Climate impacts on deglaciation and vegetation dynamics since the Last Glacial Maximum at Moossee ( Switzerland )

Since the Last Glacial Maximum (LGM, end ca. 19,000 cal BP) Central European plant communities were shaped by changing climatic and anthropogenic disturbances. Understanding long-term ecosystem 10 reorganizations in response to past environmental changes is crucial to draw conclusions about the impact of future climate change. So far, it has been difficult to address the post-deglaciation timing and ecosystem dynamics due to a lack of well-dated and continuous sediment sequences covering the entire period after the LGM. Here, we present a new palaeoecological study with exceptional chronological time control using pollen, spores and microscopic charcoal from Moossee (Swiss Plateau, 521 m a.s.l.) to reconstruct the vegetation and 15 fire history over the last ca. 19,000 years. After lake formation in response to deglaciation, five major polleninferred ecosystem rearrangements occurred at ca. 18,800 cal BP (establishment of steppe tundra), 16,000 cal BP (spread of shrub tundra), 14,600 cal BP (expansion of boreal forests), 11,600 cal BP (establishment of first temperate deciduous tree stands composed of e.g. Quercus, Ulmus, Alnus) and 8200 cal BP (first occurence of mesophilous Fagus sylvatica trees). These vegetation shifts were released by climate changes at 19,000, 16,000, 20 14,700, 11,700 and 8200 cal BP. Vegetation responses occurred with no apparent time lag to climate change, if the mutual chronological uncertainties are considered. This finding is in agreement with further evidence from Southern and Central Europe and might be explained with proximity to the refugia of boreal and temperate trees (< 400 km) and rapid species spreads. Our palynological record sets the beginning of millennial-scale land use with periodically increased fire and agricultural activities of the Neolithic period at ca. 7000 cal BP (5050 cal 25 BC). Subsequently, humans rather than climate triggered changes in vegetation composition and structure. We conclude that Fagus sylvatica forests were resilient to long-term anthropogenic and climatic impacts of the mid and the late Holocene. However, future climate warming and in particular declining moisture availability may cause unprecedented reorganizations of Central European beech-dominated forest ecosystems.


Introduction 30
In the near term, rapid climatic and environmental changes hold a substantial risk to irreversibly modify plant ecosystems in Europe (Schumacher and Bugmann, 2006;Kovats et al., 2014;Bugmann et al., 2015).
Quantifying the response or resilience of ecosystems to environmental change in the past largely improves our capacity to assess future impacts of climate and global change (Henne et al., 2015).Specifically, palaeoecological data offer the great opportunity to study long-term climate-vegetation interactions under 35 conditions that exceed the variability and duration recorded in historical archives or through measurements and experiments (Willis and Birks, 2006;Birks et al., 2016;Henne et al., 2018).During the Last Glacial Maximum (LGM), large areas in Central and Southern Europe around the Alps, in the Jura Mountains, the Black Forest, the Vosges and the Apennines were covered by ice (Ehlers and Gibbard, 2004;Bini et al., 2009;Ehlers et al., 2011;Seguinot et al., 2018).The subsequent deglaciation is generally well-40 studied, however, timing issues remain due to dating uncertainties (e.g.Wirsig et al., 2016).Recent advances in Accelerator Mass Spectrometry (AMS) radiocarbon dating offer the possibility to produce reliable results with relatively small chronological uncertainties, when using samples with extremely low carbon contents (Szidat et al., 2014;Uglietti et al., 2016).Radiocarbon dates on terrestrial plant remains extracted from the very bottom of lake sediments from sites close to the LGM glacier margins may thus help to refine the onset of deglaciation.45 However, only very few sedimentary records providing reliable deglaciation ages (i.e.no bulk dating, only terrestrial macrofossils, see Finsinger et al., 2019) are available so far from the peri-alpine lowlands (e.g.Lister, 1988;Laroque and Finsinger, 2008;Lauterbach et al., 2012).Similarly, well-dated pollen profiles covering the first two millennia of the Oldest Dryas (ca.19,000-17,000 cal BP) are almost absent and the existing chronologies are therefore inadequate (e.g.Welten, 1982;Ammann and Tobolski, 1983;Becker et al., 2006).50 Conversely, the temporal evolution of the vegetation after 17,000 cal BP is better known (e.g.Lotter, 1999;Tinner et al., 1999;Rey et al., 2017).Various sites south of the Alps indicate a first afforestation after 16,000 cal BP (e.g.Vescovi et al., 2007) and the main cause has been identified as the post Heinrich event (HE) 1 warming (Samartin et al., 2012).At around 16,000 cal BP shrub and possibly even tree birches expanded into the steppe tundra north of the Alps (Lotter, 1999;Rey et al., 2017), forming open parklands or shrub tundra.North of the 55 Alps, forests expanded after 14,700 cal BP as a consequence of the Bølling warming (see Ammann et al., 2013;van Raden et al., 2013), a process which was delayed by almost 1500 years compared to the lowlands south of the Alps (Vescovi et al., 2007).The reasons causing this long time-lag are not yet fully understood but might be related to a strong latitudinal temperature gradient and the presence of large ice masses (Heiri et al., 2014).The subsequent forested Late Glacial and Holocene vegetation history of the Swiss Plateau is best-studied and the 60 chronological framework is rather robust (e.g.Lotter et al., 1999;Wehrli et al., 2007;Rey et al., 2017).
Taken together, despite the long tradition of palaeoecological research in Switzerland with quite a high density of well-dated and highly resolved studies, a profound modern assessment of the major vegetation changes and their main causes is currently lacking.Here, the novel Moossee record has the great potential to shed new light on the timing of lake formation and on important vegetation reorganizations for the past 19,200 years in a central 65 area of the Swiss Plateau.In this study we aim (1) to reconstruct the timing of deglaciation and the establishment of first pioneer vegetation around the lake after the LGM, (2) to identify major postglacial changes in ecosystem evolution on the Swiss Plateau and to assess their causes, (3) to discuss the resilience and the vulnerability of Central European lowland forests in the past to inform the near future and (4) to emphasize the utility of exceptional temporal precision and resolution.70

Study site
Moossee is a small eutrophic lake at 512 m a.s.l.(47°1′17.0′'N,7°29′1.7′'E)located on the Swiss Plateau within the periphery of the Swiss capital Bern.The study area geologically belongs to the carbonate-rich molasse region with predominantly sandstones between the Jura Mountains in the North and the Alps in the South (Schmid et al., 2004).The lake formed after the retreat of the Rhône glacier after the LGM and has a surface area of 0.31 75 km 2 , with one main inflow in the West and one outflow in the East (Fig. 1b).The maximum water depth is 22 m, https://doi.org/10.5194/cp-2019-121Preprint.Discussion started: 7 October 2019 c Author(s) 2019.CC BY 4.0 License.
with generally anoxic waters in the hypolimnion below 12 m (Guthruf et al., 1999).The lake used to be at least ten times larger in the past, but its size has shrunk due to peat formation over the millennia and artificial lake level lowering (by 4-5 m) since the late 18 th century.Lake levels were lowered to drain the wetlands for peat exploitation and to gain agricultural land (von Büren, 1943;Guthruf et al., 1999;Harb, 2017, Fig. 1b).The 80 climate at Moossee is oceanic with mean annual temperatures of 8.8 °C and an annual rainfall of 1059 mm (data from from Bern-Zollikofen at ca. 3 km distance, MeteoSwiss, 2017).July is the warmest month with a mean temperature of 18.6 °C.The wettest months are May-August with more than 100 mm of rainfall per month (MeteoSwiss, 2017).
Alnus glutinosa and Fraxinus excelsior form fragmented stands along the lake shore, whereas mixed Fagus 85 sylvatica forests are dominant on the more elevated surrounding hills.The remaining and rather flat areas are either intensively used for agriculture or covered by settlements and infrastructure.The earliest archaeological findings around the lake date back to the Magdalenien and the Upper Palaeolithic ca.15,950-14,750 cal BP (14,000-12,800 cal BC).At that time, two reindeer hunter camps (Moosbühl I and II) were located at the former lake shore (Bullinger et al., 1997;Harb, 2017, Fig. 1b).Many finds, including preserved lake shore villages, are 90 known from the Neolithic from ca. 6450 cal BP on (4500 cal BC), impressively documenting the strong prehistoric human activities in the region (Hafner et al., 2012;Harb, 2017, Fig. 1b).The following Bronze Age and Early Iron Age is represented with scattered artifacts and grave mounds in the proximity of the lake (Harb, 2017).In the Bern area the first urban center was the oppidum Brenodor, which was built during the Late Iron Age or La Tène period (Ebnöther and Wyss, 2004).It persisted during the Roman Age as vicus Brenodurum.95 The Iron Age and Roman ruins (e.g.fortifications, bath, amphitheatre) are still visible in town on the Enge peninsula at 5 km distance from Moossee (Ebnöther and Wyss, 2004).Finally, the medieval city center, which is part of the UNESCO World Heritage "Old City of Bern", was founded in 1191 AD around the Nydegg castle that already existed before (Hofer and Meyer, 1991).

Coring and chronology
Six parallel sediment cores were retrieved at 19 m water depth with an UWITEC piston corer in the eastern part of the lake.Three cores (Moos A-C, core diameter: 60 mm, core length: 300 cm) reached coring depths of ca.17.5 m.For the other three cores (Moos F-H, core diameter: 90 mm, core length: 200 cm), due to higher friction it was only possible to recover the uppermost 7 m.A master sequence with a total length of 16.44 m was defined 105 using the Moos F-H cores for the uppermost 7 m and the Moos A-C cores for the remaining part.The sediment material below 13.5 m was not analyzed due to frequent sand layers in the lowermost part resulting in very low pollen concentrations, The chronology is based on 62 radiocarbon dates on terrestrial plant macrofossils and the Laacher See Tephra (LST; see Table 1).The radiocarbon content of terrestrial plant remains was measured at the Laboratory for 110 Radiocarbon Analysis (LARA) at the University of Bern using accelerator mass spectrometry (AMS, see Szidat et al., 2014).From 435-691 cm, additional varve counts were applied to refine the chronology.Here, the program OxCal 4.3 (V-sequence, Bronk Ramsey, 1994Ramsey, , 1995Ramsey, , 2001;;Bronk Ramsey et al., 2001) and the IntCal13 calibration curve (Reimer et al., 2013) were used to estimate the age-depth model and its 95 % (2σ) probabilities (partly published, see Rey et al., 2019b).For the remaining part (0-435 cm and 691-1335 cm), a https://doi.org/10.5194/cp-2019-121Preprint.Discussion started: 7 October 2019 c Author(s) 2019.CC BY 4.0 License.
smooth-spline curve (smoothing level = 0.3) was calculated with the program clam 2.2 (Blaauw, 2010) to assess the final age-depth model (Fig. 2).The modelled curve runs within the 95 % (2σ) probabilities of the calibrated radiocarbon ages and the 2σ confidence envelope of a generalized mixed-effect regression (GAM, Heegaard et al., 2005).

Pollen, non-pollen palynomorphs and charcoal analysis 120
A total of 514 samples for pollen and microscopic charcoal analyses were taken from the sediment core from 1296 cm to the top.The standard sampling was 1 cm 3 every 10 cm.A higher resolution was implemented for the Oldest Dryas (18,800-14,700 cal BP) and for the Neolithic-Mid Bronze Age (7400-3200 cal BP, see Rey et al., 2019a).All palynological samples were treated with HCl, KOH, HF, acetolysis, sieved with a mesh size of 0.5 mm and mounted in glycerine following standard approaches (Moore et al., 1991).Lycopodium tablets 125 (University of Lund batch no.1031 with 20,848 ± 3457 spores per tablet) were added before the chemical treatment to assess microfossil concentrations (Stockmarr, 1971).Pollen, spores and non-pollen palynomorphs (NPPs) were identified under a light microscope at 400× magnification using palynological keys (Moore et al., 1991;Beug, 2004), photo atlases (Reille, 1992) and the reference collection at the Institute of Plant Sciences (University of Bern).Betula nana and tree Betula pollen grains were separated following Birks (1968) and Clegg 130 et al. (2005).Cerealia-type pollen was identified according to size, pore diameter and annulus thickness (see Beug, 2004).
Pollen and spores were used to infer extra-local to regional vegetation dynamics (Conedera et al., 2006).A minimum pollen sum of 500 terrestrial pollen grains per sample was counted.For the lowest part of the sediment core, the minimum pollen sum was > 100 terrestrial pollen grains.Sporormiella (coprophilous fungal spore, see 135 e.g.van Geel et al., 2003) was used as a proxy for grazing activities of herbivores (e.g.Gill et al., 2013;Rey et al., 2017) and livestock farming (e.g.Rey et al., 2013;Schwörer et al., 2015).The pollen and NPP results are presented in percentages of the terrestrial pollen sum excluding Cannabis sativa pollen (due to artificial pollen input by hemp retting, see Ranalli and Venturi, 2004) and pollen of aquatic plants (Fig. 4).
We used microscopic charcoal as a proxy for regional fire activity (Tinner et al., 1998;Adolf et al., 2018).140 Particles > 10 µm and < 500 µm were analyzed and counted on the pollen slides following Tinner and Hu (2003) and Finsinger and Tinner (2005).The data are presented as microscopic charcoal influx values (particles cm -2 yr - 1 , Fig. 4).Local pollen assemblage zones (LPAZ) were identified using optimal sum-of-squares partitioning (Birks and Gordon, 1985), while statistically significant zones were determined following the broken-stick method (Bennett, 1996).All calculations were run with the program R statistics ( R Development Core Team, 145 2018).The data was plotted with the use of the programs Tilia 2.0.60 and CorelDraw.

Biodiversity estimations and ordination analysis
We first applied rarefaction analysis to calculate palynological richness (PRI), which is frequently used as a proxy for local to regional species richness (e.g.Birks and Line, 1992;Odgaard, 1999;Schwörer et al., 2015;Rey et al., 2019a).Rarefaction analysis assesses the number of taxa per sample after setting a constant minimum 150 terrestrial pollen sum (Birks and Line, 1992), which was 116 in our case.Subsequently, the probability of interspecific encounter (PIE, Hurlbert, 1971) was taken as a measure of palynological evenness (van der Knaap, 2009).To assess distortion biases related to pollen production and dispersal on PRI (e.g. through high-pollen producers such as Pinus sylvestris, Corylus avellana), evenness-detrended palynological richness (DE-PRI) was https://doi.org/10.5194/cp-2019-121Preprint.Discussion started: 7 October 2019 c Author(s) 2019.CC BY 4.0 License.calculated (see Colombaroli and Tinner, 2013).The aim of this procedure is to remove the evenness trend from 155 PRI by applying an ordinary least square regression (OLS) between PRI and PIE and adding the deriving residuals (PRI -PIE) to the original PRI values (residuals + PRI).Only if PRI and DE-PRI indicate comparable trends and changes, we suppose that PRI is uninfluenced by evenness effects and thus primarily reflects species richness.The program R statistics (R Development Core Team, 2018) was applied for the calculations.
We used ordination analysis (Birks and Gordon, 1985;Ter Braak and Prentice, 1988) to identify gradients in 160 vegetation composition over time using the program Canoco 5 (Šmilauer and Lepš, 2014).We first performed a detrended correspondence analysis (DCA, Birks and Gordon, 1985) by segments without down-weighting of rare species to assess the appropriate response model (i.e.unimodal vs. linear) for our pollen percentage data.
Since there is important turnover in the species composition as indicated by the rather long gradient length of the first DCA axis (3.41 SD), we retained the unimodal response model (DCA, Šmilauer and Lepš, 2014).165

Lithology and sedimentation
The lowermost part of the sediment sequence (1644-1300 cm, 18,900 cal BP and older), which was not analyzed for pollen and microscopic charcoal, consists of clay and sand layers.From 1300 cm (18,900 cal BP) upward, the sediment content changes to silty clay without sand layers until 980.5 cm (14,400 cal BP).According to the 170 age-depth model (Fig. 2), the sedimentation rate is steadily decreasing at the same time, suggesting the establishment of relatively stable soil conditions shortly after deglaciation.From 980.5 cm (14,400 cal BP) to the top, the sediment consists of calcite-rich fine detritus gyttja.Between 712-429 cm (7550-2650 cal BP, see Rey et al., 2019a, b) the sediment is continuously varved.Some additional partly laminated sections are present at 980.5-768 cm (14,400-8900 cal BP) and at 415−134 cm (2350-300 cal BP).The sedimentation rate becomes 175 rather stable from 980.5 cm (14,400 cal BP) onward and stays more or less constant until 429 cm (2650 cal BP).
This stabilization is probably linked to forests growing in the catchment reducing the total erosional input.The uppermost 429 cm (2650 cal BP to the present) are characterized by a steep increase of the sedimentation rate, which is most likely related to increased erosion in response to forest openings and agricultural activities in the catchment of the lake.180

Vegetation and fire history
The pollen sequence (Fig. 3) is subdivided into 21 local pollen assemblage zones (LPAZ) and five subzones 1b,20a,20b and 20c).The high number of statistically significant zones, especially in the upper part of the diagram (i.e. at 7300-2900 cal BP) is related to the exceptionally high sample resolution and the rather strong vegetation changes.For simplification and better comparison of the LPAZ, we added important climatic 185 breaks (1)(2)(3)(4)(5) that are either related to temperature changes (breaks 1-4, see compilation of chironomid-based July air temperature estimate since the LGM in Finsinger et al. (2019) or increases of moisture availability (break 4, e.g.Tinner and Lotter, 2001;Joannin et al., 2013).This climate synopsis allows for the first time a tentative regional assessment and discussion of climate amplitude variation and its impacts on vegetation for the past 19,200 years. 190 Herbaceous pollen grains from Poaceae, Artemisia, Helianthemum, Thalictrum, Chenopodiaceae, Rubiaceae and including some first pioneer dwarf shrubs such as Salix herbacea (macrofossils found, see Table 1), Betula nana and Juniperus quickly after deglaciation (not later than 19,200 cal BP as indicated by the oldest radiocarbon date 195 in Table 1).The slightly increased pollen percentages of Pinus sylvestris-type (up to 18 %) and single grains of Abies alba and Picea abies might point to reworking processes and/ or long-distance transport.High percentages of Sporormiella (up to 6 %) as a proxy for local grazing (van Geel, 2006) may be indicative of the presence of Pleistocene megafauna such as Mammuthus primigenius (wolly mammoth), Coelodonta antiquitatis (wolly rhinoceros), Rangifer tarandus (reindeer) and others that were preferentially living in the cold steppe 200 environment at that time (e.g.Nielsen, 2013).Charcoal influx values are extremely low (< 50 particles cm -2 yr -1 ), suggesting rare or almost no fire activity in the region due to scarce vegetation cover.
From 18,800-16,000 cal BP (between climatic breaks 1 and 2, LPAZ Moos-1a), the pollen record indicates that an open, species-rich (see Gypsophila repens-type and Rumex acetosa-type pollen curves), herb dominated steppe tundra persisted around Moossee.A first slight increase of shrub pollen percentages (from 7 to 14 %) 205 mainly from Betula nana, Salix, Juniperus and Ephedra fragilis-type suggest that patches of dwarf shrubs established in the region after 17,000 cal BP.This is underlined by findings of Betula nana remains (see Table 1).Sporormiella percentages remain high (at 1-5 %), pointing to prevalence of wild animals grazing at the lake.
Pollen percentages of tree Betula are markedly increasing (values up to 21 %) after the end of HE-1 (climatic 210 break 2) during LPAZ Moos-1b (16,000-14,600 cal BP), suggesting the regional establishment of tree Betula stands or woods.Shrub pollen percentages (i.e.Betula nana, Ephedra fragilis-type, Salix, Juniperus and Hippophaë rhamnoides) stay at 12-20 %, whereas non-arboreal pollen (= NAP) values decrease but remain very high (55-75 %, see pollen percentages of Artemisia, Heliathemum, Thalictrum and Chenopodiaceae).This change points to an expansion of shrub tundra with Betula nana, Salix, Juniperus and maybe even some small 215 growing Betula trees into the catchment of Moossee and is possibly related to climate warming after 16,200 cal BP (Samartin et al., 2012;Finsinger et al., 2019).The values of Sporormiella diminish and fade out at the same time, suggesting that some of the megaherbivores producing a lot of dung (e.g.Mammuthus primigenius) may have become extinct locally (see Nielsen, 2013;Cupillard et al., 2015).Charcoal influx prevail at low values (< 40 particles cm -2 yr -1 ), indicating that despite higher total biomass and fuel availability, regional fire activity did 220 not increase.
Subsequently, pollen of Pinus sylvestris-type steadily increases, suggesting the establishment of birch-pine forest around the lake (LPAZ Moos-4, 13,750-11,050 cal BP).These boreal forests prevailed through the Allerød 230 warm period (13,900-12,900 cal BP, van Raden et al., 2013), the subsequent Younger Dryas (12,700-11,700 cal BP, Finsinger et al., 2019) and the during the first centuries of the Holocene, as suggested by the pollen assemblages.However, the dominance of Pinus sylvestris-type and the decrease of Betula pollen after 12,900 cal  (Tinner and Lotter, 2001;Heiri et al., 2015).However, at 9000-8600 cal BP, Fraxinus excelsior pollen percentages increase and tree pollen indicative of mixed oak forests (Quercus,250 Ulmus, Tilia, Acer and Fraxinus excelsior, e.g.Welten, 1982) reach their highest values.Charcoal influx values remain low (Fig. 3), suggesting no significant increase of fire activity during the Early Holocene.
During LPAZ Moos-6 (8600-7250 cal BP), Corylus avellana declined, and less heliophilous trees such as Fraxinus excelsior and Tilia expanded.Likely in response to climate change around 8200 cal BP (climatic break 5, onset of mid Holocene), Abies alba and Fagus sylvatica pollen reach their empirical limit (i.e.continuous 255 curves at 8400 and 8200 cal BP respectively), suggesting the local establishment of first stands of these tree species (Birks and Tinner, 2016).At the same time, Alnus glutinosa-type pollen percentages are steadily increasing, whereas percentages values of mixed oak forest and Corylus avellana start to decline.The pollen assemblages suggest that the decline of heliophilous deciduous forests continued after 7250 cal BP (LPAZ Moos-7 to 8, 7250-6400 cal BP), when mixed beech-silver fir forests expanded massively.The general 260 prevalence of mesophilous tree species throughout LPAZ Moos-7 to LPAZ Moos-21 was likely caused by a gradual shift towards more oceanic climate conditions during the Mid-Late Holocene (Tinner andLotter 2001, 2006), as e.g.reconstructed on the basis of higher lake levels (Magny, 2013;Joannin et al., 2013).However, this long-term dominance of dark mesophilous mixed beech-silver fir forests and the replacement of the formerly widespread mixed oak-linden-elm-maple forests was also affected by agricultural activities, starting as early as 265 might point to forest management favoring oak for construction and forest pasture (acorn feeding, e.g.Gobet et al., 2000;Wick, 2015).Sporadic Sporormiella fungal spores suggest pastoral farming close to the lake.Charcoal 285 influx values generally follow the land use phases showing two major peaks at 5600 cal BP (3650 cal BC) and 700 cal BP (cal AD 1250), pointing to two phases of highest fire activity during the past 19,200 years (with up to 26,000 particles cm -2 yr -1 ).The close link to pollen of crops and weeds as well as the related declines of forests, suggest that anthropogenic burning was related to slash-and-burn activities or maintenance of open fields (Tinner et al., 2005).290

Biodiversity reconstruction and ordination
PRI (palnyological richness) and DE-PRI (eveness-detrended palynological richness) are very similar suggesting that overall, PRI is likely unaffected by evenness effects (Fig. 3).The agreement is particularly good during the periods 18,900-14,500 cal BP and 8600 cal BP to present.Here, both PRI and DE-PRI are slightly fluctuating around 15 pollen types per sample.Palynological evenness as inferred from PIE is stable (PIE at 0.8−0.9) in 295 phases where PRI and DE-PRI are in agreement.Significantly lower values (PIE down to 0.5) are recorded from 14,500-8600 cal BP when pollen grains from few taxa are dominant (either Juniperus, Betula, Pinus sylvestristype or Corylus avellana).There, palynological richness drops (PRI < 10 pollen types per sample) whereas DE-PRI stays stable at around 15 pollen types per sample.We thus assume that evenness distortions lead to underestimations of species richness during the period of strong Juniperus, Betula, Pinus (ca.14,500-11,100 cal 300 BP) and Corylus dominance (10,800-9000 cal BP, Fig. 3) and that such evenness distortions can be corrected by considering DE-PRI.
Both richness values (PRI, DE-PRI) generally increase (up to 20-25 pollen types per sample) during phases with higher human impact around 5650 cal BP (3700 cal BC), at 4650 cal BP (2700 cal BC), around 3850 cal BP (1900 cal BC) and 3500 cal BP (1550 cal BC) as well as after 2600 cal BP (650 cal BC).These increases are 305 directly linked to human induced forest openings and the introduction of cultivated plants (e.g.Cerealia-type) as well as the expansion of weeds (e.g.Plantago lanceolata), apophytes (e.g.Urtica) and heliophilous shrubs (e.g.

Corylus avellana
).Interestingly, also the tundra phase (18,900-14,600 cal BP) was rather species rich suggesting that PRI and DE-PRI are correlated with openness.A rather low share (25.6 %) of the total pollen data variance is explained by DCA axis 1.Nevertheless, the DCA scores might indicate a signal of openness as the DCA axis 310 https://doi.org/10.5194/cp-2019-121Preprint.Discussion started: 7 October 2019 c Author(s) 2019.CC BY 4.0 License.

Climate-driven deglaciation and vegetation dynamics from the Oldest Dryas to the Mid Holocene
The extent of the ice sheet around the Alpine arch during the LGM has been thoroughly studied in the past and the results are available as high-quality maps (e.g.Ehlers and Gibbard, 2004, see Fig. 1a).Radiocarbon dates on 315 terrestrial plant remains preserved in the bottom part of lake sediment sequences can be used to track the onset of deglaciation (Wirsig et al., 2016;Rey et al., 2017).At Moossee, the earliest date gives a calibrated age of 19,200 cal BP (Table 1), which is well in line with the results from other lowland locations around the Alps (Table 2).
The unpublished basal date of Lago di Monate (lab.code BE-8023.1.1)derives from a new lake sediment core that is currently under investigation.Interestingly, most of the sites (including Moossee) with calibrated ages 320 older than 18,000 cal years were located rather close to the margin of the former ice sheet (Fig. 1a, Table 2) and below < 500 m of ice (Bini et al., 2009).In contrast, sites with younger radiocarbon ages (17,270-17,770 cal BP, e.g.Lotter and Zbinden, 1989;Tinner et al., 1999;Ravazzi et al., 2014) were generally situated below a thicker ice sheet (750-1000 m) and further away from the glacier tongues.However, the relatively small age difference between all these sites suggests that the collapse of the ice sheet in the lowlands south and north of the Alps 325 occurred within 1000-1500 years, starting not later than 19,300 cal BP at the end of the LGM (23,000-19,000 cal BP for the Alps, Kaltenrieder et al., 2009;Hughes et al., 2013;Samartin et al., 2016).The huge loss of ice masses and the sudden retreat of glaciers were likely controlled by increasing summer insolation (Berger and Loutre, 1991) as well as constantly rising CO2 and CH4 concentrations in the atmosphere (Lourantou et al., 2010).330 The pollen assemblage shows that pioneer plants colonized the bare grounds around Moossee shortly after glacial retreat (ca.19,000 cal BP) to quickly form open, species rich and herb-dominated steppe tundra communities (see biodiversity estimations and DCA in Fig. 3).First arboreal plants (i.e.dwarf shrubs) could establish contemporaneously as indicated by a Salix herbacea leaf and Betula nana plant remains (see Table 1).
North of the Alps, a first important vegetation shift after the establishment of the steppe tundra occurred at ca. 345 16,000 cal BP with the expansion of shrub tundra around Moossee (see pollen of tree Betula, Betula nana, Juniperus and Salix, macrofossils of Betula nana).A similar vegetation shift has been recorded elsewhere on the Swiss Plateau (Rey et al., 2017), pointing to a regional establishment of shrub tundra with possibly even some first tree birch stands.Plant macrofossil data from Soppensee (Lotter, 1999, see Fig. 4), suggest a coeval establishment of dwarf birch thickets on the Swiss Plateau, while several records of Welten (1982) point to 350 increasing Betula abundances.Indeed, the chironomid-inferred July temperature estimates from Lago di Origlio https://doi.org/10.5194/cp-2019-121Preprint.Discussion started: 7 October 2019 c Author(s) 2019.CC BY 4.0 License.(Samartin et al., 2012, corrected to the altitude of Moossee 521 m a.s.l.assuming a constant modern temperature rate of 6 °C km -1 , Livingstone and Lotter, 1998), indicate a July air temperature warming of 2-2.5 °C reaching temperatures of 10-11.5 °C that are suitable for tree growth (Lang 1994) after 16,000 cal BP, even after considering a latitudinal temperature gradient (see Figs. 4 and 5, Samartin et al., 2012;Finsinger et al., 2019).355 Other factors than summer temperatures such as cold air extrusions from the still existing Scandinavian ice sheet in the North and a stronger latitudinal temperature gradient (Heiri et al. 2014) may have prevented the establishment of dense Betula forests north of the Alps (Rey et al., 2013).Indeed, south of the Alps, in more sheltered positions, wide-spread afforestation with Pinus cembra and Larix decidua started at around 16,500-16,000 cal BP (e.g.Tinner et al., 1999;Hofstetter et al., 2007;Vescovi et al. 2007;Pini et al., 2016, see Fig. 5).360 South of the Alps, sites above 1000 m a.s.l.became ice-free but remained unforested (Vescovi et al., 2007;Pini, 2002), with treeline positions around 800 m asl., Vescovi et al., 2007).
Juniper shrublands expanded massively at Moossee after ca.14,600 cal BP (onset of the Bølling) which is in agreement with the well-dated pollen record at Gerzensee (van Raden et al., 2013).Only ca.200 years later, Betula forests took over and completed the initial afforestation which is widely recorded in pollen assemblages across the Swiss Plateau (e.g.Ammann, 1989;Rey et al., 2017) and unambiguously confirmed by Betula 375 pubescens and Betula pendula macroremains at Soppensee (Lotter, 1999, see Fig. 4).Contemporaneously, sites up to 1800 m a.s.l. in the Alps turned ice-free (Welten, 1982;Ilyashuk, et al., 2009).South of the Alps, dense boreal forests with Pinus sylvestris and Betula established in the lowlands (Vescovi et al., 2007) and tree line reached at least 1850 m a.s.l.(Tinner andVescovi, 2007, Marta et al., 2013).This rapid deglaciation in the Alps and the forest expansion in Southern and Central Europe was caused by a sudden ca. 4 °C warming, as indicated 380 by chironomid and stable isotope records (von Grafenstein et al., 1998(von Grafenstein et al., , 1999;;North Greenland Ice Core Project members, 2004;Heiri and Millet, 2005;Larocque and Finsinger, 2008;Fleitmann et al., 2009, see Fig. 5).
Summer warming (Heiri et al., 2015, see Fig. 4) may also have triggered the increase of regional forest fires observed at Moossee and elsewhere north of the Alps (Rey et al., 2017).
Chironomid-based summer temperature reconstructions (Samartin et al., 2012;Heiri et al., 2014) suggest only a 395 marginal cooling in southern Europe, which nevertheless affected temperate trees stands (Tinner et al., 1999;Finsinger et al., 2006;Vescovi et al., 2007).This different magnitude of climate cooling may be related to the sheltered location of these lakes in the lee of the Alps, preventing direct influences from the polar ice masses and the North Atlantic during the Younger Dryas (Samartin et al., 2012;Heiri et al., 2014).
The dominance of mixed oak-linden-elm-maple forests over millennia in Central Europe (see e.g.Hadorn, 415 1992;Lotter, 1999;Litt et al., 2009;Rey et al., 2017) was most likely favored by continental climate as indicated by maximum summer and minimum winter insolation (Berger and Loutre, 1991), 1.5-2 °C warmer summers than today (e.g.Heiri et al., 2015) and generally drier conditions as reflected by lower lake levels (e.g.Magny et al., 2012, see Fig. 5).This forest type persisted until ca.8500-8000 cal BP (Figs. 3 and 5), when Abies alba and Fagus sylvatica tree stands established around Moossee.Both tree species are shade-tolerant and competitive 420 under mesophytic conditions (Tinner and Lotter, 2006;Tinner et al., 2013;Lauber et al., 2014).Similarly, Alnus and Fraxinus excelsior, both well-adapted to wet soils and moist conditions (Lauber et al., 2014;Rey et al., 2017), expanded as well.The establishment and massive spread of mesophilous mixed beech forests after 7500 cal BP (see high number of Fagus sylvatica bud scales in Table 1) is well studied on the Swiss Plateau (e.g.Lotter, 1999;Wehrli et al., 2007) and the causes for this change have been intensely discussed in the past (Tinner 425 and Lotter, 2006).Decreasing summer temperatures (Heiri et al., 2015;Finsinger et al., 2019) and increasing moisture availability (e.g.Magny et al., 2011;Magny et al., 2012;Joannin et al. 2013, see Fig. 5) suggest climate as the main trigger of this drastic change in Central European forest composition.

Vegetation and land use history during the Mid and Late Holocene
The onset of land use and agricultural activities around Moossee is documented as early as 7000 cal BP (5050 430 cal BC) by first cultural indicator pollen such as Cerealia-type and Plantago lanceolata (see Figs. 3 and 5).intensified agricultural activities caused a first dieback of the mixed beech forests.Our interpretation is in good agreement with coeval on-site archaeological evidence (e.g.log boat made of Tilia wood; Hafner et al., 2012;435 Harb, 2017).Many lowland sites south and north of the Alps indicate a contemporaneous opening of the forests.
The strong link with increasing fire activities suggests that farmers used fire to gain arable and pastoral land (i.e.slash-and-burn, e.g.Tinner et al., 1999;Kleinmann et al., 2015;Rey et al., 2017Rey et al., , 2019a)).Disruption and land use phases generated typical successional cycles starting with arboreal pioneers (Corylus avellana, tree Betula and Alnus) rapidly spreading after disturbance.These light-demanding pioneers were regularly replaced by Fraxinus 440 excelsior and Ulmus, which were in turn replaced by late-successional Fagus sylvatica and Abies alba (e.g.Kleinmann et al., 2015).Most strikingly, a recent high precision and high-resolution study was able to numerically demonstrate that land use phases and the subsequent forest successions were regionally to supraregionally synchronous (Rey et al., 2019a).The most reasonable explanation for such a striking pattern is climate and its influences on human activities.Indeed, Rey et al. (2019a) found that land use phases generally coincided 445 with warm and dry periods as indicated by lower lake levels in Western Switzerland and Eastern France (Magny, 2013) as well as higher solar irradiance (Steinhilber et al., 2009).This finding underlines that climate may have governed harvest success and through that prehistoric human population densities, an environmental effect that on the long-term was counteracted by stepwise introductions of technological innovations (e.g.metal tools, new crops; Tinner et al., 2003, Wirtz andLemmen, 2003).450 Our palaeobotanical data indicate a steady intensification of agricultural activities during the Late Holocene (from ca.3850 cal BP (1900 cal BC) onward), which is primarily evidenced in regions on the Swiss Plateau that are climatically favorable for crop production (i.e.< 550 m a.s.l., Ammann, 1989;Hadorn, 1992;Rey et al., 2017).Many tree species were strongly affected by fire disturbance, browsing and/ or overexploitation and some of them (e.g.Tilia, Taxus baccata) even collapsed completely (Rey et al., 2017).Contrarily, several taxa such as 455 Quercus and Fagus were promoted as fruit trees (Gobet et al., 2000;Wick, 2015), while others were introduced (e.g.Juglans regia, Castanea sativa) for the same reason (Tinner et al., 2003;Conedera et al., 2004).Total biodiversity increased (Fig. 3), thanks to the creation of large open habitats (Colombaroli and Tinner, 2013).As a consequence, during the Late Holocene (i.e. the past 5000 years) humans gradually replaced climate as driving factor of vegetation structure and composition, creating artificial communities (e.g.monospecific Picea, Quercus 460 or Fagus forests, hedges, pastures, fields) that are best suited for land use.However, we can show that in regard to the dominant species, beech forests were able to recover even after most intense anthropogenic disturbances (e.g.Iron Age, Roman Period, Middle Ages).Therefore, Central European beech may presumably prevail in the future if the amplitude of the anticipated climate warming remains within the Mid Holocene variability range (ca.+ 2 °C compared to the 20 th century, Heiri et al., 2015;Finsinger et al., 2019).However, if intense forestry 465 should decline, e.g. as a consequence of nature protection measures, more diverse forests may reestablish.If climate should become > 2° C warmer, possibly causing a reduction of moisture availability (Kovats et al, 2014;Henne et al. 2018), drought-sensitive beech may rapidly decline, giving way to unprecedented forest communities that will likely include drought resistant deciduous species such as Quercus pubescens, droughtresistant evergreen broadleaved species such as Quercus ilex and warm-temperate conifers such as Abies alba 470 (Bugmann et al., 2015;Henne et al., 2015;Henne et al., 2018).

Conclusions
We present a novel highly-resolved vegetation and fire history record from Central Europe that covers the entire post-LGM period.Radiocarbon dating on terrestrial plant remains (i.e.Salix herbacea leaf) resulted in a calibrated age of ca.19,200 cal BP (18,520 cal BP,95 % (2σ) probabilities) for the bottom of the 475 sediment sequence.To our knowledge, together with the novel radiocarbon date from the bottom sediments of Lago di Monate of 19,300 cal BP (18,930 cal BP,95 % (2σ) probabilities, see Table 2), this date provides the oldest age coming from peri-alpine lakes that were created by deglaciation after the ice collapse at the end of the LGM.Deglaciation was followed by the rapid establishment of pioneer steppe vegetation.After HE-1 (end ca. 16,700 cal BP, Stanford et al., 2011) shrubs (Betula nana, Betula humilis, Juniperus, Salix) and 480 possibly trees (likely Betula pubescens, Betula pendula) expanded, which is comparable to recent ecosystem changes in the Arctic in response to ongoing global warming (Pearson et al., 2013, Brugger et al., 2019).Starting points of important vegetation reorganizations at 16,000, 14,600, 11,600 and 8200 cal BP were strongly linked to climate change (temperature and/ or precipitation shifts).No apparent inertia nor lags of population establishments were detected, implying a very high sensitivity and adjustment capacity of plant communities to 485 climatic and environmental changes at decadal scales.These rapid responses without no apparent lags (due to e.g.migration) are explained by the very efficient distribution mechanisms of plants (e.g.winged fruits or bird transport of acorns, see Firbas, 1949, Tinner andLotter, 2006) and the proximity to the refugia (< 400 km) of temperate and boreal species (Kaltenrieder et al., 2009, Samartin et al., 2012, Gubler et al., 2018) https://doi.org/10.5194/cp-2019-121Preprint.Discussion started: 7 October 2019 c Author(s) 2019.CC BY 4.0 License.
Saxifraga aizoides-type are dominant at the end of the LGM (climatic break 1, ca. 19,200-18,800 cal BP, LPAZ https://doi.org/10.5194/cp-2019-121Preprint.Discussion started: 7 October 2019 c Author(s) 2019.CC BY 4.0 License.Moos 1a, Finsinger et al. 2019).This indicates the establishment of open steppe/ tundra vegetation in the region https://doi.org/10.5194/cp-2019-121Preprint.Discussion started: 7 October 2019 c Author(s) 2019.CC BY 4.0 License.BP, followed by an increase of herb pollen (≤ 20 %, see Poaceae, Artemisia) and Juniperus pollen (2-3 %), points to a transformation of closed mixed boreal forests into more open, pine-dominated parklands.We interpret 235 this change as a consequence of climate cooling during the Younger Dryas.Sporormiella fungal spores are only occurring sporadically which might indicate a low grazing activity in the close vicinity of the lake.Charcoal influx values increase temporarily (> 1000 particles cm -2 yr -1 ) at the onset of mixed birch-pine forest formation and remain at elevated but stable values (> 100 particles cm -2 yr -1 ), suggesting slightly enhanced fire activity after 13,400 cal BP. 240 The reestablishment of closed mixed birch-pine forests occurred shortly after the onset of the Holocene (climate break 4, ca.11,700 cal BP, Finsinger et al., 2019) as indicated by the increase of arboreal pollen (> 80 %).Continuous curves of Quercus, Alnus glutinosa-type, Corylus avellana and Ulmus pollen suggest the presence of first temperate forest stands likely in response to the Holocene climate warming already at ca. 11,700-11,500 cal BP.However, temperate trees and shrubs (i.e.Quercus, Corylus avellana, Ulmus, Tilia, Acer and Hedera helix) 245 expanded only after ca.11,100 cal BP (LPAZ Moos-5, 11,050-8600 cal BP) as indicated by the pollen percentages, replacing the boreal forests within ca.200-400 years (decrease of tree Betula and Pinus sylvestristype pollen).This continental open forest and shrub vegetation was advantaged by the continental climate with hot and dry summers of the Early Holocene 7000 cal BP (5050 cal BC, first pollen grains of cultural indicators such as Cerealia-type and Plantago lanceolata).The pollen stratigraphy indicates a stepwise intensification of land use over the millennia with NAP (including cultural indicator pollen) peaking at 5600 cal BP (3650 cal BC, LPAZ Moos-10, Neolithic), 3850 cal BP (1900 cal BC, LPAZ Moos-16, Early Bronze Age), 3500 cal BP (1550 cal BC, LPAZ Moos-18, Middle Bronze Age), 270 2600 cal BP (650 cal BC, LPAZ Moos 20b, Iron Age), 1800 cal BP (cal AD 150, LPAZ Moos 20b, Roman Period), 700 cal BP (cal AD 1250, LPAZ Moos 21, Middle Ages) as well as 200 cal BP (cal AD 1750, LPAZ Moos 21, Modern times).Each of these land use phases were generally accompanied by a decrease of tree pollen percentages of late successional Fagus sylvatica and Abies alba and by the expansion of light-demanding https://doi.org/10.5194/cp-2019-121Preprint.Discussion started: 7 October 2019 c Author(s) 2019.CC BY 4.0 License.pioneers such as Betula trees and Corylus avellana shrubs.Mixed beech forests were able to recover after 275 disturbances as suggested by the cyclical shape of the Fagus sylvatica curve.However, less resilient trees such as Tilia, Ulmus, Taxus baccata and the liana Hedera helix could not cope with the repeated forest disruptions mainly through logging, browsing, pollarding and massively increased fire disturbance (see Rey et al., 2019a for more details) and were strongly reduced or even disappeared after 4500 to 3500 cal BP (2550-1550 cal BC).Most striking are the massive forest openings during the Iron Age/ Roman Period (> 30 % NAP, LPAZ Moos 280 20b, 2600-1600 cal BP (650 cal BCcal AD 350)) and from the Middle Ages onward (> 60 % NAP, LPAZ Moos 21, 1050 cal BP (cal AD 900)today) which we interpret as the influence of large settlements or urban centers within close proximity (< 8 km) of the lake.The related strong increases of Quercus pollen percentages Fig. 5) initiated a recovery of cold steppe-tundra vegetation (e.g.Artemisia, https://doi.org/10.5194/cp-2019-121Preprint.Discussion started: 7 October 2019 c Author(s) 2019.CC BY 4.0 License.Chenopodiaceae).At the same time, late successional pine trees increasingly outcompeted pioneer birch trees, Increases of microscopic charcoal, which falls into a phase with fairly closed mixed beech-oak forests (Figs. 3 https://doi.org/10.5194/cp-2019-121Preprint.Discussion started: 7 October 2019 c Author(s) 2019.CC BY 4.0 License.and 5), suggest a drastic increase of fire activity at the onset of the farming.After 6500 cal BP (4550 cal BC) https://doi.org/10.5194/cp-2019-121Preprint.Discussion started: 7 October 2019 c Author(s) 2019.CC BY 4.0 License.
. The onset of varved sediments (7000 cal BP) was closely related to vegetation opening for land use.Land use gradually 490 overrode climate as the dominant factor in determining vegetation composition and structure during the Late Holocene.Present-day beech forests have been shaped by anthropogenic disturbances over millennia and were resilient to Mid and Late Holocene climate change.However, recent climate warming may exceed the Mid and Late Holocene climate variability releasing sudden collapses and unprecedented reorganizations of Central European ecosystems.495

Figure 3 .
Figure 3. Moossee sediment sequence.Presented are the dates (small black lines), the lithology, percentages of selected 945