High-resolution climate change during the Marine Isotope Stage 3 revealed by Zhouqu loess in the eastern margin of the Tibetan Plateau

Abstract

A consensus has not yet been reached on effects of climate change and driving mechanisms between the Tibetan Plateau (TP) and adjacent monsoonal areas during the Marine Isotope Stage 3 (MIS 3). Loess–paleosol sequences from the TP provide valuable information about the MIS 3 environmental history. Detailed color index and a diffuse reflectance spectral (DRS) analysis of Zhouqu (ZQ) loess from the Western Qinling Mountains were conducted to investigate climate change on the eastern margin of the TP during the MIS 3. Our results show that the variations in color index and iron oxide content in ZQ loess are mainly caused by the pedogenesis and climate conditions. The lightness (L*) value and hematite (Hm) content were used to reconstruct the precipitation history and temperature changes, respectively. The reconstructed records revealed that climate change during the MIS 3 was characterized by high frequency and large amplitude, with millennial-scale changes superimposed on orbital-scale changes. Warm–humid climate occurred in the late MIS 3, while the early climate of MIS 3 was relatively cold–dry. The Indian summer monsoon (ISM) and temperature variations during the MIS 3 mainly occurred due to obliquity and precession. The North Atlantic cooling led to the southward movement of the Intertropical Convergence Zone, and the downstream cooling of the atmosphere by the westerly jet could result in events on a millennial-scale in the eastern margin of the TP. The interhemispheric forcing may play an important role in driving the strong summer monsoon in the late MIS 3.

Keywords

Loess soil color iron oxide diffuse reflectance spectral Indian summer monsoon temperature

Introduction

The Tibetan Plateau (TP) is a key part of the Asian monsoon system, and climate change and its driving mechanisms during the Marine Isotope Stage 3 (MIS 3) are highly debated (Dutt et al., 2015; Guan et al., 2012; Han et al., 2010; Pons-Branchu et al., 2010; Shi et al., 1999; Thompson et al., 1997; Voelker, 2002; Zhang et al., 2020). The isotopic records of the Guliya ice core from the western TP indicate that the climate in the late MIS 3 (40−30 ka) was warm–humid with a strong monsoon, and the precipitation and temperature were significantly higher than those at present (Shi et al., 1999; Shi and Zhao, 2009; Thompson et al., 1997). These conditions differed from the cold–dry climate in the Chinese Loess Plateau (CLP) and Europe (Andersen et al., 2004; Antoine et al., 2001; Feng et al., 2007; Guo et al., 2009; Kang et al., 2011; Zeeden et al., 2018). The high lake levels and pollen records of the TP and its surrounding areas as well as the transgressive records of eastern China also corroborate the warm–humid climate in the late MIS 3 (Li et al., 2008; Shi and Yu, 2003; Shi and Zhao, 2009; Wang et al., 2011; Wu et al., 2004). However, the age ranges of the TP ice cores are controversial, influencing the interpretation of the past climatic and environmental conditions at different timescales (Hou et al., 2018, 2021; Thompson et al., 2001; Tian et al., 2019; Zhong et al., 2020). Recent studies have shown that the period of maximum lake expansion may have occurred in the MIS 5 rather than MIS 3 (Lai et al., 2014; Yu et al., 2019; Zhang et al., 2020a). Furthermore, the extremely warm–humid climate in the late MIS 3 is controversial. The climatic evolution, regional differences, formation mechanisms, and the orbital-scale cycles in the TP during the MIS 3 are unclear due to the uncertainties in the age model and the differences among the various proxies and geologic records (Han et al., 2010; Liu et al., 2020; Li et al., 2008; Shi and Yu, 2003; Shi and Zhao, 2009; Hou et al., 2021; Tian et al., 2019; Li et al., 2020; Zhong et al., 2020; Yang et al., 2021a). Early research on loess in the Western Sichuan Plateau and the Northeastern TP (NETP) revealed that the region suffered numerous cold–warm temperature variations during the last glacial period (Fang et al., 1996; Jiang et al., 1997; Lv et al., 2004; Pan and Wang, 1999; Qiao et al., 2006). Research on loess in the Chengdu Basin showed that the Indian summer monsoon (ISM), which originates from the subtropical Mascarene High in the Southern Hemisphere, carries large precession signals from the equatorial and tropic zones, leading to the overall intensification of the MIS 3 climate signals (Han et al., 2010). The variations of the East Asian monsoon intensity largely control the humidity record of NETP loess (Li et al., 2008). A study on loess in southeastern TP stated that the ISM intensity at 38.7, 28.3, 25.9, and 23.5 ka was notably higher than that during the Holocene, including the present day (Feng et al., 2022). According to the Ganzi loess and Maerkang loess in the west Sichuan Plateau, the conditions during MIS 3 were humid due to the intensified ISM (Chen et al., 2022; Li et al., 2023; Yang et al., 2022). The TP loess studies have significantly advanced our understanding of the paleoenvironmental changes. However, more regional and high-resolution data with good age control records are urgently needed to address the above questions.

Loess from the TP are the key for understanding climate change in the region. For the TP, high-resolution loess records, comprising a good terrestrial archive (An et al., 1991; Liu, 1985), are still lacking. Recent investigations about the loess sequences in the TP and its surrounding areas have improved our understanding on the paleoenvironment evolution and paleoclimate changes since the last interglacial (Chen et al., 2022; Han et al., 2010; Liu et al., 2021a; Yang et al., 2021a, 2021b, 2022). Aeolian loess was deposited at a high accumulation rate and was sensitive to climate change in the Western Qinling Mountains (WQM), which is the transitional zone between the CLP and the TP. This provides an opportunity for better understanding climate change during the MIS 3 (Fang et al., 1999a, 1999b; Yang et al., 2021b; Zhang, 1997; Zhang et al., 1997). Therefore, the study of loess in the WQM is crucial for comprehending the evolutionary history of the Asian monsoon and the plateau environment since the Pleistocene.

Color index and diffuse reflectance spectral (DRS) analyses are valuable proxies for reconstructing the paleoenvironment history in loess–paleosol sequences. Soil color is one of the most significant characteristics of loess–paleosol sequences (Sun et al., 2011; Yang and Ding, 2003; Yang et al., 2001). Substances that significantly affect it include organic matter, calcium carbonate, hematite (Hm), and goethite (Gt), which are closely related to pedogenesis (Balsam et al., 2004; Gunal et al., 2008; Ji et al., 2004; Sun et al., 2011; Torrent et al., 1980). The alternating reddish brown and light yellow colors of the loess–paleosol sequence reflect the soil genesis characteristics under warm–humid and cold–dry climatic conditions, respectively (An et al., 1991; Liu, 1985). Thus, soil color effectively reflects climate change of the glacial/interglacial cycle. The soil color index has been proposed to be very sensitive to high-resolution and millennium-scale climate changes (Wang et al., 2016; Yang et al., 2001). Numerous researchers have used soil color to reconstruct the monsoon and paleoenvironment evolutions over various time scales (Sprafke et al., 2020; Thompson et al., 1997; Wang et al., 2016; Yang and Ding, 2003; Yang et al., 2001). The Hm and Gt content and proportion in soil are particularly sensitive to the climate, and Hm presents the potential for temperature reconstruction (Balsam et al., 2004; Cornell and Schwertmann, 2003; Ji et al., 2004; Jia et al., 2018). A DRS analysis can effectively yield the Hm and Gt contents and has been successfully used to reconstruct paleoclimates and paleoenvironments (Balsam et al., 2004; Ji et al., 2004; Jia et al., 2018; Wang et al., 2016).

In this study, using soil color and DRS measurements from the accurately dated Zhouqu (ZQ) loess sequence in the WQM, the sensitivity of the loess color proxies and DRS-derived iron oxides (Hm and Gt) to climate change is examined. Then, these proxy data are used to reconstruct high-resolution records of climate change in the MIS 3, and the forcing mechanism in the eastern TP is investigated.

Materials and methods

Materials

The ZQ section (33°46.74°N, 104°23.94°E, 2047 m) is located on the fifth terrace of the Bailong River in the WQM. The ZQ section is 16.3 m thick and has a gravel layer at the bottom. A 2.8-m-thick dark-brown Holocene soil (S0) layer is located at the top of the ZQ section, and the upper 0.5−1.5 m of soil is affected by agricultural activity. A brown paleosol layer (L1LS1, 8.2−13.6 m) exists in the middle part of the section. The two loess layers (L1LL1 and L1LL2) are pale yellow in color and have a loose texture. An age-depth model (Figure S2) of the ZQ section was constructed based on the published results of quartz optically stimulated luminescence ages and accelerator mass spectrometry ¹⁴C ages (Yang et al., 2021b). To obtain high-resolution proxy data, the 16.3-m-thick ZQ section was sampled in 2.5 cm increments for color and DRS analyses.

The climate of the ZQ area is mainly influenced by the ISM, the TP monsoon, and the westerlies (Figure 1(a)). Regional precipitation is mainly transported by the ISM (Fang et al., 1999b; Yang et al., 2010; Zhang et al., 2011). Figure 1(b) and 3S depict the summer circulation situation. The mean annual precipitation (MAP) in ZQ is 420.60 mm, with the majority of precipitation occurring from May to September (Figure 1(d)). The mean annual temperature (MAT) is 13.4°C, and the mean annual evaporation (MAE) is 1972.4 mm.

Figure 1.

Geographical location of the ZQ section and regional records, topography of TP and its surrounding atmospheric circulation (a), average wind field of 850h Pa from July to September from 1960 to 2010 based on reanalysis data of ERA5 (b), image of ZQ section (c), and meteorological information of ZQ section (d).

Methods

The CIE L*a*b* color, DRS, and magnetic susceptibility analyses were performed at the Key Laboratory of Western China’s Environmental Systems (Ministry of Education), Lanzhou University. A total of 652 samples were measured using a portable spectrophotometer (CM-700d, Minolta, Tokyo, Japan). An agate mortar was used to grind air-dried samples into powder. Subsequently, the spectrophotometer was used to obtain three groups of diffuse reflectance spectra and color values for each sample. The standard deviation of the instrument’s chromaticity value was less than △E*AB within 0.04 *. The reflectances for violet, blue, green, yellow, orange, and red (defined as Violet%, Blue%, Green%, Yellow%, Orange%, and Redness%, respectively) were calculated from the reflectance data (Ji et al., 2001; Ji et al., 2004). The Redness% was used to determine the Hm content (Long et al., 2011). The second-derivative spectrum of the Kubelka–Munk remission function at each wavelength was calculated in accordance with Torrent and Barrón (2003). The relative variations in the mass concentrations of Gt and Hm were determined using the band amplitudes at ∼425 and ∼535 nm, which are proportional to the Gt and Hm contents, respectively (Scheinost et al., 1998). Previous studies employed the Hm/(Hm+Gt) ratio as a climate index (Ji et al., 2004; Jia et al., 2018). Low frequency (470 Hz) susceptibility (χ_lf) was measured using a Bartington MS2 susceptibility meter (MS2, Bartington Instruments, Witney, UK) (Yang et al., 2021b).

To verify the influence of Earth’s orbital parameters on the climate, an ensemble empirical mode decomposition (EEMD) and power spectrum analysis were conducted for the ZQ loess. EEMD (Huang et al., 1998) directly extracts the energy related to time scales in nonlinear fluctuations and iteratively decomposes the original composite signal with multiple time scales into a series of intrinsic model function components. This avoids the arbitrariness of frequency-band selection in multi-scale research. The EEMD method was employed to separate the various oscillation components in the ZQ loess records. Then, the period of each component was analyzed using the REDFIT spectrum. Power spectrum analyses were performed using Past software (Hammer et al., 2001) and Redfit software (Schulz and Mudelsee, 2002). The rectangular window, one overlapping (50%) segment, and eight oversampling factors with 80% and 90% confidence intervals were used to analyze the data.

Results

Color characteristics

The lightness (L*) and redness (a*) of the ZQ section considerably fluctuated, and obvious differences existed in each stratum (Figure 2(a) and (b)). The L* value in the ZQ section had an average value of 55.11 and a range of 47.70–59.01. The average L* values in S0 and L1LS1 were 54.37 and 54.14, respectively, while the average L* values in L1LL1 and L1LL2 were 55.81 and 56.43, respectively. The L* of the ZQ section had a high amplitude, low value in the paleosol, and high value in the loess layer features. The average a* value in the ZQ section was 6.49, with a range of 5.24–8.62. The a* value was higher in L1LS1 and lower in S0. From L1LS1 to S0, a* gradually decreased from the bottom to the top.

Figure 2.

Variation in soil color L* (a), a* (b), Redness% (c), Hm (d), Gt (e) and Hm/(Hm+Gt) (f) in the ZQ section. Quartz optical stimulated luminescence ages and AMS ¹⁴C ages are from Yang et al. (2021b).

DRS characteristics

The reflectance of each color band was calculated for the different strata. The results showed that each color band’s reflectance substantially varied across the different strata (Figure S1 and Table S1). Compared to L1LL1, L1LL2, and L1LS1, the average Violet%, Blue%, and Green% values of S0 were significantly higher, while the average Yellow%, Orange%, and Redness% values of S0 were significantly lower. The differences in the average reflectance in each color band of L1LL1 and L1LL2 were relatively small. However, the average Violet%, Blue%, Yellow%, and Green% values of L1LS1 were lower than those of L1LL1 and L1LL2. The Orange% value of L1LS1 was higher than that of L1LL1 and slightly lower than of L1LL2. However, the Redness% value of L1LS1 was significantly higher than that of L1LL1, L1LL2, and S0. The reflectances of each color band for L1LS1 significantly differed from those for S0.

The samples from various strata in the ZQ section essentially exhibited the same second-derivative DRS curve shape (Figure 3). The characteristic valleys for Hm and Gt of the second-derivative DRS curve of the ZQ section were situated at ∼525 and ∼410 nm, respectively. However, the band amplitude at ∼525 and ∼410 nm was clearly different in various strata. The band amplitude at ∼410 nm was the strongest in L1LL1 and L1LS1 (Figure 3(b) and (c)), while the band amplitude at ∼525 nm was the strongest in L1LS 1 (Figure 3(c)).

Figure 3.

Second-derivative curves of DRS in S0 (a), L1LL1 (b), L1LS1 (c), and L1LL2 (d) of ZQ section.

Hm and Gt contents based on DRS

In Figure 2(c), the Redness% value varies from 30.84% to 34.13%, with an average value of 32.07%. The Redness% value exhibits strong fluctuations and high amplitudes throughout the strata. The Redness% value of L1LS1 is high, with an average value of 32.60%, while that of S0 is low. The average value of Hm is 9.85 × 10⁻⁴, with a range of 5.88–13.70 × 10⁻⁴. The Hm content in L1LS1 and S0 is higher than that in L1LL1 and L1LL2, with similar fluctuations and amplitudes to the Redness% value. The average value of Gt is 2.03 × 10⁻⁴, with a range of 1.12–3.84 × 10⁻⁴. The Gt content is relatively high in L1LL1 and L1LS1, and its fluctuations significantly differ from those of the Redness% value and Hm content. The average value of the Hm/(Hm+Gt) ratio is 0.17, and the variation range is 0.10–0.34. The Hm/(Hm+Gt) ratio is relatively high in L1LS1 and S0 and relatively low in L1LL1 and L1LL2, with a significant increase from L1LL2 to L1LS1.

Discussion

Color index and iron oxide, and their environmental significance

The production and transformation of iron oxides, the production and decomposition of organic matter, and the leaching and deposition of carbonates significantly impact the soil color during environmental changes and pedogenesis. The two primary substances believed to affect the L* value of loess–paleosol sequences are carbonate and organic matter, which are closely related to regional precipitation (Balsam et al., 1999; Fernandez et al., 1988; Gunal et al., 2008; He et al., 2010; Sun et al., 2011; Yang and Ding, 2003; Yang et al., 2001). Generally, an increase in the carbonate content significantly increases the L* value, while an increase in the organic matter content significantly darkens the soil (Hammer et al., 2001; Sun et al., 2011; Yang et al., 2001). A strong correlation has been observed between the L* value and precipitation in arid semi-humid areas (Miao et al., 2013; Yang et al., 2001). Furthermore, changes in the L* value in loess–paleosol studies have been widely used to reconstruct the precipitation history, and the L* value is sensitive to millennial scale and orbital climate oscillations (Chen et al., 2002; Wang et al., 2016; Yang et al., 2001). The sedimentary environment and process of WQM loess differ from those of the CLP or Central Asian loess. According to a previous study, the loess in the investigated area is of typical aeolian origin, which is consistent with the main characteristics of loess in the CLP (Zhang et al., 1997; Fang et al., 1999a, 1999b). In terms of material sources, the TP may have provided adequate dust resources for the WQM loess (Fang et al., 1999a, 1999b; Zhang et al., 1997). However, the variation of the L* value in the ZQ section was low in the paleosol layer and high in the loess layer. The low L* value in the paleosol mainly stemmed from the increase in the organic matter content via pedogenesis. The χ_lf value reflects the pedogenic ferromagnetic mineral content, which is controlled by monsoon precipitation, and it is often used as a proxy for precipitation in loess (An et al., 1991; Kukla and An, 1989; Maher, 2016; Schaetzl et al., 2018). The changes of the L* and χ_lf values in the strata were compared (Yang et al., 2021b) and were found to be similar (Figure 2(a) and (g)). The paleosol had high χ_lf and low L* values, and the loess layer had low χ_lf and high L* values. This suggests that magnetic minerals and organic matter enriched the paleosol due to the well-developed pedogenesis related to abundant monsoon precipitation. These results depict that precipitation and pedogenesis were the main influencing factors of the L* and χ_lf values in the WQM loess. Previous loess research on the southern foothills of the Qinling Mountains showed that the L* and χ_lf values well correspond, indicating their similar response to pedogenesis and climate change (Gao et al., 2015; Mao et al., 2017; Yang et al., 2018). The latest research on the soil color of topsoil in the eastern TP showed that the L* value can be utilized to determine the precipitation changes (Chen et al., 2022). Therefore, the changes in the L* value in the ZQ section could be used to determine the changes in the precipitation in the WQM loess. Additionally, the frequency and amplitude of changes in the L* value were more sensitive to precipitation than χ_lf, as previously reported (Yang et al., 2001).

The primary minerals responsible for the color of iron oxides are Hm and Gt, which are brownish red and bright yellow in color, respectively (Cornell and Schwertmann, 2003; Ji et al., 2004; Sun et al., 2011). Changes in the different color bands reflect the selective absorption and diffuse reflection of different substances in the soil. The a* and Redness% values and Hm content reflect the spectral properties of Hm, while the Gt content reflects the spectral properties of Gt (Cornell and Schwertmann, 2003; Singh and Gilkes, 1992). The results show that the Gt content variation is low in the paleosol and high in the loess (Figure 2(e)), which could be related to the complex production of Gt in the pedogenesis (Balsam et al., 2004; Ji et al., 2004). The variation in the a* and Redness% values and the Hm content in the ZQ section were consistent in the strata (Figure 2(b)–(d)); however, the a* and Redness% values and Hm content of S0 were lower than those of L1LL1 and L1LL2. Additionally, the reflectances of each color band for S0 were significantly different from those for the other layers (Table S1 and Figure S1). Pedogenesis resulted in the weathering of iron-rich silicates, such as chlorite, and the formation of iron oxide minerals in the CLP loess under warm, moist, summer monsoon conditions (Ji et al., 2004; Wang et al., 2016). However, the enrichment of organic matter and the transformation of iron oxides in the soil caused by the human activities (Kanu et al., 2014; Schulze et al., 1993; Zhu et al., 2020) could lead to deviations in the meaning of these indicators. Holocene human activity could have led to uncertainty in the indicative significance of the Hm content with respect to climate (Kanu et al., 2014; Schulze et al., 1993; Guan et al., 2023). The high a* and Redness% values and high Hm content in the L1LS1 paleosol signify that the Hm formation was mainly related to pedogenic and climatic factors.

Hm formation is more sensitive to temperature than precipitation, with high temperature and low-moisture conditions being most favorable (Cornell and Schwertmann, 2003; Schwertmann, 1985). The changes in the a* and Redness% values and Hm content were consistent with the L * and χ_lf values in L1LL1, but their changes were significantly different from the L * and χ_lf values in L1LS1. Compared to the 860−1150 cm layer, the 1150−1360 cm layer in L1LS1 exhibited a lower χ_lf value and higher L*, Redness% value, a* value, and Hm content. These results indicate that the period represented by the 860−1150 cm layer was more favorable for the formation of ferromagnetic minerals and organic matter than the formation of Hm, whereas the period represented by the 1150−1360 cm layer was more favorable for the formation of Hm than the formation of ferromagnetic minerals and organic matter. Previous studies (Ji et al., 2001) have shown that the Luochuan loess in the CLP comprise a low magnetite content and high Hm content during the transition from L2 to S1 (MIS 5e, high-temperature period), followed by a period with a high magnetite content and low Hm content. This reflects an alternation from a dry summer monsoon environment to a warm–humid summer monsoon environment (Ji et al., 2001; Sun et al., 2011), which explains the changes in the Hm and magnetite content in the ZQ section in L1LS1. Hm formation favorably occurs under high temperatures and a limited rainfall with the rapid release of iron under neutral to slightly acidic conditions (Schwertmann, 1985; Schwertmann and Murad, 1983). A short period of highly seasonal precipitation, <350 mm/yr, under high-temperature conditions is sufficient for forming significant quantities of Hm (Maher et al., 1998). The research from Luochuan loess and Lingtai loess also support the fact that climate conditions with large evaporation and low precipitation are conducive to Hm formation (Balsam et al., 2004). Hm formation reaches the maximum when the temperature is high and the period of soil wetting is too short to produce significant magnetite (Balsam et al., 2004). Previous research on topsoil has shown that pedogenic Hm formation is preferentially dependent on MAT (Gao et al., 2018). Contemporary meteorological data shows that the ZQ area (MAP: 420.60 mm, MAT: 13.4°C, MAE: 1972.4 mm) has a semi-arid environment, with low precipitation and high evaporation. This climatic background may have resulted in the response of Hm to temperature changes being greater in the ZQ section than the other areas. The abovementioned discussions show that the Hm content is sensitive to the temperature in the ZQ section.

Previous studies have shown that the Hm and Gt ratio is sensitive to climatic factors (Ji et al., 2004; Jia et al., 2018; Wang et al., 2016). The Hm/(Hm+Gt) ratio is very sensitive to changes in the dry and wet climate (Ji et al., 2001; Wang et al., 2016; Jia et al., 2018). The study results show that the Hm/(Hm+Gt) ratio exhibits relatively good consistency, with changes corresponding to the L* and χ_lf values, and that a good correlation exists between the changes in precipitation and the Hm/(Hm+Gt) ratio. However, the changes in Gt content and magnetic susceptibility in the ZQ section significantly differ, suggesting that Gt is not simply influenced by variations in precipitation. Furthermore, the difference between Hm and Hm/(Hm+Gt) during the MIS 3 demonstrate the significant contribution of Gt. Hm production is favored by the climatic conditions of high temperatures and low precipitation, while Gt formation is mainly related to low temperatures and high seasonal precipitation (Schwertmann and Taylor, 1989). Therefore, the Gt content could be mainly influenced by a combination of low temperatures and high seasonal precipitation according to Gt variation in the ZQ section. The Hm/(Hm+Gt) ratio is believed to be more susceptible to precipitation in the study area due to the complex response of Gt to the climate and pedogenic environment (Cornell and Schwertmann, 2003; Ji et al., 2004; Wang et al., 2016).

In summary, the L* value of the ZQ loess reflected the organic matter concentration during pedogenesis. The changes in the L* value and Hm/(Hm+Gt) ratio in the ZQ loess mainly reflected the precipitation changes; the L* value was more sensitive to precipitation than the Hm/(Hm+Gt) ratio, with a more significant frequency and amplitude of change (Figure 2(a) and (f)). The a* and Redness% values and Hm content of the ZQ loess primarily indicated that the Hm content was controlled by temperature. Compared to the a* value, the Redness% value and Hm content exhibited more high-frequency and high-amplitude changes with temperature (Figure 2(c) and (d)).

Evolution of the ISM and temperature changes during the MIS 3

The above described analysis showed that the L* value of the ZQ loess could be used for reconstructing the precipitation and ISM histories, while the a* and Redness% values and Hm content could be used for reconstructing the temperature variation in the WQM. The L* value was used to reconstruct the precipitation and ISM histories (Figure 4). The temperature change in the WQM during the MIS 3 was reconstructed based on the Redness% value and Hm content (Figure 5).

Figure 4.

Comparison L* of ZQ section to the LR04 benthic δ¹⁸O record (Lisiecki and Raymo, 2005) and the atmospheric carbon dioxide records (Bereiter et al., 2015) (a), stalagmite stable δ¹⁸O isotope in China (Cheng et al., 2016) (b), stalagmite stable δ¹⁸O isotope record in Xiaobailong Cave (Cai et al., 2015) (c), pollen record of sediments in Namco lake, central TP (Wu et al., 2004) (d), annual average precipitation (MAP) reconstructed by magnetic parameters of the Ganzi loess in the ETP (Chen et al., 2022) (e), the July insolation at 30°N in the Northern Hemisphere and Earth’s obliquity (Laskar et al., 2004) (f), NGRIP δ¹⁸O record of Greenland Ice (Andersen et al., 2004) (h), EDC δD record of Antarctie (Jouzel et al., 2007) and the mid-latitude southern hemispheric SST (Mashiotta et al., 1999) (i).

Figure 5.

Comparison Redness%, Hm of ZQ section to the LR04 benthic δ¹⁸O record (Lisiecki and Raymo, 2005) (a), high-field isothermal remanence (HIRM) record of Ganzi section (Chen et al., 2022) (b), pollen records of RM core (Shen et al., 2005) (c), temperature index of Tengchong Qinghai (TCQH) core (Zhang et al., 2020b) (d), air temperature record of Mangshan loess (Peterse et al., 2014) (e), the July insolation at 30°N in the Northern Hemisphere and Earth’s obliquity (Laskar et al., 2004) (f), NGRIP δ¹⁸O record of Greenland Ice (Andersen et al., 2004) (h), EDC δD record of Antarctie (Jouzel et al., 2007) and the mid-latitude southern hemispheric SST (Mashiotta et al., 1999) (i).

The ISM gradually strengthened from the early to late MIS 3, as revealed by the ZQ loess data, with abundant precipitation in the late MIS 3. During MIS 3 (57–29 ka), the L* value gradually increased, with low precipitation in the early MIS 3 (57–49 ka), moderate precipitation in the middle MIS 3 (49−39 ka), and high precipitation in the late MIS 3 (39–29 ka) (Figure 4(g)). This differs from the composited speleothem δ¹⁸O record from eastern China by Cheng et al. (2016) (Figure 4(b)). The high precipitation in the late MIS 3 in the WQM is similar to that reported in previous studies using the Guliya ice core (Shi et al., 1999; Shi and Yao, 2002; Shi and Yu, 2003). The ISM history recorded by stalagmites in the Xiaobailong Cave in southwest China signified relatively high rainfall in the middle to late MIS 3 (Cai et al., 2015) (Figure 4(c)). An investigation of the Namco Lake sediment revealed that the early MIS 3 was relatively dry and the late MIS 3 was warm–humid in the central TP (Wu et al., 2004) (Figure 4(d)). Moreover, the high lake levels recorded by the lake sediments of the TP indicated an abnormally warm–humid climate during the late MIS 3 (Li et al., 2008; Shi et al., 2001; Zheng and Xiang, 1989). Previous studies have also recorded high lake levels for the arid lakes of northwest China in the late MIS 3 (Pachur et al., 1995; Rhodes et al., 1996; Wang et al., 2011; Zhang et al., 2002). Recently, the characteristics of the lake-level elevation and lake-surface expansion in the late MIS 3 in the western regions of China were determined using an integrated analysis of 76 lake-level records in China (Zhang et al., 2019a). According to the Tengchong Qinghai (TCQH) lake sediment record in the southeast TP, the climate during the early and middle MIS 3 was relatively cold and semi-humid and that in the late MIS 3 there was cool–humid (Zhang et al., 2019b). In the mid-MIS 3, the TP and its surrounding areas experienced a significant ice advance, possibly due to the rise in abundant summer monsoon precipitation (Cui et al., 2018; Gillespie and Molnar, 1995; Owen et al., 2002; Phillips et al., 2000; Shi and Yao, 2002). A reconstructed precipitation history from the Ganzi loess showed that precipitation in the MIS 3 was relatively low in the early stage but relatively high in the late stage in the eastern TP (Yang et al., 2022) (Figure 4(e)). A comprehensive comparison and analysis of the geological records of the TP indicated a dry climate in the early MIS 3 and high precipitation in the late MIS 3, which was recorded by the ZQ loess sequence. Additionally, studies of the loess–paleosol sequence in the Hanjiang River Basin (Mao et al., 2017; Pang et al., 2017) and the loess–paleosol sequence in the Sichuan Basin (Han et al., 2010) have shown that the climate in the late MIS 3 was unusually warm and humid. The geological records of the East Asian monsoon region revealed an increase in precipitation and forest cover as well as a rise in and transgression of the sea levels during the MIS 3 (Qiu et al., 2012; Shi et al., 2001; Tang, 1992; Yang et al., 2004). These records verify the humid climate of the MIS 3 recorded by the ZQ loess in the WQM.

The temperature of the MIS 3 in the WQM experienced significant changes at different stages. The Redness% value and Hm content of the ZQ section indicated a mild temperature in the early MIS 3 and a high temperature in the middle and late MIS 3 (Figure 5(g)). The temperature changes reflected by the Redness% value and Hm content of the ZQ section were basically consistent with the temperature changes of the MIS 3 based on the pollen record of the TCQH sediment and the RM core (Shen et al., 2005; Zhang et al., 2020b) (Figure 5(c) and (d)). The temperature reconstructed by the Mangshan loess (Peterse et al., 2014) also revealed a relatively high temperature in the middle MIS 3 (Figure 5(e)). The temperature change recorded by the high field isothermal remanence in the Ganzi loess in the eastern TP also indicated a relatively high temperature in the middle and late MIS 3 (Chen et al., 2022) (Figure 5(b)). These results support the climate characteristics of the MIS 3 recorded by the ZQ section.

The climate of the WQM in the late MIS 3 was relatively warm–humid. Compared to the MIS 2 (29–14 ka), the climate of the middle and late MIS 3 was warm and wet. The L* and Redness% values and Hm content showed a significant decrease and weak fluctuation during the MIS 2 (Figures 4 and 5), indicating a weakened ISM with a dry–cold climate during the MIS 2. The early climate of MIS 3 in the WQM recorded by the ZQ loess was relatively cold–dry, similar to MIS 2. During the Holocene, the reduced L* value signified increased precipitation and pedogenesis caused by the enhanced ISM (Figures 4 and 5), and the late MIS 3 precipitation level was similar to that of the Holocene. The reconstructed precipitation of the Ganzi loess in the MIS 1 was close to that in the late MIS 3 (Yang et al., 2022). The pollen record from the RM core suggested that the climate of the late MIS 3 was relatively warm–humid (Shen et al., 2005). The degree of warm–humid conditions in the late MIS 3 recorded by the loess–paleosol sequence in the Sichuan Basin was higher than that of the MIS 1 and similar to that of the MIS 5, but the S0 soil could have been disturbed by agricultural activities (Han et al., 2010). Climate change revealed by the ZQ section during the MIS 3 is largely consistent with the records of the TP and its surrounding areas; however, the relatively warm–wet climate in the MIS 3 recorded by the ZQ loess significantly differs from that of the CLP loess records (Guo et al., 2009; Kang et al., 2011) and European loess records (Antoine et al., 2009; Rousseau et al., 2007; Schirmer, 2000; Zeeden et al., 2018). The European region and the CLP had a relatively dry–cold climate during the MIS 3, reflecting the synchronous cooling in middle- and high-latitude areas with an increased global ice volume (Antoine et al., 2009; Guo et al., 2009; Kang et al., 2011; Rousseau et al., 2007; Zeeden et al., 2018). The differences in the climate evolution during the MIS 3 between the TP and European regions and the CLP region may be attributed to the control exerted by the different regional climate systems.

Furthermore, the ZQ loess revealed the millennial-scale changes during the MIS 3. Previous studies have shown that the climate in the MIS 3 was characterized by high instability and rapid fluctuations (Andersen et al., 2004; Bond et al., 1993; Cheng et al., 2016; Dansgaard et al., 1993; Heinrich, 1988; Sun et al., 2021). Moreover, significant worldwide differences existed in climate change amplitude during the MIS 3 (Han et al., 2010; Wetterich et al., 2014; Yu et al., 2007). The frequency and amplitude of the L*and Redness% values and Hm content recorded by the ZQ loess from the middle to late MIS 3 were very high (Figures 4 and 5), indicating that the climate rapidly fluctuated and exhibited unstable characteristics. A comparative study was performed between the L* and Redness% values and Hm content of the ZQ loess with the North Greenland Ice Core Project (NGRIP) oxygen isotope record (Andersen et al., 2004). The precipitation and temperature changes denoted by the ZQ loess were comparable with those denoted by the Heinrich and Dansgarrd–Oeschger (D–O) events (Bond et al., 1993; Dansgaard et al., 1993; Heinrich, 1988) in the NGRIP record (Figures 4 and 5). The rapid decrease in precipitation and temperature during the MIS 3 at ∼29, ∼39.5, and ∼47.5 ka well correspond to the H3, H4, and H5 Heinrich layers, respectively. Additionally, a series of rapid increases in precipitation and temperature coincided with the D–O events (Figures 4 and 5). This indicates that the WQM loess also recorded millennial-scale changes in the North Atlantic, which is consistent with the records from the TP and its surrounding areas (Feng et al., 2022; Lv et al., 2004; Zhao et al., 2020).

In summary, the results showed that the changes in the ISM dominated the precipitation history in the eastern TP. The early MIS 3 had low precipitation and temperature, and the mid-MIS 3 had high temperature and moderate precipitation, with a relatively warm–dry summer monsoon environment (Ji et al., 2001; Sun et al., 2011). A peak in precipitation occurred in the late MIS 3, with the relatively high temperature resulting in a warm–humid environment. The climate in the WQM comprised three stages from the early to late MIS3: cold–dry, warm–dry, and warm–humid. The ZQ loess clearly revealed that millennial-scale oscillations were present in the precipitation and temperature of the MIS 3.

Factors controlling climate change during the MIS 3

The changes in the ISM and temperature in the WQM were mainly driven by the combined effects of obliquity and precession. According the EEMD results and L* spectrum, the C7 component contributed 25.83% to the total variance (Figure 6(g)), and its power spectrum reflected the ∼40 ka cycle. This indicates that the obliquity cycle significantly impacted the ISM and precipitation changes. Additionally, the C5 component accounted for 20.24% of the total variance, and the power spectrum displayed a half-precession cycle (∼11.5 ka), precession (∼20 ka), and an ∼8.3 ka cycle. The ∼8.3 ka cycle could be related to the half-precession cycle (Becker, 2005; Berger and Loutre, 1997; Jian et al., 2020; Nascimento et al., 2021; Ulfers et al., 2022). Many records have demonstrated that the occurrence of half-precession is an important feature in tropical regions in terms of the amplification response of interhemispheric low-latitude insolation (Beaufort et al., 2003; Berger and Loutre, 1997; Ferretti et al., 2015; Jian et al., 2022). These results suggested that the ISM was controlled by obliquity and precession. This was also supported by the comparison of the L* values with the obliquity and precession changes (Figure 4(f) and (g)). However, the obliquity signals were stronger than the precession signals in the precipitation recorded by the ZQ loess. This is similar to the records of Asian summer monsoon precipitation (Clemens et al., 2021; Yang et al., 2022). However, it is different from the dominant ∼19–23 ka cycles recorded by the stalagmite δ¹⁸O in the Asian summer monsoon region (Cheng et al., 2016, 2021; Wang, et al., 2008), which may have stemmed from the sensitivity of different carriers and indicators (Cheng et al., 2021; Sun et al., 2021).

Figure 6.

EEMD component of L*, Hm and corresponding power spectrum in the ZQ section. The blue and red dashed lines indicate 80% and 90% levels of significance, respectively.

The EEMD results and corresponding power spectrum also showed that the temperature was affected by obliquity and precession (Figure 6(p) and (o)). The C7 component (Figure 6(p); 22.02% variance contribution) primarily reflected an obliquity cycle (∼41 ka). Figure 5(e) and (g) shows that the recorded temperature is highly consistent with the obliquity changes during the glacial stage. A recent study on the TCQH sediment stated that warm periods were probably dependent on the local annual mean insolation, controlled by Earth’s ∼41 ka obliquity cycles (Zhao et al., 2021). Obliquity may have influenced the temperature change in the WQM via changes in the local annual average insolation during the glacier stage. However, the C6 component (Figure 6(o); 19.43% variance contribution) had a strong precession cycle (∼21 ka). These findings confirm that the temperature in the WQM was influenced by insolation changes that were regulated by the obliquity and precession.

The correlation period of the millennium-scale oscillation of the North Atlantic was recorded by the ZQ loess. The power spectrum results showed that the C3 and C4 components of L* (Figure 6(c) and (d); 14.98% variance contribution) and Hm (Figure 6(l) and m; 9.72% variance contribution) support the existence of millennium-scale events. The cycles from the C3 and C4 components could be divided into two main groups: ∼7–4 ka (such as 7.5, 5.7, and 4.1 ka) and ∼3–1 ka (such as 3 and 1.5 ka) from the power spectrum. These results may correspond to the Heinrich (∼6 ka) and D–O (∼1.5 ka) cycles recorded in the North Atlantic sediments and Greenland ice core (Bond et al., 1993; Dansgaard et al., 1993; Heinrich, 1988), which correspond to the CLP loess and speleothem records (Li et al., 2015; Sun et al., 2021). The driving force of cold events on the millennial scale was associated with the weakening of the Atlantic meridional overturning circulation (AMOC), which was caused by an abrupt freshwater outflow into the North Atlantic from the Laurentide Ice Sheet (Carlson et al., 2007; McManus et al., 2004). Cold events on a millennial scale have also been recorded in the Indian Ocean sediments (Liu et al., 2021b; Mohtadi et al., 2014; Rashid et al., 2011). Furthermore, the cooling of the North Atlantic weakens the ISM by pushing the thermal equator (ITCZ) southwards through an atmospheric tele-connection and the AMOC (Arbuszewski et al., 2013; Liu et al., 2021b; Mohtadi et al., 2014; Panmei et al., 2017; Wen et al., 2016). The North Atlantic climate signal may be transmitted to East Asia through the westerlies (Song et al., 2018). A recent study suggested that the cooling of the North Atlantic led to the downstream cooling of the atmosphere by the westerly jets, which subsequently influenced the climate in the Qaidam Basin (Meng and Liu, 2018). These results suggest that cold events on a millennial scale in the WQM could have been driven by the North Atlantic cooling through the southward movement of the ITCZ and the downstream cooling of the atmosphere by the westerlies.

The results confirm that the warm–humid climate in the late MIS 3 was similar to that of the interglacial period. This could be related to the integrated response of the ISM changes due to the Earth’s orbital parameters and internal climate forcing (Clemens et al., 1991, 2021; Clemens and Prell, 2003; Kim and Wang, 2011; McGrath et al., 2021). The ISM index from the Heqing core in the southeast TP revealed that cooling in the high latitudes in the Southern Hemisphere (SH) increased the cross-equatorial pressure gradient and the ISM, which was manifested as the enhancement of the ISM with decreasing Antarctic temperatures and increasing ice volume after reaching a particular stage of the glacial period (An et al., 2011). A comparison of the LR04 Benthic δ¹⁸O record and atmospheric carbon dioxide (CO₂) record with the L* values of the ZQ loess (Figure 4(a) and (g)) revealed that the ISM gradually increased with increasing global ice volume and decreasing atmospheric CO₂ concentration during the MIS 3. A comparison of the Antarctic EPICA Dome C (EDC) δD record and the mid-latitude southern hemispheric sea surface temperature (SST) with the L* values of the ZQ loess also revealed that the ISM gradually increased with increasing Antarctic ice volume (Figure 4(i); blue) and decreasing SH SST (Figure 4(i); gray) during the MIS 3. This is similar to the findings of An et al. (2011). Additionally, the power spectrum of L* in the C6 component (Figure 6(f)) showed a ∼29 ka cycle. The ∼29 ka cycle was also present in the drill records of the Heqing Basin, where it was considered to arise from the combined influence of the Northern Hemisphere (NH) ice volume and SH temperature change (An et al., 2011). The interhemispheric forcing played an important role in driving the ISM enhancement during the MIS 3 in the WQM region. Hence, from the early to late MIS 3, the enhancement of the Mascarene high in the SH could have been caused by the decrease in the atmospheric CO₂ concentration and temperature in the SH as well as the increase in the ice volume of the Antarctic, resulting in a synchronous enhancement of the cross-equatorial pressure gradient and the ISM. Furthermore, the high insolation in the middle and late MIS 3 promoted the strengthening of the Indian low and ocean evaporation remained relatively strong. Subsequently, more warm–humid water vapor was transported to the land in South Asia, further strengthening the ISM and leading to the warm–humid climate in the late MIS 3.

In summary, the ISM and temperature in the WQM were driven by Earth’s obliquity and precession. The changes in the ISM and temperature exhibited millennial-scale climate variability during the MIS 3 in the WQM, indicating a profound impact of the high-latitude climate system of the NH. The warm–humid climate in the late MIS 3 may have jointly stemmed from obliquity, precession, and interhemispheric forcing.

Conclusion

This study analyzed the color index and DRS-derived iron oxides of the ZQ loess sequence in the eastern margin of the TP and their paleoenvironmental significance. The following conclusions were reached.

1. The soil color and iron oxide variations of the ZQ loess revealed that climate change in the MIS 3 was characterized by high frequency and large amplitude of variations as well as millennial-scale changes superimposed on orbital-scale changes in the eastern TP. The early climate of MIS 3 was relatively cold–dry, similar to the glacial time. From the early to late MIS 3, the climate in the eastern margin of the TP experienced three stages: cold–dry, warm–dry, and warm–humid.

2. The change in the ISM and temperature in the WQM was dominantly influenced by obliquity and precession. The millennial-scale changes of precipitation and temperature were mainly affected by the high-latitude climate system of the NH, the southward movement of the ITCZ, and the downstream cooling of the atmosphere by the westerlies driven by the North Atlantic cooling. The joint influence of obliquity, precession, and interhemispheric forcing may have resulted in a warm–humid climate in the late MIS 3.

Supplemental Material

Supplemental Material - High-resolution climate change during the Marine Isotope Stage 3 revealed by Zhouqu loess in the eastern margin of the Tibetan Plateau

Supplemental Material for High-resolution climate change during the MIS 3 revealed by Zhouqu loess in the eastern margin of the Tibetan Plateau by Zixuan Chen, Qiong Li, Pushuang Li, Jiantao Zhou, Yating Su, Weiming Liu, Yuanlong Luo, Chen Wen, Xuechao Xu and Shengli Yang in Progress in Physical Geography: Earth and Environment

Footnotes

Declaration of conflicting interests

The author(s) declared no potential conflicts of interest with respect to the research, authorship, and/or publication of this article.

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 Second Tibetan Plateau Scientific Expedition and Research Program (STEP) (2019QZKK0602); National Natural Science Foundation of China (42271159, 41877447), and Strategic Priority Research Program of the Chinese Academy of Sciences (XDA20090000).

ORCID iD

Zixuan Chen

Supplemental Material

Supplemental material for this article is available online.

References

Kukla

Porter

, et al. (1991) Magnetic-susceptibility evidence of monsoon variation on the Loess Plateau of central China during the last 130,000 years. Quaternary Research 36: 29–36.

Clemens

Shen

, et al. (2011) Glacial-interglacial Indian summer monsoon dynamics. Science 333: 719–723.

Andersen

Azuma

Barnola

, et al. (2004) High-resolution record of Northern Hemisphere climate extending into the last interglacial period. Nature 431: 147–151.

Antoine

Rousseau

D-D

Zöller

, et al. (2001) High-resolution record of the last Interglacial–glacial cycle in the Nussloch loess–palaeosol sequences, Upper Rhine Area, Germany. Quaternary International 76–77: 211–229.

Antoine

Rousseau

D-D

Moine

, et al. (2009) Rapid and cyclic aeolian deposition during the Last Glacial in European loess: a high-resolution record from Nussloch, Germany. Quaternary Science Reviews 28: 2955–2973.

Arbuszewski

deMenocal

Cléroux

, et al. (2013) Meridional shifts of the atlantic intertropical convergence zone since the last glacial maximum. Nature Geoscience 6: 959–962.

Balsam

Deaton

Damuth

(1999) Evaluating optical lightness as a proxy for carbonate content in marine sediment cores. Marine Geology 161: 141–153.

Balsam

Chen

(2004) Climatic interpretation of the Luochuan and Lingtai loess sections, China, based on changing iron oxide mineralogy and magnetic susceptibility. Earth and Planetary Science Letters 223: 335–348.

Beaufort

de Garidel-Thoron

Linsley

, et al. (2003) Biomass burning and oceanic primary production estimates in the Sulu Sea area over the last 380 kyr and the East Asian monsoon dynamics. Marine Geology 201: 53–65.

10.

Becker

(2005) Late Pliocene Millennial to Milankovitch-Scale Climate Variability: A Case Study of Marine Isotope Stages 101-95 in the Mediterranean and Adjacent North Atlantic. Geologica Ultraiectina, 1–164.

11.

Bereiter

Eggleston

Schmitt

, et al. (2015) Revision of the EPICA Dome C CO2 record from 800 to 600 kyr before present. Geophysical Research Letters 42: 542–549.

12.

Berger

Loutre

(1997) Intertropical latitudes and precessional and half-precessional cycles. Science 278: 1476–1478.

13.

Bond

Broecker

Johnsen

, et al. (1993) Correlations between climate records from north-atlantic sediments and Greenland ice. Nature 365: 143–147.

14.

Cai

Fung

Edwards

, et al. (2015) Variability of stalagmite-inferred Indian monsoon precipitation over the past 252,000 y. Proceedings of the National Academy of Sciences of the United States of America 112: 2954–2959.

15.

Carlson

Clark

Haley

, et al. (2007) Geochemical proxies of North American freshwater routing during the Younger Dryas cold event. Proceedings of the National Academy of Sciences of the United States of America 104: 6556–6561.

16.

Chen

Balsam

, et al. (2002) Characterization of the Chinese loess–paleosol stratigraphy by whiteness measurement. Palaeogeography, Palaeoclimatology, Palaeoecology 183: 287–297.

17.

Chen

Yang

Luo

, et al. (2022) HIRM variation in the Ganzi loess of the eastern Tibetan Plateau since the last interglacial period and its paleotemperature implications for the source region. Gondwana Research 101: 233–242.

18.

Cheng

Edwards

Sinha

, et al. (2016) The Asian monsoon over the past 640,000 years and ice age terminations. Nature 534: 640–646.

19.

Cheng

Zhang

Cai

, et al. (2021) Orbital-scale Asian summer monsoon variations: Paradox and exploration. Science China Earth Sciences 64: 529–544.

20.

Clemens

Prell

(2003) A 350,000 year summer-monsoon multi-proxy stack from the owen ridge, northern arabian sea. Marine Geology 201: 35–51.

21.

Clemens

Prell

Murray

, et al. (1991) Forcing mechanisms of the Indian-ocean monsoon. Nature 353: 720–725.

22.

Clemens

Yamamoto

Thirumalai

, et al. (2021) Remote and local drivers of Pleistocene South Asian summer monsoon precipitation: a test for future predictions. Science Advances 7: eabg3848.

23.

Cornell

Schwertmann

(2003) The Iron Oxides: Structure, Preperties, Reactions, Occurrences and Uses. Weinheim: VCH Verlagsgesellschaft Press, 1–703.

24.

Cui

Wang

, et al. (2018) Marine Isotope Stage 3 paleotemperature inferred from reconstructing the Die Shan ice cap, northeastern Tibetan Plateau. Quaternary Research 89: 494–504.

25.

Dansgaard

Johnsen

Clausen

, et al. (1993) Evidence for general instability of past climate from a 250-kyr ice-core record. Nature 364: 218–220.

26.

Dutt

Gupta

Clemens

, et al. (2015) Abrupt changes in Indian summer monsoon strength during 33,800 to 5500years BP. Geophysical Research Letters 42: 5526–5532.

27.

Fang

Chen

Shi

(1996) Ganzi loess and maximum glaciation on the Qinghai-Xizang (Tibetan) Plateau. Chin. Sci. Bull 41: 1865–1867.

28.

Fang

Van der Voo

(1999a) Age and provenance of loess in west qinling. Chinese Science Bulletin 44: 2188–2192.

29.

Fang

Van der Voo

(1999b) Rock magnetic and grain size evidence for intensified Asian atmospheric circulation since 800,000 years BP Related to Tibetan uplift. Earth and Planetary Science Letters 165: 129–144.

30.

Feng

Tang

, et al. (2007) Vegetation variations and associated environmental changes during marine isotope stage 3 in the western part of the Chinese Loess Plateau. Palaeogeography, Palaeoclimatology, Palaeoecology 246: 278–291.

31.

Feng

Wang

, et al. (2022) Amplified and suppressed regional imprints of global warming events on the southeastern Tibetan Plateau during MIS 3-2. Quaternary Science Reviews 294: 107736.

32.

Fernandez

Schulze

Coffin

, et al. (1988) Color, organic-matter, and pesticide adsorption relationships in a soil landscape. Soil Science Society of America Journal 52: 1023–1026.

33.

Ferretti

Crowhurst

Naafs

BDA

, et al. (2015) The marine isotope stage 19 in the mid-latitude North Atlantic ocean: astronomical signature and intra-interglacial variability. Quaternary Science Reviews 108: 95–110.

34.

Gao

Pang

Huang

, et al. (2015) Chroma characteristics and its significances of the Chafangcun loess-paleosol profile in Southeast Shaanxi, China (in Chinese). Acta Sedimentologica Sinica 33: 537–542.

35.

Gao

Hao

Wang

, et al. (2018) The different climatic response of pedogenic hematite and ferrimagnetic minerals: evidence from particle-sized modern soils over the Chinese Loess Plateau. Quaternary Science Reviews 179: 69–86.

36.

Gillespie

Molnar

(1995) Asynchronous maximum advances of mountain and continental glaciers. Reviews of Geophysics 33: 311–364.

37.

Guan

Gao

, et al. (2012) Modern human behaviors during the late stage of the MIS3 and the broad spectrum revolution: evidence from a Shuidonggou Late Paleolithic site. Chinese Science Bulletin 57: 379–386.

38.

Guan

, et al. (2023) Research on the cross-coloration effect of iron oxides and humus in soil. Acta Pedologica Sinica 60(1): 138–150.

39.

Gunal

Ersahin

Yetgin

, et al. (2008) Use of chromameter-measured color parameters in estimating color-related soil variables. Communications in Soil Science and Plant Analysis 39: 726–740.

40.

Guo

Berger

Yin

, et al. (2009) Strong asymmetry of hemispheric climates during MIS-13 inferred from correlating China loess and Antarctica ice records. Climate of the Past 5: 21–31.

41.

Hammer

Harper

DAT

Ryan

(2001) PAST: paleontological statistical software package for education and data analysis. Palaeontologia Electronica 4: 1–9.

42.

Han

Fang

Yang

, et al. (2010) Differences between East Asian and Indian monsoon climate records during MIS3 attributed to differences in their driving mechanisms: evidence from the loess record in the Sichuan basin, southwestern China and other continental and marine climate records. Quaternary International 218: 94–103.

43.

Sun

(2010) Changing color of Chinese loess: controlling factors and paleocliamtic significances (in Chinese). Geochimica 29: 447–455.

44.

Heinrich

(1988) Origin and consequences of cyclic ice rafting in the northeast atlantic-ocean during the past 130,000 years. Quaternary Research 29: 142–152.

45.

Hou

Jenk

Zhang

, et al. (2018) Age ranges of the Tibetan ice cores with emphasis on the Chongce ice cores, western Kunlun Mountains. The Cryosphere 12: 2341–2348.

46.

Hou

Zhang

Fang

, et al. (2021) Brief communication: new evidence further constraining Tibetan ice core chronologies to the Holocene. The Cryosphere 15(4): 2109–2114.

47.

Huang

Shen

Long

, et al. (1998) The empirical mode decomposition and the Hilbert spectrum for nonlinear and non-stationary time series analysis. Proceedings of the Royal Society of London. Series A: Mathematical, Physical and Engineering Sciences 454: 903–995.

48.

Chen

Balsam

, et al. (2004) High resolution hematite/goethite records from Chinese loess sequences for the last glacial-interglacial cycle: rapid climatic response of the East Asian Monsoon to the tropical Pacific. Geophysical Research Letters 31(3).

49.

Balsam

Chen

(2001) Mineralogic and climatic interpretations of the Luochuan loess section (China) based on diffuse reflectance spectrophotometry. Quaternary Research 56: 23–30.

50.

Jia

Wang

, et al. (2018) Variations in the iron mineralogy of a loess section in Tajikistan during the mid-pleistocene and late Pleistocene: implications for the climatic evolution in central Asia. Geochemistry Geophysics Geosystems 19: 1244–1258.

51.

Jian

Wang

Dang

, et al. (2020) Half-precessional cycle of thermocline temperature in the western equatorial Pacific and its bihemispheric dynamics. Proceedings of the National Academy of Sciences of the United States of America 117: 7044–7051.

52.

Jian

Wang

Dang

, et al. (2022) Warm pool ocean heat content regulates ocean–continent moisture transport. Nature 612: 92–99.

53.

Jiang

Xiao

, et al. (1997) The Ganzi loess stratigraphy in the west Sichuan plateau. Acta Geologica Sinica 18: 413–420.

54.

Jouzel

Masson-Delmotte

Cattani

, et al. (2007) Orbital and millennial Antarctic climate variability over the past 800,000 years. Science (New York, N.Y.) 317(5839): 793–796.

55.

Kang

Wang

(2011) Closely-spaced recuperated OSL dating of the last interglacial paleosol in the southeastern margin of the Chinese Loess Plateau. Quaternary Geochronology 6: 480–490.

56.

Kanu

Meludu

Oniku

(2014) Comparative study of top soil magnetic susceptibility variation based on some human activities. Geofisica Internacional 53: 411–423.

57.

Kim

Wang

(2011) Sensitivity of the WRF model simulation of the East Asian summer monsoon in 1993 to shortwave radiation schemes and ozone absorption. Asia-Pacific Journal of Atmospheric Sciences 47: 167–180.

58.

Kukla

(1989) Loess stratigraphy in Central China. Palaeogeography, Palaeoclimatology, Palaeoecology 72: 203–225.

59.

Lai

Mischke

Madsen

(2014) Paleoenvironmental implications of new OSL dates on the formation of the “shell bar” in the Qaidam Basin, northeastern Qinghai-Tibetan plateau. Journal of Paleolimnology 51: 197–210.

60.

Laskar

Robutel

Joutel

, et al. (2004) A long-term numerical solution for the insolation quantities of the Earth. Astronomy & Astrophysics 428: 261–285.

61.

Zhang

Shi

, et al. (2008) A high resolution MIS 3 environmental change record derived from lacustrine deposit of Tianshuihai lake, Qinghai-Tibet plateau (in Chinese). Quaternary Science 28: 122–131.

62.

Liang

, et al. (2015) Multiscale monsoon variability during the last two climatic cycles revealed by spectral signals in Chinese loess and speleothem records. Climate of the Past 11: 1067–1075.

63.

Zhang

Liu

, et al. (2020) Paleoclimatic changes and modulation of East Asian summer monsoon by high-latitude forcing over the last 130,000 years as revealed by independently dated loess-paleosol sequences on the NE Tibetan Plateau. Quaternary Science Reviews 237: 106283.

64.

Liu

, et al. (2023) Aeolian process and climatic changes in loess records from the eastern Tibetan Plateau: implications for paleoenvironmental dynamics since MIS 3. CATENA 231: 107361.

65.

Lisiecki

Raymo

(2005) A Pliocene-Pleistocene stack of 57 globally distributed benthic delta O-18 records. Paleoceanography 20(1): 1–17.

66.

Liu

(1985) Loess and the Environment. Beijing: China Ocean Press, 1–251.

67.

Liu

Chen

, et al. (2020) New insights on Chinese cave δ18O records and their paleoclimatic significance. Earth-Science Reviews 207: 103216.

68.

Liu

Yang

Cheng

, et al. (2021a) Chronology and dust mass accumulation history of the Wenchuan loess on eastern Tibetan Plateau since the last glacial. Aeolian Research 53: 100748.

69.

Liu

Chen

M-T

, et al. (2021b) Millennial-scale variability of Indian summer monsoon during the last 42 kyr: evidence based on foraminiferal Mg/Ca and oxygen isotope records from the central Bay of Bengal. Palaeogeography, Palaeoclimatology, Palaeoecology 562: 110112.

70.

Long

Balsam

(2011) Rainfall-dependent transformations of iron oxides in a tropical saprolite transect of Hainan Island, South China: spectral and magnetic measurements. Journal of Geophysical Research 116: F03015.

71.

Fang

, et al. (2004) Millennial-scale climate change since the last glaciation recorded by grain sizes of loess deposits on the northeastern Tibetan Plateau. Chinese Science Bulletin 49: 1157–1164.

72.

Maher

(2016) Palaeoclimatic records of the loess/palaeosol sequences of the Chinese Loess Plateau. Quaternary Science Reviews 154: 23–84.

73.

Mao

Pang

Jl.

Huang

, et al. (2017) Climate change during the late stages of MIS 3 period inferred from the loess record in the upper Hanjiang river valley. Quaternary Science 37: 535–547.

74.

Mashiotta

Lea

Spero

(1999) Glacial–interglacial changes in Subantarctic sea surface temperature and δ18O-water using foraminiferal Mg. Earth and Planetary Science Letters 170: 417–432.

75.

McGrath

Clemens

Huang

, et al. (2021) Greenhouse gas and ice volume drive Pleistocene Indian summer monsoon precipitation isotope variability. Geophysical Research Letters 48(4).

76.

McManus

Francois

Gherardi

, et al. (2004) Collapse and rapid resumption of Atlantic meridional circulation linked to deglacial climate changes. Nature 428: 834–837.

77.

Meng

Liu

(2018) Millennial-scale climate oscillations inferred from visible spectroscopy of a sediment core in Qarhan Salt Lake of Qaidam Basin between 40 and 10 cal ka BP. Quaternary International 464: 336–342.

78.

Miao

Yang

Zhuo

, et al. (2013) Relationship between the color of surface sediments and precipitation in arid northwest China (in Chinese). Marine Geology & Quaternary Geology 33: 77–85.

79.

Mohtadi

Prange

Oppo

, et al. (2014) North Atlantic forcing of tropical Indian Ocean climate. Nature 509: 76–80.

80.

Nascimento

Venancio

Chiessi

, et al. (2021) Tropical Atlantic stratification response to late Quaternary precessional forcing. Earth and Planetary Science Letters 568: 117030.

81.

Owen

Finkel

Caffee

(2002) A note on the extent of glaciation throughout the Himalaya during the global Last Glacial Maximum. Quaternary Science Reviews 21: 147–157.

82.

Pachur

Wunnemann

Zhang

(1995) Lake evolution in the tengger desert, northwestern China, during the last 40,000 years. Quaternary Research 44: 171–180.

83.

Pan

Wang

(1999) Loess record of Qinghai-Xizang Plateau monsoon variations in the eastern part of the plateau since the last interglacial. Quaternary Science 4: 330–335.

84.

Pang

Huang

Zhou

, et al. (2017) OSL dating of the Tuojiawan loess-palaeosol profile and climatic change during the last 55ka BP. Acta Geologica Sinica 91: 2841–2853.

85.

Panmei

Divakar Naidu

Mohtadi

(2017) Bay of bengal exhibits warming trend during the younger dryas: implications of AMOC. Geochemistry, Geophysics, Geosystems 18: 4317–4325.

86.

Peterse

Martinez-Garcia

Zhou

, et al. (2014) Molecular records of continental air temperature and monsoon precipitation variability in East Asia spanning the past 130,000 years. Quaternary Science Reviews 83: 76–82.

87.

Phillips

Sloan

Shroder

, et al. (2000) Asynchronous glaciation at nanga parbat, northwestern Himalaya Mountains, Pakistan. Geology 28: 431–434.

88.

Pons-Branchu

Hamelin

Losson

, et al. (2010) Speleothem evidence of warm episodes in northeast France during Marine Oxygen Isotope Stage 3 and implications for permafrost distribution in northern Europe. Quaternary Research 74: 246–251.

89.

Qiao

Zhao

Wang

, et al. (2006) Magnetostratigraphy and its paleoclimatic significance of a loess-soil sequence from Ganzi area, west Sichuan Plateau. Quaternary Science 26: 250–256.

90.

Qiu

Liu

Bai

(2012) Progress of the studies of paleoclimate and sea level changes in the marine oxygen isotope stage 3. Marine Geology Frontiers 28: 12–16.

91.

Rashid

England

Thompson

, et al. (2011) Late glacial to Holocene Indian summer monsoon variability based upon sediment records taken from the bay of bengal. Terrestrial Atmospheric and Oceanic Sciences 22: 215–228.

92.

Rhodes

Gasse

Lin

, et al. (1996) A late Pleistocene-Holocene lacustrine record from Lake Manas, Zunggar (northern Xinjiang, western China). Palaeogeography, Palaeoclimatology, Palaeoecology 120: 105–121.

93.

Rousseau

D-D

Sima

Antoine

, et al. (2007) Link between European and North Atlantic abrupt climate changes over the last glaciation. Geophysical Research Letters 34(22).

94.

Schaetzl

Bettis

Crouvi

, et al. (2018) Approaches and challenges to the study of loessIntroduction to the loess-fest special issue. Quaternary Research 89: 563–618.

95.

Scheinost

Chavernas

Barron

, et al. (1998) Use and limitations of second-derivative diffuse reflectance spectroscopy in the visible to near-infrared range to identify and quantify Fe oxide minerals in soils. Clays and Clay Minerals 46: 528–536.

96.

Schirmer

(2000) Rhein loess, ice cores and deep-sea cores during MIS 2–5. Journal of the German Geological Society 151(3): 309–332.

97.

Schulze

G Nagel

L Van Scoyoc

E (1993) Significance of organic matter in determining soil colors. Soil color 31: 71–90.

98.

Schwertmann

(1985) The effect of pedogenic environments on iron oxide minerals. Advances in Soil Science 1: 172–200.

99.

Schwertmann

Murad

(1983) Effect of pH on the formation of goethite and hematite from ferrihydrite. Clays and Clay Minerals 31(4): 277–284.

100.

Schwertmann

Taylor

(1989) Iron oxides. In: Dixon

Weed

(eds) Minerals in Soil Environments. 2nd edition. Madison, WI: Soil Society of America, 379–438.

101.

Shen

Tang

Wang

, et al. (2005) Pollen records and time scale for the RM core of the Zoige Basin, northeastern Qinghai-Tibetan Plateau. Chinese Science Bulletin 50: 553–562.

102.

Shi

Yao

(2002) MIS 3b(54-44 ka BP)cold period and glacial advance in middle and low latitudes (in Chinese). Journal of Glaciology and Geocryology 24: 1–9.

103.

Shi

(2003) Warm-humid climate and transgressions during 40∼30ka B.P. and their potential mechanisms (in Chinese). Quat. Sci 23: 1–11.

104.

Shi

Zhao

(2009) The special warm-humid climate and environment in China during 40∼30 ka BP: discovery and Review (in Chinese). Journal of Glaciology and Geocryology 31: 1–10.

105.

Shi

Liu

, et al., 1999. A very strong summer monsoon event during 30–40 kaBP in the Qinghai-Xizang (Tibet) Plateau and its relation to precessional cycle. Chinese Science Bulletin 44: 1851–1858.

106.

Shi

Liu

, et al. (2001) Reconstruction of the 30∼40 ka BP enhanced Indian monsoon climate based on geological records from the Tibetan Plateau (in Chinese). Palaeogeography, Palaeoclimatology, Palaeoecology 169: 69–83.

107.

Singh

Gilkes

(1992) Properties and distribution of iron-oxides and their association with minor elements in the soils of south-western Australia. Journal of Soil Science 43: 77–98.

108.

Song

Zeng

Chen

, et al. (2018) Abrupt climatic events recorded by the Ili loess during the last glaciation in Central Asia: evidence from grain-size and minerals. Journal of Asian Earth Sciences 155: 58–67.

109.

Sprafke

Schulte

Meyer-Heintze

, et al. (2020) Paleoenvironments from robust loess stratigraphy using high-resolution color and grain-size data of the last glacial Krems-Wachtberg record (NE Austria). Quaternary Science Reviews 248: 106602.

110.

Sun

Liang

, et al. (2011) Changing color of Chinese loess: geochemical constraint and paleoclimatic significance. Journal of Asian Earth Sciences 40: 1131–1138.

111.

Sun

McManus

Clemens

, et al. (2021) Persistent orbital influence on millennial climate variability through the Pleistocene. Nature Geoscience 14: 812–818.

112.

Tang

(1992) Vegetation and climate history at Menghai, Yunnan during the past 42000 years. Acta Micropalaeontologica Sinica 9: 433–455.

113.

Thompson

Yao

Davis

, et al. (1997) Tropical climate instability: the last glacial cycle from a Qinghai-Tibetan ice core. Science 276: 1821–1825.

114.

Thompson

Yao

Davis

, et al. (2001) Ice core records of climate variability on the Third Pole with emphasis on the Guliya ice cap, western Kunlun Mountains. Quaternary Science Reviews: 188: 1–14.

115.

Ulfers

Zeeden

Voigt

, et al. (2019) 81Kr dating of the Guliya ice cap, Tibetan Plateau. Geophysical Research Letters 46: 107413.

116.

Torrent

Barrón

(2003) The visible diffuse reflectance spectrum in relation to the color and crystal properties of hematite. Clays and Clay Minerals 51: 309–317.

117.

Torrent

Schwertmann

Schulze

(1980) Iron oxide mineralogy of some soils of two river terrace sequences in Spain. Geoderma 23: 191–208.

118.

Ulfers

Zeeden

Voigt

, et al. (2022) Half-precession signals in Lake Ohrid (Balkan) and their spatio-temporal relations to climate records from the European realm. Quaternary Science Reviews 280: 107413.

119.

Voelker

AHL

(2002) Global distribution of centennial-scale records for Marine Isotope Stage (MIS) 3: a database. Quaternary Science Reviews 21: 1185–1212.

120.

Wang,

Cheng

Edwards

, et al. (2008) Millennial- and orbital-scale changes in the East Asian monsoon over the past 224,000 yr. Nature 451: 1090–1093.

121.

Wang

Cheng

, et al. (2011) High lake levels on alxa plateau during the late quaternary. Chinese Science Bulletin 56: 1799–1808.

122.

Wang

Song

Zhao

, et al. (2016) Color characteristics of Chinese loess and its paleoclimatic significance during the last glacial-interglacial cycle. Journal of Asian Earth Sciences 116: 132–138.

123.

Wen

Liu

Wang

, et al. (2016) Correlation and anti-correlation of the East Asian summer and winter monsoons during the last 21,000 years. Nature Communications 7: 11999.

124.

Wetterich

Tumskoy

Rudaya

, et al. (2014) Ice complex formation in arctic East siberia during the MIS3 interstadial. Quaternary Science Reviews 84: 39–55.

125.

Zhao

, et al. (2004) Palaeovegetation, palaeoclimate and lake-level change since 120 ka BP in Nam Co, central Xizang. Acta Geologica Sinica 78: 242–252.

126.

Yang

Ding

(2003) Color reflectance of Chinese loess and its implications for climate gradient changes during the last two glacial-interglacial cycles. Geophysical Research Letters 30(20).

127.

Yang

Fang

, et al. (2001) Transformation functions of soil color and climate. Science in China Series D: Earth Sciences 44: 218–226.

128.

Yang

Chen

Shu

(2004) A preliminary study on the paleoenvironment during MIS3 in the Changjiang delta region. Quaternary Science 24: 525–530.

129.

Yang

Fang

Shi

, et al. (2010) Timing and provenance of loess in the Sichuan Basin, southwestern China. Palaeogeography Palaeoclimatology Palaeoecology 292: 144–154.

130.

Yang

Pang

Zhou

, et al. (2018) Chroma characteristics and its significance of the Junwangcun loess-paleosol profile in the Hanzhong basin Journal of Natural Sciences University 57(1): 93–101.

131.

Yang

Liu

Chen

, et al. (2021a) Variability and environmental significance of organic carbon isotopes in Ganzi loess since the last interglacial on the eastern Tibetan Plateau. CATENA 196: 104866.

132.

Yang

Liu

, et al. (2021b) Quartz OSL chronology of the loess deposits in the Western Qinling Mountains, China, and their palaeoenvironmental implications since the Last Glacial period. Boreas 50: 294–307.

133.

Yang

Chen

, et al. (2022) Magnetic properties of the Ganzi loess and their implications for precipitation history in the eastern Tibetan plateau since the last interglacial. Paleoceanography and Paleoclimatology 37.

134.

Gui

Shi

, et al. (2007) Late marine isotope stage 3 palaeoclimate for East Asia: a data–model comparison. Palaeogeography, Palaeoclimatology, Palaeoecology 250: 167–183.

135.

S-Y

Colman

Lai

Z-P

(2019) Late-Quaternary history of ‘great lakes’ on the Tibetan Plateau and palaeoclimatic implications – a review. Boreas 48: 1–19.

136.

Zeeden

Hambach

Veres

, et al. (2018) Millennial scale climate oscillations recorded in the Lower Danube loess over the last glacial period. Palaeogeography, Palaeoclimatology, Palaeoecology 509: 164–181.

137.

Zhang

(1997) A study on loess sedimentation and environmental change since late Pleistocene in Wudu Area. Journal of Lanzhou University (Natural Sciences) 33: 105–114.

138.

Zhang

Shi

Dai

(1997) The climate records of deep sea oxygen isotope stage 5: a detailed correlation between Wudu Loess, ice cores from polar glaciers and deep sea (in Chinese). Journal of Lanzhou University (Natural Sciences) 33: 107–115.

139.

Zhang

Wunnemann

, et al. (2002) Lake level and climate changes between 42,000 and 18,000 ¹⁴C yr B.P. in the tengger desert, northwestern China. Quaternary Research 58: 62–72.

140.

Zhang

Chen

Holmes

, et al. (2011) Holocene monsoon climate documented by oxygen and carbon isotopes from lake sediments and peat bogs in China: a review and synthesis. Quaternary Science Reviews 30: 1973–1987.

141.

Zhang

Xue

Yao

(2019a) The lake status records and paleoclimatic changes of China since the Last Interstadial. Quaternary International 527: 12–18.

142.

Zhang

Gosling

, et al. (2019b) Vegetation and climate evolution during the last glaciation at Tengchong in yunnan province, southwest China. Palaeogeography, Palaeoclimatology, Palaeoecology 514: 441–452.

143.

Zhang

Zhao

, et al. (2020a) Spatiotemporal complexity of the “greatest lake period” in the Tibetan plateau. Science Bulletin 65: 1317–1319.

144.

Zhang

Zheng

Huang

, et al. (2020b) Sensitivity of altitudinal vegetation in southwest China to changes in the Indian summer monsoon during the past 68000 years. Quaternary Science Reviews 239: 106359.

145.

Zhao

Tzedakis

, et al. (2020) Evolution of vegetation and climate variability on the Tibetan Plateau over the past 1.74 million years. Science Advances 6: eaay6193.

146.

Zhao

Rohling

Liu

, et al. (2021) Possible obliquity-forced warmth in southern Asia during the last glacial stage. Science Bulletin 66: 1136–1145.

147.

Zheng

Xiang

(1989) Saline Lakes in the Qinghai-Xizang Plateau. Beijing: Beijing Science and Technology Publishing House.

148.

Zhong

Solonenko

, et al. (2020) Glacier ice archives fifteen-thousand-year-old viruses. bioRxiv 9: 160.

149.

Zhu

Yuan

Song

, et al. (2020) Advance in dissimilatory iron reduction in wetland soils and sediments. Wetland Science 18(1): 122–128

Supplementary Material

Please find the following supplemental material available below.

For Open Access articles published under a Creative Commons License, all supplemental material carries the same license as the article it is associated with.

For non-Open Access articles published, all supplemental material carries a non-exclusive license, and permission requests for re-use of supplemental material or any part of supplemental material shall be sent directly to the copyright owner as specified in the copyright notice associated with the article.

0.00 MB

0.23 MB