Near-surface permafrost aggradation in Northern Hemisphere peatlands shows regional and global trends during the past 6000 years

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1 752858HOL / The HoloceneTreat and Jones research-article2018 Research paper Near-surface permafrost aggradation in Northern Hemisphere peatlands shows regional and global trends during the past 6000 years The Holocene 2018, Vol. 28(6) The Author(s) 2018 Reprints and permissions: sagepub.co.uk/journalspermissions.nav DOI: journals.sagepub.com/home/hol Claire C Treat 1,2 and Miriam C Jones 3 Abstract The history of permafrost aggradation and thaw in northern peatlands can serve as an indicator of regional climatic history in regions where records are sparse. We infer regional trends in the timing of permafrost aggradation and thaw in North American and Eurasian peatland ecosystems based on plant macrofossils and peat properties using existing peat core records from more than 250 cores. Results indicate that permafrost was continuously present in peatlands during the last 6000 years in some present-day continuous permafrost zones and formed after 6000 BP in peatlands in the isolated to discontinuous permafrost regions. Rates of permafrost aggradation in peatlands generally increased after 3000 BP and were greatest between 750 and 0 BP, corresponding with neoglacial cooling and the Little Ice Age (LIA), respectively. Peak periods of permafrost thaw occurred after 250 BP, when permafrost aggradation in peatlands reached its maximum extent and as temperatures began warming after the LIA, suggesting that permafrost thaw is likely to continue in the future. The broader correlation of permafrost aggradation in peatlands with known climatic trends and other proxies such as pollen records suggests that this record can be a valuable addition to regional climate reconstructions. Keywords bog, boreal, fen, macrofossils, peatlands, permafrost, synthesis, taiga, tundra Received 7 September 2017; revised manuscript accepted 9 December 2017 Introduction Widespread permafrost aggradation and degradation in highlatitude peatlands highlight the importance of long-term permafrost dynamics in northern peatlands. The recent permafrost degradation and thaw that has been observed in peatlands in Alaska (Jones et al., 2016a; Jorgenson et al., 2006; Wickland et al., 2006), Canada (Payette et al., 2004; Sannel and Brown, 2010; Vitt et al., 2000), and Scandinavia (Hodgkins et al., 2014; Johansson et al., 2006) is predicted to continue because of climatic warming (Lawrence et al., 2012). The timing of near-surface permafrost aggradation and degradation within the northern high-latitude permafrost zone, where peat formation initiated often millennia after the Last Glacial Maximum (LGM) (MacDonald et al., 2006), remains unevaluated at a broad, regional scale. Similarly, whether widespread permafrost thaw is a new phenomenon in peatlands or has occurred previously during the Holocene in response to warmer temperatures remains largely unknown. Hence, an important gap remains in understanding the timing and rates of permafrost aggradation and degradation in peatlands across northern high latitudes, which has important implications for carbon (C) storage in permafrost peatlands as climate warms and permafrost thaws (Jones et al., 2016b; O Donnell et al., 2012; Schneider von Deimling et al., 2012). Permafrost is ground that remains frozen for more than two consecutive years and occurs both where mean annual air temperatures (MAAT) are less than 0 C, although it remains in some regions where the present-day MAAT are greater than 0 C and the permafrost is insulated by vegetation or peat (Halsey et al., 1995; Shur and Jorgenson, 2007). Climate plays a role in permafrost aggradation because air temperatures must be cold enough to result in perennially frozen ground (Shur and Jorgenson, 2007). However, MAAT alone are not enough to determine whether permafrost will aggrade or thaw. Permafrost aggradation and thaw have been linked to a range of local and regional factors including the colonization of peatland surfaces by Sphagnum species (spp.), microtopography, tree and shrub cover, snow thickness, snow distribution, and disturbance (Allard and Seguin, 1987; Camill, 2000, 2005; Camill and Clark, 1998; Johansson et al., 2013; Payette et al., 2004; Zoltai and Tarnocai, 1975). These local factors cause changes in the soil thermal regime and can result in decreased thermal conductivity during the summer or increased exposure to cold winter temperatures (Halsey et al., 1995; Oksanen et al., 2003; Seppälä, 1994, 2011; Zoltai, 1993, 1 University of Alaska Fairbanks, USA 2 Department of Environmental and Biological Sciences, University of Eastern Finland, Finland 3 U.S. Geological Survey, USA Corresponding author: Claire C Treat, Department of Environmental and Biological Sciences, University of Eastern Finland, P.O. Box 1627, Kuopio, Finland. Claire.treat@uef.fi

2 Treat and Jones ; Zoltai and Tarnocai, 1975), which can ultimately result in permafrost formation given sufficiently cold temperatures. Permafrost thaw in peatlands can be associated with climate warming (Halsey et al., 1995) and local factors such as hydrologic changes, disturbance (including wildfire), and increased snow cover (Camill, 2005; Johansson et al., 2006; Payette et al., 2004; Zoltai, 1993). The combination of site factors and climate factors is key to determining both the history of permafrost dynamics and future responses to climate change with warming temperatures (e.g. Seppälä, 2011). Northern peatlands have experienced both warmer and cooler temperatures during the Holocene. Whether permafrost aggradation or degradation occurred broadly in peatlands as a result of these differing climatic conditions is unknown. Temperatures were warmer than present during the time-transgressive Holocene Thermal Maximum (HTM) (Kaufman et al., 2004). Peatland initiation, accumulation, and expansion rose sharply on the landscape during the HTM (Jones and Yu, 2010; MacDonald et al., 2006; Smith et al., 2004). In nearly all regions, the HTM preceded the period of study; the latest occurrence of the HTM in North America was in the middle Holocene in Eastern Canada (Kaufman et al., 2004). Whether permafrost aggradation and/or permafrost thaw in peatlands occurred under the warmer conditions during the HTM is unknown but is highly relevant given future climatic warming. Subsequent neoglacial cooling may have set the stage for permafrost aggradation in peatlands. Decreasing summer insolation across the northern high latitudes contributed to cooler-than-present temperatures following the HTM (Berger and Loutre, 1991), in particular during neoglaciation (Sharp, 1960) and the Little Ice Age (LIA) (Alley, 2000; Marcott et al., 2013). Colder temperatures resulted in permafrost aggradation during the Holocene across northern high latitudes (Mann et al., 2010, 2002). For example, detailed plant macrofossil analyses from peat cores in Canada show permafrost aggradation occurred in numerous peatland sites coinciding with cooling at the end of the HTM around 4000 BP as evidenced by vegetation changes from wet fen vegetation to drier, forested bog species (Zoltai, 1995). The timing of near-surface permafrost aggradation in peatlands (hereafter permafrost aggradation ) is an important controlling factor for determining the potential for C loss with permafrost thaw (Jones et al., 2016b; Treat et al., 2014). If permafrost aggradation results in the incorporation of relatively undecomposed material into permafrost (i.e. syngenetic permafrost), then the potential C losses found in soil incubations are similar to C losses from surface soils, whereas thawing of highly decomposed material in peat deposits with epigenetic permafrost results in small C losses (Lee et al., 2012; Treat et al., 2014). Given that in the present-day permafrost zone, permafrost peatlands (histels) comprise ~19% of the area but contain 40% of the soil organic C (top 3 m of soil; Hugelius et al., 2014; Tarnocai et al., 2009), understanding how permafrost aggradation and thaw impacts biogeochemical cycling of C is important for understanding feedbacks to warming. The timing of permafrost aggradation and thaw in peatlands can be inferred from a combination of detailed plant macrofossil analysis, physicochemical peat properties, and detailed chronologies (Camill et al., 2009; Oksanen et al., 2003; Treat et al., 2016). Here, we use a dataset of peat properties and peatland vegetation community reconstructions to identify the timing of peatland initiation, permafrost aggradation, and permafrost thaw in 266 cores from across the northern hemisphere for the last 6000 years (Treat et al., 2016). Specifically, we ask whether there are coherent, regional trends in the timing of permafrost aggradation and thaw in northern hemisphere peatlands and how they relate to paleoclimate. Methods Dataset development: Ecosystem classification, age depth models, and synthesis Permafrost aggradation can result in changes in vegetation that can be preserved in organic soil horizons, which generally contain both plant macrofossil records and means to obtain good chronologic constraints (Zoltai and Tarnocai, 1971, 1975). Therefore, this study focused on the aggradation of permafrost in peatlands. We compiled records of plant macrofossils, radiocarbon dates, lithologies, and peat properties from cores from 441 peatland cores within the boreal and tundra regions of North America, Europe, and Asia using methods described in detail in a previous study (Treat et al., 2016). We selected cores that were well-described using the criteria below. We were interested in the development of permafrost during the Holocene, and given the climatic variability during the Holocene, we selected peat cores from the regions that contained permafrost during the LGM (Vandenberghe et al., 2014). While there were numerous records and even permafrost aggradation prior to 6000 BP (Figure 1; Table S1, available online), we focused on the period after 6000 BP in this analysis in order to have a higher data density to conduct an analysis of regional trends in permafrost aggradation. We selected cores from the larger dataset that met the following criteria: (1) organic soils >30 cm thick that contained a minimum of 65% organic matter); (2) available plant macrofossil assemblages to classify the peatland environmental type, including the presence/absence of permafrost; (3) chronologic control of one or more dates for every 2000 years; (4) location within the present-day permafrost zone (Brown et al., 2001 [1998]) or within the zone of permafrost at the Last Permafrost Maximum (LPM), ka BP (Figure 1) (Lindgren et al., 2015; Vandenberghe et al., 2014). This resulted in the inclusion of 266 cores from 214 sites across the northern hemisphere (Figure 1; Table S1, available online). These data are summarized in Table S1, available online; the complete dataset including site information, peat properties, plant macrofossils, and chronologic information is available through PANGAEA ( PANGAEA ). Sections of peat cores were classified into wetland classes including permafrost-free fens, permafrost-free bogs, permafrost peatlands (including peat plateaus, palsas, bogs and fens with permafrost, polygonal peat plateaus, high- and low-center polygons, and tundra with >30 cm of organic soils in present day), thawed permafrost (including collapse-scar fens, bogs, and thaw ponds), and other (peatland pools, marshes, swamps, and ponds, lakes, and upland forests that later develop peatlands) using the classification scheme described by Treat et al. (2016) and in more detail below regarding the delineation of permafrost dynamics. Briefly, plant macrofossil assemblages, detailed descriptions of lithology, and peat properties were used to classify peat core sections into wetland classes based on the Canadian Wetland Classification system (National Wetlands Working Group, 1988; Treat et al., 2016). This approach relied heavily on the original authors interpretation of the plant macrofossil data because there is no single indicator species of permafrost formation (see discussion below; Oksanen and Väliranta, 2006). In this study, we derived new age depth models for each core based on the reported chronology using BACON (Blaauw and Christen, 2011) and IntCal13 (Reimer et al., 2013). We assumed that peat core surfaces were from the year of sampling, unless specified otherwise in the original dataset. The wetland classification of each peat core was converted from the depth-scale to a time scale using the age depth model for each core and subsequently binned into 250-year age bins for analysis and plotted

3 1000 The Holocene 28(6) Figure 1. The estimated timing of permafrost aggradation in peatland sites located within the present-day permafrost zone (hatched lines) and the Last Permafrost Maximum zone (beige shading). The noncontinuous permafrost zone includes zones of discontinuous, isolated, and sporadic permafrost. Permafrost aggradation: Colors indicate the timing of permafrost aggradation, with warmer colors indicating older permafrost and cooler colors indicating more recent permafrost formation. Peatlands without evidence of permafrost in the past or present are classified as none. Regions correspond with regions discussed in the text. The extent of present-day peatlands is shown in yellow (Yu et al., 2010). using bin midpoints. We use calibrated 14 C ages throughout the text and abbreviate cal yr BP as simply BP. Determination of potential permafrost aggradation The presence of permafrost in a peatland site at the time of sampling considerably simplifies the identification of permafrost aggradation. In boreal regions, surface uplift from ice expansion within the permafrost results in a vegetation shift to species indicative of dry conditions (Seppälä, 2011; Zoltai and Tarnocai, 1975; Zoltai et al., 1988). Plant macrofossil analysis can be used to identify the vegetation shifts to dry, forested communities associated with permafrost aggradation from relatively wet vegetation communities associated with fens and bogs (Zoltai and Tarnocai, 1975). While this is generally evident in peat cores from a transition from wet, sedge-dominated peat to forest peat or Sphagnumwoody peat (Camill et al., 2009; Jones et al., 2013; Kuhry, 2008; Zoltai and Tarnocai, 1975), more detailed analysis is required because of the similarities between dry bogs and permafrost peatland species (Camill et al., 2009; Jones et al., 2013; Oksanen, 2006). Moss species associated with permafrost aggradation include Polytrichum spp., Pleurozium spp., Tomenthypnum nitens, Dicranum elongatum, and some hummock-forming Sphagnum mosses, generally Sphagnum sect. Acutifolia (Camill et al., 2009; Jones et al., 2013; Kuhry, 2008; Oksanen, 2006; Sannel and Kuhry, 2009; Zoltai, 1993). In near-surface peat, an increase in lichen abundance (Cladina spp. and Cladonia spp.) and fungal sclerotia is also commonly associated with permafrost aggradation (Camill et al., 2009; Kuhry, 2008; Oksanen, 2006; Zoltai and Tarnocai, 1975; Zoltai et al., 1988). The alternation of Sphagnum fuscum peat and rootlet layers has also been used to identify permafrost aggradation and persistence in both Western Canada and European Russia (Oksanen et al., 2003; Sannel and Kuhry, 2008, 2009). In continuous permafrost zones and tundra sites, indicators of permafrost aggradation take somewhat different forms. In many places, permafrost may have formed during or prior to the LGM (e.g. Kanevskiy et al., 2014; Vandenberghe et al., 2014) and may be present in the mineral soil underlying at the time of peat inception (Zoltai and Tarnocai, 1975). In this case, a mixing of peat and mineral soils indicative of cryoturbation at the peat-mineral soil interface has been used to indicate permafrost presence at the time of peat formation (Zoltai and Tarnocai, 1975). The presence of permafrost in the years to centuries following peat inception was unlikely for peats formed in drained thaw lake basins, a relatively common occurrence in Arctic regions of Alaska and Siberia (Bockheim et al., 2004; De Klerk et al., 2011; Jones et al., 2012; Walter Anthony et al., 2014). In these cases, the timing of permafrost aggradation has been identified using species composition shifts toward drier conditions, including from Carex-dominated tundra fen vegetation to dry tundra vegetation (Jones et al., 2012) or from the development of dry microforms associated with patterned ground permafrost features (Davis, 2001). In low-center polygons formed in drained thaw lake basins, permafrost aggradation can be identified by the development of patterned ground features such as polygon rims or ridges. Some species associated with polygon rims or ridges include mosses such as Tomenthypnum nitens, Hylocomnium splendens, Sphagnum cf. subsecundum, Sphagnum teres, Sphagnum warnstorfia, and shrubs including Salix spp. (De Klerk et al., 2011). The polygon ridges are easily distinguishable from the wetter, Carex- and Eriophorumdominated polygon low centers with mosses such as Calliergon giganteum and Drepanocladus revolvens (De Klerk et al., 2011; Tarnocai and Zoltai, 1988; Zoltai and Tarnocai, 1975), which also occur in wet, permafrost-free tundra fen sites. The delineation of the timing of permafrost aggradation and thaw also becomes more difficult with a more complex site history, including instances of partial or complete thaw resulting in

4 Treat and Jones 1001 the absence of permafrost in the present day. In the discontinuous permafrost region of North America, specifically, in Alaska and Western Canada, the clearest indicator of past permafrost is a sequence indicative of permafrost aggradation followed by thaw, similar to the transitions between species observed on the margins of and lawns of collapse-scar features in the present day (Jones et al., 2013). In these records, species in the macrofossil record frequently indicated dry, treed peat plateaus with Picea mariana, lichens, and fungal sclerotia, followed by a distinct, abrupt transition to wetter conditions with species such as Sphagnum riparium, Carex spp., and Eriophorum spp., and the decrease/ disappearance of Picea mariana, all indicating permafrost thaw (e.g. Jones et al., 2013, 2016b; Kuhry, 2008; Oksanen et al., 2003; Sannel and Kuhry, 2008; Zoltai, 1993). Permafrost aggradation can also result in changes in peat properties. Because of the drier conditions associated with surface uplift because of ice expansion, the degree of peat decomposition shifts to a more or highly humified peat (e.g. Camill et al., 2009; Kuhry, 2008; Oksanen et al., 2003; Sannel and Kuhry, 2008, 2009) and apparent C accumulation rates generally slow (Jones et al., 2012; Oksanen, 2006; Sannel and Kuhry, 2009; Treat et al., 2016; Zoltai et al., 1988). Using a large synthesis dataset from the permafrost region, Treat et al. (2016) found higher C/N ratios in peats that were likely deposited after permafrost aggradation, suggesting that C/N ratios could provide an additional evidence for the timing of permafrost aggradation. More advanced chemical analysis has also been used to characterize the timing of permafrost aggradation with varying degrees of success, but these newly developed approaches have not yet been applied widely (Ronkainen et al., 2015; Routh et al., 2014). It is difficult, if not impossible, to determine the exact rather than the potential timing of permafrost aggradation. Using detailed vegetation surveys in adjacent areas with and without permafrost, previous studies have shown that there is no single indicator species of permafrost formation (Oksanen and Väliranta, 2006). For example, in the boreal region, the species composition of boreal dry bogs and permafrost peat plateaus is similar. Both are generally distinguished by the transition to drier species, often making a determination of the timing of permafrost aggradation imprecise (Camill et al., 2009). Slow peat accumulation rates or erosion of surface peat because of windscour or other factors can remove important parts of the peat record (Peteet et al., 1998; Ronkainen et al., 2015), introducing difficulties for determining the timing of permafrost aggradation or other environmental changes precisely. Still, a careful multi-proxy analysis of plant macrofossils, peat accumulation rates, degree of decomposition, and peat chemistry (including carbon, nitrogen, and hydrogen content) offers the best chance at identifying this important aspect of ecosystem history. Regional trends We aggregated the peat cores regionally in order to analyze trends in permafrost aggradation and thaw. The regional analysis was based on administrative unit boundaries and geographic positions where the cores were located (Figure 1). Regions in North America included Eastern North America, Central North America (Eastern Canadian Rockies, Hudson Bay Lowlands, the Great Lakes, and Ontario), Alaska and British Columbia, and the Canadian Arctic (Yukon Territory, Northwest Territories, and Nunavut). Regions in Eurasia included the Fennoscandia (including the Kola Peninsula in Russia), European Russia, West and Central Siberia (including the Lake Baikal Region, n = 1), and Eastern and Far Eastern Siberia. The number of cores in each wetland classification were summed within each region for each time bin in order to understand regional trends in fen-to-bog transitions (an increase in the number of bog cores and a decrease in fen cores), permafrost aggradation (an increase in the number of cores with permafrost), and permafrost thaw (a decrease in the number of cores with permafrost). When discussing the regional trends in rates of permafrost aggradation and thaw, we use the normalized rate of transition. The normalized rate of transition (dna/dt) for a given peatland type a is calculated using Equation 1, where N is the number of peatland cores. dna Nat Nat = dt N 1 0 This approach reduces bias introduced by varying data density (number of cores) within each region, and specifically looks at changes occurring relative to the existing peat cores. The mean normalized rates and standard errors of rates from Equation 1 are calculated for the time periods of interest. Both approaches assume that an increase in the number of cores or percentage of cores correlates with an increasing peatland area and that cores in this study are representative of peatland dynamics as a whole. However, without accurate maps and areal estimates of each type of peatland area regionally, we cannot evaluate whether the number of samples from each peatland type is truly representative. For comparison between the records of permafrost aggradation and thaw in peatlands to regional and hemispheric climatic trends (e.g. Figure 2), we used existing paleoclimatology records. We used the modeled summer solar insolation at 65 N (Berger and Loutre, 1991) because of the correlation between radiation and peatland C accumulation at shorter time scales (Charman et al., 2013). We used percent melt data from the Agassiz Ice Cap in the Canadian Arctic, which is indicative of summer temperatures routinely increasing to above freezing (Fisher et al., 1995) and correlated well with sea ice records from this region (Vare et al., 2009). While it is likely that regional temperature trends diverged from the local trends in the Canadian Arctic, given its high-latitude location, continuous nature, proximity to sites in North America, and lack of other continuous regional records, it is included in this analysis. Results Northern high latitudes Over the past 6000 years, the number of peatland sites increased by 63% from 164 cores in 6000 BP to 266 cores in 0 BP (Figure 2g). During this time, the percentage of permafrost-free fens decreased from 62% to 23%, despite occurring in ~100 sites until 750 BP (Figure 2c). After 750 BP, the number of permafrost-free fens decreased rapidly by 40%, from 96 cores to 62 cores at 0 BP. The number of permafrost-free bogs increased from 20% of cores (Figure 2d, n = 35) to 30% of cores (n = 79) over this same period. Permafrost aggradation in peatlands increased from less than 10% of peatland cores (n = 15) in 6000 BP to 40% of peatland cores at 0 BP (Figure 2e, n = 111). Overall, rates of permafrost aggradation were greatest between 750 BP and 0 BP, developing in 55 new cores during this period, or at a rate of 2.2% ± 0.4% of cores per century. From 6000 BP to 1000 BP, permafrost aggraded at a rate of 0.4% ± 0.1% of cores per century, although rates were slightly higher between 2750 and 2250 BP. Permafrost thaw was greatest between 250 BP and present, occurring in 10 of 266 cores, or less than 4% of cores. Rates of permafrost thaw were low, averaging 0.07% ± 0.05% of cores per century. The location of sites was important for determining the timing of permafrost aggradation. Permafrost aggraded earliest in sites in the continuous permafrost zone (3500 ± 500 BP, where error is standard error among sites), and on average 1730 ± 470 years t0 (1)

5 1002 The Holocene 28(6) Figure 2. Changes in Northern Hemisphere climate and peatlands for the past 6000 years. (a) Modeled July solar insolation at 65 N (W m -2 ; Berger and Loutre, 1991); (b) melt percentage of ice layers from the Agassiz Ice Cap in the Canadian Arctic, averaged over 50 years, where higher melt percentage indicates warmer summer temperatures (Fisher et al., 1995); (c) the number of cores classified as permafrost-free fens (solid line) and percent of fen cores (dashed line); (d) the number of cores classified as permafrost-free bogs (solid line) and percent of bog cores (dashed line); (e) the number of cores classified as permafrost peatlands, including peat plateaus, palsas, permafrost fens, permafrost bogs, polygonal peat plateaus, high- and low-center polygon peatlands, and tundra with >30 cm of organic soils in present day (solid line) and percent of permafrost cores (dashed line); (f) the number of cores classified as thawed permafrost, including collapse-scar fens, bogs, and thaw ponds (solid line) and percent of permafrost thaw cores (dashed line); and (g) the total number of peat core records in each time step. The ecosystem classification was based on plant macrofossil analysis and physicochemical properties from individual cores. The ages were based on core chronology and BACON-derived age depth models.

6 Treat and Jones 1003 Figure 3. Aggregated regional records of ecosystem transitions, permafrost aggradation, and permafrost thaw in North America. The ecosystem classification is shown by the number of cores in each class for each region: (a) Eastern North America (n = 51); (b) Central North America, including the Eastern Canadian Rockies, Hudson Bay Lowlands, the Great Lakes, and Ontario (n = 60); (c) Canadian Arctic including Yukon Territory, Northwest Territories, and Nunavut (n = 29); (d) Alaska and British Columbia (n = 28). The increasing number of cores indicates new peatland initiation. after peatland inception. In discontinuous permafrost, permafrost aggraded 1500 years later at 2000 ± 250 BP in sites that were older (4320 ± 320 years). Permafrost aggraded more frequently in bog cores than fen cores in the boreal region (42/62 cores) and discontinuous permafrost zones (40/70 cores), but more frequently in fens than bogs in the tundra region (28/33 cores) and continuous permafrost zone (20/28 cores). North America In North America, the number of peatland sites increased during the past 6000 years (Figure 3). The majority of cores in North America were from Eastern and Central North America (more than 50 cores each, Figure 3a and b), whereas Alaska and British Columbia combined had the fewest records (28 sites, Figure 3d) followed by the Canadian Arctic (29 cores, Figure 3c). In Eastern North America, the number of fens peaked around 3125 BP (n = 21 cores) and decreased to a minimum of less than 5% of cores in the present from a maximum of 65% at 6000 BP (Figure 3a). During this period, the percentage of bogs increased from 20% at 6000 BP (n = 6 cores) to 50% in the present (n = 26 cores). The first appearance of permafrost occurred between 5500 and 5250 BP (Payette, 1988) and increased to 43% of cores in the present day (n = 22). Peak periods of permafrost aggradation in Eastern North America occurred between 2250 and 2000 BP and between 1250 and 1000 BP (Figure 3a). In Central North America, the fraction of fens decreased from 50% in 6000 BP to 20% in present day, while the fraction of bogs reached a maximum between 3000 and 1750 BP. Bogs occurred in ~50% of cores before decreasing to 37% of sites at 0 BP (Figure 3b). Permafrost aggradation in Central North America was small prior to 750 BP, when rates increased from 0.2% ± 0.1% of cores per century to 3.3% ± 0.8% of cores per century between 750 and 500 BP (Figure 3b). The occurrence of permafrost in Central North America peaked at 0 BP (n = 25, 42%, Figure 3b). Some permafrost thaw also occurred in Central North America, with a maximum occurring between 1500 BP and 500 BP in two to three cores (Figure 3b) (Kuhry, 2008; Zoltai, 1993). In many sites in both the Canadian Arctic and Northern Alaska, permafrost aggradation occurred in peats prior to 5000 BP (Figures 1, 3c, and d). In the Canadian Arctic, permafrost formed in 30% of cores prior to 5000 BP (n = 6). Most peatland initiation occurred in this region before 5000 BP and from 4000 to 1250 BP. Permafrost aggradation increased between 4000 BP and 3000 BP and again from 1250 BP to 0 BP, resulting in 90% of cores developing permafrost by 0 BP (n = 26, Figure 3c). In Alaska and British Columbia, the fraction of fens and bogs remained relatively constant until 250 BP, when the number of fens decreased and the number of bogs and permafrost cores increased (Figure 3d). Between 750 BP and present, permafrost thaw occurred and reached a maximum of eight cores in present day (2015 CE; Figure 3d). Eurasia Trends in peatland transitions differed between North America and Eurasia and were generally more stable in Eurasia (Figures 3 and 4). In Fennoscandia, fens were the most common peatland type in cores at 6000 BP (73%), but were found in less than 50% of records after 1250 BP as the number of bogs increased (Figure 4b). Permafrost aggradation in Fennoscandian peatlands occurred after 1250 BP in very few sites (Figure 4b, Table S1, available online) (Kokfelt et al., 2010; Van der Knaap, 1989). In continental Europe, the proportion of fens and bogs was relatively stable during the past 6000 years at 57% fen (n = 7) and 36% bog, (n = 6) respectively; no permafrost formed in this region.

7 1004 The Holocene 28(6) Figure 4. Aggregated regional records of ecosystem transitions, permafrost aggradation, and permafrost thaw in Eurasia from 6000 BP to present. The ecosystem classification is shown by the number of cores in each class for each region for (a) Arctic Europe, European Russia, and Russian Coastal Plain (n = 17); (b) Fennoscandia, including Iceland, Norway, Sweden, Finland, Kola Peninsula (Russia), and Spitsbergen (n = 13); (c) Western and Central Siberia, including the Lake Baikal region (n = 25); (d) Eastern and Far Eastern Siberia (n = 24). The increasing number of cores indicates new peatland initiation. Apart from Fennoscandia, European records do not contain permafrost in peatlands. In the Russian and European Arctic, permafrost consistently aggraded in peatlands after 1000 BP and led to a sharp decrease in the number of permafrost-free fens (Figure 4a). The occurrence of permafrost increased from 31% between 1000 and 750 BP (n = 5 cores) to 59% in the present day (n = 10, Figure 4a). Several cores in the Russian and European Arctic experienced permafrost thaw and re-aggradation in the past 6000 years (Table S1, available online; Oksanen et al., 2001, 2003). In West and Central Siberia, fraction of permafrost-free fens decreased slightly from 63% to 44% of cores and the fraction of permafrost-free bogs increased from 11% to 28% between 6000 BP and present. Permafrost-free fens were the dominant peatland type in these records between 6000 BP and 4750 BP and again 3750 BP and 750 BP (Figure 4c). From 4750 BP to 3750 BP and again between 750 BP and present, permafrost-free bogs and permafrost assemblages increased (Figure 4c). Present-day permafrost extent was reached in West and Central Siberia before 500 BP and occurred in 24% of cores (Figure 4c). In East and Far East Siberia, permafrost-free fens were more common than permafrost-free bogs, which were relatively rare in our records (Figure 4d). In Eastern and Far East Siberia, permafrost aggradation accelerated after 1000 BP and peaked after 250 BP, occurring in 58% of cores (n = 14, Figure 4d). Many peatlands formed in thermokarst lake basins in this region, resulting in a relatively high number of permafrost thaw records in these cores (Figure 4d, Table S1, available online; Walter Anthony et al., 2014). Discussion Regional trends in permafrost aggradation One advantage to the reconstructions of permafrost history using peat cores is the broad spatial coverage of this dataset (Figure 1), which allows global and regional comparisons to other climate proxies, as well as information from relatively data-sparse regions, such as much of Siberia (Kaufman et al., 2009; Mann et al., 2009; Marcott et al., 2013). The increased aggradation of permafrost in peatlands over the past 6000 years is in general agreement with other regional proxies of these northern regions showing a general cooling trend (Andreev et al., 2011; Fisher et al., 1995; Kaufman et al., 2004, 2016; MacDonald et al., 2000; Marcott et al., 2013). Permafrost aggradation in peatlands increased since 6000 BP, with the most widespread evidence for permafrost aggradation occurring in the middle to late-holocene after 4000 BP and more strongly after 1000 BP (Figure 2e), which generally followed the decrease in summer solar insolation (Figure 2a) and decreasing melt recorded on the Greenland icesheet (Figure 2b; Fisher et al., 1995). The increased rate of permafrost aggradation across the northern high latitudes after 4000 BP suggests a climatic response to neoglacial cooling. Both decreased Greenland melt layers (Figure 2b) and evidence for increased Arctic sea ice (Vare et al., 2009) corroborate the trend of a cooling climate. Today, more than 40% of the peat cores studied are underlain by permafrost. We observe regionally coherent patterns in permafrost aggradation in both North America and Eurasia, potentially indicating climatic cooling such that MAAT dropped below 0 C during the time periods of high permafrost aggradation. Prior to neoglaciation (>4000 BP). Our data show that less than 10% of the northern high-latitude peatlands studied contained permafrost prior to 6000 BP (14 of 153 cores), owing to warmerthan-present temperatures during the HTM (Fisher et al., 1995; Kaufman et al., 2004). This study shows early (prior to 6000 BP) permafrost aggradation in peatlands of northern Alaska, similar to early permafrost aggradation in Arctic Canada and parts of Siberia (Figure 1, Table S1, available online). At 6000 BP, scattered permafrost existed in peatlands within the continuous permafrost

8 Treat and Jones 1005 zone of Arctic Canada but did not exist in peatlands in other parts of Canada (Zoltai, 1995), indicating warmer temperatures than present day in much of Canada (Kaufman et al., 2004). Peat records in Alaska and British Colombia remain nearly unchanged throughout the past 6000 years (Figure 3d), which suggests that these peatland ecosystems are either resilient to environmental change, or that there has been relatively little environmental change over the last 6000 years that affects temperature or effective moisture, or both. The largest temperature changes in Alaska and Yukon occurred prior to 6000 years ago, when cold deglacial temperatures followed by warmer-than-present temperatures driven by insolation-driven changes in the early Holocene (Kaufman et al., 2004). Summer temperature reconstructions from two Alaskan lakes are stable for the past 6000 years (Kurek et al., 2009). Sea ice extent in the Canadian Arctic showed only small increases from 6000 to 4000 BP and more abrupt increases after 4000 BP (Vare et al., 2009), indicating stable conditions during the former period. A recent study by Kaufman et al. (2016) showed high spatial variability in climate in Alaska and the Yukon territories over the Holocene, likely driven by orographic complexity in the region, masking trends in temperature and moisture beyond the identification of a mid- HTM from 7000 to 5000 BP and neoglacial cooling after 4000 BP. Alternatively, the resiliency and stability of permafrost peatlands in the Alaska/Western Canada region (Figure 3d) could be a reflection of the age of the peatlands and subsequent peat thickness (Table S1, available online). These peatlands largely developed in the early Holocene (Jones and Yu, 2010), several millennia before the other peatland regions in this study (Gorham et al., 2007), and accumulated peat rapidly in their initial stages (Jones et al., 2009, 2014). In several Arctic and taiga regions, the peatland permafrost aggradation followed a previous cycle of thaw of surface permafrost (yedoma) that formed on this unglaciated landscape, forming thermokarst lakes, which later drained to create flat basins ideal for peatland accumulation (Jones et al., 2012; Jorgenson et al., 2013; Kanevskiy et al., 2014). These processes are common drivers of peatland initiation in Western Alaska, including the Seward Peninsula, Koyukuk National Wildlife Refuge, and Innoko National Wildlife Refuge (Jones et al., 2012, 2016b; Kanevskiy et al., 2014; O Donnell et al., 2012) as well as in Eastern Siberia, where peat also accumulated in drained lake basins (De Klerk et al., 2011; Walter Anthony et al., 2014). Therefore, Holocene permafrost aggradation in these regions were not unique, first-time events, but instead part of a more complex evolution of deglacial climate and geomorphology. The formation of peatlands in response to permafrost thaw occurred in several other sites, but the initial cause of permafrost thaw is not described (Table S1, available online). In Siberia, permafrost aggradation occurs prior to 6000 BP in many sites, indicating continued cold temperatures throughout much of the Holocene (Figure 1). Much like Alaska, most of Siberia remained unglaciated and peat initiation was driven by warming and increased moisture availability at the end of the LGM (Alexandrov et al., 2016; MacDonald et al., 2006; Smith et al., 2004). A relatively stable number of permafrost peatlands existed in western and central Siberia, consistent with the stability of the unglaciated Alaska sites (Figures 3d and 4c). These sites also have in common with Alaska an early HTM, followed by a cooler and more stable mid- to late-holocene climate (Kaufman et al., 2004). Permafrost also aggraded prior to 5500 BP in other Arctic sites in eastern European Russia (Hugelius et al., 2012; Ronkainen et al., 2015; Väliranta et al., 2003). Neoglaciation ( BP). The records of permafrost aggradation in peatlands (Figure 2e) show general agreement with the climate trends of neoglacial cooling beginning BP (Alley, 2000; Fisher et al., 1995; Kaufman et al., 2004; Payette and Lavoie, 1994). The widespread presence of permafrost in peatlands in both Eurasia and North America began around 4000 BP and accelerated after 1000 BP (Figures 2 4). Regionally, permafrost in peatlands steadily increased in arctic Canada, Eastern North America, and Arctic and European Russia after ~3000 BP (Figures 3a, c, and 4a). Records of glacial expansions from both Alaska (Barclay et al., 2009) and Europe (Holzhauser et al., 2005) during neoglaciation also show increased prevalence of glacial advance after 3000 BP, with multiple glacier expansion events (3000 BP, 1500 BP, 250 BP; Barclay et al., 2009), roughly coinciding with the timing of increases in permafrost occurrence (Figure 2e). A remarkable increase in sea ice in the Canadian Arctic Archipelago occurs after 3000 BP (Belt et al., 2010; Vare et al., 2009), suggesting that insolation-driven neoglacial cooling increased persistence of sea ice, which would have resulted in cooler and drier conditions on land. This is consistent with the small number of new peat sites (one) in the Canadian Arctic between 3000 BP and 2250 BP (Table S1, available online; Ellis and Rochefort, 2006) but is not reflected in the timing of the aggradation of permafrost, which mainly aggraded either prior to 3000 BP or after 1000 BP (Figure 1). The regional trends in permafrost aggradation show spatial and temporal differences that suggest regional differences in timing of post-htm neoglacial cooling across North America. Permafrost aggradation increased steadily after 3000 BP in eastern North America (Figure 3a), which agrees with other records showing a steady cooling in eastern North America (Allard and Seguin, 1987). The cooling, driven by decreasing summer insolation (Berger and Loutre, 1991), shortened growing seasons and resulted in shifts in vegetation, as evidenced by pollen records (Viau et al., 2006) and tree-line limits retreated in eastern Canada (Payette and Gagnon, 1985). However, in arctic Canada, permafrost occurrence increased earlier, beginning around 4000 BP (Figure 3c; Zoltai, 1995), which coincides with the onset of cooler summer temperatures (Figure 2b; Fisher et al., 1995). However, the most rapid rates of permafrost aggradation occurred later in Central North America than in Eastern North America or arctic Canada, beginning after 1000 BP (Figure 3b). The regional coherence of permafrost aggradation in Central North America (Figure 3b) suggests a strong climatic driver of permafrost aggradation that interacted with site factors such as peat thickness, Sphagnum presence, or snow cover to control whether permafrost aggraded in individual sites. The contrast in the timing of permafrost aggradation between the more continental climate of Central North America (later, Figure 3b) and more humid, maritime Eastern North America (earlier, Figure 3a) suggests that in addition to neoglacial cooling, other climatic factors may have affected the timing of permafrost aggradation in Central North America, such as changes in winter precipitation or snow cover. Across Eurasia, the strongest trend in permafrost aggradation occurred in the Russian and European Arctic after 3250 BP (Figure 4a). In the Laptev Sea region, Betula nana pollen disappeared after 3300 BP, indicating that conditions became colder and more similar to today after that time (Andreev et al., 2011), and tree lines retreated southward (MacDonald et al., 2000; Oksanen et al., 2001). Permafrost aggradation in all Eurasian peat cores increased between ~3000 and 1000 BP, with the increase occurring earlier in Arctic Europe, European Russian, and Russian Coastal Plain (Figure 4a) (Oksanen, 2006; Väliranta et al., 2003) and transgressing eastward across central Siberia (Figure 4c) and eastern and far eastern Siberia (Figure 4d), where permafrost increased markedly after 1000 BP. A transgressive west to east cooling is also suggested by pollenderived temperature reconstructions that show similar lags between western and eastern Siberia (Anderson et al., 2002). Evidence for permafrost aggradation from peatland sites in the eastern part of

9 1006 The Holocene 28(6) Eurasia that were not included in this analysis both agree and disagree with this interpretation. Permafrost aggradation was indicated around 2550 BP at Vaisjeäggi palsa mire in northern Finland (Oksanen, 2006) and around 3100 BP in one of two Sub-Arctic cores from Ortino in far northeastern European Russia (Väliranta et al., 2003). However, permafrost aggraded in the other core from Ortino significantly earlier, around 5550 BP (Väliranta et al., 2003), again suggesting that local factors also play a significant role in the timing of permafrost aggradation. Generally, Eurasian neoglacial cooling results in a decrease of permafrost-free fens in our records and increase in permafrost in the cores after 2000 BP (Figure 4a c), suggesting a colder and potentially drier shifted fens to drier, ombrotrophic bogs and permafrost peatlands. Pollen and macrofossil data from a peatland in the Pur Taz region of Siberia indicate dry conditions with abundant charcoal, and either little peat accumulation or oxidation of peat following deposition after ~4500 BP, or both (Peteet et al., 1998). Late-Holocene ( BP). Several Arctic regions showed an increase in permafrost aggradation beginning around 1500 BP. Other regions continued to aggrade permafrost (Eastern North America) or showed little change in permafrost conditions. The large increase in permafrost presence after ~1500 BP in eastern Siberia (Figure 4d) agrees with the timing of permafrost aggradation in Fennoscandia, Arctic Europe, and the Russian Coastal Plain (Figure 4a, b, and d), as well as Arctic Canada (Figure 3c), suggesting much colder temperatures in the Arctic at that time. Tree growth rings also indicated cool conditions in this region during this period (Naurzbaev and Vaganov, 2000) and reconstructed winter sea-surface temperatures also decline after 1500 BP (De Vernal et al., 2013; Voronina et al., 2001). From 1250 to 1000 BP, the number of cores with fen vegetation increased in Central North America (Figure 3b), as does the number of cores with thawed permafrost peat (Figure 3b). This pattern suggests slightly warmer temperatures, which coincides with the Roman Warm Period (RWP, BP) in North America (Viau et al., 2006), or that permafrost temperatures remained close to 0 C and may have increased above the freezing point interannually, increasing their susceptibility to thaw. Medieval Climate Anomaly and LIA (after 1000 BP). The peak in permafrost aggradation observed in this study occurs between 750 and 0 BP (Figure 2e), the beginning of which broadly corresponds with the end of the Medieval Climate Anomaly (MCA) and encompasses the LIA (Mann et al., 2009). This trend was spatially widespread and occurred throughout the study region after 750 BP (Figures 1, 3, and 4) (Kultti et al., 2004; Oksanen, 2006), with the exception of Western and Central Siberia, where the rate of permafrost aggradation increased earlier (Figures 1 and 4c). Our data show that the maximum modern permafrost extent was reached ~125 BP (Figure 2e), corresponding with the end of the LIA, which slightly lags the coolest LIA temperatures that occurred ~200 yr BP (Kaufman et al., 2004; Marcott et al., 2013). In North America, the pattern of greater permafrost aggradation earlier in Western North America (Arctic Canada and Alaska, British Columbia) (Figures 1, 3c, and 3d) than in Eastern North America (Figures 1, 3a, and 3b) agrees broadly with the spatial patterns of cool and warm anomalies (relative to CE) during the MCA. During the MCA, warm anomalies persisted over the North Atlantic and eastern Canada and cool anomalies persisted over western Canada and Alaska (Mann et al., 2009). This likely resulted in climatic conditions that were not favorable to permafrost aggradation in Eastern North America, and thus, permafrost aggradation commenced when conditions cooled after 1000 BP. At the end of the MCA, permafrost aggradation increased sharply in Central North America after 750 BP at a rate of more than 3% of cores per century, especially compared with previous rate in the Holocene of 0.2% of cores per century. The increased permafrost aggradation in Central North America, especially in the present-day discontinuous and sporadic permafrost zone of Manitoba and Alberta (Figures 1 and 3b), culminated in the greatest number of permafrost sites after 250 BP (Figure 3b). This suggests that MAAT decreased below 0 C in this region, and agrees with previous studies that found much of the permafrost in the discontinuous permafrost zone of Canada formed during the LIA (Halsey et al., 1995; Zoltai, 1995). Eastern North America was the only region experiencing an increasing fraction of permafrost-free bogs relative to other peatland classes during the LIA as indicated by the presence of ombrotrophic bog vegetation. This suggests either autogenic shifts toward ombrotrophy as aging peatlands accumulated peat above the water table or drier conditions resulting in a decrease in water level, or both, that would have resulted in ombrotrophic conditions and vegetation changes (Figure 3a) (e.g. Magnan and Garneau, 2014). A recent late-holocene hydroclimate synthesis of multi-proxy records found drying in many records from Quebec and Nova Scotia during this time (Rodysill et al., in press). Using lake-level reconstructions, Payette and Delwaide (2004) found increased tree mortality at the onset of the LIA ( BP) in Northern Quebec because of a combination of peatland permafrost aggradation and a drier climate. Taken together, peatland succession and permafrost aggradation does not appear to be the result of autogenic peatland processes alone, but due in part to climatic drivers, including both temperature and moisture. Permafrost thaw Unlike permafrost aggradation, permafrost thaw is rarely recorded in the peat cores studied despite warmer late-holocene climatic periods associated with the RWP and the MCA. The highest rates of permafrost thaw in this study occurred as a result of both widespread permafrost extent in peatlands and warmer conditions since the end of the LIA (Figure 2b and f). Given the relatively widespread occurrence of permafrost in peatlands today (Figure 2e) combined with warming temperatures, it seems likely that the thawing trend will continue, especially in regions where permafrost has thawed before. Understanding the resilience of permafrost under a warming climate is important given the potential for losses of peatland carbon stocks with permafrost thaw is significant (Jones et al., 2016b; O Donnell et al., 2012). The long-term response of peatland C stocks to permafrost thaw is important for predicting carbon cycle feedbacks to future warming. Permafrost thaw requires that permafrost was established at the site previously. This study examines permafrost aggradation since 6000 BP, as permafrost was only present in a few cores prior to this time (Figures 2 4). Many peatlands formed in formerly glaciated regions (Gorham et al., 2007), where in lieu of permafrost, ice sheets covered much of the present-day permafrost peatland region (Vandenberghe et al., 2014). Rapid warming following the LGM was likely unfavorable to permafrost aggradation, as HTM occurred early in the Holocene (Fisher et al., 1995; Kaufman et al., 2004). However, where permafrost was present prior to 6000 BP, evidence for permafrost thaw prior to peat formation ( BP) occurred in Arctic Canada, Alaska, and Eastern Siberia (Table S1, available online; Figure 4d), as peatlands developed in drained thaw lake basins following permafrost thaw (De Klerk et al., 2011; Geurts, 1985; Jones et al., 2012; Walter Anthony et al., 2014). While permafrost aggradation generally increases over the 6000-year period (Figure 2e) in agreement with