3.2. Holocene Variations in Soot and Char
 Long-term trends in ECsoot and ECchar showed large variations during the Holocene (Figures 2a and 2b). ECsoot fluxes varied from 0.002 to 0.71 mg cm−2 yr−1 and ECchar varied from 0.002 to 2.63 mg cm−2 yr−1 since 10 ka. In general, mean ECsoot and ECchar values were low prior to 8 ka and higher after, although a marked decline occurred in the past 2000 years. ECsoot levels increased from very low values prior to 8 ka to relatively high values ca. 5 ka. A brief decline in ECsootoccurs from 5 - 4 ka, and then levels stabilize until about 1.5 ka. After 1.5 ka, ECsoot declines sharply to a local minima ca. 600 and 300 years ago, and then increases rapidly toward present. ECcharremains low during the early Holocene prior to 8 ka, and then increases gradually from 8 - 1.5 ka. After 1.5 ka, ECchar declines rapidly (like ECsoot) to a local minima between ca. 600 and 300 years ago and then increases sharply in the past few centuries.
Figure 2. Comparison of (a) standardized ECsoot values with (b) standardized ECcharvalues from Lake Daihai; (c) total organic carbon; (d) the composite standardized and smoothed (250- and 500-yr windows) biomass burning record from 36 sites in eastern Asia [Marlon et al., 2012]; (e) GISP2 δ18O data [Stuiver et al., 1995] used to infer northern hemisphere temperature variations; (f) δ18O data from Dongge Cave [Dykoski et al., 2005] and (g) Sanbao Cave [Dong et al., 2010] used to infer regional moisture availability; (h) estimated changes in population and (i) crop and pastureland from the HYDE data set [Klein Goldewijk et al., 2010] for 10°–45°N latitude, 65°–150°E longitude; (j) tree pollen from [Xiao et al., 2004]; (k) changes in dynasties [Zhang et al., 2008]; and (l) deposition times (yr cm−1) based on the age-depth model. High-resolution data (i.e., the EC, Dongge Cave and GISP2 records) were smoothed with a lowess curve using a 500-year moving window.
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 High-frequency variations in ECsoot and ECchar also show several shifts during the Holocene. Variability is low, for example, prior to 8 ka, but this is likely due in part to low sedimentation rates (high deposition times) during this period (Figure 2l). High-frequency changes in both ECsoot and ECcharare relatively consistent from 8 - 2 ka, with the exception of a reduction in variability from about 5 - 4 ka, and an abrupt decline in overall EC values ca. 2.9 ka. Variability in ECsoot and ECchar appears to increase after 2 ka, and modern ECsoot and ECchar values in particular are among the highest observed during the Holocene. Holocene ECsoot levels surpassed modern values only twice–once ca. 1750 years ago (AD 200) and once ca. 600 years ago (AD 1350) (Figure 2a). The modern ECchar value is the highest in the Holocene and surpasses the mean by more than six standard deviations (Figure 2b).
 ECchar accounts for the largest component of total EC concentrations (86%) and is strongly correlated with total EC (r = 0.99; S3 in SI). ECsoot is less correlated with total EC (r = 0.49; S3 in SI), but ECsoot and ECchar fluxes are strongly positively correlated (r = 0.89). Despite this high correlation, several noticeable differences exist between the standardized ECsoot and ECchar flux data (Figures 2a and 2b). For example, after 8 ka, ECsoot increases immediately, whereas ECchar remains relatively low until 7 ka. ECcharalso shows a steady increase in its mean values from the mid- to late-Holocene (e.g., between 6 and 2 ka), whereas mean ECsoot values remain similar during these two periods. ECcharalso shows a stronger decline between 2 - 1 ka. Thus, although ECsoot and ECchar appear to reflect a similar underlying process (e.g., biomass burning), differences in ECsoot and ECchar production, transportation or deposition may account for the observed differences [Elmquist et al., 2006; Han et al., 2010]. A negative correlation between ECchar and TOC (r = −0.64; P < 0.001) and ECsoot and TOC (r = −0.42; P < 0.05) suggests that charring in the analysis has little impact on the EC and EC fractions, and non-pyrogenic matter is not included in the analysis processes using the thermal optical method. ECchar and ECsoot, which are indicators of biomass burning in this region, are mainly associated with dry conditions (see details of the comparison of EC with the speleothem record below).
3.3. Controls on Holocene Trends in Biomass Burning
 Biomass burning is controlled predominantly by climate on regional- to global-scales [Carmona-Moreno et al., 2005]. Seasonal changes in temperature and precipitation, for example, produce a distinct cycle of burning in the northern and southern hemispheres each year [van der Werf et al., 2006]. Vegetation changes also have a strong influence on millennial-scale trends in fire through their effects on fuel characteristics (e.g., abundance, distribution, and flammability). Human activities may have contributed to carbon accumulations in Lake Daihai, but the extent and timing of such human impacts during the Holocene are largely unknown. The charcoal-based biomass burning index as well as data for each of the primary controls on biomass burning–climate, vegetation and human activities–show distinct trends during the Holocene that may help explain the observed changes in EC at Lake Daihai.
 The regional biomass burning index (Figure 2d) shows large variations prior to 6 ka, and an upward trend from 6 - 1.6 ka. Biomass burning declines sharply from 1.6 - 1 ka, then rises to AD 1950, and finally declines again to present. Monsoon intensity inferred from the Dongge and Sanbao caveδ18O values (Figures 2f and 2g) is relatively strong prior to 6 ka, decreases from 6 - 3 ka, and subsequently remains relatively low. There are no regional temperature reconstructions for eastern Asia for the Holocene. Theδ18O data from Greenland indicate a slight and gradual decline in temperatures since about 8 ka at high latitudes, but may have varied from midlatitude temperature trends [Grafenstein et al., 1999]. Daniau et al. , for example, suggest that midlatitude temperatures increased during the Holocene and that such changes are consistent with increased biomass burning. Aside from the long-term temperature trends, however, theδ18O data from Greenland indicate several large short-term changes that are known to be widespread, including a decline in temperatures associated with the “8.2 ka event” [Alley and Agustsdottir, 2005] and the Little Ice Age ca. 550 - 250 years ago [Mann et al., 2009].
 Vegetation changes near Lake Daihai were reconstructed from changes in the relative abundance of pollen taxa [Xiao et al., 2004] and are representative of broader trends in the region [Zhao et al., 2009]. The primary vegetation types shift from steppe in the early Holocene to steppe forest in the mid-Holocene (8 - 3 ka), to desert steppe in the late Holocene (after ∼3 ka). Specifically, arid steppe (high Artemisia, Chenopodiaceae and Ephedra pollen and low arboreal pollen [AP] values) existed from ca. 10 - 8 ka due to a colder-than-present climate and low effective moisture. There is an apparent contradiction between the climate inferred from the vegetation at this time and lowδ18O values in the speleothem records thought to reflect increased monsoon intensity. This discrepancy is widely recognized and may be attributable to low North Atlantic sea-surface temperatures (SSTs) and high-latitude air temperatures that affect the availability, amount and transport of water vapor [Chen et al., 2008]. Between 8 - 3 ka, tree pollen is relatively high and regional climate was generally warmer and wetter than previously. The transition from the mid- to late-Holocene was marked by a shift toward drier conditions, evidenced both by the Dongge Caveδ18O record and the decline of tree pollen (Figures 2c and 2f). Xiao et al. infer that forest steppe shifted to steppe vegetation during this time due to effectively drier conditions. Today the region is characterized by desert steppe vegetation due to its transitional semi-arid/semi-humid climate.
 The population estimates from the HYDE data set show a steady increase during the Holocene (Figure 2h) with the exception of two important events: the Mongol Invasions starting in AD 1211, when China lost about a third of its population, and the fall of the Ming Dynasty around AD 1644, when it lost about a sixth of its population [McEvedy and Jones, 1978; Pongratz et al., 2008]. The land-use area estimates from HYDE show an initial increase beginning 7000 years ago, with variations during the past 2000 years that reflect the changes in population estimates.
 Analyses of correlations between the EC data and its potential controls, including the biomass burning and monsoon indexes, as well as temperature proxies and vegetation changes indicate that both ECsoot and ECchar fluxes have significant positive correlations with biomass burning and negative correlations (although not significantly) with monsoon intensity, TOC content, tree pollen and the Greenland temperature record (Table 1). In particular, ECsoot and ECchar are both positively correlated with the regional biomass burning index, and the relationship is slightly stronger for ECchar (r = 0.51, P < 0.01) than for ECsoot (r = 0.41, P < 0.05). ECchar is also more strongly correlated with the δ18O values from Dongge Cave (r = 0.88, P < 0.001) than is ECsoot (r = 0.66, P < 0.001) (higher δ18O values reflect a weaker monsoon, so there is a negative relationship between biomass burning and monsoon intensity). Finally, ECchar (but not ECsoot) is negatively correlated with the tree pollen data (r = −0.35, P < 0.05).
Table 1. Statistical Correlations Between EC and Potential Explanatory Variables
|EC Data||Potential Explanatory Variable||Correlation|
|ECsoot||Regional charcoal composite||r = 0.41; P < 0.05|
|ECsoot||Dongge Cave δ18O||r = 0.66; P < 0.001|
|ECsoot||Arboreal Pollen %||r = −0.04; P > 0.05|
|ECsoot||GISP2 δ18O||r = −0.16; P > 0.05|
|ECchar||Regional charcoal composite||r = 0.51; P < 0.01|
|ECchar||Dongge Cave δ18O||r = 0.88; P < 0.001|
|ECchar||Arboreal Pollen %||r = −0.35; P < 0.05|
|ECchar||GISP2 δ18O||r = −0.23; P > 0.05|
 Low ECsoot and ECchar levels prior to 8 ka imply low biomass burning and are consistent with cool, dry climate conditions (and potentially reduced convection and lightning). Limited vegetation productivity and fuel abundance inferred from the pollen data also likely reduced fire spread (Figure 2j). High charcoal levels from 10 - 9 ka primarily reflect increased burning to the northeast of Lake Daihai [Li et al., 2005], although high burning is common elsewhere in arid steppe environments during the late glacial; this may be partly reflected by moderately high EC values ca. 10 ka (Figures 2a and 2b).
 A comparison of ECsoot with the GISP2 δ18O record (Figures 2a and 2e) indicate that both were reduced during the abrupt and widespread cooling that occurred during the 8.2 ka event [Alley and Agustsdottir, 2005]. The maximum likely age of the EC samples currently dated to 8 ka is 8.35 ka, and thus it is possible that the minimum in the EC data was synchronous with the 8.2 ka event. Because the independent regional biomass burning record also reaches a minimum ca. 8.2 we argue that the climate changes reduced fire and thus ECsoot and ECchar levels at 8.2 ka (Figures 2a, 2b, 2d, and 2e). Elsewhere in the world, climate changes often resulted in an opposite effect (i.e., increased burning), however [Marlon et al., 2012].
 Human activities may be expected to have had an increasing influence on ECsoot and ECchar trends and/or variability as the Holocene progressed, particularly in the late Holocene with the advent of dynasties in China (Figure 2k) [Expert Group of the Xia-Shang-Zhou Project, 2000]. The location of Lake Daihai near the northern limit of intensive agriculture also made it an area of frequent conflict between nomadic tribes and settled farming societies [Huang and Su, 2009]. Early Holocene hunter gatherer populations were low, however, ECsoot and ECchar were low, and charcoal levels were declining, suggesting very limited ecological effects from people on fire prior to 8 ka (Figures 2a, 2b, 2d, and 2h). The earliest large increase in ECsoot, ECchar, and charcoal occurs between 8 - 7 ka and is associated with a large increase in AP (Figures 2a, 2b, 2d, and 2j), which is also not consistent with widespread human impacts on burning.
 Two short-term increases in ECchar at 7 ka and 4.1 ka coincide with declines in AP that may reflect localized human impacts on burning through deforestation, but the first local evidence for agricultural activity does not occur until 6.3 ka [Tian, 2000] and there is no major shift in ECsoot or ECchar at that time. The second increase in ECchar and decline in tree pollen ca 4.1 ka occurs during a period of declining farming and cultural activity following the demise of the Laohushan culture [Tian, 2000]. The gradual increase in ECcharfrom 7 - 2 ka parallels a gradual increase in regional human activity (Figures 2h and 2i), but increases in Pinus (pine), Quercus (oak) and Ostryopsis(birch) during the mid-Holocene indicate an expansion of forest steppe vegetation and a shift to generally warm, wet conditions [Xiao et al., 2004] that do not support a strong human impact on burning then. It is possible that fire gradually increased during more frequent but brief intervals of drought during this interval.
 A sharp decline in both ECsoot and ECchar occurs ca. 3 ka coincident with a reduction in regional biomass burning and also with increased effective moisture (Figures 2a, 2b, 2d, and 2g). Subsequently, available moisture declines, AP declines sharply, regional biomass burning increases, and both ECsoot and ECcharshow some of their highest Holocene values; all of this occurs despite a reduction in estimated population growth and land-use (Figures 2a, 2b, 2d, 2f, 2g, 2h, 2i, and 2j). Reduced burning during wet periods and increased fire during dry periods suggests that changes in monsoon intensity or effective moisture continued to exert a strong influence on ECsoot, ECchar and biomass burning throughout the late Holocene.
 The large decline in ECsoot and ECcharfrom ca. 2 - 1 ka is matched by a similar decline in biomass burning (Figures 2a, 2b, and 2d) and a final decline in AP (Figure 2j). The lowest late-Holocene EC levels occur around the Little Ice Age (LIA, 550 - 250 years ago [Mann et al., 2009]), but the regional fire decline starts well before the LIA. In central eastern Inner Mongolia [Huang et al., 2005] and the southern loess plateau, burning is thought to be primarily human-caused since 3.1 ka [Huang et al., 2006]. A large fire decline during the past ∼600 years in the loess plateau is thought to result from the complete transformation of the landscape to agricultural land and to a lack of vegetation left to burn [Huang et al., 2006], but the HYDE data at least are not consistent with this idea. There is not a unique shift in EC, for example during the Mongol invasions, and changing dynasties also do not coincide with major EC fluctuations (Figures 2a, 2b, and 2k). Shifts in moisture availability seem a more plausible explanation for the large changes in late-Holocene biomass burning, which are similar across the northern hemisphere [Marlon et al., 2012].
 The large variations in biomass burning of the past millennium have implications for debates beyond eastern Asia. The EC minimum 500 years ago in particular parallels a global decline in fire activity at that time, which is roughly coincident with both the onset of the LIA and with the arrival of Europeans in America that led to widespread indigenous population collapse. The fire decline in the tropical Americas is often argued to have reduced global atmospheric CH4 concentrations from biomass burning [Ferretti et al., 2005; Houweling et al., 2008; Mischler et al., 2009; Finkelstein and Cowling, 2011], but whether the decline was triggered by population collapse or reduced temperatures is debated [Marlon et al., 2008; Nevle and Bird, 2008; Dull et al., 2010; Power et al., 2012]. The occurrence of a marked regional decline in soot and char in eastern Asia provides important evidence that the fire decline was not limited to the Americas, and thus pandemics associated with European colonization alone cannot account for the decline in Asia. Rather, climate changes or other shifts in human activities must be invoked to explain the widespread fire decline.
 In the past 50 years, ECsoot and ECchar reach very high levels (Figures 2a and 2b). The sources of EC during this interval shift from biomass burning to a combination of biomass and fossil-fuel combustion products, partly reflecting the industrialization of Asia during the latter half of the 20th century. As the Holocene EC records show, similar ECsoot levels were only attained about 800 and 2000 years ago, and current ECchar levels are higher than at any other time in the Holocene.