How dry was the Younger Dryas? Evidence from a coupled δH-δO biomarker paleohygrometer applied to the Lake Gemündener Maar sediments, Western Eifel, Germany

Chair of Geomorphology and BayCEER, University of Bayreuth, Universitätsstrasse 30, D95440 Bayreuth, Germany Institute of Agronomy and Nutritional Sciences, Soil Biogeochemistry, Martin-LutherUniversity Halle-Wittenberg, Von-Seckendorff-Platz 3, D-06120 Halle, Germany 10 Institute of Geography and Oeschger Center for Climate Change Research, University of Bern, Hallerstrasse 12, CH-3012 Bern, Switzerland Faculty of Physics and Applied Computer Science, AGH University of Science and Technology, Al. Mickiewicza 30, 30-059 Kraków, Poland Institute of Geosciences, Group of Climate and Sediments, Johannes Gutenberg University of 15 Mainz, J.-J.-Becher-Weg 21, D-55128 Mainz, Germany Institute of Geography, Dresden University of Technology, Helmholtzstrasse 10, D-01062 Dresden, Germany


Abstract
The Late Glacial to Early Holocene transition phase and particularly the Younger Dryas period, i.e. the major last cold spell in Central Europe during the Late Glacial, are considered crucial for understanding rapid natural climate change in the past. The sediments from Maar lakes in the Eifel, Germany, have turned out to be valuable archives for recording such 5 paleoenvironmental changes.
For this study, we investigated a Late Glacial to Early Holocene sediment core that was retrieved from the Gemündener Maar in the Western Eifel, Germany. We analysed the hydrogen (δ 2 H) and oxygen (δ 18 O) stable isotope composition of leaf wax-derived lipid biomarkers (n-alkanes C27 and C29) and a hemicellulose-derived sugar biomarker (arabinose), 10 respectively. Both δ 2 Hn-alkane and δ 18 Osugar are suggested to reflect mainly leaf water of vegetation growing in the catchment of the Gemündener Maar. Leaf water reflects δ 2 H and δ 18 O of precipitation (primarily temperature-dependent) modified by evapotranspirative enrichment of leaf water due to transpiration. Based on the notion that the evapotranspirative enrichment depends primarily on relative humidity (RH), we apply a previously introduced 'coupled δ 2 Hn-15 alkane-δ 18 Osugar paleohygrometer approach' to reconstruct the deuterium-excess of leaf water and in turn Late Glacial-Early Holocene RH changes from our Gemündener Maar record.
Our results do not provide evidence for overall markedly dry climatic conditions having prevailed during the Younger Dryas. Rather, a two-phasing of the Younger Dryas is supported with moderate wet conditions on Allerød level during the first half and drier conditions during 20 the second half of the Younger Dryas. Moreover, our results suggest that the amplitude of RH changes during the Early Holocene was more pronounced than during the Younger Dryas. This included the occurrence of a 'Preboreal Humid Phase'. One possible explanation for this unexpected finding could be that solar activity is a hitherto underestimated driver of Central European RH changes in the past. 25

Introduction
In order to evaluate the relevance of man-made climate change in the future, it is of great importance to study and understand large and rapid climate fluctuations in the past. Many studies have focused on the Late Glacial to Early Holocene transition phase, a period with various expressions in temperature, atmospheric circulation and hydrology worldwide (Alley, 5 2000;Brauer et al., 2008;Denton et al., 2010;Partin et al., 2015;Wagner et al., 1999).
Particularly the Younger Dryas (YD) period, i.e. the major last cold spell in Central Europe during the Late Glacial just before the onset of the Holocene warm period (Denton et al., 2010;Heiri et al., 2014;Isarin and Bohncke, 1999), has long been considered crucial for understanding rapid natural climate change in the past (Alley, 2000). The sediments from maar 10 lakes in the Eifel, Germany, have turned out to be valuable archives for paleoenvironmental reconstructions providing high resolution palynological, sedimentological and geochemical records for climate, vegetation and landscape history (Brauer et al., 2008;Brunck et al., 2015;Litt et al., 2003;Litt and Stebich, 1999;Sirocko et al., 2013;Zolitschka, 1998).
Within this study we aimed at applying the coupled δ 2 Hn-alkane-δ 18 Osugar paleohygrometer 10 approach to the Late Glacial-Early Holocene sediment cores of the Gemündener Maar. More specifically, we addressed the following objectives: (i) source identification of the sedimentary organic matter and the investigated n-alkanes and sugars (aquatic vs. terrestrial), (ii) reconstructing leaf water isotope composition based on compound-specific δ 2 H and δ 18 O values of the n-alkane and sugar biomarkers, (iii) reconstructing RH changes using the coupled δ 2 Hn-15 alkane-δ 18 Osugar paleohygrometer approach and (iv) inferring implications for middle European paleoclimate history from the established Gemündener Maar RH record.
The maar was formed during a phreatomagmatic explosion within the local Devonian siltstone (greywacke) around 20-25 ka BP (Büchel, 1993). The lake is 39 m deep at maximum and has a diameter of roughly 300 m. Due to its formation conditions the lake is almost circular with a 25 lake surface area of 75,000 m 2 and is surrounded by a small catchment (Fig. 1B), with an area of 430,000 m 2 (Scharf and Menn, 1992). The lake is fed by precipitation and groundwater (no surface inflow and outflow present). The sediments are, accordingly, not affected by fluvial sediment input. The catchment area is furthermore steep and densely vegetated with broadleaved trees (Fig. 1C). The investigated samples were taken from the 8 m Gemündener Maar 30 core (GM1), which was taken at approximately 20 m water depth near the center of the maar ( Fig. 1B) with a Livingston piston corer (UWITEC, Mondsee, Austria). The GM1 core was retrieved from a terrace on the steep slope of the maar exactly in a fan of an underwater erosion gully structure. The core is part of the Eifel Laminated Sediment Archive Project of the Institute for Geoscience at Johannes Gutenberg University Mainz (Sirocko et al., 2013(Sirocko et al., , 2016).

5
Bulk analyses were carried out on 112 samples, covering a section of 606 cm to 727 cm depth of the Gemündener Maar GM1 core. Total carbon (TC) and nitrogen (N), bulk carbon isotope composition (δ 13 CTC) and nitrogen isotope composition (δ 15 N) were determined at the Institute of Agronomy and Nutritional Sciences, Soil Biogeochemistry, Martin-Luther-University Halle-Wittenberg, using an EuroVector EA 3000 elemental analyzer (Hekatech, 10 Wegberg, Germany) coupled via a Conflo III Interface to a Delta V Advantage isotope ratio mass spectrometer (IRMS; both from Thermo Fisher Scientific, Bremen, Germany).
Additionally, total organic carbon (TOC) and bulk δ 13 C of the total organic carbon (δ 13 CTOC) were assessed after removal of carbonate with 32% hydrochloric acid (HCl) fumigation followed by a neutralization step with moist sodium hydroxide, both for 24 h under 60°C and 15 vacuum conditions. This allows calculating TOC/N ratios. Laboratory standards from the International Atomic Energy Agency (IAEA) as well as from United States Geological Survey (USGS) with known total carbon, nitrogen, 13 C and 15 N contents (IAEA N2, IAEA CH6, IAEA NO3, IAEA CH7, IAEA 305A, USGS 41) were used for calibration. The 13 C and 15 N contents are expressed in the common δ-notation as relative to an international standard (δ 13 C: Vienna 20 Pee Dee Belemnite, VPDB; δ 15 N: atmospheric N2, Air).
For pollen analysis, 16 samples were investigated covering the relevant depth section.
Each sample covered a depth range of 1 cm. Preparation was conducted by F. Dreher according to standard procedures at the laboratory of the Group of Climate and Sediments, Institute of Geosciences, Johannes Gutenberg University of Mainz, using potassium hydroxide, HCl and 25 hydrofluoric acid (Sirocko et al., 2016). Afterwards, acetic acid and a mixture of acetic anhydride and sulphuric acid (9:1) were used for acetolysis. The samples were then centrifuged at 3000 to 3500 rounds per minute for 5 min and then sieved over 200 µm and a 10 µm sieve.
Afterwards, the samples were fixed with anhydrous glycerol for reliable identification and a maximum magnification of 600 was used for counting the remains. Pollen results are reported 30 in relative percentages (%).

Age control
The investigated sediments are partially laminated. The first tie point to establish a chronology for the Gemündener Maar core is a radiocarbon-dated piece of charcoal in 727 cm core depth, which dated to 13,800 ± 110 cal a BP (Fig. 2D). This age is derived from a 14 C age of 11,950 ± 65 a BP as part of the supplement material of Sirocko et al. (2013), calibrated using 5 CalPal-software (Weninger and Jöris, 2008) calculated with the IntCal13 calibration curve (Reimer et al., 2013). The uncertainty of the calibrated 14 C age represents the 68% probability range. The second tie point is the clearly visible Laacher See Tephra between 673 to 680 cm core depth ( Fig. 2A and D). The latter can be used as chronological marker due to the varve counted age of 12,880 a BP in the adjacent Meerfelder Maar (Fig. A and D;Brauer et al., 1999). 10 The onset of the Younger Dryas period was set to 12,680 a BP (varve counted in Lake Meerfelder Maar sediments; Brauer et al., 1999;Litt et al., 2009) identified in a depth of 670 cm in the GM1 core due to a clear color change ( Fig. 2A and D). The onset of the Preboreal (Holocene) was found to date to 11,590 a BP in Lake Meerfelder Maar by varve counting (Litt et al., 2009). This was used to wiggle match the distinct changes in the pollen spectra 15 (decreasing Poaceae, peaking Artemisia, increasing Pinus and Betula; Fig. 2B and C), the clear rise in TOC (Fig. 3A) and the color change ( Fig. 2A), which were identified in 643 cm depth ( Fig. 2D). The Late Glacial to Preboreal (Holocene) transition is commonly well recorded in maar sediments from the Eifel region, i.e. clear changes in deposition as well as pollen pattern (Brauer et al., 1999;Litt et al., 2001Litt et al., , 2003Litt and Stebich, 1999), dated e.g. in Lake Holzmaar 20 to 11,600 a BP by a combination of varve counting and 14 C dating (Zolitschka, 1998). The last time marker used to constrain the age model is the middle of the sharp increase in Corylus (hazel) pollen at 622 cm depth (Fig. B, C and D). We used this sharp increase as marker for the Preboreal to Boreal transition, which is varve counted by Litt et al. (2009) to 10,740 a BP in the Meerfelder Maar sediments. The offset of 60 years to the varve counted Holzmaar record of 25 Zolitschka (1998), as it is presented by Litt et al. (2009), is within the uncertainty of placing the onset of the Preboreal in the Gemündener Maar Corylus curve.
The investigated core section from 607 cm to 694 cm depth covers therefore the time between ~ 13,150 a and 10,140 a BP, i.e. the Allerød, the Younger Dryas, the Preboreal, and the beginning of the Boreal, with regard to the biomarkers ( Fig. 2A and D). Assuming constant 30 sedimentation rates between the markers, an average resolution of 51 a/cm can be calculated; the minimum and maximum resolution are 19 and 124 a/cm, respectively. The part above the Laacher See Tephra reveals a lower mean resolution (55 a/cm) than the section below (30 a/cm).

Biomarker and compound-specific isotope analysis
For δ 2 H analyses of n-alkanes as well as δ 18 O analyses of sugars, 59 samples were prepared from 607 cm to 694 cm depth of the Gemündener Maar GM1 core, in order to cover the core section with already high TOC content and the Late Glacial to Holocene transition 5 ( Fig. 2 and 3A). n-Alkanes were extracted from 1-6 g freeze-dried and grinded samples by microwave extraction at 100°C for 1 h, using 15 ml of solvent (dichloromethane and methanol, in a ratio 9:1). The resultant total lipid extracts were separated over aminopropyl silica gel (Supelco 45 µm) filled pipette columns. Nonpolar compounds (including n-alkanes) were eluted with n-hexane. The fraction was spiked with a known amount of 5α-androstanone, used 10 as internal standard. Identification and quantification was carried out on an Agilent MS 5975 (EI) interfaced with an Agilent 7890 GC equipped with a 30 m fused silica capillary column (HP5-MS 0.25 mm i.d., 0.25 µm film thickness), and a split/splitless injector operating in splitless mode at 320°C. Carrier gas was helium and the temperature program was 1 min at 50°C, from 50 to 200°C at 30°C min -1 , from 200 to 320°C at 7°C min -1 , 5 min at 320°C. Data 15 recording comprised the total ion count (scan mode from m/z 40 to m/z 600) and single ion monitoring (m/z 57, 71, 85 and 99). Concentrations were calculated relative to the internal standard and to an external standard (n-C21 to n-C40 alkane mixture, Supelco), injected in different concentrations (40 ng/µl, 4 ng/µl, 1 ng/µl, 0.4 ng/µl).
Prior to compound-specific isotope analyses, the n-alkanes were further purified. The 20 nonpolar fractions were passed over a pipette column filled with activated AgNO3 impregnated silica gel and a pipette column filled with zeolite (Geokleen). After drying, the zeolite was removed using hydrofluoric acid and the n-alkanes were recovered by liquid-liquid extraction with hexane. The purified n-alkane fractions were measured for their compound-specific stable hydrogen isotope composition (δ 2 H). The measurements were performed at the Institute of 25 Geography, University of Bern on an IsoPrime 100 IRMS, coupled to an Agilent 7890A GC via a GC5 pyrolysis/combustion interface operating in pyrolysis modus with a Cr (ChromeHD) reactor at 1000°C. Samples were injected with a split/splitless injector. The GC was equipped with 30 m fused silica column (HP5-MS, 0.32 mm inner diameter, 0.25 µm film thickness).
The precision was checked by co-analyzing a standard alkane mixture (n-C27, n-C29, n-C33) with 30 known isotope composition (A. Schimmelmann, University of Indiana), injected twice every six runs. The samples were analyzed in threefold repetitions (except from the samples in 622 and 672 cm depth), and the analytical precision was generally better than 5‰. The stable hydrogen isotope compositions are given in the δ-notation (δ 2 Hn-alkane) versus Vienna Standard Mean Ocean Water (VSMOW). The H3 + -correction factor was checked every two days and stayed stable over the course of measurements at 3.14. The δ 2 Hn-alkane values refer to the area weighted mean of the δ 2 H values of n-alkanes with 27 and 29 carbon atoms (n-C27, n-C29), respectively, because of their relatively high abundance in the samples (Fig. 4A).
The sample preparation for δ 18 O analyses of hemicellulose/polysaccharide-derived 5 sugars followed standard procedures at the Institute of Agronomy and Nutritional Sciences, Soil Biogeochemistry, Martin-Luther-University Halle-Wittenberg, according to the method of Zech and Glaser (2009). The monosaccharide sugars were hydrolytically extracted from samples containing approximately 10 mg total organic carbon with 10 ml of 4 M trifluoroacetic acid at 105°C for 4 h, applying the method described by Amelung et al. (1996). After filtration 10 over glass fibre filters, the extracted sugars were cleaned using XAD-7 (to remove humic-like substances) and Dowex 50WX8 columns (to remove interfering cations). Afterwards, the purified samples were freeze-dried and derivatized by adding methylboronic acid (4 mg in 400 µl pyridine) for 1 h at 60°C.  4B). Rhamnose areas, respectively concentrations, were too low for reliable isotope measurements in most samples.

30
The coupled δ 2 Hn-alkane-δ 18 Osugar paleohygrometer approach was described in detail by Tuthorn et al. (2015) and Zech et al. (2013a). The most fundamental assumption of the approach is that the isotope composition of leaf water can be reconstructed by applying biosynthetic fractionation factors on the measured δ 2 Hn-alkane and δ 18 Osugar values (Fig. 5). The concept is furthermore based on the observation that the isotope composition of global precipitation plots typically close to the global meteoric water line (GMWL; δ 2 Hprec = 8 • δ 18 Oprec + 10;Dansgaard, 1964). In Germany, a local meteoric water line (LMWLGermany) slightly deviating from GMWL was described by Stumpp et al. (2014) (δ 2 Hprec = 7.72 ± 0.13 • δ 18 Oprec + 4.90 ± 0.01; Fig. 5), 5 which we used as the baseline for our calculations. The quite similar LMWLs for Trier (δ 2 Hprec = 7.81 ± 0.08 • δ 18 Oprec + 5.06 ± 0.60) and Koblenz (δ 2 Hprec = 7.80 ± 0.07 • δ 18 Oprec + 2.68 ± 0.53) as well as the GMWL are additionally displayed in Fig. 5 for comparison. The local precipitation is the source for soil water and shallow groundwater, which in turn acts as source water for plants. During daytime, however, leaf water is typically 2 H-and 18 O-enriched 10 compared to the source water due to evapotranspiration through the stomata ( Fig. 5; Allison et al., 1985;Bariac et al., 1994;Walker and Brunel, 1990). The leaf water reservoir at the evaporative sites achieves quickly steady-state conditions (Allison et al., 1985;Bariac et al., 1994;Gat et al., 2007;Walker and Brunel, 1990). Thus, the isotope composition of the transpired water vapor is equal to the isotope composition of the source water utilized by the 15 plants during the transpiration process. The evaporative enrichment of leaf water under steadystate conditions, can be described via a Craig-Gordon model (e.g. Flanagan et al., 1991;Roden and Ehleringer, 1999) by the following expression (e.g. Barbour et al., 2004): where δe, δs and δa are the hydrogen and oxygen isotope compositions of leaf water at the evaporative sites, source water and atmospheric water vapor, respectively, ε* are the 20 equilibrium enrichment expressed as (1-1/αL/V) • 10 3 where αL/V is the equilibrium fractionation between liquid and vapor in ‰, εk are the kinetic fractionation parameters for water vapor diffusion from intracellular air space through the stomata and the boundary layer, both for 2 H and 18 O, respectively, and ea/ei is the ratio of atmospheric vapor pressure to intracellular vapor pressure. When leaf temperature is equal to air temperature, the ea/ei ratio represents RH of the 25 local atmosphere. If the plant source water and the local atmospheric water vapor are in isotopic equilibrium, the term δa -δs can approximated by -ε*. Thus. Eq. (1) can be reduced to: The kinetic fractionation parameters (εk) are typically related to stomatal and boundary layer resistances with respect to water flux (Farquhar et al., 1989). Since direct measurements of those plant physiological parameters can be hardly assessed in a paleo application we used the 30 kinetic enrichment parameters Ck instead, derived from a more generalized form of the Craig-Gordon model, for describing the kinetic isotope enrichment for 2 H and 18 O, respectively, which leads to Eq. (3) (Craig and Gordon, 1965;Gat and Bowser, 1991): In a δ 2 H-δ 18 O diagram, the hydrogen and oxygen isotope composition of leaf and source water can be described as a local deuterium-(d-)excess = δ 2 H -7.72 • δ 18 O (Stumpp et al., 2014) in one equation by using the slope of the LMWLGermany (Eq. 4). This approach is comparable to 5 the d-excess definition from Dansgaard (1964), who used the equation d = δ 2 H -8 • δ 18 O for a measure of the parallel deviation between a given point in the δ 2 H-δ 18 O diagram to the GMWL.
Where de and ds are the d-excess values of the leaf water at the evaporative sites and the source water, respectively, and the equilibrium ( 2 * and 18 * ) and kinetic enrichment parameters (C k 2 and C k 18 ) are expressed for both isotopes. From Eq.
(1) to Eq. (4) the primary control of RH on the 10 isotope composition of the leaf water is demonstrated when stomata are open through transpiration. If de can be derived from compound specific δ 2 H and δ 18 O measurements of the n-alkane and sugar biomarkers, which derive δ 2 He and δ 18 Oe values for the purpose of calculating de values via the equation de = δ 2 He -7.72 • δ 18 Oe, and the ds can be approximated also from the d-excess of the LMWLGermany (= 4.9). Accordingly, Eq. (4) can be rearranged in 15 order to calculate RH of the local atmosphere normalized to leaf temperature as given by Eq.
(5) (Hepp et al., 2017;Tuthorn et al., 2015;Zech et al., 2013a): where Δd is the distance between de and ds, calculated as Δd = de -ds. Equilibrium fractionation parameters ( 2 * and 18 * ) are derived from empirical equations of Horita and Wesolowski (1994) with mean daytime growth-period temperature of 14.8°C (from 6 a.m. to 7 p.m. and April to 20 October, derived from the Nürburg-Barweiler station, approx. 25 Merlivat (1978), who reported maximum values during the molecular diffusion 25 process of water through a stagnant boundary layer. The assumption that maximum kinetic fractionation occurs seems to be most suitable for sedimentological application where a signal averaging over decades can be assumed (see above and discussion in Zech et al., 2013a). It should be also noted that εk values of broadleaf trees and shrubs over broad climatic conditions are well in the range with used C k 2 and C k 18 values, revealing 23.9 (±0.9) and 26.7‰ (±1.0) for δ 2 H and δ 18 O, respectively (derived from supplementary data of Cernusak et al., 2016).
The nominator of Eq. (5) describes the parallel distance between the d-excesses of LMWL and leaf water at the evaporative sites which is converted into RH values, while the denominator is a combination of the slopes of LMWL and the local evaporation line (LEL). 5 This means in turn that the quantification with Eq. (5) is done by obtaining the distance between the source water points, calculated via the intersects between the individual LELs and the LMWLGermany, and the leaf water points. The underlying slope of those LELs can be derived from Eq. (6) via the Craig-Gordon model using the same assumptions as outlined above in a rearranged form (Eq. 6; Zech et al., 2013a). When using the fractionation parameters from 10 above, the slope of the LEL is constant over time, independent from RH and equal to ~ 2.8 (Eq. 6). This is well in agreement with field and laboratory studies (Allison et al., 1985;Bariac et al., 1994;Gat et al., 2007;Tipple et al., 2013;Walker and Brunel, 1990).
In order to provide an uncertainty interval in terms of measurement precision covering the Gemündener Maar RH record, we calculated an error propagation for de values according 15 to Eq. 7, by using the analytical standard errors (SE). Maximum and minimum values were then applied to Eq. (5) resulting in a lower and upper RH limit (blue shaded area in Fig. 7A for biomass production and possibly increasing minerogenic sedimentation. Interestingly, the Late-Glacial-Early Holocene TOC patterns seem not to be the same for all maar lakes, because the Meerfelder Maar shows a clear TOC two-phasing during the Younger Dryas (Brauer et al., 1999) and the Holzmaar is lacking an Allerød TOC maximum (Lücke et al., 2003). The δ 15 N values of the Gemündener Maar record range from 0 to 5‰, showing the maximum and minimum within the Allerød period. δ 13 CTC and δ 13 CTOC reveal values between -31 to -17‰ 5 and -36 to -24‰, respectively. While δ 13 CTC show maximum values in 703 cm depth, δ 13 CTOC is decreasing continuously from the beginning to the end of the Allerød, followed by increasing values during the Younger Dryas and the Preboreal/Boreal, interrupted by a short decrease around the beginning of the Holocene. δ 13 CTC clearly show the presence of carbonate between 690 cm and 727 cm depth with more positive δ 13 CTC values compared to δ 13 CTOC values. 10 TOC/N ratios range from 5 to 16 with the end of the Allerød revealing increasing ratios, while the late Younger Dryas shows slightly decreasing ratios and the Preboreal is marked by the highest ratios.
( Fig. 3) The source of organic matter in lacustrine sediments of small lakes, as one of the most 15 crucial questions and challenges when interpreting organic proxies from lacustrine sedimentary records (Meyers and Ishiwatari, 1993), can either be autochthon (aquatic origin) or allochthon (terrestrial origin). The TOC/N ratio and δ 13 C values are most common proxies for sedimentary source determination. While non-vascular aquatic organisms often reveal C/N ratios between 4 and 10 (due to low amounts of cellulose and lignin), vascular plants show commonly C/N ratios 20 of 20 and higher (Meyers and Ishiwatari, 1993). After Meyers (2003), a TOC/N ratio of 10 is often used as threshold for identifying aquatic versus terrestrial input (Fig. 3E). Accordingly, input of terrestrial organic matter increased during the Allerød, decreased during the Younger Dryas and was highest during the Holocene. The Gemündener Maar δ 13 CTOC values (Fig. 3D) are well in range of C3 land plants and lacustrine algae (Meyers and Lallier-Vergés, 1999); 25 evidence for the occurrence of C4 land plants is missing. Overall, no clear additional information about the sedimentary organic matter origin of the Gemündener Maar sediments can be inferred neither from δ 13 CTOC alone (c.f. Lücke et al., 2003), nor by combining δ 13 CTOC with TOC/N ratios (c.f. Meyers and Lallier-Vergés, 1999). When considering that both δ 13 CTOC and TOC/N values of terrestrial organic matter are additionally affected by mineralization and 30 degradation, resulting in more positive δ 13 CTOC values and lower TOC/N ratios (e.g. Zech et al., 2007), a straightforward interpretation of those proxies seems to be challenging. Similarly, δ 15 N has been investigated as proxy for sedimentary organic matter origin (Meyers and Ishiwatari, 1993;Meyers and Lallier-Vergés, 1999;Wolfe et al., 1999). However, numerous possibly influencing processes like nitrogen uptake by plants, various nitrogen sources, discrimination during denitrification and diagenesis complicate the use of δ 15 N as direct source determination proxy.
Despite the above presented uncertainties, concerning the origin of bulk sedimentary organic matter in the Gemündener Maar, the origin of the sedimentary biomarkers, namely n-5 alkanes and sugars, needs to be addressed. This is crucial because aquatic biomarkers incorporate the isotope composition of lake water, whereas terrestrial biomarkers incorporate the isotope composition of leaf water (Huang et al., 2004;Kahmen et al., 2013;Mügler et al., 2008;Sachse et al., 2004Sachse et al., , 2012Sauer et al., 2001;Tuthorn et al., 2014;Zech et al., 2013bZech et al., , 2014b. With regard to the n-alkane biomarkers, high amounts of the chain lengths n-C27 and 10 n-C29 are characteristic for the Gemündener Maar sediments. Such patterns are typical for epicuticular leaf wax layers of higher terrestrial plants (e.g. Eglinton and Hamilton, 1967). With regard to the sugar biomarkers, they were previously studied in detail by Hepp et al. (2016).
According to the authors' own results and a compilation from the literature (including e.g. Jia et al., 2008;Prietzel et al., 2013;Zech et al., 2012Zech et al., , 2014b, relatively high amounts of arabinose 15 are a good indicator for a primarily terrestrial origin (higher vascular plants) of the sugars. This interpretation is in agreement with the Gemündener Maar being a small lake with densely forested steep crater walls (Fig. 1C). We therefore conclude and suggest that arabinose as well as n-C27 and n-C29 in our Gemündener Maar record are primarily of terrestrial rather than aquatic origin and thus reflect δ 2 H/δ 18 Oleafwater rather than δ 2 H/δ 18 Olake-water. 20
From the study of tree rings, it is known that stem cellulose does not show the full leaf water 18 O enrichment signal. Barbour and Farquhar (2000) related this signal damping to the proportion of unenriched source water contributing to the local synthesis water (px) and to the proportion of exchangeable oxygen during cellulose synthesis (pex). The latter is often assumed 25 to be rather constant around 0.40, as estimated from leaf and wood cellulose of Eucalyptus globulus and values compiled from literature (Cernusak et al., 2005), meaning that around 40% of the oxygens in the stem cellulose exchanged. Already Helliker and Ehleringer (2002) compared the signal transfer from leaf water to the cellulose of tree stems with the signal transfer occurring in grasses. And Liu et al. (2016) reported on signal damping in the range 30 between 34 and 53% for the C4 grass Cleistogenes squarrosa. δ 18 Ol = (δ 18 Oarabinose -ε 18 bio)/(1 + ε 18 bio/1000) (Eq. 8) δ 2 Hl = (δ 2 Hn-alkane -ε 2 bio)/(1 + ε 2 bio/1000) (Eq. 9) (Eq. 10) The correction presented in Eq. (10) is based on assumptions that 40% (0.4) of the leaf water enrichment is lost during hemicellulose biosynthesis of grass leaves, which is well in range with values presented in the literature for cellulose synthesis in tree rings and grasses, respectively (Cernusak et al., 2005;Liu et al., 2016). Furthermore, the Poaceae pollen concentration in percentage is used to calculate the fraction of non-grassy pollen [fnon-grasses = (100 -10 Poaceae)/100] corresponding to the non-grassy biomarker contribution, which may serve as rough approximation. For a paleo application, δ 18 Os remains a priori unknown. Therefore, the intercept between the individual LEL´s (Eq. 6) and the LMWL of Germany were used to generate δ 18 Os values. Note that the signal damping effect described here for cellulose synthesis is likely not fully applicable to our approach using the sugar biomarker arabinose. In fact, 15 pentoses like arabinose are biosynthesized via decarboxylation of the carbon at position six (C6) from glucose (Altermatt and Neish, 1956;Burget et al., 2003;Harper and Bar-Peled, 2002). Waterhouse et al. (2013) showed that the oxygens at C6 position in glucose moieties are most strongly affected by the exchange with local water medium of 80%, as indicated by heterotrophic cellulose synthesis. Thus, arabinose has lost a strongly exchanged (dampened) 20 oxygen and the remaining pentose shows less 18 O signal dampening.
With regard to the ε 2 bio value of −160‰, this biosynthetic fractionation factor is confirmed by climate chamber studies of dicotyledonous plants (Kahmen et al., 2011(Kahmen et al., , 2013Tipple et al., 2015). However, the latter studies reveal also a range of ~ 35‰, interpreted as species-specific effects during n-alkane biosynthesis. Much more pronounced is the difference 25 between dicotyledonous and monocotyledonous C3 plants regarding the degree to which the leaf water isotope enrichment is transferred into leaf n-alkanes (Gamarra et al., 2016;Kahmen et al., 2013). While dicotyledonous plants show signal transfer rates of 96% on average (Kahmen et al., 2013), a larger range between 38 and 61% is found for monocotyledonous plants ( Gamarra et al., 2016). The latter imply that 39 to 62% of the 2 H leaf water enrichment 30 is not recorded by the n-alkanes of grasses. Hence, like for δ 18 O a correction may be requested to account for grass-derived n-alkanes: δ 2 Hl # = {(δ 2 Hl -δ 2 Hs)/[fnon-grasses + (1 -0.5) -fnon-grasses • (1 -0.5)]} + δ 2 Hs, (Eq. 11) where δ 2 Hl* are the grass corrected δ 2 Hl values. The δ 2 Hs values and the non-grassy pollen fraction are defined as in Eq. (10). The mass balance correction presented in Eq. (11) is based on assumptions that only 50% of the leaf water enrichment is incorporated by the n-alkanes during biosynthesis in grass leaves.

Reconstructing relative humidity based on the coupled δ 2 Hn-alkane-δ 18 Osugar paleohygrometer approach
The biomarker-based leaf water values (δl = δ 2 Hl, δ 18 Ol via Eq. 8 and 9) result in dexcess values of leaf water (dl) ranging between -125 and -30‰ ( Fig. 5 and 6A). This is well within the range as it can be expected. the middle Preboreal reveals the most pronounced RH maximum. The mean reconstructed RHdv value is 53% (mean RHdv upper limit = 45%; mean RHdv lower limit = 62%; see section 2.5). 30 For comparison, the modern RHdv value (6 a.m. to 7 p.m. from April to October) from the adjacent meteorological station Nürburg-Barweiler (approx. 25 km northeast of Lake GM; hourly data from 1995 to 2015 from Deutscher Wetterdienst, 2016) is 67% (Fig. 6B). In addition, the range of the reconstructed RHdv values of 50% is well in agreement with the modern RHdv variability of 45%, within a range of 48% to 93% (definition and meteorological station details as above). As proposed in the previous chapter, three correction scenarios can be 5 applied when reconstructing dl and RHdv values in order to account for 2 H and 18 O signal dampening occurring in grasses.
( Fig. 6) Accordingly, the full correction for grass-derived alkane and sugar biomarkers (scenario 4 in Tab. 1) results in 0.0 to 6.3% (mean 1.8%) lower RHdv values (RHdv # * in Fig. 6B). This 10 corresponds to dl decreases of 0.0 to -12.0‰ (dl # * in Fig. 6A). Such small changes are still far below the pooled analytical standard errors . When only correcting for the 18 O signal dampening (scenario 3 in Tab. 1), dl values decrease by 0.0 to -22.7‰, corresponding to RH decreases of 0.0 to -12.0% (dl # and RHdv # in Fig. 6A and B, respectively). By contrast, when only correcting for the 2 H signal dampening (scenario 2 in Tab. 1), this leads to 0.0 to 10.6‰ more positive and 15 0.0 to 5.6% higher RHdv values (dl* and RHdv* in Fig. 6A and B). Overall, these results suggest that the reconstructed RHdv values are not strongly affected by 2 H and 18 O signal dampening of grasses.
We are aware, that microclimatic conditions with higher RH values often develop in lower canopy levels of forests (Graham et al., 2014;Parker, 1995). This may result in RH 20 overestimations when applying the coupled δ 2 Hn-alkane-δ 18 Osugar paleohygrometer approach.
However, most leaf biomass is produced in higher canopy levels being exposed to sunlight and free-air RH values. This is in agreement with a study of Zech et al. (2015), who investigated nalkanes in soils of the tropical montane rainforest of Mt. Kilimanjaro. There, n-alkanes reflect δ 2 Hleaf-water as calculated from free-air RH rather than as calculated from nearly saturated 25 ground-level RH.
A basic assumption of our coupled δ 2 Hn-alkane-δ 18 Osugar paleohygrometer approach is isotopic equilibrium between plant source water and water vapour. In order to test the robustness of this assumption and respective effects on reconstructed RH values, we used data of Jacob and Sonntag (1991), who measured the isotope composition of precipitation and of 30 atmospheric water vapour in Heidelberg, Germany, during the period 1981 to 1989. The mean difference between annual weighted means of precipitation (≈ plant source water) and the water vapour, which was calculated as average over the vegetation period (April-October). The derived apparent fractionation (ap) amounts to 18.3 and 1.57‰ in average for 2 H and 18 O, respectively. We used this ap in Eq. (1) instead of the difference δa -δs and recalculated the RH values. This recalculation leads to an average RH change of only -1.7% (±0.9), which is far below the analytical errors of the d-excess of leaf water.
Finally, the stability of the d-excess and slope of the LMWLGermany through the past 5 needs to be discussed. According to Stumpp et al. (2014), the long-term d-excess of precipitation from 28 sites in Germany does not show pronounced relationships to local climate conditions of the site. All reported values are close to 10‰, which indicates that Atlantic air masses are the main moisture source for Germany (e.g. Rozanski et al., 1993). In addition, the Europe suggest that the atmospheric circulation patterns over the European continent were rather constant throughout the past 35,000 years, which also implies that long-term d-excess of precipitation did not deviate substantially from the modern values (Rozanski, 1985). In 20 summary, the variations in the slope of the LMWL of Germany are assumed to be rather small over longer timescales.
The detailed discussions in the above three chapters address numerous uncertainties when using the coupled δ 2 Hn-alkane-δ 18 Osugar paleohygrometer approach. Conclusively, the reconstructed RHdv history of the Gemündener Maar seems, however, robust enough to infer 25 reliable paleoclimatic/-hydrologic implications.

How dry was the Younger Dryas in Western Europe?
While it is well known that the Younger Dryas was a cold spell occurring in the Northern Hemisphere during the Late Glacial (Denton et al., 2010;Heiri et al., 2014;Isarin and Bohncke, 30 1999), there is much less clear evidence concerning moisture supply/availability and RH changes during the Younger Dryas. The Gemündener Maar RHdv record suggests quite some variability but on average moderate RHdv conditions of ~ 56% during the end of the Allerød and the first half of the Younger Dryas. This is within the range of modern RHdv values (Fig.   6B). In the second half of the Younger Dryas, a clear RHdv decrease of ~ 11% occurred (Fig.   7A). Such a two phasing of the Younger Dryas has been suggested previously based on multiproxy climate data for western Europe (Isarin et al., 1998). In more detail, Isarin et al. (1998) reported on a cold and humid first phase being followed by drier and warmer conditions. 5 It is moreover speculated that a shift of the mean sea-ice margin during winter in the North Atlantic Ocean slightly to the north could have caused this two phasing. Reduced cyclonic activity and precipitation affected thereby primarily Western Europe because this region was situated at the southern margin of the main storm tracks during the first Younger Dryas period (Isarin et al., 1998). The authors presented also evidence for the strengthening of the westerly 10 winds in Western Europe as consequence of northwards shifted North Atlantic Ocean sea-ice margin during the late Younger Dryas period. This contradicts, however, with the interpretation of the Meerfelder Maar sedimentary record. Here, the thicker varves during the early Younger Dryas (between 12,680 and 12,240 varve a BP) are used along with geochemical results as indicator for stronger winter winds (Brauer et al., 2008). In line with this, Brauer et al. (1999) 15 interpreted high biogenic opal contents and Pediastrum remain concentrations during the early Younger Dryas as enhanced aquatic productivity due to an increased nutrient supply caused by soil erosion and the reworking of littoral sediments. The varve formation throughout the second Younger Dryas period (between 12,240 and 11,590 varve a BP) is interpreted to be mainly controlled by snowmelt driven surface runoff (Brauer et al., 1999). Moreover, the authors 20 speculate if during that time the Meerbach creek began to drain into the Meerfelder Maar, which could be possibly linked to enhanced precipitation amounts. In summary, the interpretations derived from the Younger Dryas sediments of the Meefelder Maar by Brauer et al. (2008Brauer et al. ( , 1999 seem neither to be in accordance with the results of Isarin et al. (1998) Maar RH record not corroborating an overall dry Younger Dryas. While the uncertainties of the 30 coupled δ 2 Hn-alkane-δ 18 Osugar paleohygrometer approach were discussed in detail in the previous chapters, according to our opinion the most important uncertainties affecting the dual biomarker approach of Rach et al. (2014Rach et al. ( , 2017 are the following. First, lake water is assumed to reflect δ 2 H of precipitation. Indeed, Lake Holzmaar, which seems to be comparable to the Meefelder Maar at least for the drainage conditions via one creek, shows a difference of 7.4‰ in δ 2 H between inflow and lake water (Sachse et al., 2004). This lake water enrichment is likely to have been variable in the past, especially when including the speculation concerning the drainage of the Meerbach creek during the Younger Dryas (Brauer et al., 1999). Second, n-C23 is interpreted to be of aquatic origin (from Potamogeton) and used for reconstructing δ 2 Hlake-5 water. However, there is increasing evidence that n-C23 is also of terrestrial origin (Rao et al., 2014). For instance, Aichner et al. (2018) have recently shown for a lake in Poland that n-C23 is a variable mixture of aquatic and terrestrial origin in those Late Glacial and Early Holocene sediments. And birch as pioneering and one of the dominant tree species during Late Glacial reforestation of Central Europe is known to produce considerable amounts of mid-chain n-10 alkanes (Tarasov et al., 2013). Albeit not included in the latter publication, n-C23 concentrations of Betula exilis and Betula pendula reached 653 and even 2323 μg/g in that study. This is highly relevant, because the biosynthetic fractionation factor of aquatic n-alkanes is much smaller than the one of terrestrial n-alkanes. Minor changes in the contribution of terrestrial vs. aquatic nalkanes will thus have a considerable impact on the reconstructed δ 2 H n-C23 record and in turn 15 on reconstructed RH values when applying the dual biomarker approach. Finally, it may be noteworthy to acknowledge, that Sachse et al. (2004) found no significant correlation for δ 2 H of n-C23 and lake water and precipitation along a European lake surface transect.
Also recently and also applying the dual biomarker approach, Muschitiello et al. (2015) studied Younger Dryas lake sediments from Hässeldala Port in Southern Sweden. Here, the 20 authors used δ 2 H of n-C21 as proxy for lake water, respectively summer precipitation.. The calculated difference between terrestrial and aquatic n-alkane δ 2 H values suggests more humid conditions at the beginning of the Younger Dryas followed by a more or less steady trend towards drier conditions, peaking around 11,700 a BP (Muschitiello et al., 2015). This would be, within age uncertainties, in line with Gemündener Maar RHdv minimum between ~ 11,700 25 and 11,900 a BP. Last but not least, Gázquez et al. (2018) analyzed triple oxygen and hydrogen isotopes of gypsum in the Southern Pyrenees and thus reconstructed RH changes. Again, more humid conditions are reported for the beginning of the Younger Dryas.
( Fig. 7) In search of possible drivers/mechanisms for the observed Gemündener Maar RHdv 30 record we came across the 14 C production and 10 Be flux rates (Fig. 7B), derived from IntCal13 and the Greenland ice-cores (GRIP, GISP2), respectively (Muscheler et al., 2014). These records are commonly interpreted in terms of solar activity (and thus insolation) changes (Stuiver and Braziunas, 1988;Vonmoos et al., 2006) and reveal striking similarities with our Gemündener Maar RHdv record. For instance, all three records reveal quite high centennialscale variability during the Allerød and the first half of the Younger Dryas. Generally low RHdv values during the second half of the Younger Dryas and the Early Preboreal coincide with high solar activity, whereas the pronounced RHdv maximum from 11,260 to 11,050 a BP coincides within age uncertainties with a pronounced solar activity minimum ( Fig. 7A and B). We dub 5 this wet period the 'Preboreal Humid Phase', which should not be confused with the Preboreal Oscillation (Björck et al., 1997). The Preboreal Oscillation is a short cold event recorded in Greenland ice-cores ~11,400 ka (Rasmussen et al., 2007) and led to more arid conditions at least in the Netherlands according to palynological results (Bos et al., 2007;van der Plicht et al., 2004). These pollen records also show the existence of a pronounced humid phase 10 thereafter, thus corroborating the 'Preboreal Humid Phase'. Widespread glacial advances in the Alps are also attributed to the Preboreal Oscillation (Moran et al., 2016). However, given the dating uncertainties they may actually rather reflect increased precipitation during the Preboreal Humid Phase.
It should be emphasized, that the described similarities between the Gemündener Maar 15 RHdv record and the solar activity records do not allow an a priori causality interpretation. It is widely accepted that the Younger Dryas and the Preboreal Oscillation are related to freshwater forcing in the North Atlantic (e.g. Fisher et al., 2002;Murton et al., 2010;Muschitiello et al., 2015). However, the causes and mechanisms responsible for climate and environmental changes during the rest of Holocene remain vague, and more research including paleoclimate 20 modelling is clearly needed and encouraged to investigate the possible influence of solar activity (Renssen et al., 2007;Rind, 2002). We propose that both the North Atlantic Ocean temperature and solar activity, which triggered solar insolation, were the two main drivers for the RHdv variability in Central Europe. A key example might be the Preboreal Humid Phase. It can be expected that the North Atlantic Ocean, the main moisture source for Central Europe, 25 revealed already considerably higher temperatures during the Preboreal Humid Phase compared to the Younger Dryas, as indicated by a consistent ~ 2°C increase in planktonic foraminifera (Globorotalia inflata, Globorotalia bulloides and Neogloboquadrina pachyderma) derived Mg/Ca temperatures from a marine sediment core south of Iceland (Fig. 7B, Thornalley et al., 2009Thornalley et al., , 2010Thornalley et al., , 2011. This led to an enhanced moisture content of the atmosphere. When these 30 wet air masses were transported onto continental Europe, where low solar insolation inhibited warming up and drying of these air masses, pronounced humid climate conditions were established.

Conclusions
Referring to the underlying research questions and based on the presented results and the outlined discussion (including the cited literature) the following conclusions have to be drawn:  The terrestrial vs. aquatic origin of bulk sedimentary organic matter cannot be 5 determined unambiguously for the Gemündener Maar. This is caused by the bulk proxies (TOC/N, δ 13 C and δ 15 N) being not straightforward interpretable. By contrast, the alkane biomarkers with the chain-length n-C27 and n-C29 and the sugar biomarker arabinose can be most likely associated with the epicuticular leaf-wax layers and the hemicellulose structures of higher terrestrial plants, respectively. Therefore, they are 10 interpreted to originate primarily from leaf material of the Gemündener Maar catchment.
 δ 2 H/δ 18 Oleaf-water could be reconstructed from δ 2 Hn-alkane (n-C27 and n-C29) and δ 18 Osugar occurring from ~11,260 to 11,050 a BP. We propose North Atlantic Ocean temperature and solar activity (and thus insolation) to be the main drivers for Late Glacial-Early 30 Holocene RH changes in Central Europe and encourage respective paleoclimate modelling studies in order to validate or falsify our proposition.