winter over the last 100 years. Therefore, a second criterion was used to confirm our winter snow selection. This criterion used succinic acid, a light dicarboxylic acid with an observed strong summer maximum; and a winter level is almost non-existent in the current atmosphere in Europe (Legrand et al., 2007b). The very low winter levels are related to the absence of source as this species is mainly photochemically produced from biogenic precursors. The concentration distribution of succinate values was inspected in recent firn layers (0–12 mw.e.), and the 50 % concentration limit of 5 ppb separates snow deposition arriving from summer and winter precipitation at Elbrus. Winter snow and ice layers were identified when both ammonium and succinate criteria were fulfilled for more than two successive samples.

Figure 8a, shows the result of this data dissection over the uppermost 12 mw.e. along with the δ18O record (Fig. 8e). The mean δ18O level of selected winter data is -19.6 ‰, and as demonstrated in Fig. 8a and c, the winter season selected from ammonium and succinate concentrations matches with winter sections deduced from the δ18O profile quite well. However, when examining the δ18O variability compared to the major ions, it appears that sometimes the spring season or even the beginning of the summer season may be included. For dating by annual layer counting, this shortcoming is not critical; however, if the data set is inspected at seasonal resolution this definition of the spring season might be a handicap. In this case a stronger criteria (NH4+ < 50 ppb and succinate < 3 ppb) may be applied in addition to ensure that only deposition corresponding to winter precipitation and associated atmospheric background conditions are selected within the winter period. The mean δ18O level of winter data selected using this criteria is -21.1 ‰, whereas seen in Fig. 8a, c, and e, this selection may omit some winter seasons.

Examination of NH4+ and succinate minima below a depth of 12 m contrasts with results from the European Alps, where in Elbrus the NH4+ winter level decreases significantly from near the surface to around 70 mw.e. depth (see Fig. 8). Therefore, the NH4+ winter and background criteria were adjusted using a winter (background) threshold of 50 ppb (30 ppb) from 52 to 62 mw.e. of the core and 30 ppb (20 ppb) from below a core depth of 62 mw.e. In contrast, the succinate winter levels did not change and the 5 ppb criterion applied in recent times was also applied in deeper layers. Figure 8b, d, and f showed a comparison of NH4+, succinate with the δ18O record between 75 and 78 mw.e. (i.e., from AD 1931 to 1928). The winter criteria match well with recent winter deposition as deduced from the stable isotope content, although the stable isotope record tends to already be a bit smoothed compared to the uppermost firn layers. As observed for the uppermost core section, we cannot exclude that the winter criteria include parts of an intermediate season, whereas the background criteria select only deposition arriving from the coldest precipitation.

Figure 9 shows the result of the dating of the Elbrus ice core. In addition to model calculations detailed in Sect. 3.4, the depth–age scale obtained by annual layer counting using the NH4+ and succinate criteria is reported down to 122 mw.e. Annual layer counting was achieved as described above down to 85 mw.e. Below 85 mw.e., winter levels became rather thin due to annual layer thinning but also likely due to upstream effects as commonly encountered on such small-scale glaciers (Preunkert et al., 2000). Therefore, below 85 mw.e. ice-core layers, in which fewer than three samples had reached the winter criteria were considered as winter seasons, and from 113 to 122 mw.e. winter layers were also assigned when only one of the two criteria was fulfilled for at least one sample, while the other sample showed only a relative minimum that sometimes exceeded the fixed threshold. This lack of sharp minima could be either due to the fact that winter sections become smaller than our sampling resolution of 2–3 cm applied to core depths below 90 mw.e., and/or could be the result of an incomplete precipitation preservation due to wind erosion upstream of the borehole as observed on other small-scale Alpine glacier sites (e.g., Preunkert et al., 2000). In this latter case a systematic lack of winter snow accumulation would occur in the deeper ice-core layers.

Dating based on annual layer counting of the chemical stratigraphy agrees fairly well with the AD 1963 tritium time horizon located at the core depth of 50.7 mw.e. and which is dated at AD 1965 using the ammonium stratigraphy; Fig. 10a. The layer counting results fit well with the dating achieved to 106.7 m based on the seasonal stratigraphy of the stable isotope profile. Whereas stable isotopes predict the year AD 1924 at a core depth of 106.7 m, the chemical stratigraphy leads to an estimate of the year AD 1926 at this depth.

To anchor the depth–age relationship with further absolute time horizons, we inspected the sulfate profile to identify volcanic horizons such as found in other northern hemispheric ice cores between AD 1912 (Katmai) and AD 1783 (Laki) in Greenland (Legrand et al., 1997; Clausen et al., 1997) and at Colle Gnifetti (Bohleber 2008). However, since Elbrus is a volcanic crater, it is sometimes difficult to attribute a peak either to a well-known global eruption or to a local event. Furthermore, numerous sulfate peaks in the Elbrus ice core originate from terrestrial input as suggested by the presence of concomitant calcium peaks. The Katmai eruption in AD 1912 could be clearly identified at 87.7 mw.e. (dated at AD 1911 using the ammonium stratigraphy), with several neighboring samples showing relatively high sulfate levels (up to 1200 ppb, i.e., 25 µEqL-1) compared to sulfate peaks generally present in summer layers in the early 20th century. Furthermore, in contrast to neighboring summer sulfate peaks located at 87.2, 87.4, 88.0, and 89.3 mw.e. that are alkaline (see Fig. 10b), the acidity of samples of the 87.7 mw.e. sulfate peak reaches 8 µEqL-1. Furthermore, samples located at the top part of the 87.7 mw.e. sulfate peak remain neutral in spite of a large presence of calcium. Figure 11 demonstrates that the 1-year uncertainty in dating this horizon is in excellent agreement with our annual counting.

We were still able to easily examine annual counting down to 113 mw.e. resulting in a date of AD 1860, but below this depth the dating becomes more uncertain (see the green line in Fig. 9). Below 88 mw.e., seven significant potential volcano horizons are suspected based on their ionic balance and sulfate levels (not shown), from which at least one horizon is of local origin as suggested by small stones up to 1–2 mm in the corresponding ice layer. Nevertheless, a series of three narrow ionic spikes occur at 118–120 m w.e. (dated at ∼ AD 1840–1833), where two of the spikes are characterized by an increase of sulfate and acidity (up to 7.8 µEqL-1, not shown) that may be related to the well-known eruptions observed in Greenland that are dated to AD 1840±2 years. One of these eruptions may possibly be due to the Coseguina eruption in AD 1835 (Legrand et al., 1997).

We calculated the depth–age relationship from the depths and thicknesses of the counted annual layers (Fig. 9). Despite high variability in the annual layers' thickness, the data demonstrate layer thinning with depth related to the ice flow. Applying the thickness-age relationship developed by Nye (Dansgaard and Johnsen, 1969) to the actual annual layer data (Fig. 9) provides the mean accumulation over the whole studied time period in the ice core resulting in 1.583 m of ice equivalent per year. The Nye curve corresponds to the depth–age relationship from Nye's model with a best-fit (constant over time) accumulation rate and the glacier thickness at the drilling site (Fig. 9). The green line is the depth–age relationship as suggested by Salamatin's model (Salamatin et al., 2000) with the same best-fit accumulation rate and drill site basal and surface descriptions as assumed when applying Nye's model.

Dating based on annual layer counting of the chemical stratigraphy agrees well with the AD 1963 time horizon that is located at the core depth of 50.7 mw.e. (dated at 1965 using the ammonium stratigraphy, Fig. 10a). In addition it fits very well with the dating achieved so far (i.e., core down to 106.7 m) on the base of the seasonal stratigraphy of the stable isotope profile. Whereas stable isotopes predict the year 1924 at a core depth of 106.7 m, the chemical stratigraphy leads to an estimation of the year 1926 in this depth.

To anchor the depth age relation with further absolute time horizons, a first inspection of the sulfate profile was made with the view to identify volcanic horizons as found in other northern hemispheric ice cores between 1912 (Katmai) and 1783 (Laki eruption) in Greenland (Legrand et al., 1997; Clausen et al., 1997) and at Colle Gnifetti (Bohleber, 2008). However since the Elbrus is an in active volcanic crater, it is sometimes difficult to attribute a peak either to a well-known global eruption or to a local event. Furthermore, numerous sulfate peaks in the Elbrus ice core originate from terrestrial inputs as suggested by the presence of concomitant calcium peaks. So far, the Katmai eruption in 1912 could be clearly identified at 87.7 mw.e. (dated at 1911 using the ammonium stratigraphy) with several neighboring samples showing relatively high sulfate levels (up to 1200 ppb, i.e., 25 µEqL-1) compared to those seen in sulfate peaks generally present in summer layers of the early 20th century. Furthermore, as seen in Fig. 10b, in contrast to neighboring summer sulfate peaks located at 87.2, 87.4, 88.0, and 89.3 mw.e., that are alkaline (see Fig. 10b), the acidity of samples of the 87.7 mw.e. sulfate peak reaches 8 µEqL-1 at the bottom part of the sulfate peak. Furthermore, samples located at the top part of the 87.7 mw.e. sulfate peak remain neutral in spite of a large presence of calcium (similar to those seen in neighboring summer sulfate peaks). As seen in Fig. 9 it appears that within 1-year uncertainty, this horizon is in excellent agreement with our annual counting.

Below 88 mw.e., we were still able to easily proceed annual counting down to 113 (1860); whereas further down, the dating become more uncertain (see the green line in Fig. 9). Below 88 mw.e., seven significant potential volcano horizons can be suspected on the basis of the ionic balance and sulfate levels (not shown), from which however at least one are of local origin (as suggested by small stones with size of up to 1–2 mm were found in the corresponding layer). Nevertheless, a series of three narrow spikes was located at 118–120 mw.e. (dated at around 1840–1833) among which two that are characterized by an increase of sulfate and acidity (up to 7.8 µEqL-1, not shown) may be related to the well-known eruptions observed in Greenland in a time distance of 2 years around 1840 (one of them being possibly due to the Coseguina eruption in 1835) (Legrand et al., 1997).