Lake level evidence for a mid-Holocene East Asian summer monsoon maximum and the impact of an abrupt late-Holocene drought event on prehistoric cultures in north-central China

Abstract

Uncertainty regarding the timing of the highest Holocene water level of enclosed Dali Lake in northern China has led to controversy about whether the East Asian summer monsoon (EASM) peaked in the early Holocene or the mid-Holocene. Therefore, a record combining a reliable chronology with unambiguous lake level indicators is essential to resolve the issue. In this study, we established a temporal sequence of well-preserved paleolake shorelines at different elevations around Dali Lake using quartz optically stimulated luminescence (OSL) dating. Combining the new OSL-based chronological data with the previously published ages of beach ridges, we constructed an integrated lake level record for Dali Lake since the last deglaciation; the record is chronologically well-constrained and provides a clearer representation of lake level changes than previous studies. The main findings are as follows: (1) the level of Dali Lake rose gradually during 14.5–8.0 ka and reached a highest level during 8.0–6.0 ka that was ~61 m higher than today, before gradually declining after ~6.0 ka; (2) although the short-lived highstand of Dali Lake during the early Holocene was caused by both monsoon precipitation and snow/ice meltwater influx, the mid-Holocene lake level maximum was mainly the result of increased monsoon precipitation. These findings indicate that the EASM maximum in the region occurred during the mid-Holocene, which is supported by precipitation-proxy records from the EASM margin; (3) a major decline (~30 m) of the level of Dali Lake occurred at ~4.2 ka, reflecting a regional-scale drought event in northern China. Combined with near-contemporaneous frequent floods in the lower Yellow River, we propose that the pattern of “drought in northern China, flooding in central China” occurred during ~4.2–3.8 ka, triggering the migration of the prehistoric human population of the area to the central Plain of China. This population migration may have destabilized the existing social order and promoted the emergence of more complex societies, leading to the development of early civilization in north-central China.

Keywords

Abrupt lake level decline early civilization highstand Holocene Paleolake shorelines quartz optically stimulated luminescence dating

Introduction

The East Asian summer monsoon (EASM) is a major component of the global climate system, and its variability directly impacts the lives of more than one billion people (An, 2000; Chen et al., 2015, 2021; Liu et al., 2015a). Understanding the past variability of the EASM is critically important for predicting future changes (Chen et al., 2015). However, there is an ongoing and vigorous debate about whether the Holocene highest water level of Dali Lake, in northern China, occurred in the early Holocene (11.7–8.0 ka) or the mid-Holocene (8.0–4.0 ka) (Goldsmith et al., 2017a, 2017b; Jiang et al., 2020; Liu et al., 2017a). It is also disputed whether replenishment of the lake water by snow/ice meltwater from the Great Khingan Mountains affected the level of Dali Lake during the early Holocene (Fan et al., 2017, 2019; Goldsmith et al., 2017a, 2017b; Jiang et al., 2020; Liu et al., 2017a; Xiao et al., 2008, 2009; Yue et al., 2021). Although the history and the factors driving the lake level changes of Dali Lake have been of major scientific interest for several years, the lack of a resolution of these issues has led to much controversy about whether the Holocene EASM maximum occurred in the early Holocene or the mid-Holocene (Goldsmith et al., 2017a, 2017b; Jiang et al., 2020; Liu et al., 2017a; Xiao et al., 2008).

Based on the dating of lake beach ridges, lacustrine sediment outcrops and alluvium samples, Goldsmith et al. (2017a, 2017b) reconstructed the Holocene history of lake level fluctuations at Dali Lake. These studies concluded that the level was highest during the early Holocene, and therefore that the EASM reached its Holocene maximum at that time. In contrast, Xiao et al. (2008, 2015) and Liu et al. (2017a) demonstrated that the early Holocene highstand of Dali Lake was caused by snow/ice meltwater influx rather than by monsoon precipitation. In addition, Jiang et al. (2020) pointed out that only four beach ridge samples were used by Goldsmith et al. (2017a, 2017b) in their reconstruction, and that most of the materials used were from lacustrine sediment outcrops and alluvium, which may be ambiguous lake level indicators. Also, Han and Li (2020) proposed that the radiocarbon dating results of Goldsmith et al. (2017a) were affected by sediment reworking and a carbon reservoir effect and demonstrated that several of their dating results were problematic. Furthermore, Liu et al. (2017a) pointed out that even if the lake level reconstruction of Goldsmith et al. (2017a) was correct, the data still indicated that the highest lake level would have occurred in the mid-Holocene if Dali Lake had not been in an outflow state at that time. Recently, Jiang et al. (2020) studied 13 beach ridges at different elevations at Dali Lake to reconstruct the Holocene history of lake level fluctuations. They concluded that the highest level of Dali Lake occurred during the mid-Holocene and that this represented the Holocene EASM maximum. Nevertheless, the reality of the early Holocene lake level changes of Dali Lake proposed by Jiang et al. (2020) are uncertain, given that there was only a single data point during the interval of 12.0–7.7 ka. Additionally, Goldsmith and Xu (2020) suggested that the 3 out of the 13 beach ridges used in Jiang et al. (2020) are not in stratigraphic order, refuting the conclusion of Jiang et al. (2020) that the early Holocene was relatively dry.

In addition, the abrupt termination of the Holocene humid period, as indicated by an abrupt lake level change at Dali Lake, may have had major impacts on the development of prehistoric cultures (Goldsmith et al., 2017a; Magny, 2004; Sun et al., 2019; Yan et al., 2020a). Evidence of several Neolithic cultures has been found in and around the Dali Lake catchment, such as the Hongshan (Yang et al., 2015) and Laohushan cultures (Tian and Akiyama, 2001). Therefore, more detailed, integrated, and well-dated evidence of paleolake shorelines is needed to reconstruct the lake level history of Dali Lake, which can potentially resolve the controversy regarding the timing of the EASM maximum. Moreover, a high-quality lake level record can also be used to explore the impact of abrupt climate change on the succession of Chinese cultures in the region.

In this study, we excavated and sampled a series of trenches in well-preserved paleolake shorelines at different elevations around Dali Lake. We then used optically stimulated luminescence (OSL) dating of quartz grains extracted from the sediments to provide a chronological framework for these sequences. Combined with previous results from the site, we used the chronology to provide a robust lake level history of Dali Lake and to determine whether the EASM peaked in the early Holocene or mid-Holocene. We also investigated the influence of the abrupt lake level decline in the middle-late Holocene and associated ecosystem change on the development of prehistoric cultures in the area. Overall, the findings improve our understanding of past EASM variations and prehistoric human-environment interactions in the EASM region.

Regional setting

Dali Lake is the second largest lake in Inner Mongolia, China. It lies on the northern margin of the EASM region within the ranges of 43.2°–43.4°N and 116.5°–116.7°E (Figure 1). The modern area and elevation of Dali Lake are ~180 km² and 1223 m a.s.l., respectively. The lake basin is hydrologically closed and had a catchment area of ~5900 km² before 4.2 ka, and ~4700 km² afterward (see Section 5.3). A basalt platform delimits the lake catchment in the north and west, and lake cliffs have developed on the margin of the basalt platform. Several band-shaped paleolake shorelines are distributed on the now terrestrialized lacustrine plain between the lake cliff and the modern lake shoreline in the northern part of the catchment (Figure 2). The Gongger grassland neighboring the Great Khingan Mountains is located to the east of Dali Lake. To the south of Dali Lake is the Otindag Sandy Land, with an area of 21,400 km², where the geomorphology has been transformed to varying degrees by intensive wind erosion, making the distribution of the paleolake shorelines in the area unclear.

Figure 1.

Inset (a) shows the modern northern limit of the Asian summer monsoon (green dashed line, Chen et al., 2008) and the location of Dali Lake and other lakes in the EASM margin (Chagan Nur Lake, Daihai Lake, Gonghai Lake, Baijian Lake and Qinghai Lake). The main modern trajectories of the westerlies and the EASM are indicated by arrows. Modern catchment of Dali Lake (b).

Figure 2.

Remote sensing image of the Dali Lake catchment and the localities of sections at the paleolake shorelines. Inset squares show the locations of a–h.

Dali Lake is fed by four rivers (Figure 1): two perennial rivers, the Gongger and Salin Rivers, enter the lake from the northeast, and two intermittent rivers, the Holai and Liangzi Rivers, enter from the southwest. The Gongger River, which originates in the Great Khingan Mountains, is the most important water source and comprises 57% of the total annual river supply, reaching 3.2 × 10⁷ m³ per year (Management Office of the Dalinuoer National Nature Reserve in the Inner Mongolia, 2015). According to the records of the closest meteorological station to Dali Lake (Hexigten station) (Supplemental Figure S1), the region has a mean annual temperature of 3.2°C and mean annual precipitation of 392 mm, 66% of which is from intense summer precipitation during June–August.

Materials and methods

Elevation determination and lithology of the paleolake shorelines at Dali Lake

A paleolake shoreline is generally a landform formed by the accumulation of sediment near the shore during the stable period of the ancient lake, and its elevation is approximately equal to the lake level at that time (Jones and Jordan, 2007). As shown in the satellite image in Figure 2, a series of paleolake shorelines are preserved in the Dali Lake catchment. However, most of these shorelines are discontinuous segments that are visible in remotely sensed images. Furthermore, several built roads parallel to the lakeshore in the Dali Lake catchment appear very similar to the paleolake shorelines in the satellite images. We used the following methodology to determine the genuine paleolake shorelines of Dali Lake. First, we analyzed the satellite image of Dali Lake catchment in detail through visual interpretation, identified the possible paleolake shorelines with the characteristics of ridge-like landform along the contour line in the catchment, and then recorded the corresponding position and elevation information. Second, we used a semi-automated paleolake-recovery tool (PaleoLakeR—details of the operational procedures are given in Sheng (2009)), to determine whether the paleolake shorelines assessed by visual interpretation could form a closed lake surface. Third, the shorelines determined by PaleoLakeR were analyzed by field investigation to ensure that each paleolake shoreline was formed by accumulation, so that the OSL samples could be obtained from the sedimentary profile. Correspondingly, we identified a total of 11 paleolake shorelines (designated DL19-1 to DL19-11, in order of increasing elevation) that developed within the Dali Lake catchment.

The longitude and latitude of each sampling point on the paleolake shorelines were recorded in a satellite image during fieldwork and were subsequently matched to a digital elevation model (DEM) obtained from a topographic map (1:50,000) with a contour interval of 10 m, in order to obtain the absolute elevation of each shoreline. The area of the DEM pixel size is 25 × 25 m and the elevation error of the DEM measurements is <5 m. In addition, the relative elevation difference between the paleolake shorelines and the modern lake level was further evaluated using a Differential Global Positioning System (DGPS) with an elevation error of ~2–3 m. Comparison of the relative elevation difference between the paleolake shorelines and the modern lake level obtained using the two methods (DEM and DGPS) showed that the results were very consistent within the error range (Supplemental Table S1), which demonstrates that the elevations retrieved from the DEM used in this study are reliable. Moreover, using the DEM to determine the elevation of each paleolake shoreline ensured that the calculations of other hydrological parameters (e.g. lake area, water volume) of Dali Lake are consistent, further reducing the uncertainty associated with instrumental measurements.

An excavator and hand shovels were used to dig trenches in the paleolake shorelines, penetrating uncompacted beach sands and gravels to reach former fine-grained lacustrine sediments. Ten out of eleven paleolake shorelines are indicated by beach ridges, while DL19-9 is a wetland sediment-like paleolake shoreline. The lithology of the paleolake shorelines can typically be divided into two units (Figure 3): (1) An upper horizontal layer of soil or loess; and (2) a lower obliquely-laminated layer of well-sorted and rounded sand and gravel, with occasional fragments of mollusk shells (Radix auricularia, Supplemental Figure S2), and/or a lacustrine layer consisting mainly of silt and clay. Profile DL19-9 is a 210 cm thick wetland sedimentary sequence, located next to the Holai River which connects Dali Lake and Duolun Lake (Figure 2). The upper part of the profile (0–100 cm) is a black fine silt-clay, intercalated with a grayish-yellow layer of clayey and fine silty-sand, with an irregular and disturbed stratification. The lower part of this profile (100–210 cm) consists of sand and lacustrine clay. Detailed information about the paleolake shorelines is listed in Table 1, the stratigraphy and sampling locations are shown in Figure 3 and marked-up photographs of the paleolake shorelines taken in the field are shown in Supplemental Figure S3. In addition, to increase the sample size of the relative paleolake level indicators, we included the lowest elevation of Luwang City (1253 m a.s.l.), a fortification built in 1270 AD (680 a), and the lowest section of the Jin Wall (1231 m a.s.l.), built in 1198 AD (752 a). These supplemental elevations were also obtained by DEM data and they provide an upper limit for the level of Dali Lake.

Table 1.

Information about the elevation and location of paleolake shorelines in the Dali Lake catchment.

Dali Lake shoreline section	Longitude/E	Latitude/N	Elevation/m	Above modern lake level/m	Lakeshore profile height/m	Number of samples
DL19-1	116°38′22.47″	43°23′21.23″	1231.1	8.1	1.4	3
DL19-2	116°37′42.27″	43°23′45.98″	1240.6	17.6	1.8	1
DL19-3	116°43′14.32″	43°28′10.72″	1263.5	40.5	2.55	5
DL19-4	116°43′15.01″	43°28′17.15″	1267.2	44.2	2.35	1
DL19-5	116°43′14.67″	43°28′14.48″	1267.3	44.3	1.86	3
DL19-6	116°55′58.94″	43°22′21.11″	1274.0	51.0	2.23	5
DL19-7	116°39′58.01″	43°25′38.08″	1275.0	52.0	6.8	2
DL19-8	117°02′44.21″	43°26′53.92″	1276.6	53.6	1.56	1
DL19-9	116°26′16.96″	43°15′11.54″	1280.4	57.4	2.1	2
DL19-10	117°05′38.47″	43°19′34.56″	1283.2	60.2	1.6	1
DL19-11	116°49′2.16″	43°27′18.83″	1287.5	64.5	4.35	2

Figure 3.

Lithological units and luminescence ages of samples from paleolake shorelines DL19-1 to DL19-11 in the Dali Lake catchment. The paleolake shoreline sections are arranged in clockwise order around Dali Lake, starting with DL19-1. The detailed locations are shown in Figure 2.

OSL dating

Sampling and pretreatment

All OSL samples were obtained by hammering steel tubes (33 cm long, 6 cm diameter) into the freshly excavated vertical section. The tubes were covered and sealed with aluminum foil, and then wrapped in black plastic bags and taped to avoid light exposure and moisture loss. In total, 26 samples were collected from shorelines at various elevations within the Dali Lake catchment. Detailed information about the samples is given in Table 2.

Table 2.

Quartz OSL dating results for samples from paleolake shorelines in the Dali Lake catchment.

Sample No.	Depth (m)	Aliquots (Accepted/Measured)	Grain size (μm)	De (Gy)	Over-dispersion (%)	U (ppm)	Th (ppm)	K (%)	W.C. (%)	Cosmic dose rate (Gy ka⁻¹)	Dose rate (Gy ka⁻¹)	Age (ka)	Age model
DL19-1-35	0.35	14 (16)	90–300	1.45 ± 0.07	6.90 ± 4.88	0.46 ± 0.01	1.80 ± 0.04	1.64 ± 0.03	5 ± 2.5	0.26	2.10 ± 0.06	0.69 ± 0.04	CAM
DL19-1-95	0.95	16 (16)	90–180	34.78 ± 1.70	17.51 ± 3.53	1.49 ± 0.07	7.17 ± 0.33	2.41 ± 0.02	10 ± 5	0.23	3.18 ± 0.13	10.94 ± 0.70	CAM
DL19-1-120	1.20	15 (16)	90–180	49.01 ± 2.14	14.61 ± 3.48	1.76 ± 0.06	8.67 ± 0.29	2.50 ± 0.05	10 ± 5	0.22	3.41 ± 0.14	14.39 ± 0.87	CAM
DL19-2-55	0.55	16 (16)	90–125	36.90 ± 1.35	12.55 ± 2.69	1.66 ± 0.03	8.07 ± 0.13	2.33 ± 0.06	10 ± 5	0.25	3.25 ± 0.14	11.35 ± 0.64	CAM
DL19-3-30	0.30	16 (16)	63–180	5.91 ± 0.72	53.84 ± 9.62	0.83 ± 0.02	3.52 ± 0.08	2.04 ± 0.07	5 ± 2.5	0.27	2.70 ± 0.09	2.19 ± 0.28	CAM
DL19-3-77	0.77	15 (16)	90–180	11.75 ± 0.24	6.54 ± 1.56	0.34 ± 0.01	1.66 ± 0.05	1.93 ± 0.07	5 ± 2.5	0.23	2.32 ± 0.08	5.07 ± 0.23	CAM
DL19-3-110	1.10	16 (16)	90–300	10.13 ± 0.21	6.85 ± 1.65	0.51 ± 0.04	2.24 ± 0.11	1.41 ± 0.01	5 ± 2.5	0.22	1.81 ± 0.06	5.59 ± 0.22	CAM
DL19-3-197	1.97	16 (16)	90–125	68.77 ± 2.87	15.50 ± 3.31	1.66 ± 0.06	8.47 ± 0.36	2.35 ± 0.04	5 ± 2.5	0.21	3.44 ± 0.10	19.98 ± 1.09	CAM
DL19-3-225	2.25	13 (16)	90–125	61.69 ± 1.52	5.88 ± 2.13	1.56 ± 0.03	8.27 ± 0.12	2.33 ± 0.06	10 ± 5	0.20	3.21 ± 0.17	19.20 ± 1.18	CAM
DL19-4-115	1.15	13 (15)	63–125	33.77 ± 1.54	15.07 ± 3.34	1.64 ± 0.06	8.15 ± 0.33	2.13 ± 0.05	10 ± 5	0.22	3.05 ± 0.13	11.07 ± 0.68	CAM
DL19-5-66	0.66	13 (16)	90–180	53.26 ± 3.24	20.28 ± 4.30	1.39 ± 0.03	7.40 ± 0.16	2.15 ± 0.03	10 ± 5	0.24	2.97 ± 0.15	17.94 ± 1.46	CAM
DL19-5-91	0.91	16 (16)	90–125	36.91 ± 0.91	8.18 ± 1.75	0.98 ± 0.03	3.77 ± 0.07	2.08 ± 0.05	5 ± 2.5	0.23	2.87 ± 0.09	12.84 ± 0.56	CAM
DL19-5-136	1.36	15 (16)	90–125	210.03 ± 7.50	10.61 ± 2.65	1.47 ± 0.02	7.67 ± 0.11	2.43 ± 0.05	10 ± 5	0.22	3.25 ± 0.14	64.70 ± 3.57	CAM
DL19-6-46	0.46	15 (17)	63–300	7.99 ± 0.37	15.61 ± 3.19	0.67 ± 0.03	2.81 ± 0.09	1.73 ± 0.04	5 ± 2.5	0.27	2.24 ± 0.08	3.57 ± 0.21	CAM
DL19-6-99	0.99	16 (18)	63–300	7.54 ± 0.30	16.65 ± 3.17	0.65 ± 0.02	2.37 ± 0.04	1.34 ± 0.04	5 ± 2.5	0.22	1.82 ± 0.06	4.13 ± 0.22	CAM
DL19-6-119	1.19	17 (17)	90–300	10.45 ± 0.45	17.24 ± 3.10	0.60 ± 0.01	2.14 ± 0.07	1.63 ± 0.02	5 ± 2.5	0.22	2.03 ± 0.07	5.15 ± 0.28	CAM
DL19-6-148	1.48	16 (16)	90–180	12.65 ± 0.45	13.11 ± 2.51	0.64 ± 0.01	2.48 ± 0.05	1.81 ± 0.02	5 ± 2.5	0.21	2.25 ± 0.06	5.62 ± 0.26	CAM
DL19-6-203	2.03	16 (16)	90–125	16.98 ± 0.38	7.42 ± 1.66	0.78 ± 0.04	3.59 ± 0.20	2.53 ± 0.05	5 ± 2.5	0.20	3.05 ± 0.10	5.56 ± 0.22	CAM
DL19-7-300	3.00	16 (16)	90–125	13.74 ± 0.32	8.62 ± 1.72	1.26 ± 0.08	5.79 ± 0.37	2.31 ± 0.06	5 ± 2.5	0.17	3.07 ± 0.10	4.47 ± 0.18	CAM
DL19-7-590	5.90	16 (16)	90–125	35.54 ± 1.33	13.47 ± 2.56	1.57 ± 0.03	5.30 ± 0.11	2.50 ± 0.04	10 ± 5	0.13	3.08 ± 0.14	11.53 ± 0.67	CAM
DL19-8-96	0.96	16 (16)	90–300	12.71 ± 0.36	9.62 ± 2.13	0.29 ± 0.01	1.15 ± 0.03	1.71 ± 0.04	5 ± 2.5	0.23	1.97 ± 0.07	6.45 ± 0.31	CAM
DL19-9-70	0.70	16 (16)	63–180	23.10 ± 0.78	12.77 ± 2.43	0.78 ± 0.02	3.01 ± 0.02	1.79 ± 0.05	10 ± 5	0.24	2.23 ± 0.11	10.36 ± 0.62	CAM
DL19-9-150	1.50	14 (16)	90–125	33.42 ± 1.03	10.88 ± 2.28	1.44 ± 0.04	5.40 ± 0.01	2.50 ± 0.05	10 ± 5	0.21	3.18 ± 0.16	10.52 ± 0.67	CAM
DL19-10-120	1.20	16 (16)	90–300	11.45 ± 0.70	21.66 ± 4.01	0.31 ± 0.01	1.37 ± 0.02	1.20 ± 0.04	5 ± 2.5	0.22	1.52 ± 0.06	7.55 ± 0.56	CAM
DL19-11-210	2.10	16 (16)	38–90	20.09 ± 0.48	8.56 ± 1.72	1.65 ± 0.07	8.34 ± 0.42	2.23 ± 0.08	5 ± 2.5	0.20	3.33 ± 0.11	6.04 ± 0.25	CAM
DL19-11-395	3.95	13 (16)	90–125	513.52 ± 20.74	6.08 ± 4.48	0.65 ± 0.03	2.57 ± 0.16	2.22 ± 0.07	5 ± 2.5	0.16	2.62 ± 0.08	195.80 ± 9.97	CAM

OSL sample pretreatment followed the procedures described by Aitken (1998) and was conducted under subdued red light. The sediments at each end of the sampling tube (2–3 cm) were removed and used for measurements of dose rate and water content. The unexposed materials in the middle part of the tube were pretreated for luminescence measurements, which consisted of the addition of 10% HCl and then 30% H₂O₂ to remove carbonates and organic matter, respectively. The samples were then sieved using water. It should be noted that when we sieved the samples, we prioritized the extraction of particles in the grain size range of 90–125 μm. When the quantity of extracted particles was too low, we increased the grain size range according to the grain size distribution of the sample, until the amount of sieved sample fulfilled the requirements for OSL dating. Heavy liquids with densities of 2.62 and 2.75 g/cm³ were then used to separate the quartz fraction of each sample. Quartz grains were treated with 40% HF for 60 min to remove the outer layer irradiated by alpha particles and any remaining feldspar grains. Then, all samples were treated with 1 mol/L HCl for 10 min to remove fluorides produced during HF etching. Finally, the separated quartz grains were mounted as mono-layers on 6 mm diameter stainless steel discs using silicone oil adhesive, with an average of ~2000 quartz grains on each disc.

OSL measurement procedures

OSL measurements were performed using an automated Risø TL/OSL DA-20 reader in the Luminescence Dating Laboratory of Lanzhou University, China. Irradiation was carried out using a 90 Sr/90Y beta source built into the reader. Stimulation was carried out with a blue LED (λ = 470 ± 20 nm) stimulation source. The OSL signal was detected through a 7.5 mm thick U-340 filter. The samples were bleached at 125°C for 40 s. Quartz OSL signals were calculated using an integral of the first 0.32 s of the decay curve, minus an integral of the following 0.48–0.96 s in order to minimize any influence of the slow and medium components (Cunningham and Wallinga, 2010). The purity of the quartz extracts was verified using an infrared (IR) depletion ratio method (Duller, 2003). The equivalent dose (De) of each sample was determined using the standard single aliquot regenerative dose (SAR) protocol (Murray and Wintle, 2003, Supplemental Table S2). To select an appropriate preheat temperature and to check the suitability of the SAR sequence for dating the quartz OSL samples (Murray and Wintle, 2003), preheat plateau-dose recovery tests under different preheat temperatures were conducted on representative samples DL19-5-91 and DL19-10-120. In the dose recovery test, the representative samples DL19-5-91 and DL19-10-120 were first bleached using two room temperature optical stimulations with the blue LEDs switched on for 250 s, with 10⁴ s room temperature storage between the two stimulations (Murray and Wintle, 2003). Doses of 35.14 and 9.96 Gy were then given to samples DL19-5-91 and DL19-10-120, respectively. The preheating temperature was set from 180°C to 280°C with a 20°C increment and was applied to six groups (three aliquots in each group, 18 aliquots in total) for each sample. The cut-heat temperatures were set to the range of 160–220°C. As illustrated in Figure 4a to c, measured/given dose ratios for these two samples are broadly consistent (unity ± 10%) with the application of preheat temperatures between 200°C and 240°C, and the calculated recycling ratios for 200°C and 240°C fall within the range of 0.9–1.1. Thus, a 240°C preheat and 200°C cut-heat temperature were selected for SAR De determination. For each sample, 16–18 large aliquots (6 mm diameter) were measured. Aliquots with recycling ratios between 0.85 and 1.15 and the recuperation within ±5% were accepted. Finally, 13–17 aliquots of each sample were used to calculate De. Any partial bleaching problem or mixture of grains with different ages would be masked by the bright signals from the large aliquot sizes. Therefore, five samples with typical over-dispersion (OD) values (<10%, 10–20%, 20–30%, >40%) were selected for further De measurement by using small aliquot sizes (~1–2 mm diameter stainless steel discs, with at least 24 aliquots for each sample). Based on an analysis of the characteristics of the De distributions for each sample, the central age model (CAM, Galbraith et al., 1999) was applied to well-bleached samples, the minimum age model (MAM, Galbraith et al., 1999) was applied to incomplete-bleached samples, and the finite mixture model (FMM, Roberts et al., 2000) was applied to the samples affected by a mixture of young grains.

Figure 4.

Dose recovery test results of measured/given dose ratio (a), recycling ratio (b), and recuperation of the natural signal (c) for quartz samples DL19-5-91 from shoreline section DL19-5 and DL19-10-120 from shoreline section DL19-10; (d and e) show the decay curves for the natural, regenerative dose, zero dose and test dose signals for samples DL19-5-91 and DL19-10-120, respectively (the inset is the growth curve); (f) recycling ratio histogram for all measured quartz aliquots of samples from paleolake shorelines in the Dali Lake catchment (the inset shows the histogram of the recuperation values for all measured quartz aliquots).

U and Th concentrations were measured using Inductively Coupled Plasma Mass Spectrometry (ICP-MS) and K concentrations using inductively coupled plasma optical emission spectroscopy (ICP-OES). Beta and gamma dose rates were determined based on the U, Th, and K concentrations according to the conversion factors of Guérin et al. (2011). In addition, to eliminate the effect of the gamma ray contributions from neighboring layers, the dose rate of the samples near a major boundary, or located within stratigraphic units less than 30 cm thick, was corrected following Aitken (1985) (Supplemental Figure S4). The cosmic ray dose rate was estimated for each sample as a function of depth, altitude, and geomagnetic latitude (Prescott and Hutton, 1994). Given the variability of the water content of the sediments during the burial interval, we assumed an average burial-interval water content of 5 ± 2.5% for samples of beach sand, top soil and loess, and 10 ± 5% for samples of lacustrine sediments. For quartz, although the grains may have a slight alpha activity, it is usually a good approximation to assume that the internal dose rate is negligible (Aitken, 1998). Thus, we did not consider an internal alpha dose rate. The annual dose rate and OSL ages were calculated using LDAC (V1.0), a Microsoft Excel Visual Basic Application (VBA) package (Liang and Forman, 2019). Calculation parameters and results are summarized in Table 2.

Results

Internal checks of the luminescence characteristics of quartz OSL signals

The reliability of quartz OSL ages for samples from the paleolake shorelines of Dali Lake was monitored by internal checks of the luminescence characteristics of the quartz OSL signals. The OSL signals of the natural, regenerative and test doses for the representative samples DL19-5-91 and DL19-10-120 all decreased rapidly to background within 2 s (Figure 4d and e), indicating that the OSL signals are dominated by fast components (Smith and Rhodes, 1994). The growth curves are both fitted by a saturating exponential model I/I₀ = (1 – exp^–D/D0) in which the D₀ value represents the saturation level (Figure 4d and e) (Wintle and Murray, 2006). The De values of all the samples (except sample DL19-11-395) are less than the respective 2D₀ values (e.g. Figure 4d and e), suggesting that the quartz signal has not reached saturation. In addition, the mean recycling ratio of all accepted aliquots of the samples is 1.005 ± 0.003, and the recuperation of the natural signal for all accepted aliquots of samples is within ±5% (Figure 4f). In summary, the SAR protocol was deeded suitable for our samples.

The De distributions of the representative samples DL19-5-91 and DL19-10-120 are illustrated using an Abanico plot (Dietze et al., 2016, Figure 5). For sample DL19-5-91, the De distribution is normally distributed. Most aliquots are clustered around the ±2 SD region, which is centered at the mean and the OD value is 8.18 ± 1.75%. In contrast, the De distribution of sample DL19-10-120 has a positively skewed distribution and the OD value is 21.66 ± 4.01%. Almost all the samples have concentrated De distributions (Supplemental Figure S5), with OD values <20%. A few samples have scattered De distributions with OD values up to 50%. According to their OD values (Table 2), the samples can be classified into four types: (a) OD < 10%; (b) 10% < OD < 20%; (c) 20% < OD < 30%; (d) OD > 40%. In order to minimize the quartz signal being masked by the bright signals from the large aliquot sizes, one sample of type (a) (DL19-3-77), type (c) (DL19-10-120), type (d) (DL19-3-30), and two samples of type (b) (DL19-1-95; DL19-6-99) were selected for further De measurement by using small aliquot sizes (~1–2 mm diameter stainless steel discs, with at least 24 aliquots for each sample). The results show that the De distribution of samples DL19-3-77, DL19-1-95, and DL19-6-99 (types (a) and (b)) generally have a normal distribution (Supplemental Figure S6) and their OD values are low (<20%) (Supplemental Table S3), which is like the samples measured using the large aliquot sizes. This suggests that the OSL signals of samples of type (a) and (b) are well-bleached (Bailey and Arnold, 2006). In contrast, the De distribution of sample DL19-10-120 (type (c)) is positively skewed (Supplemental Figure S6) and the OD value is over 40% and higher than that measured using the large aliquot sizes (Supplemental Table S3), suggesting that the OSL signals of samples of type (c) are incompletely bleached (Bailey and Arnold, 2006). For sample DL19-3-30 (type (d)), the De distribution is significantly scattered and negatively skewed (Supplemental Figure S6), suggesting that samples of type (d) may have been mixed with young quartz particles from the upper layers as a result of bioturbation (Jankowski et al., 2020). Therefore, we conclude that samples of types (a) and (b) are well-bleached, type (c) is incompletely bleached, and type (d) is mixed with young grains. We then calculated the De value for each sample according to its De distribution characteristics; that is, the CAM was used for samples of type (a) and (b), the MAM was used for samples of type (c), and the FMM was used for samples of type (d). For the FMM, σ_b = 20%, four groups were defined according to the principle of maximum log likelihood and the group with the highest proportion was selected for age calculation (Roberts et al., 2000) (Supplemental Table S4).

Figure 5.

Abanico plots showing the De distribution for samples DL19-5-91 (a) and DL19-10-120 (b).

Disequilibrium in the U decay chain

Disequilibrium in the U decay chain is caused mainly by leaching or depositional processes. In this study, the dose rate was consistent with depth within the same sedimentary facies and ranges from 1.52 ± 0.06 Gy/ka to 3.41 ± 0.14 Gy/ka, with an average of 2.66 ± 0.12 Gy/ka (Supplemental Figure S4). This means that the U, Th, K contents are not distorted (Supplemental Figure S4) (Wang et al., 2019; Wei et al., 2021; Zhang et al., 2018b). In addition, the radioactive elements of thorium and uranium have a relatively stable ratio in various sediment types (Allègre et al., 1986). Here, the thorium and uranium (Th/U) ratio was determined to be within the range of 3–6 (Supplemental Figure S4), which indicates that the effect of disequilibrium in the U decay chain on the dose rate is negligible.

Chronostratigraphy of the paleolake shorelines of Dali Lake

The quartz OSL dating results for the depositional sequences at different shoreline elevations are summarized in Figure 3, and we now describe the results for each profile.

Profile DL19-1 (H = 1231.1 m a.s.l.)

The chronology shows that the layer of littoral sediments was formed at ~0.7 ± 0.04 ka, indicating that the lake level reached 1231.1 m a.s.l. at that time. Below this layer are deep-water sediments deposited during the terminal Pleistocene–early Holocene. Although the deep-water sediments do not directly reflect the lake level, we infer that the lacustrine sediments represented by samples DL19-1-95 and DL19-1-120 belong to the deep-water lacustrine facies, as revealed by our modern process study of the relationship between the water depth and the grain size distribution of the surface lacustrine sediments of Dali Lake (Supplemental Figure S7). This indicates that the level of Dali Lake was much higher than the elevation of the deep-water sedimentary layer of samples DL19-1-95 and DL19-1-120 at that time.

Profiles DL19-2 (H = 1240.6 m a.s.l.) and DL19-4 (H = 1267.2 m a.s.l.)

Although the satellite image reveals obvious geomorphological characteristics of paleolake shorelines (Figure 2), the littoral sediments in these two cases were not preserved. This may have been the result of erosion by wave action or wind deflation. Despite this, these shorelines can extend up to several kilometers in length at almost the same elevation (Figure 2). The OSL dating results reveal that these two paleolake shorelines overlie deep-water sediments of early Holocene age. In addition, the grain size components of these lacustrine sediment samples are mainly clay and silt, indicative of a deep lake depositional environment (Supplemental Figure S7). Therefore, the level of Dali Lake was higher than the elevation of the lacustrine sedimentary layers of samples DL19-2-55 and DL19-4-115, at ~11.4 ± 0.64–11.1 ± 0.68 ka.

Profile DL19-3 (H = 1263.5 m a.s.l.)

The littoral sedimentary layer indicates that this paleolake shoreline was formed at ~5.6 ± 0.22–4.8 ± 0.21 ka, and the layer incorporates shell fragments (Supplemental Figure S2). The littoral sedimentary layer was covered by a soil layer and the age of ~3.1 ± 0.26 ka indicates that the lake level did not subsequently exceed this height. Beneath the littoral sedimentary layer there were late Pleistocene alluvial and deep-water sediments, for which the OSL ages are consistent within the error range (Figure 3).

Profile DL19-5 (H = 1267.3 m a.s.l.)

The age of the littoral sedimentary layer, which is intercalated in between two deep-water sediment layers, is ~12.8 ± 0.56 ka. The age of the upper deep-water sediment layer is not in stratigraphic order, but there is a slightly greater OD of the corresponding OSL sample (OD = 20.28 ± 4.30%); thus, it may have failed to achieve sufficient sunlight exposure before deposition and the age may be overestimated.

Profile DL19-6 (H = 1274 m a.s.l.)

The OSL ages are in stratigraphic order, and the depositional age of the base of the littoral sedimentary layer is like that of profile DL19-3, but their elevations differ by nearly 10 m. This implies that the lake level fluctuated frequently during this interval (5.6 ± 0.26–4.8 ± 0.21 ka). In addition, a littoral sedimentary layer was formed at ~4.6 ± 0.25 ka. The top of this profile is covered by a soil layer with an age of ~3.6 ± 0.21 ka, which indicates that the lake level did not subsequently exceed this height.

Profile DL19-7 (H = 1275 m a.s.l.)

A gravel layer is intercalated between an upper loess layer and the basal deep-water sediment layer. It is noteworthy that profile DL19-7 was covered by very thick loess layer, mainly due to the existence of a mountain (Zhenzi Mountain) within the vicinity of the profile, which blocked dust storms and promoted the accumulation of loess. The age of the loess layer is 4.5 ± 0.18 ka and that of the deep-water sediment layer is 11.5 ± 0.67 ka. The deep-water sediment layer contains shells (Supplemental Figure S2) and its grain size is slightly coarser (Supplemental Figure S7); this indicates deposition in a nearshore environment and that the lake level at that time was relatively close to the elevation of this layer, compared with the finer grained deep-water sediments of similar age in profiles DL19-1, DL19-2, and DL19-4.

Profiles DL19-8 (H = 1276.6 m a.s.l.) and DL19-10 (H = 1283.2 m a.s.l.)

The littoral sedimentary layer is homogeneous without bedding; the ages are 6.5 ± 0.31 ka and 6.1 ± 0.51 ka, respectively. There is a small quantity of gravel incorporated in the sediments of both profiles, which indicates that the lake level reached the elevation of these two paleolake shorelines during the mid-Holocene.

Profile DL19-9 (H = 1280.4 m a.s.l.)

The sedimentary facies characteristics and location of this profile indicate that the top layer (0–100 cm) was deposited in a wetland/floodplain environment during the gradual retreat of paleo–Dali Lake. Beneath is a deep-water sediment layer with the age of ~10.5 ± 0.67 ka. The grain size of these two layers is slightly coarser than the deep-water lacustrine muds (Supplemental Figure S7), which indicates that they were deposited in shallow water and thus that the lake level during the early Holocene was close to the elevation of the layer.

Profile DL19-11 (H = 1287.5 m a.s.l.)

A gravel layer is intercalated between the loess and a littoral sediment layer. Like profile DL19-7, loess accumulation was promoted by the existence of a mountain in the vicinity of the profile. The age of the loess layer is 6.0 ± 0.25 ka, indicating that the lake level did not exceed the elevation of this layer at this time. The OSL signal of the quartz from the littoral sedimentary layer was saturated as the result of the De values exceeding 2D₀ (D₀ = 197.30 ± 4.01 Gy), and its minimum age is estimated as 195.8 ± 9.97 ka, which is beyond both the age range of quartz OSL determinations at local dose rates and the scope of the present study.

Almost all the OSL ages of the paleolake shorelines of Dali Lake are in stratigraphic order. The shifts in sedimentary facies within the profiles clearly demonstrate the processes of lake level advance, retreat, and shoreline formation, all of which provide a basis for establishing a better constrained history of lake level fluctuations at Dali Lake since the last deglaciation.

Lake level fluctuations of Dali Lake since the last deglaciation

The reconstructed lake level record of Dali Lake obtained in this study is presented in Figure 6a. The results indicate that the lake level gradually rose since the beginning of the Bølling/Allerød (B/A) warming period (~15.0 ka), and that at ~12.8 ka the level was 43 m higher than today. The lake level may have experienced a low stage during the Younger Dryas (YD) cold period, and subsequently began to rise and reached a peak at ~10.4 ka, when the lake level was ~57 m higher than today. The lake level during the interval of 10.0–7.0 ka is unclear because no lakeshore relics were found at Dali Lake during this interval. The lake level reached its maximum since the last deglaciation during the mid-Holocene (~6.1 ka). Notably, when the level of Dali Lake exceeds 1285 m a.s.l., an outflow is generated and the water will flow out through the pass to the north of the catchment (Supplemental Figure S8), and therefore the actual level may have been higher than the reconstruction. After ~6.1 ka, the lake level continued to decrease and the elevation of the modern level is the lowest throughout the Holocene.

Figure 6.

Comparison of lake level records from the EASM margin proposed by different studies. (a) Variation in the lake level of Dali Lake proposed by this study (blue circles are beach ridges, green circles are deep-water sediments, brown circles are topsoil or loess). The lowest section of the Jin Wall, the elevation of Luwang City and the modern lake level are indicated by the red stars. (b) Lake level reconstruction of Dali Lake proposed by Goldsmith et al. (2017a) (triangles represent alluvial deposits, squares represent lacustrine sediments, and diamonds represent beach ridges). (c) Lake level reconstruction of Dali Lake proposed by Jiang et al. (2020). (d) Lake level reconstruction of Dali Lake proposed by Liu et al. (2016). (e) Integrated lake level record for Dali Lake. Circles are samples obtained in this study, triangles are the beach ridges from Jiang et al. (2020), squares are the beach ridges from Goldsmith et al. (2017a) (the open square represents a ¹⁴C date from charcoal, and the solid square represents an OSL date). (f) Lake depth reconstruction of Daihai Lake proposed by Sun et al. (2009). (g) Lake level reconstruction of Baijian Lake proposed by Long et al. (2012). (h) Lake level reconstruction of Qinghai Lake proposed by Liu et al. (2015b). (i) Lake level reconstruction of Chagan Nur Lake proposed by Li et al. (2020). In Figure 6, EH and MH represent the early Holocene and mid-Holocene, respectively; the black vertical dashed lines in curve Figure 6e represent two abrupt falls in the level of Dali Lake, at 5.5and 4.2 ka (the same hereinafter); the dark gray shaded area in the curve in Figure 6e indicates that the lake level of Dali Lake fluctuated throughout this interval. In the curve in Figure 6e, the error bars of the OSL chronology are represented by 1σ, and the error bars of the radiocarbon (¹⁴C) chronology are represented by 2σ.

Discussion

Comparison of records of lake level fluctuations for Dali Lake reconstructed by different studies

Although the history of lake level fluctuations at Dali Lake has been reconstructed in previous studies (e.g. Goldsmith et al., 2017a, 2017b; Jiang et al., 2020; Liu et al., 2016), the results have been intensively debated. To provide an unambiguous history of lake level fluctuations at Dali Lake, and to resolve the discrepancies between different studies, we compared our new lake level record with those obtained in previous studies of Dali Lake, reconstructed mainly by paleolake shorelines (Figure 6). The results of the systematic comparison are given below.

Stage 1: Last deglaciation (>11.7 ka). Comparison of elevation of the lake level at ~13.0 ka proposed by Goldsmith et al. (2017a) (Figure 6b) with our record (Figure 6a) shows that the elevation difference exceeds 10 m. This may be because the lacustrine sediments and alluvial deposits used by Goldsmith et al. (2017a) cannot directly indicate the lake level (Jiang et al., 2020). In addition, the minor highstand (Figure 6d) during the YD proposed by Liu et al. (2016) may be a chronological artifact produced by the reservoir effect and sediment re-deposition (Han and Li, 2020; Jiang et al., 2020).

Stage 2: ~11.7–8.0 ka. Although there were no OSL-dating data from beach ridges fall within this interval, based on the relationship between water depth and the grain size component of the surface lacustrine sediments (Supplemental Figure S7), the lacustrine sediments (DL19-7-590, DL19-9-70, DL19-9-150) from the paleolake shorelines can be generally regarded as nearshore sediments (with a water depth around 1–2 m). These lacustrine sediments reveal a highstand at Dali Lake at the beginning of the early Holocene (~11.7–10.0 ka), which conflicts with the relatively low level at Dali Lake during the early Holocene reconstructed by Jiang et al. (2020) (Figure 6c). This may be because the lake level record of Jiang et al. (2020) has only a single data point during 12.0–7.7 ka, which cannot reflect the lake level fluctuations during the early Holocene. Compared with our lake level record, the record of Goldsmith et al. (2017a) provides only qualitative evidence of a short-lived highstand at Dali Lake during the early Holocene, which is based on a series of samples of lacustrine sediments and alluvial deposits. Moreover, the lake level record of Liu et al. (2016) does not show a short-lived early Holocene highstand, which may be caused by the limited number of data points in the early Holocene, or by the influence of the reservoir effect and sediment redeposition (Han and Li, 2020; Jiang et al., 2020). During the interval of 10.0–8.0 ka, none of the studies of the paleolake shorelines of Dali Lake found lakeshore relics (Figure 6a–d).

Stage 3: 8.0–6.0 ka. Although the early Holocene lake level reconstruction at Dali Lake by Jiang et al. (2020) has been challenged (Goldsmith and Xu, 2020), Jiang et al. (2020) still reported a high elevation (1282.3 m a.s.l.) paleolake shoreline dated to ~6.0 ka. Moreover, the few beach ridge samples retrieved by Goldsmith et al. (2017a) also indicate that the highest level occurred in the mid-Holocene. Therefore, as shown in our record, there is robust evidence demonstrating that the level of Dali Lake reached a maximum during the mid-Holocene.

Stage 4: 6.0 ka–present-day. From ~6.0 ka onward, all the records of the lake level fluctuations for Dali Lake reconstructed by different studies show that the level of Dali Lake declined continuously (Figure 6a–d).

In order to produce a complete history of the lake level fluctuations at Dali Lake since the last deglaciation, we integrated published paleolake shoreline records with a reliable chronology and unambiguous lake level indicators with our new OSL-based chronology, with 21 age control points (Figure 6e). They include four paleolake shoreline records from Goldsmith et al. (2017a), and 13 paleolake shoreline records from Jiang et al. (2020). The integrated lake level record for Dali Lake shows a gradually increasing lake level during ~15.0–10.0 ka, a highstand existed during both the early Holocene and the mid-Holocene (8.0–6.0 ka), but the lake level was higher during the mid-Holocene than that during the early Holocene (~10.5 ka). Subsequently, after ~6.0 ka, the lake level gradually fell and the modern level is the lowest throughout the Holocene (Figure 6e).

It is important to note that the elevation of the integrated lake level record differs by ~20 m from that of the lake level from Jiang et al. (2020) at ~7.7 ka. This may be because the paleolake shoreline labeled R6 (with the age of 7.7 ka) was affected by reworking (Goldsmith and Xu, 2020), and thus the chronology of this paleolake shoreline may not be reliable. In addition, our integrated record has 13 age control points for the interval of 6.0–4.0 ka, which enables us to detect abrupt episodes of lake level change during this period. The integrated record shows two abrupt lake level declines, at ~5.5 ka and ~4.2 ka, respectively (Figure 6e); however, there are several differences between them. During the 5.5 ka event, the lake level fell at ~5.5 ka and then rapidly rebounded with frequent fluctuations (dark gray shaded area in Figure 6e). In contrast, during the 4.2 ka event, the level fell from 1273 to 1249 m a.s.l. within a short interval and failed to recover. It may be inappropriate to try to identify an abrupt lake level decline in Dali Lake at 4.2 ka using only two age data and more evidence is needed to confirm the authenticity of this event in future study. However, for the moment at least, we propose that an abrupt lake level decline at 4.2 ka in Dali Lake is possible for the following reasons. First, the chronology of sample DL19-6-99 was measured using two different aliquot sizes and both are dated to ~4.2 ka within the dating uncertainty (Supplemental Table S3). This ensures the reliability of the dating result. Second, two age data that defined the 4.2 ka lake level decline both are based on quartz OSL dating, which would reduce the dating uncertainty due to different dating methods. Also, the elevation difference of the paleolake shoreline (DL19-6-99; R8 (Jiang et al., 2020)) is ~30 m, and considering the unlikelihood of such a large elevation error, this elevation difference may reflect the rapid lake level decline at this stage. Third, numerous paleoclimatic records from northern China show that a significant aridification event occurred at ~4.2 ka, leading to abrupt changes in the regional hydrology (e.g. Liu and Feng, 2012; Ran and Chen, 2019; Sun et al., 2019; Xiao et al., 2018, 2019). The abrupt change of the lake level of Dali Lake during this period is therefore consistent with the pattern of regional climate change.

In summary, compared with results of previous studies, our integrated lake level record has more age control points and provides a more complete history of the lake level fluctuations at Dali Lake since the last deglaciation.

Comparison of the lake level fluctuations of Dali Lake since the last deglaciation with records from other lakes in the EASM margin

We now compare the lake level records from a large area within the EASM margin to differentiate between regional climate change and local non-climatic factors as possible causes of the lake level changes at Dali Lake, and thus to determine the possible forcing mechanisms.

Stage 1: Last deglaciation (>11.7 ka). A comparison of the lake level records from lakes in the EASM margin (see Figure 1 for the locations) shows that the pattern of lake level evolution of Dali Lake during the last deglaciation (~15.0–11.7 ka) is roughly representative of much of the EASM margin (Figure 6). It is noteworthy that, due to the limited influence of the EASM in the northeastern margin of the Qinghai-Tibet Plateau during the last deglaciation (An et al., 2012), the lake level of Qinghai Lake during this period was thought to be mainly controlled by meltwater from glaciers and permafrost (Liu et al., 2015b; Yan et al., 2020b). Thus, the pattern of lake level change of Qinghai Lake is somewhat inconsistent with that recorded at Dali Lake during the last deglaciation (Figure 6).

Stage 2: 11.7–8.0 ka. At the beginning of the early Holocene (~11.7–10.0 ka), the pattern of variation of lake level change at Dali Lake is inconsistent with that of other lakes in the EASM margin (Figure 6f–h). A conflict can also be seen between the lake level of Dali Lake and the qualitative chironomid-based lake water-depth reconstruction for Gonghai Lake in northern China (Wang et al., 2016) during the early Holocene. This suggests that the occurrence of a brief highstand at Dali Lake during ~11.7–10.0 ka was a local phenomenon, which can be attributed mainly to local non-climatic factors (see Section 5.3). During 10.0–8.0 ka, no lakeshore sediment deposition occurred at Dali Lake, which is consistent with paleolake shoreline records from other lakes in the EASM margin (Liu et al., 2015b; Long et al., 2012). This may indicate that the sediments of the major lake shorelines during this interval were destroyed by erosion caused by the frequent lake level fluctuations. In addition, climate proxy records from sediment cores from other lakes located on the northwestern margin of the EASM region also indicated that the water depth fluctuated frequently during the early Holocene (Yan et al., 2020b).

Stage 3: 8.0–6.0 ka. During the mid-Holocene (8.0–6.0 ka), a lake level maximum was recorded at Dali Lake, which is also consistently observed in other lakes in the EASM margin (Figure 6). The mid-Holocene lake level maximum is also supported by the water depth reconstruction for Gonghai Lake (Wang et al., 2016) and by climate proxy records from Kuhai Lake in the Northeast Tibetan Plateau (Wünnemann et al., 2018; Yan et al., 2018). Although a lake level reconstruction based on seven paleolake shorelines from Chagan Nur Lake, located close to the Dali Lake catchment, indicates that the lake level has gradually decreased since the beginning of the Holocene (Figure 6i), Li et al. (2020) proposed that this may have been because evidence of several millennial-timescale high lake level events during the mid-Holocene was destroyed by erosion. Therefore, the occurrence of a mid-Holocene lake level maximum is characteristic of the lakes in the EASM margin, clearly indicating a strong climatic control.

Stage 4: 6.0 ka–present-day. After ~6.0 ka, the level of Dali Lake and that of other lakes in the EASM margin all declined continuously. However, a rise in the level of Qinghai Lake at ~2.0 ka has been reported, which may reflect the influence of the westerly jet stream (Liu et al., 2015b).

Possible forcing mechanisms of lake level change in Dali Lake and its implications for a mid-Holocene EASM maximum

The lake level changes at Dali Lake were not controlled solely by monsoon precipitation, and the role of non-monsoon precipitation-related factors such as snow/ice meltwater input from the surrounding mountains may also have been significant (e.g. Liu et al., 2017a; Xiao et al., 2008, 2015; Yue et al., 2021). We now discuss the possible forcing mechanisms of lake level change in Dali Lake during different stages.

Stage 1: Last deglaciation (>11.7 ka). A comparison with records from other sites in the EASM margin shows a synchronous pattern of lake level variations, which reveals that the level of Dali Lake during the last deglaciation was controlled by regional climate change. With the rising temperature since the beginning of the BA warming period (Figure 7e), the precipitation in northern China would have increased (Figure 7c), which in turn caused a rise in the level of Dali Lake (Figure 7a). During the YD cold period, however, the Atlantic Meridional Overturning Circulation (AMOC) weakened (McManus et al., 2004; Praetorius et al., 2008) and temperatures decreased significantly (Figure 7e). This would have reduced the precipitation in northern China (Chen et al., 2015; Zhang et al., 2018a; Figure 7c), which led to a fall in the level of Dali Lake (Figure 7a).

Stage 2: 11.7–8.0 ka. Although a pollen-based precipitation record from Gonghai Lake shows an increase at the beginning of the early Holocene (~11.7–10.0 ka) (Figure 7c), the mean precipitation in this interval was still lower than that during the mid-Holocene. In addition, an integrated pollen-based quantitative precipitation record from northern China (Li et al., 2018) also suggests a relatively dry climate during ~11.7–10.0 ka, which is supported by the low paleosol density in the Loess Plateau at the same time (Wang et al., 2014). These observations further suggest that the highstand of Dali Lake at ~11.7–10.0 ka was mainly controlled by local factors superimposed on the monsoon precipitation. Two possible factors are discussed below. The first potential factor is vegetation cover. Vegetation cover plays an important role in the partitioning of precipitation into evapotranspiration and runoff at the catchment scale (Gan et al., 2021). The tree coverage in the Dali Lake catchment was relatively low during the early Holocene (Figure 7d). Evapotranspiration from the sparse vegetation was likely relatively low, which would have increased the runoff and caused a rise in the level of Dali Lake. The second potential factor is meltwater supply to the lake basin, which is based on the modern observation that the discharge of the Gongger River, a major inflow to Dali Lake, is greater during spring floods than during summer floods, owing to the meltwater runoff from snow/ice packs that cover the mountainous areas in the upper reaches of the river in winter and spring (Compilatory Commission of Annals of Hexigten Banner, 1993). Given that the Greater Khingan Mountains region may have been covered by thick snow during the last deglaciation (Pang et al., 2021), the snow/ice meltwater from the surrounding mountains, generated in response to the persistent warming in the EASM margin induced by the high Northern Hemisphere summer insolation during the early Holocene, may have made an important contribution to the lake level change in Dali Lake. This inference is supported by other studies (Fan et al., 2019; Li et al., 1990; Xiao et al., 2008; Yue et al., 2021). In addition, the low lake water temperature caused by the cold meltwater influx during the early Holocene (Yue et al., 2021) may also have reduced the evaporation of the lake water. Therefore, we propose that the highstand of Dali Lake during ~11.7–10.0 ka was not solely the result of enhanced EASM precipitation (Figure 7b–d) but was also the result of the influx of large amounts of snow/ice meltwater, together with increased catchment runoff caused by the absence of tree coverage and low lake evaporation. These additional factors may explain the seeming inconsistency between the highstand of Dali Lake and the relatively weak EASM during the early Holocene (Chen et al., 2015; Goldsmith and Xu, 2020; Goldsmith et al., 2017a, 2017b; Han and Li, 2020; Jiang et al., 2020; Liu et al., 2017a). Admittedly, more direct and valid evidence to support the early Holocene highstand of Dali Lake caused by snow/ice meltwater is still needed. A paleoclimate-vegetation-hydrological coupling model to quantify the relationship between precipitation and lake level change in the Dali Lake catchment could potentially resolve the issue in a future study. As for the frequent lake level fluctuations during 10.0–8.0 ka can be attributed to an unstable EASM climate during this interval (Yang et al., 2019; Zhang et al., 2021).

Stage 3: 8.0–6.0 ka. During the mid-Holocene (8.0–6.0 ka), the synchronous lake level maximum at Dali Lake and other lakes in the EASM margin indicates that the lake level during this interval was mainly controlled by regional monsoon precipitation. Although the mid-Holocene megathermal maximum (Shi et al., 1994), with high temperatures (Figure 7e), may have increased the rate of lake water evaporation (E) compared to the early Holocene, the level of Dali Lake exceeded that in the early Holocene and reached its maximum at this time (Figure 7a). This clearly indicates a positive water balance (P > E) and that the monsoon precipitation in the EASM margin increased significantly. The peak lake level at Dali Lake implies that the EASM reached its maximum during the mid-Holocene, which is supported by a peak in water balance model-based precipitation from Dali Lake (Li et al., 1990) and pollen-based precipitation from the EASM margin (Figure 7b and c), and by increased tree pollen percentages (Figure 7d). Furthermore, the high degree of paleosol development in the Chinese Loess Plateau at the time (Wang et al., 2014) strongly supports our inference of a lake level maximum at Dali Lake, and thus a mid-Holocene EASM maximum. Notably, there was a short interval of lake level decline at ~6.5 ka, which may be related to the mid-Holocene drought event reported in the semi-arid region of China (e.g. Chen et al., 2003; Jiang et al., 2006; Yin et al., 2013).

Stage 4: 6.0 ka–present-day. After 6.0 ka, although the temperature gradually decreased (Figure 7e), the significant decrease in precipitation in the EASM margin (Figure 7b and c) was so great that it may have caused a negative water balance (P < E) and thus the fall of the lake level of Dali Lake (Figure 7a). Against the background of gradually decreasing lake level since ~6.0 ka, the two prominent lake level falls, at ~5.5 ka and ~4.2 ka (Figure 7a), may be attributed to a change in the regional hydroclimate and catchment hydrology (Bai et al., 2017; Yang et al., 2015). At ~5.5 ka, the pollen-based quantitative precipitation reconstruction from Dali Lake shows a significant decrease (Figure 7b), which is consistent with a significant regional decrease in tree pollen percentages (Figure 7d). Although several paleoclimatic records did not register rapid climate changes during this period, likely due to the lack of resolution and sensitivity of the proxy records, several high-resolution and climatically sensitive East Asian stalagmite records (Liu et al., 2020) clearly documented the rapid retreat of the Asian summer monsoon at ~5.5 ka (e.g. Bai et al., 2017; Dong et al., 2010; Dykoski et al., 2005; Jiang et al., 2013). In addition, Magny and Haas (2004) integrated 44 multiproxy global climate records and they also detected a precipitation decrease during ~5.6–5.0 ka in the region of 20°–40°N, especially in the Asian monsoon region where the climate was predominantly cold and dry. During 5.5–4.2 ka the regional humidity increased slightly (Figure 7b–d) and the lake level of Dali Lake recovered somewhat (Figure 7a), which indicates that the lake system was able to recover from the event. The lake level fell at Dali Lake at ~4.2 ka may also have been caused by an abrupt and sustained decrease in monsoon precipitation in the EASM margin (Figure 7b–d). However, it is important to note that, although the decrease in monsoon precipitation at ~4.2 ka was similar to or even less than that at ~5.5 ka, there was still a large and irreversible fall (by ~30 m) in the level of Dali Lake within a short period of time (Figure 7a). This implies that an additional factor is needed to explain the magnitude and irreversibility of the lake level decline. Interestingly, Yang et al. (2015) proposed that at ~4.2 ka an abrupt hydrological shift in the catchment of Dali Lake occurred as a result of a large decrease in monsoon precipitation. This large-scale hydrological event breached multiple drainage divides, permanently diverting water resources into the east-flowing Xilamulun River (Yang et al., 2015). Based on hydrological analysis we have reconstructed the drainage network of Dali Lake prior to capture by the east-flowing Xilamulun River via groundwater sapping, as proposed by Yang et al. (2015). The results show that, before 4.2 ka, the catchment area of Dali Lake was ~5900 km² and that a channel connected Dali Lake with the modern Otindag Sandy Land, which is consistent with a relict channel observed in the field (Supplemental Figure S9). Additionally, based on a comprehensive multidisciplinary analysis of topography, tectonics, hydrogeology, groundwater dynamics, and stable isotopes, a recent study also proposed that a channel existed that connected Dali Lake and the modern Otindag Sandy Land (Li et al., 2021a). Therefore, we infer that against the background of rapid and widespread monsoon precipitation recession at ~4.2 ka, the steep decline in precipitation and the resulting groundwater capture by the Xilamulun River in the Otindag Sandy Land jointly led to the catastrophic and irreversible lake level decline in Dali Lake at ~4.2 ka.

Figure 7.

(a) Lake level fluctuations of Dali Lake since the last deglaciation. (b) Pollen-based precipitation reconstruction from Dali Lake (Xiao et al., 2015). (c) Pollen-based precipitation reconstruction from Gonghai Lake (Chen et al., 2015). (d) Tree pollen percentages from Dali Lake (Wen et al., 2017). (e) Synthesized Northern Hemisphere (30°–90°N) temperature record during the last deglaciation (Shakun et al., 2012) and the Holocene (Marcott et al., 2013).

Influence of the abrupt lake level decline on the development of prehistoric cultures in north-central China

Changes in ecological systems caused by large-scale and abrupt hydroclimatic fluctuations have a major impact on the growth and reproduction of animals and plants, and thus on the development of prehistoric cultures (Sun et al., 2019; Zhang et al., 2021), especially in northern China where the freshwater resources are limited (Liu et al., 2017b). Our results show for the first time that the lake level of Dali Lake decreased rapidly at ~5.5 and 4.2 ka, demonstrating the occurrence of two abrupt EASM-induced drought events. However, these two events were substantially different and are thought to have had different effects on the development of the regional prehistoric cultures.

Goldsmith et al. (2017a) suggested that the abrupt termination of the Holocene humid period at ~5.5 ka recorded at Dali Lake was responsible for the disappearance of the Hongshan culture in northern China. However, the reduction in precipitation in northern China at this time was insufficient to substantially affect the local dryland agricultural system, because the late Hongshan culture in northern China continued to develop until ~4.2 ka (Xia et al., 2000; Xu et al., 2019; Yang et al., 2015). At the same time, the scale of population migration from northern China to the south was relatively minor (Wu and Liu, 2002). We speculate that in the case of the 5.5 ka event, the relatively short-lived climatic fluctuations may in fact have increased the resilience of the local human population to this environmental change (Dong et al., 2021; Xu et al., 2017), and contributed to the evolution of the Hongshan culture which continued without major hiatuses.

In contrast, the abrupt and sustained decrease in monsoon precipitation at ~4.2 ka, and the resulting groundwater capture of the Xilamulun River in the Otindag Sandy Land, intensified the already relatively dry environment of the Dali Lake catchment, which underwent a catastrophic and irreversible change. The prehistoric inhabitants were unable to adapt to the much less favorable environment and were forced to respond by emigrating from the region. In an archeological investigation, Yang et al. (2015) determined that there was an absence of artifacts of the post-Hongshan culture in the Otindag Sandy Land after 4.2 ka. This temporal absence of archeological evidence suggests rapid abandonment and migration at the onset of desertification in the Otindag Sandy Land (Yang et al., 2015), which is also indicated by the absence of artifacts in the vicinity of the modern headwater of the Xilamulun River, from ~4.2 ka (Liu and Feng, 2012; Wagner et al., 2013). Therefore, the profound climatic deterioration in northern China, represented by the irreversible lake level decline of Dali Lake at 4.2 ka, is thought to have triggered the demise of the Hongshan culture in the region, which was displaced southward to the Yellow River Basin (Yang et al., 2015).

The widespread drought conditions in northern China caused by the recession of monsoon precipitation at ~4.2 ka (Xiao et al., 2018, 2019) likely also had a profound impact on the prehistoric cultures across a large area of north-central China (Figure 8). For example, Han (2017, 2020) suggested that the prehistoric cultures in northern China gradually moved southward from ~4.2 ka. Additionally, the clearest evidence of the southward migration of prehistoric cultures is the appearance in the Linfen Basin, in southern Shanxi Province, of large quantities of ceramics with typical characteristics of the Laohushan culture (Figure 8), which led to the transformation of the Taosi culture to the late Taosi culture (Han, 2017, 2020). Subsequently, the influence of the Laohushan culture extended southward, shown by the appearance of ceramics, microlithic arrowheads and oracle bones in the vast areas of the middle and lower reaches of the Yellow River, and the infiltration of the Wangwan Phase III culture by the Laohushan culture (Han, 2017, 2020; Figure 8). This cultural evidence points to a large-scale catastrophic and irreversible drought event at 4.2 ka in northern China, which likely exacerbated regional water shortages and reduced agricultural productivity. This would have rendered the region uninhabitable and forced the prehistoric population of northern China to migrate southward to western Henan Province (the central Plain of China) (Han, 2007).

Figure 8.

(a) Schematic diagram of the possible direction of prehistoric population migration. The illustrations of ceramic styles around the base map are representative of prehistoric cultures of the region, and the two stratigraphic sections indicate the occurrence of floods in eastern and western Henan Province, respectively (see Yu et al., 2020; Zhang and Xia, 2011 for details). (b) Lake level fluctuations of Dali Lake (this study) and wet-dry climatic index from Daihai Lake (Xiao et al., 2019), demonstrating a regional-scale drought event in northern China at ~4.2 ka. The blue and brown curves represent the frequency of extreme floods in the middle and lower reaches of the Yellow River, respectively (Yu et al., 2020).

Within the limits of the dating error, frequent floods occurred in central China at almost the same time as the catastrophic ~4.2 ka drought in northern China (Chen et al., 2020; Li and Gao, 2021; Li et al., 2021b; Yu et al., 2020; Zhang and Xia, 2011). For example, the lithology of the Erlitou section in the Luoyang Basin, western Henan province, revealed the occurrence of a major flood at ~4.0 ± 0.3–3.8 ± 0.3 ka, within the age uncertainty (Zhang and Xia, 2011) (Figure 8). Subsequently, a vast floodplain developed in the middle of the Luoyang Basin (Zhang and Xia, 2011), which comprised productive agricultural land and promoted cultural development (Wei, 2020). The favorable post-flooding environment would have been attractive to the southward-migrating population from northern China. The lithology of the Shilipu section in the lower Yellow River also revealed the occurrence of a major flood at ~3.8 ± 0.4 ka, within the age uncertainty. However, owing to the low-lying terrain in eastern Henan Province and the frequent shifting of the channel of the Yellow River (Figure 8), the frequent floods in eastern Henan province may have had a disastrous impact on the indigenous population (Chen et al., 2020; Cui and Zhou, 2003; Yu et al., 2020), forcing a migration westward to the higher altitude western Henan Province. For example, the Zaolvtai culture, which was distributed mainly in eastern Henan province, was affected by the floods (except for a component of the population that migrated southward to develop the Guangfulin culture (Zhao, 2014)) and migrated westward to high-elevation Zhengzhou, which is the modern capital city of Henan Province (Wei, 2021). This migrant population merged with the local Wangwan Phase III culture, forming the Xinzhai culture (Wei, 2021). Almost 79% of the characteristics of the ceramics of the Xinzhai culture were inherited from the Zaolvtai culture (Wei, 2018). Subsequently, the Erlitou culture, which belongs to the “Proto-Shang dynasty,” developed from the Xinzhai culture (Wei, 2021), which is supported by the fact that almost 78% of the characteristics of the ceramics of the Erlitou culture were inherited from the Xinzhai culture (Wei, 2018, Figure 8).

From the foregoing, we speculate that the climatic pattern of “drought in northern China, flooding in central China” at ~4.2 ka had a far-reaching impact on the prehistoric cultures in north-central China. That is, it caused a large influx of population to the central Plain of China, and the resulting increased regional population and competition for resources caused an increase in the frequency of wars. The consequence was a major reorganization of the Chinese civilization. The core position of the civilization in the central Plain of China was highlighted, and the Erlitou bronze cultures arose, which in turn may have affected the political situation in most parts of China, causing Chinese civilization to enter a more mature “Kingdom” stage (Han, 2010, 2020; Ran and Chen, 2019; Xu, 2009).

Conclusions

We have conducted an OSL dating–based study of the paleolake shorelines at Dali Lake. Comparison of this record with lake level records from elsewhere along the EASM margin and with EASM precipitation reconstructions, together with archeological evidence, leads to the following main conclusions.

(1) On the sub-orbital timescale, the lake level of Dali Lake rose from 14.5 ka onward, reached a maximum during the mid-Holocene (8.0–6.0 ka), and declined afterward.

(2) The established lake level record of Dali Lake indicates that the EASM intensity reached a maximum in the mid-Holocene (8.0–6.0 ka), while a short-lived early Holocene highstand of Dali Lake was likely caused by the influx of snow/ice meltwater, combined with enhanced monsoon precipitation, the weakening of runoff interception, and low lake evaporation.

(3) The climatic pattern of “drought in northern China, flooding in the central China” at ~4.2 ka forced the migration of the prehistoric population of the region to the central Plain of China, which promoted the development of early civilization in north-central China.

Supplemental Material

sj-docx-1-hol-10.1177_09596836221145362 – Supplemental material for Lake level evidence for a mid-Holocene East Asian summer monsoon maximum and the impact of an abrupt late-Holocene drought event on prehistoric cultures in north-central China

Supplemental material, sj-docx-1-hol-10.1177_09596836221145362 for Lake level evidence for a mid-Holocene East Asian summer monsoon maximum and the impact of an abrupt late-Holocene drought event on prehistoric cultures in north-central China by Zhiping Zhang, Zhongwei Shen, Shanjia Zhang, Jie Chen, Shengqian Chen, Dongxue Li, Shuai Zhang, Xiangjun Liu, Duo Wu, Yongwei Sheng, Qiuhong Tang, Fahu Chen and Jianbao Liu in The Holocene

Footnotes

Acknowledgements

We thank Leibin Wang, Zhong Wang, Xiaokang Liu for valuable suggestions; Jan Bloemendal for English improvement and discussion.

Funding

The author(s) disclosed receipt of the following financial support for the research, authorship, and/or publication of this article: This work was supported by the National Natural Science Foundation of China (41790421).

ORCID iD

Zhongwei Shen

Supplemental material

Supplemental material for this article is available online.

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