Environmental imprints of climate changes and anthropogenic activities in the Ore Mountains of Bohemia (Central Europe) since 13 cal. kyr BP

Abstract

We have measured pollen, aluminum, and lead abundances in a Czech peat bog at Bozi Dar in order to investigate the environmental impact of Holocene climate changes and mining activities in Central Europe. The pollen record shows a continuous vegetal cover since 13 kyr BP. Aluminum and pollen deposition characterize the Younger Dryas (YD) cold event at 12.3–11.0 kyr BP, within its expected time frame in Northern Europe. Aluminum fluxes reveal four other significant dust events with significant Asian and Saharan sources, as shown with variations in stable ²⁰⁴Pb, ²⁰⁶Pb, and ²⁰⁷Pb isotopes. Recurrent Asian incursions are seen at Bozi Dar during the YD and Mid-Holocene dust events (at 7.1, 6.5, and 5.9 kyr BP). These Asian imprints are also observed by Le Roux and colleagues during the YD in the nearest Holocene-long geochemical peat record from Switzerland (Etang de la Gruère). We do not see at Bozi Dar the distinct western European volcanic inputs identified in the Swiss peat by Shotyk and colleagues. However, isotopic ratios characteristics of Mediterranean airflows are evident at both locations during the Mid- and Late Holocene. The occurrence of these various air mass sources in Switzerland and the Czech Republic substantiates Holocene atmospheric circulation models in Central Europe with dominating zonal airflows during the Early Holocene, followed by a governing meridian dispersion during the Mid- and Late Holocene. Lead inputs at 5.1–4.3 and 1.9 kyr BP cannot be explained by increased dust deposition only. The 1.9 kyr event fits local Bohemian ore isotopic imprints of which isotopic signatures are seen as early as 2.2 kyr BP in our core, hence revealing the oldest ever environmental record of mining-related activities in the Czech Republic. Relatively elevated Pb deposition at 5.1 and 4.3 kyr BP are tentatively attributed to contamination that may be from local sources and/or long-range transport.

Keywords

Asia Central Europe dust Holocene lead isotopes Mediterranean mining peat bog pollen

Introduction

Continental climate changes and environmental imprints of human activities during the Holocene are generally identified with long-term sediment, ice, or speleothem records. Among those, ice and peat bog cores, when available, provide more direct and specific records of airborne trace elements, and their stratigraphy is less perturbed. The duration and regional/global extension of these natural or anthropogenic perturbations are best characterized with a multi-proxy approach. In particular, the combination of high-resolution stratigraphy and biogeochemical markers improves characterization of abrupt climate changes, whose duration and threshold intensity are marked, versus other parts of the record that indicate more constant conditions (Alley et al., 2003; Clark et al., 2002; Trenberth, 1997; Trenberth et al., 2004). The former includes one of the most intensively chronicled climate crises during the Upper Pleistocene, the Younger Dryas (YD) cold event, which occurred between 13.0 and 11.5 kyr BP and extended throughout continental Europe (e.g. Alley et al., 1993; Berger, 1990; Broecker, 2003; Clark et al., 2001, 2002; Dansgaard et al., 1989; Peteet, 1992, 1995; Severinghaus and Brook, 1999; Wright, 1989; Wunsch, 2006). That crisis was triggered by meltwater discharge from the Laurentide ice sheet drainage, which caused a partial shutdown of the Meridional Overturning Circulation (MOC; Hughen et al., 1998; Lie and Paasche, 2006; Muscheler et al., 2000) that substantially controls the thermal mass balance of the North Atlantic regions (e.g. Bond et al., 1997; Broecker, 2003; Broecker and Denton, 1989; Fairbanks, 1989; Keigwin et al., 1991; Lehman et al., 1991). Vegetation cover, as characterized by relative pollen abundances, allows identification of other continental past climate episodes, with some being much shorter than the YD event, like the abrupt 8.2 kyr extreme cold episode that has generated drier and colder weather conditions for 200–400 years in Europe (e.g. Allen et al., 1999; Genty et al., 2003, 2005; Veski et al., 2004). Pollen also provide indices for early human activities related to deforestation and/or agricultural practices (e.g. Davis et al., 2003; Heiri et al., 2003; Magny and Begeot, 2004; Magny et al., 2003; Prasad et al., 2009; Rosén et al., 2001; Sarmaja-Korjonen and Seppä, 2007; Seppä and Poska, 2004; Seppä et al., 2005, 2007; Snowball et al., 2002; Spurk et al., 2002; Tinner and Lotter, 2001). Still there are uncertainties in the spatial and temporal distribution of these Holocene climate and early anthropogenic activities as recorded with pollen analyses. Here, we combine pollen abundances along with aluminum (Al) concentrations and lead (Pb) isotopic compositions that are efficient proxies for atmospheric dust and aerosol sources, respectively. Aluminum is a major crustal component that provides a measure of atmospheric dust loads, whose variations between cold and warm spikes are indicative of atmospheric conditions in connection with the extension of desert (lack of vegetation cover) or emerged lands caused by sea-level changes during cold events. Variations in ratios of stable Pb isotopes (²⁰⁴Pb, ²⁰⁶Pb, ²⁰⁷Pb, ²⁰⁸Pb; the last three are end members of uranium–thorium decay chains) can be used to discriminate between various natural (e.g. diverse wind-eroded crustal bedrocks and volcanic emissions) and anthropogenic (e.g. different metallogenous mines) sources because of the specific U–Th content of different crustal reservoirs and ores. Isotopic imprints are preserved during smelting and/or the erosion of geological materials, hence providing an accurate tool to identify the geographical origin of airborne deposited dust and anthropogenic imprints associated with mining/smelting activities in continental reservoirs (Abouchami and Zabel, 2003; Alfonso et al., 2001; Baron et al., 2005, 2009; Brännvall et al., 1997, 2001; Chambers et al., 2012; Chow and Patterson, 1962; De Vleeschouweret al., 2007, 2010; Dunlap et al., 1999; Jones et al., 2000; Jouffroy-Bapicot et al., 2007; Kempter and Frenzel, 2000; Kylander et al., 2005, 2010; Le Roux et al., 2004, 2012; Marcantonio et al., 2002; Martinez-Cortizas et al., 2002; Monna et al., 2004; Novak et al., 2003; Renberg et al., 2001, 2002; Shirahata et al., 1980; Shotyk et al., 1998, 2001; Thevenon et al., 2011).

We have measured pollen abundances, Al concentrations, Pb concentrations and corresponding isotopic compositions in a Holocene-long peat bog from the Czech Republic (Central Europe) in order to assess (1) the inland continental signature of ocean-driven climate events and associated atmospheric circulation patterns; (2) the imprints of volcanic episodes possibly deposited in Central Europe from the main known sources in the Faroe Islands, Iceland, and the French Massif Central; and (3) the earliest discernible environmental imprint of mining/smelting activities in the northwestern Czech Republic, which is one of the richest Ag–Pb European mining districts and where mining has been well documented since the 14th century ce.

Two of the main difficulties in this type of study are to ascertain the geogenic provenance of dust from different sources and whether or not the environmental disturbances recorded within the peat bogs originate from local or distant sources. Here, we shall make use of the most discriminating of Pb, ²⁰⁴Pb, along with ²⁰⁶Pb and ²⁰⁷Pb. This approach is often precluded in peat bog investigations because of the difficulty in accurately measuring ²⁰⁴Pb in environmental samples with low Pb concentrations (Doe, 1970; Kylander et al., 2005, 2010). ²⁰⁴Pb, along with ²⁰⁶Pb and ²⁰⁷Pb, has been commonly used to characterize volcanic and crustal imprints, as well as specific Pb ore signatures in Europe (see references in Allan et al., 2013; Baron et al., 2006; Biscaye et al., 1997, 2002; Garçon et al., 2012; Holm et al., 2001; Kylander et al., 2005, 2010; Stos-Gale et al., 1995, 1996; Thirlwall et al., 2004; Wittig et al., 2007).

This long peat core also provides the opportunity to compare our Czech Republic geochemical record with the other nearest Holocene-long Pb and dust record from Etang de la Gruère in the Jura Mountains, Switzerland (Le Roux et al., 2012; Shotyk et al., 1998; Figure 1). We attempt to corroborate the two data sets and possibly extend findings regarding the occurrence of climate episodes along with the geographical source of deposited aerosols in Central Europe, in particular the occurrence of volcanic events, incursions of Mediterranean and Asian air masses, and the early geochemical record of mining and smelting activities.

Figure 1.

(a) Localization of the Etang de la Gruère (Shotyk et al., 1998) and (b) the Bozi Dar (this study) peat cores. Various sized and gray circles represent the oldest Holocene peat cores analyzed for lead (Pb) and its stable isotopes in Europe (12–6 kyr BP according to the caption in the figure). This figure has been modified after De Vleeschouwer et al. (2010).

Materials and methods

Study site and sampling

The Bozi Dar peat bog (15°23′N, 12°54′E) is a rain-fed sphagnum-dominated, mountain-top peatland situated in Ore Mountains (Krusné hory) of Bohemia in the northern Czech Republic (Figure 1). It is a large upland moor located on a watershed that covers an area of 105 ha at the elevation of 993–1035 m. It lies in a moderately warm to cold region with a mean annual temperature of 4.1°C and an average annual rainfall of 1250 mm (Dohnal et al., 1965). The bedrock is composed of micashists and paragneisses, with Neogene nephelinite extrusive rocks covering part of the area. The total thickness of the peat is 3.0 m. The vegetation is dominated by Norway spruce (Picea abies), dwarf pine (Pinus mugo spp. rotundata), European beech (Fagus sylvatica), and dwarf birch (Betula nana). Large sphagnum mats are found in open segments of the bog. The most abundant sphagnum species are Sphagnum riparium Angstr., Sphagnum fallax v. Klinggr., and Sphagnum girgensonii Russ. The herbaceous layer includes Carex limosa, Oxycoccus palustris, Vaccinium myrtillus, Sedum villosum, Pedicularis palustris, Drosera rotundifolia, Arnica montana, and Swertia perennis. The peat is sphagnum-dominated down to 143 cm, and then eriophorum-carex-dominated (Brizova, 1993). This peat was mined between the 18th century and the early 1960s. In 1540 ce, a 13-km-long channel was excavated to drain bog surface water, and it was used until 1945 to provide water to an ore processing plant in Horni Blatna. In the vicinity of Bozi Dar, spoil from small mining Sn–Ag operation has been dated back from the 17th to the 19th centuries.

A vertical peat profile was sampled in an intact segment of the peat bog outside the area of historical peat mining in July 1992. A 15-cm-thick layer of relatively dry peat was removed from a 3-m-high cut, and discarded. Compact, dark-colored peat was sampled from the cleaned profile in an upward direction using a plastic shovel for Pb determinations. Approximately 10 g of peat was taken at 5-cm intervals (43 samples in all). Since the uppermost 25 cm of the peat profile may have been disturbed, it was not sampled. Peat samples for Pb analysis were stored in zip-lock™ bags. A full 3-m-deep peat profile for pollen analysis was sampled using metal boxes (10 × 10 × 50 cm³; no peat monoliths could be collected from the depth of 0–80 cm because the substrate was too dry). Sub-samples for pollen analysis were collected with a spatula from peat monoliths at the same intervals as those collected for Pb analysis (43 samples).

Analyses

Pollen was extracted from the peat by HF treatment and subsequent acetolysis (Erdtman, 1954; Overbeck, 1958). The resulting pollen mixtures were placed on a smear glass along with glycerin, ethanol, and water. Over 500 arboreal pollen grains and usually a lower number of non-arboreal pollen grains were counted from each peat section. Pollen abundances are shown in Figure 2.

Figure 2.

General vegetation-type phases (Firbas, 1949/1952) and main climate periods (A to E) according to the detailed percentage pollen diagram at Bozi Dar.

Radiocarbon dating of seven depth levels (0.90, 0.95, 1.55, 1.75, 2.0, 2.35, 2.6, and 3.0 m below surface) was made at the Gliwice and Hannover Geochronological Laboratories. The data were calibrated for variable initial radiocarbon concentrations using the OxCal v4.1 and Clam programs (Bronk Ramsey, 2009; Reimer et al., 2009). The median calibrated ages are listed in Table 1, with 2σ calibrated BP ages and relative areas for each ¹⁴C analyzed sample. Calibrated ages show a linear correlation (R² = 0.97) with depth (Figure 3) suggesting a fairly continuous deposition with no observable hiatus. This correlation can be used to calculate accumulation rates (ARs) that are necessary for Al and Pb deposition flux calculations. However, we do not have ¹⁴C age determination between 100 and 150 cm (Figure 3) to outline accurate AR. Therefore, we established an age model curve on the basis of calibrated ¹⁴C dates in order to assess proper ages and AR thorough the length of the core using the CLAM program (under R programming language, version 2.10.0, copyright^©2009, The R Foundation for Statistical Computing; Blaauw, 2010).

Table 1.

Calibrated ¹⁴C ages for the Bozi Dar peat core.

Depth (cm)	Age (14C yr BP)	2σ cal. BP age (relative area)	Median cal. Age (cal. BP)
300	11,240 ± 290	12,668–13,748 (100%)	13,155
260	10,340 ± 180	11,598–12,772 (96.6%) 11,473–11,555 (2.4%) 1407–11,443 (1%)	12,167
235	8400 ± 90	9200–9539 (96.1%) 9138–9179 (3.9%)	9401
200	7870 ± 110	8452–8998 (100%)	8718
175	6830 ± 100	7553–7867 (94.4%) 7508–7545 (3.7%) 7899–7924 (1.9%)	7680
155	5860 ± 95	6443–6902 (100%)	6675
95	2140 ± 80	1966–2332 (98.2%) 1950–1964 (1.8%)	2134
90	1800 ± 30	1690–1820 (86.9%) 1626–1668 (13.1%)	1736

Figure 3.

Linear correlation of calibrated ¹⁴C ages versus depth in Bozi Dar peat core.

Modeled ARs vary from 0.013 to 0.033 cm/yr (with the exception of the top, where ¹⁴C date determination may be hampered because of removal of a dried section, mining of the peat for the past 300 years, and surface runoff) with a mean AR of 0.020 ± 0.006 cm/yr (n = 40, 2σ). The lower AR are encountered between 100 and 155 cm (0.014 ± 0.001 cm/yr, n = 10, 2σ) but are close to the mean calculated AR with no noticeable hiatus. Numerical ages are presented in Table 2 along with geochemical data for each sampled depth. They are used in the discussion to define the chronology of encountered pollen, dust, and anthropogenic episodes.

Table 2.

Al, Pb concentrations (ppm), and corresponding deposition fluxes (µg/cm² and ng/cm²/yr for Al and Pb, respectively), Pb isotopic signatures, as a function of depth, calculated ages (see age model in the method section), and ash content (%) in the Bozi Dar peat core.

Depth (cm)	Age (kyr BP)	Ash (wt %)	Al (ppm)	Pb (ppm)	Al flux (µg/cm²/yr)	Pb flux (ng/cm²/yr)	²⁰⁶Pb/²⁰⁷Pb	²⁰⁶Pb/²⁰⁴Pb	²⁰⁷Pb/²⁰⁴Pb
30	0.14	1.34	1510	0.66	13.68	5.98	1.1860	18.553	15.644
50	0.44	0.99	nd	0.94	nd	4.34	1.1835	18.515	15.646
60	0.67	1.40	6021	1.38	20.82	4.77	1.1854	18.545	15.645
80	1.39	1.12	1585	1.90	3.34	4.01	nd	nd	nd
85	1.63	4.64	1219	2.40	2.31	4.54	1.1820	18.429	15.592
90	1.89	4.41	774	3.48	1.32	5.94	1.1734	18.298	15.591
95	2.18	3.62	1124	1.35	1.75	2.10	1.1771	18.415	15.644
100	2.50	2.79	1940	1.02	2.81	1.48	1.1929	18.608	15.600
105	2.84	4.77	4516	2.75	6.19	3.77	1.2046	18.886	15.678
110	3.20	3.17	2761	1.35	3.63	1.77	1.1983	18.800	15.689
115	3.57	2.14	1728	0.83	2.20	1.06	nd	nd	nd
120	3.95	1.72	1611	nd	2.02	nd	1.2029	18.855	15.677
125	4.33	1.94	1675	2.54	2.08	3.15	1.1952	18.641	15.599
130	4.72	2.50	1489	0.96	1.85	1.20	nd	nd	nd
135	5.10	3.10	1696	3.70	2.14	4.67	1.1944	18.731	15.683
145	5.85	1.90	4585	0.72	6.15	0.97	1.1991	18.769	15.649
155	6.53	2.92	843	0.70	1.27	1.05	1.1991	18.740	15.629
160	6.84	4.05	816	nd	1.32	nd	1.1899	18.621	15.644
165	7.13	1.87	774	0.58	1.36	1.02	1.2041	18.837	15.644
170	7.39	1.85	726	0.64	1.38	1.22	1.1971	18.784	15.692
175	7.64	2.19	737	0.64	1.53	1.33	1.1965	18.764	15.683
180	7.86	2.99	673	nd	1.53	nd	1.2025	18.848	15.676
185	8.06	0.96	662	0.55	1.64	1.36	1.1909	18.640	15.652
190	8.25	1.11	726	nd	1.95	nd	1.1871	18.581	15.654
195	8.42	1.06	700	0.36	2.01	1.04	1.1962	18.725	15.655
200	8.59	1.33	620	nd	1.89	nd	nd	nd	nd
205	8.74	1.57	784	0.41	2.48	1.29	1.1995	18.778	15.660
210	8.89	1.45	853	0.44	2.68	1.38	1.2050	18.874	15.662
215	9.05	1.30	954	0.42	2.86	1.26	1.1893	18.616	15.655
220	9.21	1.72	912	0.58	2.53	1.61	1.1977	18.763	15.667
225	9.39	2.68	954	nd	2.38	nd	1.1923	18.653	15.645
230	9.60	1.96	1537	1.28	3.40	2.83	1.1924	18.604	15.602
235	9.82	3.74	2931	1.60	5.67	3.09	1.1908	18.642	15.659
240	10.08	3.43	2666	1.63	4.64	2.84	1.1927	18.643	15.632
245	10.37	3.06	2390	1.28	3.90	2.09	1.1981	18.780	15.675
250	10.67	3.95	2237	2.33	3.55	3.70	1.1939	18.722	15.682
255	10.97	11.3	16547	nd	26.35	nd	1.1933	18.684	15.658
260	11.27	30.5	28175	18.40	46.51	30.4	1.2006	18.803	15.662
265	11.55	28.9	42146	15.28	72.98	26.5	1.2013	18.836	15.680
270	11.82	43.6	46317	21.04	83.72	38.0	1.2002	18.780	15.647
275	12.09	35.6	34090	18.40	63.99	34.5	1.2013	18.804	15.654
280	12.34	46.0	47695	19.36	92.44	37.5	1.1968	18.718	15.640
290	12.82	84.9	1108	4.62	2.24	9.35	1.1948	18.801	nd
300	13.29	nd	nd	nd	nd	nd	nd	nd	nd

‘nd’ means ‘not determined’; statistically significant dust episodes are highlighted in bold.

Pb and Al analyses were performed by Inductively Coupled Plasma Atomic Emission Spectrometry (ICP-AES, Jobin Yvon Ultima C, CEREGE) after acid digestion (HCl–HNO₃–HF) of peat material using trace metal clean techniques in a HEPA (class 1000) filtered trace metal clean laboratory. Reproducibility and accuracy for our ICP-AES analyses of major and trace elements have been previously reported (Germanique, 1994; Ollivier et al., 2010). Procedure and analytical blanks were below detection limits. Overall uncertainties on Al and Pb concentrations were generally below 5%. Concentrations are shown in Table 2 and Figure 4. Al and Pb deposition fluxes (Table 2) are calculated using ARs from the CLAM age model and a measured mean bulk density of 0.096 g/cm³.

Figure 4.

Aluminum (Al, ppm) and Pb/Al ratios (×10⁴) in Bozi Dar peat core.

A fraction of acid digested peat was purified on ×G1 × 8 anionic resin to extract Pb and measure the ratios of the four stable Pb isotopes (²⁰⁴Pb, ²⁰⁶Pb, ²⁰⁷Pb, ²⁰⁸Pb) by Thermo-Ionization Mass Spectrometry (TIMS, Finnigan MAT 262, CEREGE). Analytical mass discrimination was corrected using concurrent and repeated analyses of a National Institute of Standard and Technology (NIST) standard reference material for common lead (SRM981). Standard deviations of the ²⁰⁶Pb/²⁰⁷Pb ratios did not exceed 0.05%. Uncertainties on ²⁰⁶Pb/²⁰⁴Pb and ²⁰⁷Pb/²⁰⁴Pb ratios were generally below 0.2% and 0.3%, respectively. Lead isotope ratios are presented in Table 2.

Reliability of the atmospheric record

We have to ensure that Al and Pb accumulated within our peat derive from atmospheric deposition only (as expected with ombrotrophic peats) and not from lateral advection or upward migration from runoff and bedrock weathering, respectively. Ash analyses were performed to help identify the ombrotrophic and mineral horizons. Acid-insoluble ashes were determined on separate peat fractions that were dry-ashed at 550°C prior to aqua regia digestion (HCl–HNO₃), then filtered through Whatman 42 ashless filter paper and heated again at 550°C according to Vile et al. (1995, 2000). The ash content is below 5% down to 250 cm (Table 2 and Figure 5), corroborating the ombrotrophic character of the peat (Tolonen, 1984). Below, in a fen-dominated section, the ash content reaches 26–92% (Table 2 and Figure 5). This increase could be an indication of the contribution of bedrock weathering to the peat that would interfere with atmospherically deposited Al and Pb of which concentrations increase by a factor 10 or more in this bottom section of the core (Table 2 and Figure 4). We tentatively assess a potential mineral contribution within the peat using the ratio of the crustal element Al to that of ashes. We determine a mean Al/Ash (×10) ratio of 0.8 ± 0.6 (n = 35, 2σ) in the ombrotrophic sphagnum-dominated fraction of the peat (Figure 5). In the enriched ash section of the peat (below 250 cm), we calculate a ratio of 1.1 ± 0.3 (n = 7, 2σ; Figure 5). The ratios are not statistically different (p > 0.1, student t-test) and suggest that the potential minerotrophic nature of the bottom peat does not significantly influence trace metal content in this fraction. The Al/Ash ratios are significantly divergent from these means (p < 0.0001, student t-test) within the very surface (Al/Ash = 4.3) and the very bottom (Al/Ash = 0.01) layers of the peat core (Table 2) likely owing to runoff and drainage at the top (Brizova, 1993) and a possible mineralogical contribution from the bedrock at the very base of the core. The bedrock contribution can also be discerned using the variations of the Pb/Al ratio with depth in the peat. Indeed, any significant deviation of the Pb/Al ratios within the bottom layers of the peat may indicate a contribution in addition to the atmosphere. Here, Pb/Al ratios are generally uniform with the exception of a peak (at 80–95 cm, Figure 4) and three outliers (at 125, 135, and 290 cm; Figure 4). The mean Pb/Al ratios below 250 cm (4.8 ± 1.2 10⁴, n = 5, 2σ) are not significantly different (p > 0.05, student t-test) from the mean Pb/Al ratio in the rest of the core (6.4 ± 2.8 10⁴, n = 24, 2σ), suggesting no interference with the underlying bedrock. The three elevated ratios determined in the ombrotrophic section of the core may reflect the occurrence of pollutant Pb. Indeed, higher Pb/Al ratios generally reveal an anthropogenic contribution to Pb accumulated in sediment cores (Angeledis et al., 2011; García-Alix et al., 2013). These potential pollutant Pb imprints are discussed later. The outlier at 290 cm (Figure 4), corroborates ash results regarding a possible bedrock contribution to accumulated trace elements in the peat and is not considered in our discussion of atmospheric delivered imprints.

Figure 5.

Ash (%) and Al/ash ratios (×10) in the Bozi Dar peat core. The vertical dotted line shows the 5% ash limit above which the peat may be considered minerotrophic.

These results suggest that the minerotrophic feature of the bottom peat (below 250 cm) does not seem to prevent the record of atmospheric markers with no noticeable contribution from upward migration of bedrock-derived Pb and Al, with the exception of the very bottom layer at 290 cm. Shotyk et al. (2002) and Monna et al. (2004) show that minerotrophic peats also can preserve atmospherically deposited metal records at the millennia time scale. Meanwhile, this minerotrophic feature, along with compaction, contributes to enhance the bulk density of the sampled material below 250 cm, up to a factor 2 according to Weiss et al. (2002). Because no density was measured in this section, we assume that our flux calculations, using a mean density determined in the ombrotrophic section of the core, likely are underestimated below 250 cm and therefore are not used to calculate a mean deposition flux in the core and for comparison with other published deposition during the same period.

Results and discussion

Pollen abundances

In our description of vegetative phases (I–X), we refer to Firbas (1949/1952), whose classification is most commonly used in Central European regions. The oldest segment of the profile (vegetative phases II to mid-IV) exhibits the lowest proportion of arboreal species. The base of the profile (at 290 cm) is associated with 50% of arboreal species, while mere 30% of arboreal species are found at 260 cm. In younger sections of the profile, pollen of arboreal species predominate (around 75% in vegetative phases V to mid-VIII, and around 85% in vegetative phases mid-VIII to IX). Figure 2 shows an abbreviated BD pollen diagram. In the Late Glacial, the most abundant tree species are Pinus, Betula, Salix, and Juniperus, while herbaceous species (Cyperaceae) prevail. Solely in the Late Glacial, pollen of Hippophae is noted. In vegetative phase IV (Preboreal), the diagram records mixed forests, with Pinus and Betula, and an understory consisting of Juniperus and shrub-like Salix. During the Boreal (phase V), maximum relative abundance of Corylus is recorded. Simultaneously, the abundance of herbaceous species drops significantly. During the Atlantic (phases VI and VII), the relative representation of Pinus is at its lowest. Since the Boreal (V), Quercetum mixtum (QM) peaks in the Atlantic (VI–VII). This group of woody species includes Quercus, Ulmus, Tilia, Acer, and Fraxinus. At the same time, Corylus becomes less common. Also, during the Atlantic, Picea first appears. Alnus and Fraxinus build up from the end of the Boreal through the Atlantic. Unlike Alnus, which remains present throughout the upper parts of the profile, Fraxinus disappears at the end of phase VIII. The onset of the Sub-Boreal (VIII) is marked by a larger representation of Fagus, Abies, and Carpinus. These species remain in the area until the Sub-Atlantic (IX). During the Sub-Atlantic, QM abundance decreases, while Pinus once again becomes most common. The profile records a continuous forested landscape until around 2.3 kyr BP.

Climate-related vegetation changes

In the vertical Bozi Dar profile, vegetative phase I (Firbas, 1949/1952) is missing because no pollen is preserved in the non-organic eluvial sediment. The Late Pleistocene is represented by the Allerod (II) and the YD (III). As seen in Figure 2, Late Glacial is characterized by an open landscape, with grasses-dominated vegetation. The advent of pine and birch coincides with the Early-Holocene climatic warming (Preboreal, IV, 11.8–9.6 kyr BP). Hippophae coincide with the continuing presence of an inorganic admixture in the substrate. These species often suggest periodical colder climatic phases throughout the Holocene. The Boreal (V, 9.6–8.1 kyr BP) is characterized by a further temperature increase. Shade-intolerant shrubs, such as Corylus, are present at the margins of Pinus-/Betula-dominated woods. The Atlantic (VI–VII, 8.1–4.0 kyr BP) represents the warmest period of the Holocene. Pines become less common, restricted to extreme sites, such as rock outcrops and wetlands. We note that pollen does not provide a tool to distinguish between pines and dwarf pines. The climatically demanding group of QM trees peaks in the warmest period (Atlantic). At the same time, the abundance of Corylus declines because of its poor shade tolerance and decreasing nutrient availability. Also, in the Atlantic, Pinus becomes more common in upland areas, at elevations of around 1000 m. The Boreal/Atlantic boundary (here, c. 8.0 kyr BP) is commonly characterized by increasing precipitation (see Brizova, 1999; Mentlík et al., 2010) as evidenced by the advent of moisture-demanding Alnus and Fraxinus. Both species populate numerous riparian zones and also the area of the nearby Komorany Lake. The beginning of the Sub-Boreal (at 4.0 kyr BP) is drier and colder than the previous Atlantic period as shown by the high occurrence of Fagus, Abies, and Carpinus that are temperature-sensitive species (QM). Fagus, Abies, and Carpinus also are frequent with soil maturation and may not be directly related to climate changes. Pollen of Picea and Abies, unlike pollen of most other woody species, is capable of long-distance transport. The abundance of these two species is not necessarily related to local abundances.

According to the observed local pollen arrangements and the Firbas vegetative phases, we infer five local climate trends (A to E) that are shown in Figure 2. They comprise the very cold Late Glacial dominated by Cyperaceae grass land (A, 12.5–11.0 kyr BP), the Early-Holocene warming and the onset of Pinus (B, 11.0–9.0 kyr BP), the temperature increase during the Boreal (V) with a sharp increase of the shade non-tolerant Corylus and concurrent decline of Pinus (C, 9.0–8.1 kyr BP), the warmest and wettest period of the Holocene (Atlantic VI–VII) characterized by the onset of moisture-demanding species Alnus and Fraxinus and the occurrence of matured-soil demanding Alnus and Picea (D, 8.1–4.3 kyr BP), and the drier and possibly colder Sub-Boreal with the decline of temperature-sensitive species (E, <4.3 kyr BP). This climate zonation is used when discussing geochemical data. Pollen records from Bozi Dar and Etang de la Gruère in Switzerland show the same general trends (Roos-Barraclough et al., 2004). Meanwhile, the cold Sub-Boreal period ends earlier in Etang de la Gruère (c. 10 kyr BP) as shown by the occurrence of Corylus, Ulmus, and Quercus. Warmer trends (D) and (E) identified in Bozi Dar with the succession of Alnus–Picea and Fagus–Abies (c. 8.1 kyr BP and 4.3 kyr BP, respectively) are comprised within the same trend in Etang de la Gruère starting c. 7.0 kyr BP.

Dust deposition and climate shifts

We calculate a mean Al deposition flux of 2.66 ± 1.35 µg/cm²/yr (n = 33, 2σ; Table 2) that excludes the minerotrophic layer (below 250 cm) as well as the upper layer where Al deposition is above 20 µg/cm²/yr and could be hampered by lateral advection. According to a student t-test, any flux above 3.5 µg/cm²/yr is significantly higher (with p < 0.001, student t-test) than this mean deposition and is considered a dust episode. Consequently, we can identify five individual dust spikes (D₁–D₅; Table 2 and Figure 6) with deposition fluxes significantly higher than the mean input. The oldest one (D₁) is by far the largest (Al fluxes reach 84 µg/cm²/yr; Table 2) and extends from 280 to 255 cm (i.e. 12.3–11.0 kyr BP). It is comprised within the Late Glacial climate zone A and exhibits the characteristics of the expected YD cold event during the Late Glacial stage with enhanced eolian activity (De Boer, 1995; Kaiser et al., 2007). It is in agreement with the YD interval identified by Shotyk et al. (1998) in Etang de la Gruère (12.5–10.5 kyr). The YD episode at Bozi Dar matches other YD occurrences between 12.9 and 11.5 kyr BP from Greenland ice and Central European lakes and peats (Alley et al., 1993; Björck, 2006; Dansgaard et al., 1989; Enters et al., 2010; Le Roux et al., 2012; Peteet, 1995; Wolters, 2002), thus validating our age model. The second Al peak (D₂; Table 2 and Figure 6) is less pronounced with a maximum Al deposition of 5.7 µg/cm²/yr and extends from 10.7 to 9.8 kyr BP (250–235 cm). It covers the warming transition climate zone B (Figure 6). The following maxima D₃ and D₄ show similar maxima (6.2 µg/cm²/yr) at 5.9 and 3.2–2.8 kyr BP, respectively (Table 2). They both occur during the warm wet and dry climate zones D and E (Figure 6) and therefore are less likely to be associated with cold episodes, but rather northward flows from southern regions. The last Al peak (D₅) reaches 20.8 µg/cm²/yr (Table 2). It starts at 60 cm and is post-Medieval.

Figure 6.

Calculated Pb and Al deposition fluxes, crustal enrichment factors, and ²⁰⁶Pb/²⁰⁷Pb ratios from the Bozi Dar peat core. We identify (1) various dust events (Di, light gray) defined as peaks (several data points) or outliers (single data point) and (2) anthropogenic events (Ai, dark gray) that are shown with pollen-based climate zones (A to E). Single dates correspond to ¹⁴C calibrated ages (Table 1 and Figure 3).

We have identified several dust deposition maxima in our core using Al. While D₁ can easily be associated with the YD cold episode, there is no obvious well-known climate event that could be invoked to explain the other peaks or outliers. For example, there is no visible imprint of the 8.2 kyr cold event. Contrary to the YD, the 8.2 kyr episode is not as widely evidenced in continental Europe, possibly because of its shorter duration and intensity (Alley et al., 2003; Wunsch, 2006). For instance, it is not observed north of 60°N in Europe (Bakke et al., 2005; Bigler et al., 2002, 2003; Jones et al., 2004; Korhola et al., 2002; Larocque and Hall, 2003, 2004; Paasche et al., 2004; Rosén et al., 2001; Seppä et al., 2002, 2007), likely owing to a seasonal emphasis of the cooling phases during spring and winter that are not well recorded in septentrional regions. An insufficient sampling resolution may also cause this abrupt event not to be seen in the Bozi Dar record. In order to improve and complement the characteristics of these dust episodes, we use Pb and its stable isotopes that are efficient geochemical tracers of aerosol deposition sources.

Dust sources

The ²⁰⁶Pb/²⁰⁷Pb ratios measured within the peat core are above 1.190, except at 190 cm (8.3 kyr BP) and in the top 95 cm that corresponds to aerosols younger than 2.2 kyr BP (Table 2) and may be associated with pollutant Pb as discussed below. ²⁰⁶Pb/²⁰⁷Pb ratios of dust maxima D1–D4 are not statistically different from the mean ²⁰⁶Pb/²⁰⁷Pb ratio below 95 cm (1.197 ± 0.005, n = 33, 2σ; Table 2). This isotopic imprint is within the range of that measured in other Mid-Holocene (MH) peat bogs from Central and Northern Europe (Klaminder et al., 2003; Kylander et al., 2010; Shotyk et al., 1998) and Greenland ice cores (Biscaye et al., 1997; Burton et al., 2007; Rosman et al., 1997). The D₅ dust spike is too young (less than 700 years; Table 2) to be accurately defined owing to uncertainties on the age model. On the basis of ²⁰⁶Pb/²⁰⁷Pb ratios alone, none of the dust episodes can be isotopically identified with certainty. The isotope outlier at 8.3 kyr BP could be associated with specific aerosol sources, but shows no Al deposition maximum that could be related to a dust episode associated with the 8.2 kyr event. In order to resolve the determination of isotopic sources, we use ²⁰⁷Pb/²⁰⁴Pb versus ²⁰⁶Pb/²⁰⁴Pb plot types that allow to distinguish between well-known soil sources from various geographical origins, including local bedrock and long-distance sources from Asia (China, Gobi) and the Mediterranean (Sahara; Figure 7). The local Bohemian Pb ores are also shown to identify local metallurgical activities (Figure 7). Dust isotopic imprints are well scattered within the ²⁰⁷Pb/²⁰⁴Pb versus ²⁰⁶Pb/²⁰⁴Pb plot (Figure 7) that may help resolve specific source imprints for the various dust peaks (identified on the basis of Al deposition fluxes).

Figure 7.

²⁰⁶Pb/²⁰⁴Pb versus ²⁰⁷Pb/²⁰⁴Pb plot with dust sources from the Sahara (Abouchami and Zabel, 2003; Hamelin et al., 1989, 1997), the Gobi desert and Chinese loess (Biscaye et al., 1997; Jones et al., 2000), and local bedrock and imprints from Bohemian ores (Vanecek et al., 1985; Veron, unpublished data). The mean and standard deviations (sd) for both ²⁰⁶Pb/²⁰⁴Pb and ²⁰⁷Pb/²⁰⁴Pb ratios are shown in the figure for closed gray squares. Samples with higher standard deviations are either highlighted (sd_6/4 < 0.05 and sd_7/4 < 0.03) or crossed (sd_6/4 < 0.07 and sd_7/4 < 0.05) gray squares.

The YD episode (D₁) displays a clear Asian imprint that is particularly enhanced at the beginning of this event (12.3–11.7 kyr BP), when dust deposition reaches a maximum (Table 2 and Figure 8). These results corroborate a similar Asian imprint as evoked by Nd isotopes in the Etang de la Gruère (Jura Mountains, Switzerland) high-resolution sampling during the same period (Le Roux et al., 2012). These findings suggest that significant Asian-originating dust plumes invaded all of Central Europe during the early YD period. The local bedrock does not seem to contribute much during the YD, likely owing to remaining ice cover during this post-Glacial period. A possible Mediterranean imprint may also be recorded at Bozi Dar at 11.6 kyr BP. This Mediterranean imprint is recorded further at 10.7 kyr BP and 10.4 kyr BP, at the beginning of the following D₂ peak (Figure 8). The strongest flux recorded during this event (5.67 µg/cm²/yr at 9.8 kyr BP) may also retain a significant Saharan component given its closeness to this source imprint. Meanwhile, D₂ also reflects some mixing with other source imprints, as seen at 10.1 kyr BP (Figure 8). If we assume a single event for D₂, then a mixture of the Mediterranean source with local aerosols (as defined by the bedrock imprint) is more likely to explain the distribution of the D₂ isotopic signatures than a mixture with the Asian source that would imply significant wind shifts during the D₂ episode. This dust event is also observed in the Etang de la Gruère peat core, where it is partly assigned a volcanic origin on the basis of Nd isotopes (Le Roux et al., 2012). No such source is evident in our Czech core. Indeed, all of the potential volcanic isotopic signatures fall far outside of the observed range of Pb isotopes in Bozi Dar with ²⁰⁷Pb/²⁰⁴Pb ratios below 15.55 for the Icelandic ridges (Elliott et al., 1991; Prestvik et al., 2001; Sun et al., 1975; Thirlwall et al., 2004) and the Faroe Island (Gariépy et al., 1983; Holm et al., 2001), and ²⁰⁶Pb/²⁰⁴Pb ratios above 19.2 for the Massif Central alkali basalts (Pilet et al., 2004; Wilson and Downes, 1991; Wittig et al., 2007). The absence of volcanic signature in our core possibly arises from the far-east location of the Bozi Dar sampling site in Central Europe (as compared with the Jura Mountains) that prevents western European volcanic eruptions from reaching our peat core. For the same reason, we do not detect the dust episodes observed in the Swiss core at 8.4 kyr BP (on the basis of Nd isotopes; Le Roux et al., 2012) and the so-called 8.2 kyr event (from ²⁰⁶Pb/²⁰⁷Pb ratios; Shotyk et al., 1998) in the Bozi Dar geochemical record. An insufficient sampling resolution may also cause these sharp deposition episodes not to be detected at Bozi Dar. The MH period (8.6–6.5 kyr BP) is characterized by distinct Mediterranean and Asian dust imprints at 7.9–7.4 kyr BP and at 7.1 and 6.5 kyr BP, respectively (Table 2 and Figure 8). D₃ dust event at 5.9 kyr BP can also be assigned an Asian origin (Figure 6). This imprint is defined on the basis of 14 published Pb isotope analyses that are intertwined between the Saharan and the local bedrock on the ²⁰⁷Pb/²⁰⁴Pb versus ²⁰⁶Pb/²⁰⁴Pb plot (Figure 7). Therefore, any data point positioned within the Asian imprint source may also be interpreted as a mixture between the Saharan and the bedrock imprints and should be considered with caution. The D₃ event can be related to the climate Bond Event 4 (Bond et al., 2001). This outlier is also found in the Etang de la Gruère peat (Jura Mountains, Switzerland), where it could not be assigned any specific source origin on the basis of Nd isotopes and ²⁰⁶Pb/²⁰⁷Pb ratios alone (Le Roux et al., 2012; Shotyk et al., 1998). The Mediterranean signature is pronounced during the following D₄ episode between 3.2 and 2.8 kyr BP. These results suggest frequent Mediterranean dust incursions reaching Central Europe since 7.9 kyr BP at the onset of the warmer climate zones D and E (Figure 6) corroborating similar findings from Etang de la Gruère (Le Roux et al., 2012; Shotyk et al., 1998, 2001). The discrepancy between the geochemical records in Czech Republic and Switzerland for the westerly (marine 8.2 kyr event, volcanism) and easterly (Asian dust) source imprints suggest a less widely and efficient dispersion of zonal airflows in Central Europe, while similar occurrences of Mediterranean incursions in both peat cores suggest a more efficient meridian transport. This assumption is consistent with circulation patterns determined during the Holocene in Northern Europe from reconstructed lake levels, pollen–climate calibration, and General Circulation Models that propose an efficient zonal circulation during the Early-Holocene followed by more intense meridian airflows during the MH and Late Holocene (see Harrison et al., 1991, 1992; Seppä and Poska, 2004; Yu and Harrison, 1995). Most particularly, prevalent summer time anticyclonic conditions would explain the frequent recurrence of meridian warm Mediterranean incursions that reach Northern European regions during the MH (Antonsson et al., 2008; Seppä and Birks, 2001). This past atmospheric circulation contrasts with the present zonal-dominated airflows established in Northern Europe for the past two millennia. The post-Medieval D₅ dust imprint likely shows the contribution of local Bohemian ores with Pb deposition fluxes three to four times as high than during the preceding period. This pollutant imprint shall be further discussed below in the section on ‘Anthropogenic imprints’.

Figure 8.

²⁰⁶Pb/²⁰⁴Pb versus ²⁰⁷Pb/²⁰⁴Pb source imprint plot with identified dust (YD, Di) and anthropogenic events (Ai) from Bozi Dar. MH represents the Mid-Holocene period.

Calculated Pb deposition fluxes in our peat core show maxima that generally match dust episodes and are statistically different (p < 0.001, student t-test) from the mean Pb deposition flux (1.96 ± 1.08 ng/cm²/yr, n = 23, 2σ, Table 2) calculated below 95 cm (with ²⁰⁶Pb/²⁰⁷Pb > 1.190) and above the minerotrophic layer (at 250 cm). This mean deposition flux is within the range of the other lowest pre-anthropogenic background measured in European peat bogs (Klaminder et al., 2003; Kylander et al., 2005; Shotyk et al., 2001). Most Pb deposition maxima are clearly from crustal origin as shown by the calculated enrichment factors (EFs; Figure 6). EFs are determined as the ratio of Pb/Al in the sample to that of Pb/Al from a crustal reference chosen here as a mean of continental loess, soils, and upper crust sediments (McLennan, 1995; Wedepohl, 1995). Uncertainties of the calculated EFs reflect the concentration disparities of the chosen mean crustal reference ratio in various continental reservoirs (Figure 6). Our EFs are defined according to Shotyk et al. (1998) for peat material, using Al instead of scandium as a crustal marker for our study. When close to unity, one can assume that the sample is not enriched in Pb compared with a natural ‘pre-anthropogenic’ background, and therefore, increased Pb deposition can be explained by increased dust fluxes. Some EFs clearly extend beyond unity for two Pb deposition maxima at 5.1–4.3 kyr BP (A₀) and 2.2–1.4 kyr BP (A₁; Figure 6), suggesting the deposition of non-crustal Pb from anthropogenic origin. A₀ and A₁ show EFs that range from 2.5 to 9.4 (Table 2). These enriched layers are discussed according to Pb isotope imprints in order to corroborate and explain the source of the pollutant aerosols deposited in the peat.

Anthropogenic imprints

The ²⁰⁶Pb/²⁰⁷Pb ratios for A₁ (1.173–1.182; Table 2) are significantly (p < 0.0001, student t-test) less radiogenic than the mean pre-anthropogenic ²⁰⁶Pb/²⁰⁷Pb ratio measured below 95 cm (1.197 ± 0.005, n = 33, 2σ). However, ²⁰⁶Pb/²⁰⁷Pb ratios may be occasionally misleading owing to low radiogenic sources from local soil or volcanic emissions (Kylander et al., 2005, 2007, 2010; Rosman et al., 2003; Zheng et al., 2007). The use of the ²⁰⁷Pb/²⁰⁴Pb versus ²⁰⁶Pb/²⁰⁴Pb plot clearly links A₁ to the local Bohemian ore source (Figure 8). We may therefore infer that A₁ shows evidences of local metallurgical activities as early as 2.2 kyr BP. It is the oldest geochemical evidence for such industry in Czech Republic. Pollen records during this episode display a decrease of QM species that can be partly attributed to mining-related activities and associated deforestation. The other oldest strong geochemical features of metallurgical activity recorded from European peat bogs have been found in Spain at 3.9 kyr BP (García-Alix et al., 2013), in Germany at 3.5 kyr BP (Monna et al., 2000), in France at 3.2 kyr BP (Jouffroy-Bapicot et al., 2007; Monna et al., 2004), and in Switzerland and Sweden at 3.0 kyr BP (Klaminder et al., 2003; Shotyk et al., 1998).

It is difficult to ascertain the geographical origin and extent of anthropogenic disturbances recorded in sediments during the Bronze Age in remote continental regions owing to the fact that (1) most of the anthropogenic markers originate from aerosol transported over long distances, hence inducing weak signatures; (2) most cultivated plants are self-pollinating and do not readily emit pollen in the atmosphere, hence remaining invisible in the regional pollen rain; and (3) a disturbance extending at least 100 ha is needed to generate a visible signal in Arborum Pollen (AP) decline. In spite of these difficulties, environmental imprints of early human activity have been seen in continental Europe from the combination of pollen and Pb records in peat bogs and lakes, in Switzerland at 6.1 and 4.0 kyr BP (Thevenon et al., 2011; Weiss et al., 1997), in Germany at 5.5–3.5 kyr BP (Enters et al., 2010; Monna et al., 2000), in Sweden at 5.4 kyr BP (Klaminder et al., 2003), in France at 5.0 kyr BP (Jouffroy-Bapicot et al., 2007; Monna et al., 2004), and in Spain at 3.9 kyr and 3.2 kyr BP (García-Alix et al., 2013; Kylander et al., 2005). Some uncertainties can be resolved, thanks to the use of Pb isotope when local soil and ore imprints are well known (Monna et al., 2000, 2004). The A₀ Pb peaks display ²⁰⁶Pb/²⁰⁷Pb ratios (1.194–1.195; Table 2) that are within the range of the pre-anthropogenic background (1.196 ± 0.006), and therefore, no conclusion could be drawn from ²⁰⁶Pb/²⁰⁷Pb ratios alone. When plotted in the ²⁰⁷Pb/²⁰⁴Pb versus ²⁰⁶Pb/²⁰⁴Pb diagram (Figure 8), they clearly indicate Mediterranean (at 5.1 kyr BP) and local bedrock (at 4.3 kyr BP) sources. Because of a significant increase of EF in these samples, we may infer (1) long-range transport of contaminated aerosols from a distant anthropogenic signal in Mediterranean regions and (2) local disturbances associated with local agricultural activity and/or tentative mining activities. The Mediterranean imprint is also encountered in the layer immediately above A₀, at 4.0 kyr BP (Table 2) corroborating the frequent Mediterranean airflow incursions during this period. Our pollen record shows no specific trend that could confirm human activities, neither Cerealia-type species nor strong deforestation evidences. Two ²⁰⁶Pb/²⁰⁷Pb ratios (1.187 and 1.190) that are within the lowest range of the mean ²⁰⁶Pb/²⁰⁷Pb imprint below 95 cm (1.197 ± 0.005, n = 33, 2σ) are measured during the MH at 8.3 and 6.8 kyr BP (MH₀; Table 2). The ²⁰⁷Pb/²⁰⁴Pb versus ²⁰⁶Pb/²⁰⁴Pb plot indicates that MH₀ imprint tends toward the Bohemian ore source (Figure 8) and may be indicative of some contamination. Unfortunately, no Pb concentration, and corresponding fluxes or EFs, are available at these depths, and therefore, no conclusion can be drawn from the MH₀ isotopic imprints alone that could signify the contribution of another non-identified geogenic source. It should be noted that the isotopic imprints measured during the post-Medieval dust event D₅ (60–30 cm) likely include some Bohemian ore component (Figure 8).

Conclusion

The most remarkable features from the Bozi Dar peat core are (1) the well-known YD event with specific aerosol transport from Asia at 12.3–11.7 kyr BP, that is also observed at 7.1, 6.5, and 5.9 kyr BP; (2) the recurrent imprint of Mediterranean originating aerosols at 7.9–7.4 kyr BP (before aridification of the Sahara estimated c. 6.0–3.5 kyr BP; Jung et al., 2004; Renssen et al., 2006) and at 4.0–2.9 kyr BP; and (3) indisputable local mining activity starting at 2.2 kyr BP, the oldest ever encountered in Central Europe. The YD event and the Mediterranean airflow incursions are recorded during the same periods in the peat core collected by Shotyk et al. (1998) in the Jura Mountains, Switzerland. Asian sources are suggested by Nd isotopes in the Swiss core, but solely during the strong YD dust deposition episode (Le Roux et al., 2012). More frequent Asian-originating events are encountered in Bozi Dar during dust events of lesser importance and the MH, possibly owing to the eastern inland location of our sampling site compared with the Swiss peat. The inland eastward location of the Bozi Dar peat may also explain why we do not see the 8.2 kyr event as well as volcanic incursions that are detected in the Jura Mountains (Le Roux et al., 2012; Shotyk et al., 1998). Our geochemical markers clearly respond to atmospheric circulation patterns determined from pollen, hydrological, and physical models during the Holocene with (1) a dominant zonal circulation during the Early Holocene with the occurrence of Asian dust plumes in both Switzerland and the Czech Republic and (2) a more efficient meridian dispersion of Mediterranean aerosols during the MH and Late Holocene in Central Europe. Further high-resolution geochemical records from Holocene-long peat cores are needed in central and eastern Europe to corroborate these results, in particular the recurrences of Mediterranean and Asian airflow incursions, the occurrence of abrupt episodes (the 8.4 and 8.2 kyr BP events, volcanism), as well as the environmental imprint of mining activities during the Bronze Age.

Footnotes

Acknowledgements

We are very grateful to G Le Roux for his help with the completion of the ¹⁴C age model curve. We thank F Oldfield, JK Cochran, AR Flegal, and CE Lambert for useful editorial remarks. Three reviewers provided very constructive comments that helped improve the manuscript. We acknowledge the support of the OSU-Institut Pythéas at Aix Marseille University.

Funding

This work was supported by the European Commission (grant 244118 SoilTrEC) and the Czech Geological Survey (grants 335600 and 323000).

References

Abouchami

Zabel

(2003) Climate forcing of the Pb isotope record of terrigenous input into the Equatorial Atlantic. Earth and Planetary Science Letters 213: 221–234.

Alfonso

Grousset

Massé

. (2001) A European lead isotope signal recorded from 6000 to 3000 years BP in coastal marshes (SW France). Atmospheric Environment 35: 3595–3605.

Allan

Le Roux

De Vleeschouwer

. (2013) High-resolution reconstruction of atmospheric deposition of trace metals and metalloids since ad 1400 recorded by ombrotrophic peat cores in Hautes-Fagnes, Belgium. Environmental Pollution 178: 381–394.

Allen

JRM

Brandt

Brauer

. (1999) Rapid environmental fluctuations in southern Europe during the last glacial period. Nature 400: 740–743.

Alley

Marotzke

Nordhaus

. (2003) Abrupt climate change. Science 299: 2005–2010.

Alley

Meese

Shuman

. (1993) Abrupt increase in snow accumulation at the end of the Younger Dryas event. Nature 362: 527–529.

Angeledis

Radakovitch

Veron

. (2011) Anthropogenic metal contamination and sapropel imprints in deep Mediterranean sediments. Marine Pollution Bulletin 62: 1041–1052.

Antonsson

Chen

Seppä

(2008) Anticyclonic atmospheric circulation as an analogue for the warm and dry mid-Holocene summer climate in central Scandinavia. Climate of the Past 4: 215–224.

Bakke

Dahl

Paasche

. (2005) Glacier fluctuations, equilibrium-line altitudes and palaeoclimate in Lyngen, northern Norway, during Lateglacial and Holocene. The Holocene 15: 518–540.

10.

Baron

Carignan

Laurent

. (2006) Medieval lead making on Mont-Lozère Massif (Cévennes-France): Tracing ore sources by using Pb isotopes. Applied Geochemistry 21: 241–252.

11.

Baron

Lavoie

Ploquin

. (2005) Record of metal workshops in peat deposits: History and environmental impact on the Mont-Lozère Massif (France). Environmental Science and Technology 39: 5131–5140.

12.

Baron

Mahé-Le-Carlier

Carignan

. (2009) Archaeological reconstruction of medieval lead production: Implications for ancient metal provenance studies and paleopollution tracing by Pb isotopes. Applied Geochemistry 24: 2093–2101.

13.

Berger

(1990) The Younger Dryas cold spell – A quest for causes. Global and Planetary Change 3(3): 219–237.

14.

Bigler

Grahn

Larocque

. (2003) Holocene environmental change at Lake Njulla (999 m a.s.l.), northern Sweden: A comparison with four small nearby lakes along an altitudinal gradient. Journal of Paleolimnology 29: 13–29.

15.

Bigler

Larocque

Peglar

(2002) Quantitative multiproxy assessment of long-term patterns of Holocene environmental change from a small lake near Abisko, northern Sweden. The Holocene 12 (4): 481–496.

16.

Biscaye

Grousset

Revel

. (1997) Asian provenance of Last Glacial Maximum dust in the GISP-2 ice core, Summit, Greenland. Journal of Geophysical Research 102: 26765–26781.

17.

Biscaye

Grousset

Revel

. (2002) Quantitative multiproxy assessment of long-term patterns of Holocene environmental change from a small lake near Abisko, northern Sweden. The Holocene 12: 481–496.

18.

Björck

(2006) Younger Dryas oscillation, global evidence. In: Elias

(ed.) Encyclopedia of Quaternary Science, vol. 3. Oxford: Elsevier B.V., pp. 1987–1994.

19.

Blaauw

(2010) Methods and code for ‘classical’ age-modeling of radiocarbon sequences. Quaternary Geochronology 5: 512–518.

20.

Bond

Kromer

Beer

. (2001) Persistent solar influence on North Atlantic climate during the Holocene. Science 294: 2130–2136.

21.

Bond

Showers

Cheseby

. (1997) A pervasive millennial-scale cycle in North Atlantic Holocene and glacial climates. Science 278: 1257–1266.

22.

Brännvall

Bindler

Emteryd

. (1997) Stable isotope and concentration records of atmospheric lead pollution in peat and lake sediments in Sweden. Water, Air and Soil Pollution 100: 243–252.

23.

Brännvall

Bindler

Emteryd

. (2001) Four thousand years of atmospheric lead pollution in northern Europe: A summary from Swedish lake sediments. Journal of Paleolimnology 25: 421–435.

24.

Brizova

(1993) Rekonstrukce vyvoje vegetace raseliniste Bozi Dar na zaklade pylove analyzy. Praha: Czech Geological Survey, 36 pp.

25.

Brizova

(1999) Late Glacial and Holocene development of the vegetation in the Labe (Elbe) River flood-plain (Central Bohemia, Czech Republic). Acta Paleobotanica (Suppl. 2): 549–554 (also in: Proceedings of the 5th EPPC, Kraków).

26.

Broecker

(2003) Does the trigger for abrupt climate change reside in the ocean or in the atmosphere? Science 300: 1519–1522.

27.

Broecker

Denton

(1989) The role of ocean–atmosphere reorganization in glacial cycles. Geochimica et Cosmochimica Acta 53: 2465–2501.

28.

Bronk Ramsey

(2009) Bayesian analysis of radiocarbon dates. Radiocarbon 51(1): 337–360.

29.

Burton

Rosman

KJR

Candelone

. (2007) The impact of climatic conditions on Pb and Sr isotopic ratios found in Greenland ice, 7–150 ky BP. Earth and Planetary Science Letters 259: 557–566.

30.

Chambers

Booth

De Vleeschouwer

. (2012) Development and refinement of proxy-climate indicators from peats. Quaternary International 268: 21–33.

31.

Chow

Patterson

(1962) The occurrence and significance of lead isotopes in pelagic sediments. Geochimica et Cosmochimica Acta 26: 263–308.

32.

Clark

Marshall

Clarke

GKC

. (2001) Freshwater forcing of abrupt climate change during the last glaciation. Science 293: 283–287.

33.

Clark

Pisias

Stocker

. (2002) The role of the thermohaline circulation in abrupt climate change. Nature 415: 863–869.

34.

Dansgaard

White

JWC

Johnsen

(1989) The abrupt termination of the Younger Dryas climate event. Nature 339: 532–534.

35.

Davis

BAS

Brewer

Stevenson

. (2003) The temperature of Europe during the Holocene reconstructed from pollen data. Quaternary Science Reviews 22: 1701–1716.

36.

De Boer

(1995) A olische Prozesse und Landschaftsformen im mittleren Baruther Urstromtal seit dem Hochglazial der Weichselkaltzeit. Berlin: Fachbereich Geographie der Humboldt-Universität Berlin, 215 pp.

37.

De Vleeschouwer

Le Roux

Shotyk

(2010) Peat as an archive of atmospheric pollution and environmental change: A case study of lead in Europe. PAGES Newsletter 18(1): 21–22.

38.

De Vleeschouwer

Gérard

Goormaghtigh

. (2007) Atmospheric lead and heavy metal pollution records from a Belgian peat bog spanning the last two millennia: Human impact on a regional to global scale. Science of the Total Environment 377: 282–295.

39.

Doe

(1970) Lead Isotopes. Berlin: Springer Verlag.

40.

Dohnal

Kunst

Mejstřík

. (1965) Československá rašeliniště a slatiniště. Praha: Academia.

41.

Dunlap

Steinnes

Flegal

(1999) A synthesis of lead isotopes in two millennia of European air. Earth and Planetary Science Letters 167: 81–88.

42.

Elliott

Hawkesworth

Grönvold

(1991) Dynamic melting of the Iceland plume. Nature 351: 201–206.

43.

Enters

Kirilova

Lotter

. (2010) Climate change and human impact at Sacrower See (NE Germany) during the past 13,000 years: A geochemical record. Journal of Paleolimnology 43: 719–737.

44.

Erdtman

(1954) An Introduction to Pollen Analysis. 2nd Edition. Waltham, MA: Chronica Botanica, 239 pp.

45.

Fairbanks

(1989) A 17,000-year glacio-eustatic sea level record: Influence of glacial melting rates on the Younger Dryas event and deep-ocean circulation. Nature 342: 637–642.

46.

Firbas

(1949/1952) Spät und nacheiszeitliche Waldgeschichte Mitteleuropas nördlich der Alpen. Jena: Gustav Fischer.

47.

García-Alix

Jimenez-Espejo

Lozano

. (2013) Anthropogenic impact and lead pollution throughout the Holocene in Southern Iberia. Science of the Total Environment 449: 451–460.

48.

Garçon

Chauvel

Chapron

. (2012) Silver and lead in high-altitude lake sediments: Proxies for climate changes and human activities. Applied Geochemistry 27: 760–773.

49.

Gariépy

Ludden

Brooks

(1983) Isotopic and trace element constraints on the gensesis of the Faeroe lava pile. Earth and Planetary Science Letters 63: 257–272.

50.

Genty

Blamart

Ouahdi

. (2003) Precise dating of Dansgaard-Oeschger climate oscillations in western Europe from stalagmite data. Nature 421: 833–837.

51.

Genty

Combourieu-Nebout

Hatté

. (2005) Rapid climatic changes of the last 90 kyr recorded on the European continent. Comptes Rendus de l’Académie des Sciences, Paris, Géosciences 337: 970–982.

52.

Germanique

(1994) Major, trace and rare-earth elements in fourteen reference samples. Determination by X-Ray fluorescence spectrometry and inductively coupled plasma optical emission spectrometry. Geostandards Newsletter 18(1): 91–100.

53.

Hamelin

Ferrand

Alleman

. (1997) Isotopic evidence of pollutant lead transport from North America to the subtropical North Atlantic gyre. Geochimica et Cosmochimica Acta 61(20): 4423–4428.

54.

Hamelin

Grousset

Biscaye

. (1989) Lead isotopes in trade wind aerosols at Barbados: The influence of European emissions over the North Atlantic. Journal of Geophysical Research 94(C11): 16243–16250.

55.

Harrison

Prentice

Bartlein

(1991) What climate models can tell us about the Holocene palaeoclimates? Paläoklimaforschung – Palaeoclimate Research 6: 285–299.

56.

Harrison

Prentice

Bartlein

(1992) Influence of isolation and glaciation on atmospheric circulation in the North Atlantic sector: Implications of general circulation model experiments for the late Quaternary climatology of Europe. Quaternary Science Reviews 11: 283–299.

57.

Heiri

Lotter

Hausmann

. (2003) A chironomid-based Holocene summer air temperature reconstruction from the Swiss Alps. The Holocene 13: 477–484.

58.

Holm

Hald

Waagstein

(2001) Geochemical and Pb–Sr–Nd isotopic evidence for separate hot depleted and Iceland plume mantle sources for the Paleogene basalts of the Faroe Islands. Chemical Geology 178: 95–125.

59.

Hughen

Overpeck

Lehman

. (1998). Deglacial changes in ocean circulation from an extended radiocarbon calibration. Nature 391: 65–68.

60.

Jones

Halliday

Rea

. (2000) Eolian inputs of lead to the North Pacific. Geochimica et Cosmochimica Acta 64: 1405–1416.

61.

Jones

Leng

Solovieva

. (2004) Holocene climate of the Kola Peninsula: Evidence from the oxygen isotope record of diatom silica. Quaternary Science Reviews 23: 833–839.

62.

Jouffroy-Bapicot

Pulido

Baron

. (2007) Environmental impact of early palaeometallurgy: Pollen and geochemical analysis. Vegetation History and Archaeobotany 16: 251–258.

63.

Jung

SJA

Davies

Ganssen

. (2004) Stepwise Holocene aridification in NE Africa deduced from dust-borne radiogenic isotope records. Earth and Planetary Science Letters 221: 27–37.

64.

Kaiser

Rother

Lorenz

. (2007) Geomorphic evolution of small river–lake-systems in northeast Germany during the Late Quaternary. Earth Surface Processes and Landforms 32: 1516–1532.

65.

Keigwin

Jones

Lehman

. (1991) Deglacial meltwater discharge, North-Atlantic deep circulation, and abrupt climate change. Journal of Geophysical Research 96(C9): 16811–16826.

66.

Kempter

Frenzel

(2000) The impact of early mining and smelting on the local tropospheric aerosol detected in ombrotrophic peat bogs in the Harz, Germany. Water, Air and Soil Pollution 121: 93–108.

67.

Klaminder

Renberg

Bindler

. (2003) Isotopic trends and background fluxes of atmospheric lead in northern Europe: Analyses of three ombrotrophic bogs from south Sweden. Global Biogeochemical Cycles 17: GB1019.

68.

Korhola

Vasko

Toivonen

. (2002) Holocene temperature changes in northern Fennoscandia reconstructed from chironomids using Bayesian modeling. Quaternary Science Reviews 21: 1841–1860.

69.

Kylander

Klaminder

Bindler

. (2010) Natural lead isotope variations in the atmosphere. Earth and Planetary Science Letters 290: 44–53.

70.

Kylander

Muller

Wuüst

RAJ

. (2007) Rare earth element and Pb isotope variations in a 52 kyr peat core from Lynch’s Crater (NE Queensland, Australia): Proxy development and application to paleoclimate in the southern Hemisphere. Geochimica et Cosmochimica Acta 71: 942–960.

71.

Kylander

Weiss

Martinez-Cortizas

. (2005) Refining the pre-industrial atmospheric Pb isotope evolution curve in Europe using an 8000 year old peat core from NW Spain. Earth and Planetary Science Letters 240: 467–485.

72.

Larocque

Hall

(2003) Chironomids as quantitative indicators of mean July air temperature, validation by comparison with century-long meteorological records from northern Sweden. Journal of Paleolimnology 29: 475–493.

73.

Larocque

Hall

(2004) Holocene temperature estimates and chironomid community composition in the Abisko Valley, northern Sweden. Quaternary Science Reviews 23: 2453–2465.

74.

Le Roux

Fagel

De Vleeschouwer

. (2012) Volcano- and climate-driven changes in atmospheric dust sources and fluxes since the Late Glacial in Central Europe. Geology 40(4): 335–338.

75.

Le Roux

Weiss

Grattan

. (2004) Identifying the sources and timing of ancient and medieval atmospheric lead pollution in England using a peat profile from Lindow bog, Manchester. Journal of Environmental Monitoring 6(5): 502–510.

76.

Lehman

Jones

Keigwin

. (1991) Direct evidence for early retreat of the Fennoscandian ice sheet during the last deglaciation. Nature 349: 513–516.

77.

Lie

Paasche

(2006) How extreme was northern hemisphere seasonality during the Younger Dryas? Quaternary Science Reviews 25(5–6): 404–407.

78.

McLennan

(1995) Sediments and soils: Chemistry and abundances. In: Ahrens

(ed.) A Handbook of Physical Constants: Rock Physics and Phase Relations (AGU Ref. Shelf Ser., vol. 3). Washington, DC: AGU, pp. 8–19.

79.

Magny

Begeot

(2004) Hydrological changes in the European midlatitudes associated with freshwater outbursts from Lake Agassiz during the Younger Dryas event and the early Holocene. Quaternary Research 61: 181–192.

80.

Magny

Begeot

Guiot

. (2003) Contrasting patterns of hydrological changes in Europe in response to Holocene climate cooling phases. Quaternary Science Reviews 22: 1589–1596.

81.

Marcantonio

Zimmerman

. (2002) A Pb isotope record of mid-Atlantic US atmospheric Pb emissions in Chesapeake Bay sediments. Marine Chemistry 77: 123–132.

82.

Martinez-Cortizas

Garcia-Rodeja

Pontevedra-Pombal

. (2002) Atmospheric Pb deposition in Spain during the last 4600 years recorded by two ombrotrophic peat bogs and implications for the use of peat as archive. Science of the Total Environment 292: 33–44.

83.

Mentlík

Minár

Brizova

. (2010) Glaciation in the surrounding of Prášilské Lake (Bohemian Forest, Czech Republic). Geomorphology 117: 181–194.

84.

Monna

Galop

Carozza

. (2004) Environmental impact of early Basque mining and smelting recorded in a high ash minerogenic peat deposit. Science of the Total Environment 327: 197–214.

85.

Monna

Hamer

Lévêque

. (2000) Pb isotopes as a reliable marker of early mining and smelting in the Northern Harz province (Lower Saxony, Germany). Journal of Geochemical Exploration 68: 201–210.

86.

Muscheler

Beer

Wagner

. (2000) Changes in deep-water formation during the Younger Dryas event inferred from Be-10 and C-14 records. Nature408: 567–570.

87.

Novak

Emmanuel

Vile

. (2003) Origin of lead in eight Central European peat bogs determined from isotope ratios, strengths and operation times of regional pollution sources. Environmental Science and Technology 37: 437–445.

88.

Ollivier

Hamelin

Radakovitch

(2010) Seasonal variations of physical and chemical erosion: A three-year survey of the Rhone river (France). Geochimica et Cosmochimica Acta 74: 907–927.

89.

Overbeck

(1958) Pollenanalyse quartärer Bildungen. In: Freund

Grehn

(eds) Handbuch der Mikroskopie in der Technik. Frankfurt/Main: Umschau Verlag, pp. 325–410.

90.

Paasche

Løvlie

Dahl

. (2004) Bacterial magnetite in lake sediments: Late glacial to Holocene climate and sedimentary changes in northern Norway. Earth and Planetary Science Letters 223: 319–333.

91.

Peteet

(1992) The palynological expression and timing of the Younger Dryas event – Europe versus eastern North America. In: Bard

Broecker

(eds) The Last Deglaciation: Absolute and Radiocarbon Chronologies (NATO ASI Series I, vol. 2). Berlin/Heidelberg: Springer-Verlag, pp. 327–344.

92.

Peteet

(1995) Global Younger-Dryas. Quaternary International 28: 93–104.

93.

Pilet

Hernandez

Bussy

. (2004) Short-term metasomatic control of Nb/Th ratios in the mantle sources of intraplate basalts. Geology 32: 113–116.

94.

Prasad

Witt

Kienel

. (2009) The 8.2 ka event: Evidence for seasonal differences and the rate of climate change in Western Europe. Global and Planetary Change 67: 218–226.

95.

Prestvik

Goldberg

Karlsson

. (2001) Anomalous strontium and lead isotope signatures in the off-rift Öræfajökull central volcano in south-east Iceland. Evidence for enriched end member(s) of the Iceland mantle plume? Earth and Planetary Science Letters 190: 211–220.

96.

Reimer

Baillie

MGL

Bard

. (2009) IntCal09 and Marine09 radiocarbon age calibration curves, 0–50,000 years cal. BP. Radiocarbon 51: 1111–1150.

97.

Renberg

Bindler

Brännvall

(2001) Using the historical atmospheric lead-deposition record as chronological marker in sediment deposits in Europe. The Holocene 11(5): 511–516.

98.

Renberg

Brännvall

Bindler

. (2002) Sable lead isotopes and lake sediments – A useful combination for the study of atmospheric lead pollution history. Science of the Total Environment 292(1): 45–54.

99.

Renssen

Brovkin

Fichefet

. (2006) Simulation of the Holocene climate evolution in northern Africa: The termination of the African Humid period. Quaternary International 150: 95–102.

100.

Roos-Barraclough

van der Knaap

van Leeuwen

JFN

. (2004) A Late Glacial and Holocene record of climate and possible modern anthropogenic influences from a Swiss peat humification profile. The Holocene 14(1): 7–19.

101.

Rosén

Segerström

Eriksson

. (2001) Holocene climate change reconstructed from diatoms, chironomids, pollen and near-infrared spectroscopy at an alpine lake (Sjuodjijaure) in northern Sweden. The Holocene 11: 551–562.

102.

Rosman

KJR

Chisholm

Hong

. (1997) Lead from Carthaginian and Roman Spanish mines isotopically identified in Greenland ice dated from 600 bc to 300 ad. Environmental Science and Technology 31: 3413–3416.

103.

Rosman

KJR

Hong

Burton

. (2003) Pb and Sr isotopes from an ice-core provides evidence for the changing atmospheric conditions at the Sajama icecap, South America. Journal of Physics IV 107: 1157–1160.

104.

Sarmaja-Korjonen

Seppä

(2007) Abrupt and consistent responses of aquatic and terrestrial ecosystems during the 8200 cal yr cold event – A lacustrine record from Lake Arapisto, Finland. The Holocene. 17: 457–467.

105.

Seppä

Birks

HJB

(2001) July mean temperature and annual precipitation trends during the Holocene in the Fennoscandian tree-line area: Pollen-based climate reconstructions. The Holocene 11(5): 527–539.

106.

Seppä

Poska

(2004) Holocene annual mean temperature changes in Estonia and their relationship to solar insolation and atmospheric circulation patterns. Quaternary Research 61: 22–31.

107.

Seppä

Hammarlund

Antonsson

(2005) Low- and high frequency changes of temperature and effective humidity during the Holocene in south central Sweden: Implications for atmospheric and oceanic forcings of climate. Climate Dynamics 25: 285–297.

108.

Seppä

Birks

HJB

Giesecke

. (2007) Spatial structure of the 8200 cal yr BP event in northern Europe. Climate of the Past 3: 225–236.

109.

Seppä

Nyman

Korhola

. (2002) Changes in treelines and alpine vegetation in relation to post-glacial climate dynamics in northern Fennoscandia based on pollen and chironomid records. Journal of Quaternary Science 17: 287–301.

110.

Severinghaus

Brook

(1999) Abrupt climate change at the end of the last glacial period inferred from trapped air in polar ice. Science 286: 930–934.

111.

Shirahata

Elias

Patterson

. (1980) Chronological variations in concentrations and isotopic compositions of anthropogenic atmospheric lead in sediments of a remote subalpine pond. Geochimica et Cosmochimica Acta 44: 149–162.

112.

Shotyk

Krachler

Martinez-Cortizas

. (2002) A peat bog record of natural, pre-anthropogenic enrichments of trace elements in atmospheric aerosols since 12370 ¹⁴C yr BP, and their variation with Holocene climate change. Earth and Planetary Science Letters 202(2): 21–37.

113.

Shotyk

Weiss

Appleby

. (1998) History of atmospheric lead deposition since 12,370 ¹⁴C yr BP recorded in a peat bog profile, Jura Mountains, Switzerland. Science 281: 1635–1640.

114.

Shotyk

Weiss

Kramers

. (2001) Geochemistry of the peat bog at Etang de la Gruère, Jura Mountains, Switzerland, and its record of atmospheric Pb and lithogenic trace elements (Sc, Ti, Y, Zr, Hf and REE) since 12,370 14C yr BP. Geochimica et Cosmochimica Acta 65(14): 2337–2360.

115.

Snowball

Zillen

Gaillard

(2002) Rapid early-Holocene environmental changes in northern Sweden based on studies of two varved lake-sediment sequences. The Holocene 12: 7–16.

116.

Spurk

Leuschner

Baillie

MGL

. (2002) Depositional frequency of German fossil oaks: Climatically and non-climatically induced fluctuations in the Holocene. The Holocene 12: 707–715.

117.

Stos-Gale

Gale

Annetts

(1996) Lead isotope data from the Isotrace Laboratory, Oxford: Archaeometry data base 3, ores from the Aegean, part 1. Archaeometry 38: 381–390.

118.

Stos-Gale

Gale

Houghton

. (1995) Lead isotope data from the Isotrace Laboratory, Oxford: Archaeometry data base 1, ores from the Western Mediterranean. Archaeometry 37: 407–415.

119.

Sun

Tatsumoto

Schilling

(1975) Mantle plume mixing along the Reykjanes Ridge axis, lead isotope evidence. Science 190: 143–147.

120.

Thevenon

Graham

Chiaradia

. (2011) Local to regional scale industrial heavy metal pollution recorded in sediments of large freshwater lakes in central Europe (lakes Geneva and Lucerne) over the last centuries. Science of the Total Environment 412–413: 239–247.

121.

Thirlwall

Gee

MAM

Taylor

. (2004) Mantle components in Iceland and adjacent ridges investigated using double-spike Pb isotope ratios. Geochimica et Cosmochimica Acta 68: 361–386.

122.

Tinner

Lotter

(2001) Central European vegetation response to abrupt climate change at 8.2 ka. Geology 29: 551–554.

123.

Tolonen

(1984) Interpretation of changes in the ash content of ombrotrophic peat layers. Bulletin of the Geological Society of Finland 56: 207–219.

124.

Trenberth

(1997) Short-term climate variations: Recent accomplishments and issues for future progress. Bulletin of the American Meteorological Society 78: 1081–1096.

125.

Trenberth

Overpeck

Solomon

(2004) Exploring drought and its implications for the future. EOS 85: 27.

126.

Vanecek

Patocka

Posmourny

. (1985) The Use of Isotopic Composition of Ore Lead in Metallogenic Analysis of the Bohemian Massif. Prague: Academia.

127.

Veski

Seppä

Ojala

AEK

(2004) Cold event at 8200 yr BP recorded in annually laminated lake sediments in Eastern Europe. Geology 32: 681–684.

128.

Vile

Wieder

Novak

(2000) 200 years of Pb deposition throughout the Czech Republic: Patterns and sources. Environmental Science and Technology 34(1): 12–21.

129.

Vile

Novák

MJV

Brizova

. (1995) Historical rates of atmospheric metal deposition using 210Pb dates Sphagnum peat cores: Corroboration, computation, and interpretation. Water, Air and Soil Pollution 79(1–4): 89–106.

130.

Wedepohl

(1995) The composition of the continental crust. Geochimica et Cosmochimica Acta 59(7): 1217–1232.

131.

Weiss

Shotyk

Cherburkin

. (1997) Atmospheric lead deposition from 12 400 to ca. 2000 yrs BP in a peat bog profile, Jura Mountains, Switzerland. Water, Air and Soil Pollution 100: 311–324.

132.

Weiss

Shotyk

Rieley

. (2002) The geochemistry of major and selected trace elements in a forested peat bog, Kalimantan, SE Asia, and its implication for past atmospheric dust deposition. Geochimica et Cosmochimica Acta 66(13): 2307–2323.

133.

Wilson

Downes

(1991) Tertiary-Quaternary extension-related alkaline magmatism in Western and Central Europe. Journal of Petrology 32: 811–849.

134.

Wittig

Baker

Downes

(2007) U–Th–Pb and Lu–Hf isotopic constraints on the evolution of the sub-continental lithospheric mantle, French Massif Central. Geochimica et Cosmochimica Acta 71: 1290–1311.

135.

Wolters

(2002) Vegetationsgeschichtliche Untersuchungen zur spätglazialen und holozänen Landschaftsentwicklung in der Döberitzer Heide. Berlin: J. Cramer, 157 pp.

136.

Wright

Jr (1989) The Amphi-Atlantic distribution of the Younger Dryas paleoclimatic oscillation. Quaternary Science Reviews 8: 295–306.

137.

Wunsch

(2006) Abrupt climate change: An alternative view. Quaternary Research 65: 191–203.

138.

Harrison

(1995) Holocene changes in atmospheric circulation patterns as shown by lake status changes in northern Europe. Boreas 24(3): 260–268.

139.

Zheng

Shotyk

Krachler

. (2007) A 15,800-year record of atmospheric lead deposition on the Devon Island Ice Cap, Nunavut, Canada: Natural and anthropogenic enrichments, isotopic composition, and predominant sources. Global Biogeochemical Cycles 21(2): GB2027.