Early-Holocene to present palaeoenvironmental shifts and short climate events from the tropical wetland and lake sediments,Kukkal Lake,Southern India: Geochemistry and palynology

Abstract

The Kukkal basin, Tamil Nadu, India, receives most of its rain from the southwest monsoon (SWM). A sediment core from Kukkal Lake preserves a continuous sediment record from the early-Holocene to present (9000 yr BP to present). The present lake is situated at an elevation of ~1887 m a.s.l., in a small basin that appears to have alternated between a and wetland depositional environment. Climate proxies, including sediment texture, total organic carbon (TOC), total nitrogen (TN), C/N, pollen and geochemical composition indicate a steady progression to wetter conditions, with two stepwise changes at about 8000, and between 3200 and 1800 yr BP. The change at 8000 yr BP appears to correspond to a brief (100–150 years) dry spell recorded elsewhere in India. The change at 3200–1800 yr BP consisted in a rapid intensification of the SWM, and may correlate with the initiation of the ‘Roman Warm Period’. There is no clear evidence of changes at the times of the ‘Medieval Warm Period’ (‘MWP’) and the ‘Little Ice Age’ (‘LIA’). The C/N ratio of the sediments ranges from 14.02 to 8.31, indicating that the organic matter originated from a mixture of lacustrine algae, vascular and terrestrial plants. Chemical weathering indices (Chemical Index of Alteration (CIA), Chemical Index of Weathering (CIW), and Plagioclase Index of Alteration (PIA)) are consistent with extreme silicate weathering. Pollen data show a development from savanna vegetation prior to about 8000 yr BP, followed by grassland with palms, the appearance of ferns just prior to 3200 yr BP and the establishment of the tropical humid forest between 3200 and about 1800 yr BP.

Keywords

chemical indices of weathering climate change ferns Holocene India radiocarbon dating southwest monsoon Tamil Nadu tropical lake tropical wetland

Introduction

The Indian subcontinent is one of the largest monsoon-dominated regions of the world, and the climate is controlled primarily by the southwest monsoon (SWM). The term ‘monsoon’ refers to seasonal moisture-rich winds, and it is a unique climatic factor that differentiates south Asian countries – particularly India and parts of Asia, Australia and Africa – from the rest of the world. The Indian subcontinent is predominantly affected by the changes in these winds, which have a southerly and a westerly component. Combined together, they constitute the SWM, the sign of life and the much needed rains. The influence of the SWM is felt as far north as China’s Xinjiang region. India gets most of its rainfall by the SWM from June to September, and only the coastal tract of the Tamil Nadu state receives the Northeast rains or the return monsoon during mid-September to mid-December. Around September, as the sun retreats south, the northern land mass of the Indian subcontinent cools fast, and with this, the air pressure rises over northern India, while the surrounding atmosphere of the Indian Ocean holds the heat. This causes the cold wind to move down from the Himalayas and Indo-Gangetic Plain towards the vast spans of the Indian Ocean south of the Deccan peninsula. This is known as the northeast monsoon (NEM) or Retreating Monsoon. However, the amount and intensity of both the monsoonal rains have fluctuated since the late Quaternary period (Clift and Plumb, 2008; Gupta et al., 2006; Singhvi et al., 2010, 2012). During the last two decades, scientific interest in the reconstruction of the Asian monsoon variations, particularly the SWM, since the late Quaternary to Holocene period has increased significantly. In particular, in the Indian subcontinent, many types of natural archives such as lake and marine sediment cores, palynology, tree rings and corals (Mayewski et al., 2004) have been used to restructure late Quaternary–Holocene palaeoclimatic changes. Such records include fluvial archives of the Gujarat region (Chamyal et al., 2002), deep marine sediment cores studied from the Arabian Sea and the Bay of Bengal (Achyuthan et al., 2014; Chauhan et al., 2010; Gupta et al., 2003, 2005; Nagasundaram, 2014; Nagasundaram et al., 2014), lacustral and aeolian sediments from the Thar desert (Achyuthan et al., 2007a, 2007b; Enzel et al., 1999), lacustrine sediments from Kerala (Farooqui et al., 2010; Farooqui and Sekhar, 2011; Kumaran et al., 2008; Veena et al., 2014), peat deposits from the foreland basin of the Himalayas (Rawat et al., 2015; Sharma et al., 2006), glacial moraines and proglacial deposits of the Ladakh region (Sant et al., 2011), speleothems (Yadava and Ramesh, 2005), stalagmites (Sinha et al., 2007) and tree rings (Esper et al., 2002a, 2002b, 2003; Shah and Bhattacharya, 2009; Yadav and Singh, 2002).

Lakes and wetlands in this region function as repositories of sediment deposited in a closed system and act as major sinks for organic carbon. Therefore, sediment cores from such sites in the area affected by the SWM can provide a continuous, high resolution of palaeoenvironmental record of the evolution of the SWM (Verschuren, 2003; Verschuren et al., 2000). It is important to reconstruct past climate and to study palaeoclimate records with high resolution especially since the early-Holocene period because the past performance of the monsoon may provide vital input for projecting future climate change and predictions. Recent high-resolution studies from the northern Andes show that vegetation has responded to climate change within decades to centuries (Gonzalez-Carranza et al., 2012). The growth and impacts of the human species worldwide had profound implications for human development, advance of major river valley civilizations took place because of rain-fed agriculture and were established during this time interval (Allchin, 1997; Enzel et al., 1998; Gupta, 2004; Ponton et al., 2012; Possehl, 1996) and that from the past records the early-Holocene period was wetter than today (Allchin, 1997; Enzel et al., 1998; Gupta, 2004; Kaufman et al., 2004; Possehl, 1996). Hence, it is important to reconstruct past climate changes with higher resolution and greater degree of accuracy.

The Holocene Megathermal, also known as Hypsithermal, Altithermal or Holocene Climatic Optimum (HCO) or Holocene Optimum or Holocene Thermal Maximum, was a warmer phase during roughly the interval 9000–5000 yr BP in areas that received East Asian monsoon (EAM) (An et al., 2000; Sun and Li, 2012; Wang et al., 2005). In the Asian region, palaeoclimate records based on lake sediments indicate that both southwest summer and northeast winter rains were stronger in the early-Holocene period (10,000–7000 yr BP) than today, resulting in a wetter and cooler climate over the Asian continent. Warming started around 8000 yr BP, and the Holocene Megathermal Maximum (HMM) that occurred between 7200 and 6000 yr BP observed particularly change in vegetation zones, increase in temperatures, precipitation and sea-level fluctuations. Annual mean temperature during this period was estimated at 1°C higher than today for South China; 2°C higher for the Changjiang River Valley; about 3°C higher for North China, Northeast China and Northwest China; and up to 4–5°C higher for southern Qinghai-Xizang Plateau (Shi et al., 1993). After 3000 yr BP, climate in China area returned to a cold and dry state (Sun and Li, 2012). Peat deposits from Sanjiang Plain, China, revealed a dry and warm climate from 1900 to 1000 yr BP, a warm and humid phase from 1000 to 500 yr BP, and a cool and dry climate since 500 yr BP (Xia and Wang, 2001). A period of warming in Oman coeval with that in China was observed from the index fossil Globigerina bulloides, which indicates wind-induced upwelling of ocean water, during the interval 1200–600 yr BP (Gupta et al., 2003). These conclusions are based on a wide variety of proxies, including pollen and diatoms (Prell and Kutzbach, 1987; Van Campo et al., 1996), oxygen isotopes from marine sediments (Sirocko et al., 1993, 1996) and Arabian Sea upwelling intensities (Overpeck et al., 1996; Prell and Van Campo, 1986). Staubwasser et al. (2002, 2003) provided some insight into the climate variability in the SWM region since the mid-Holocene.

A history of flow volume of the Indus River over the past 5000 years has been reconstructed by Von Rad et al. (1999a, 1999b). The Indus has a large catchment in a region that is now partly semi-arid. Being fed by melting Himalayan ice as well as precipitation in the semi-arid hinterland, its discharge is governed by a complex interplay of phenomena. Information obtained to date shows a stronger SWM since about 1500 yr BP. Temperature indicators are less clear, except around the ‘Medieval Warm Period’ (‘MWP’) (about 1000–600 yr BP) and the ‘Little Ice Age’ (‘LIA’) (Cronin et al., 2003; Crowley and Lowery, 2000). Chauhan and Vogelsang (2006) reported that the ‘LIA’ had a large temporal and spatial variability because of the variations in the intensity of SWM and had its impact across the globe. However, in the northern hemisphere, the ‘LIA’ was a mild cooling event (Anderson et al., 2002; Cronin et al., 2003; Crowley, 2000; Crowley and Lowery, 2000; Singhvi and Kale, 2008). Summarizing these studies, it is evident that a significant variation of the SWM, encompassing several cycles of fluctuations, occurred during the Holocene.

To date, the marine record from the Arabian Sea has proved to be the most consistent and continuous record of the SWM fluctuations (Kudrass et al., 2001; Lückge et al., 2001; Sirocko et al., 1993, 1996, 2000; Thamban et al., 2002; Tiwari et al., 2005, 2006), whereas limited studies have been carried out in the Bay of Bengal (Achyuthan et al., 2014; Chauhan et al., 2004; Chauhan and Vogelsang, 2006; Duplessy, 1982). However, lakes of peninsular India preserve a great wealth of palaeoenvironmental and palaeoclimate data spanning the last 6000 years as described by Limaye et al. (2010) and Veena et al. (2014). The lake and wetland archives of Tamil Nadu have received so far little attention. Kukkal Lake, in the Palani Hills, Dindigul District, Tamil Nadu State, lies in a basin affected by both the SWM and NEM rains. Hence, the main objectives of this study are to reconstruct the Holocene palaeomonsoon and palaeoenvironmental shifts using the Kukkal Lake sediments.

Study area

Kukkal Lake (10°16′N: 77°22′E) is a small, irregular, fresh water body covering an area of 18 ha, and situated at an altitude of 1887 m a.s.l. (Figure 1) about 33 km from Kodaikanal, Dindigul district, Tamil Nadu, India. The present lake is impounded by a low dam at the northern end, and floods an area that appears to have been a wetland similar to one about 300 m south of the lake. Both features may have been lakes at times in the past. Presently, the lake has a catchment area of 2–3 km² that is forested and relatively undisturbed by human activities; some abandoned paddy terraces on the east side of the lake are the only evidence of agriculture in the catchment. Sediment in the basin beneath the lake is therefore suitable for carrying out palaeoenvironment research. Bedrock underlying the catchment is Precambrian charnockite associated with hornblende-biotite gneiss, granite and quartzite, and is covered by a thin veneer of red vertisol and lateritic soil (Bagyaraj and Gurugnanam, 2011). The geology and geomorphic observations in and around the lakes reveal that the Kukkal Lake basin is structurally controlled by faults and lineaments, as are the main tributary watercourses within the basin. The catchment receives rainfall from both SWM and NEM, the NEM being dominant (Meher-Homji and Gupta, 1999). The mean annual rainfall for Kodaikanal over a 4-year period (2001–2004) was 1690 mm (Bagyaraj and Gurugnanam, 2011). The dry season is from December through March, and there is also a short dry period in July. The climate is tropical humid for most of the year, with summer temperatures reaching 19.8°C (max.) and 11.3°C (min.) and winter temperatures between 17.3°C (max.) and 8.3°C (min.). Meher-Homji and Gupta (1999) state that Kodaikanal experiences a bixeric climate regime receiving high rainfall from October to December (i.e. NEM). The lake is surrounded by the ‘Shola’ forests (the word ‘Shola’ derived from the Tamil language word ‘cÕlai’ meaning grove or thick vegetation cover (Meher-Homji, 1967). The Sholas are patches of stunted tropical montane forest found in the valleys in the midst of rolling grassland in the higher montane regions of the Nilgiri Hills, South India. The soil texture varies from clay to silty clay; soils are acidic in nature and contain a high percentage of iron oxide and alumina. In the ‘Shola’ forests, accumulation of humus in the top layers of the soil renders it a black colour (Bera and Farooqui, 2000; Meher-Homji, 1967). The dominant trees in the Shola forest are represented by Michelia nilagirica, Bischofia javanica (bishop wood), Calophyllum tomentosum, Cedrela toona (Indian mahogany), Eugenia (myrtle) spp., Ficus glomerata and Mallotus spp.; and small trees such as Pygeum gardneri, Schefflera racemosa, Linociera ramiflora, Syzygium spp., Rhododendron nilgiricum, Mahonia nepalensis, Elaeocarpus recurvatus, Ilex denticulata, Michelia nilagirica, Actinodaphne bourdillonii and Litsea wightiana. These forests are interspersed with grasslands, characterized by frost and fire-resistant grass species like the Chrysopogon zeylanicus, Cymbopogon flexuosus, Arundinella ciliata, Arundinella mesophylla, Arundinella tuberculata, Themeda tremula and Sehima nervosum.

Figure 1.

Map showing the location of the Kukkal Lake and the sediment coring site.

Materials and methods

Field methods

The field map was prepared using Survey of India toposheets, a Google Earth map and Satellite imagery data products IRS LISS III for delineating faults, lineaments and channels draining into the lake. Faults and the rock types exposed around the lake surroundings were also noted. A piston core of 1 m in length was used to retrieve the sediment core from lake floor (Figure 1). The cored sampled location was marked using a Garmin III Global Positioning System (GPS) (Figure 1). The sediment core was then subsampled at 2-cm intervals, collected in clean zip-locked polythene bag, marked, described (Table 1) and stored for further analyses.

Table 1.

Lithology of the Kukkal Lake sediment core.

Lithological unit (depth, cm)	Description	Munsells colour
III (28-0)	Brownish black (10 Y 3/1) in colour and the sediment type was sandy silt and silty sand laminae, peaty layers, with the coarse organic carbon pellets with rootlets and root pores.	Brownish black 10 YR 3/1
II (47-28)	Brownish black (10 Y 2/3) in colour moderately sorted sandy silt and silty sand quartz, fine clear coarse quartz, angular to sub-angular with ferruginous nodules, and peat 3.5 mm in length	Brownish black 10 YR 2/3
I (100-47)	Peaty layer, dark black, clay with few roots, organic rich layer, fine sandy silt, clayey silt and silty sand layers	Black 10 YR 2/1

Radiocarbon dating

Four organic carbon rich samples from the core were subjected to acid-base-acid pre-treatment, and dated by liquid scintillation counting on benzene at the Department of Geosciences, University of Arizona, Tucson, USA. The radiocarbon ages were then calibrated using Calib 5.2.0 (Stuiver et al., 2010; Stuiver and Reimer, 1993) (Figure 2, Table 2).

Figure 2.

Sediment texture, radiocarbon ages, sedimentation rate (mm/a), CaO/MgO and CWI variation with depth.

Table 2.

Radiocarbon dates of the lake sediment core collected from the Kukkal Lake.

S.no.	Lab ID	Depth (cm)	¹⁴C age (yr BP)	Calibrated date	δ ¹³C‰
1					−22.6
2	A16345	16–18	99.7 ± 1.1 pMC	AD 1650–1955	−23.3
3					−22.8
4	A16346	42–44	3215 ± 120	1862–1131 BC	−20.6
5	A16347	70–72	7640 ± 165	7028–6110 BC	−17.7
6	A16311	98–100	9105 ± 345/−330	9255–7526 BC	−16.1

Chemical and physical analyses

For the sediment texture analyses, about of 5 g of the dried sediment was washed following Carver (1971) through a 230 American Standard Testing Machine (ASTM) sieve mesh opening of 0.063 mm until the water passing through was clear (Tables 1 and 2), and the washed water was collected in a large porcelain basin. This sieving separated the sand particles from the fine fractions of silt and clay particles. The sediments collected in the 230 sieve mesh were calculated as the sand content. Furthermore, the silt and clay percentages were determined using the measuring jar and settling velocity procedure put forward by Carver (1971). Procedures of Loring and Rantala (1992) and Piper (1947) titration methods were adopted for the determination of calcium carbonate content. The organic matter (OM) in the sediment sample was determined following the methods of Gaudette et al. (1974). The total organic carbon (TOC) and total nitrogen (TN) were determined using about 1 g of sediment sample, decarbonated with a 1M solution of hydrochloric acid, washed three times with deionized water, freeze-dried and then analysed in a Thermo Scientific FLASH 2000 CHNS-O Organic Elemental Analyser (OEA) at the Department of Geology, Anna University, Chennai (Figure 3, Table 3). Major oxide and trace element analyses were carried out at the Central Instrumentation Facility (CIF) at Pondicherry University, Pondicherry, using a wavelength dispersive x-ray fluorescence (WD-XRF) spectrometer with samples in powder form. The samples were analysed for SiO₂, Al₂O₃, Na₂O, CaO, Fe₂O₃, K₂O, TiO₂, P₂O₅, MnO, MgO, Sr, Cr, Cu, Ni, Co, Pb, Zn, V, Ba and Zr (Tables 4 and 5). Results were normalized to Al (Figure 4).

Figure 3.

Down core variations of texture, OM, TOC, TN percentage and C/N ratio showing increase in sand from 7640 to 3215 yr BP because of sediment flux from the catchment area.

Table 3.

Sediment texture, OM, TOC, TN content, C/N ratio and age of Kukkal sediment core.

S.no.	Depth (cm)	Age	Sand %	Silt %	Clay %	CaCO₃%	OM %	TOC %	TN %	C/N	Sediment type
1	2–4	17.58	45.75	50.72	3.53	2.25	42.34	15.79	1.13	14.02	Sandy silt
2	4–6	29.30	40.11	55.61	4.27	0.48	40.86	12.13	1.04	11.62	Sandy silt
3	6–8	41.02	50.61	45.49	3.90	1.32	47.02	12.34	1.09	11.32	Silty sand
4	8–10	52.74	41.98	53.57	4.46	1.10	32.70	12.54	1.14	11.01	Sandy silt
5	10–12	64.46	42.23	53.39	4.39	1.17	34.77	10.29	1.19	8.73	Sandy silt
6	12–14	76.18	36.98	58.00	5.02	0.93	34.89	8.05	1.25	6.45	Sandy silt
7	14–16	87.90	41.00	54.50	4.50	0.90	36.64	13.19	1.18	11.15	Sandy silt
8	16–18	99.70	40.83	53.99	5.17	1.40	35.11	13.20	1.09	12.21	Sandy silt
9	18–20	339.34	40.87	54.33	4.81	1.25	36.30	13.21	1.00	13.28	Sandy silt
10	20–22	578.98	44.20	51.30	4.50	1.16	36.17	10.53	1.04	10.25	Sandy silt
11	22–24	818.62	39.02	56.40	4.58	1.19	35.38	7.85	1.09	7.22	Sandy silt
12	24–26	1058.25	39.88	55.26	4.86	1.29	37.15	12.54	1.12	11.19	Sandy silt
13	26–28	1297.89	42.66	53.02	4.31	1.59	36.45	13.49	1.18	11.39	Sandy silt
14	28–30	1537.53	51.73	44.26	4.01	1.38	38.93	14.45	1.25	11.59	Silty sand
15	30–32	1777.17	46.24	49.67	4.09	0.76	32.95	8.60	0.93	8.02	Sandy silt
16	32–34	2016.81	39.73	56.40	3.87	1.16	22.48	2.76	0.62	4.45	Sandy silt
17	34–36	2256.45	40.75	55.42	3.83	1.57	23.97	5.35	0.51	10.52	Sandy silt
18	36–38	2496.08	41.26	55.47	3.27	0.19	24.58	7.83	0.63	12.12	Sandy silt
19	38–40	2735.72	39.49	57.15	3.36	0.57	26.81	10.30	0.75	13.72	Sandy silt
20	40–42	2975.36	50.66	46.01	3.33	0.53	30.97	7.94	0.89	9.56	Silty sand
21	42–44	3215.00	64.97	32.41	2.62	0.55	34.74	5.58	1.03	5.41	Silty sand
22	44–46	3531.07	78.47	19.55	1.98	0.72	37.81	11.50	0.97	11.86	Sand
23	46–48	3847.14	91.92	7.56	0.52	0.63	40.04	9.86	0.88	11.13	Sand
24	48–50	4163.21	91.29	7.97	0.74	0.36	41.56	8.22	0.79	10.40	Sand
25	50–52	4479.29	82.66	14.65	2.69	0.57	42.72	8.57	0.71	12.23	Sand
26	52–54	4795.36	77.90	18.83	3.27	0.59	42.18	8.92	0.63	14.07	Sand
27	54–56	5111.43	78.02	18.76	3.22	0.36	41.68	8.61	0.69	12.43	Sand
28	56–58	5427.50	78.91	17.91	3.18	1.00	42.20	8.13	0.63	12.84	Sand
29	58–60	5743.57	81.84	14.80	3.36	0.78	42.02	7.64	0.58	13.25	Sand
30	60–62	6059.64	82.38	13.63	4.00	0.76	40.65	7.57	0.55	13.66	Sand
31	62–64	6375.71	57.92	34.06	8.02	0.77	38.85	7.50	0.53	14.07	Silty sand
32	64–66	6691.79	35.73	54.99	9.28	1.48	37.55	6.96	0.73	9.50	Sandy silt
33	66–68	7007.86	21.26	67.64	11.10	0.98	36.20	6.39	0.60	11.10	Sandy silt
34	68–70	7323.93	12.25	74.56	13.20	1.00	36.25	5.82	0.46	12.70	Clayey silt
35	70–72	7640.00	17.91	69.41	12.68	1.67	31.80	5.49	0.47	11.61	Sandy silt
36	72–74	7744.64	20.52	67.08	12.40	1.00	31.67	5.16	0.49	10.52	Sandy silt
37	74–76	7849.29	19.68	68.12	12.20	1.64	30.06	4.25	0.47	9.03	Sandy silt
38	76–78	7953.93	24.78	61.25	13.96	2.44	30.96	4.38	0.51	8.69	Sandy silt
39	78–80	8058.57	36.44	50.75	12.81	2.28	30.71	4.51	0.54	8.34	Sandy silt
40	80–82	8163.21	23.12	60.79	16.09	2.38	32.14	4.51	0.55	8.28	Sandy silt
41	82–84	8267.86	22.61	63.39	14.01	2.79	32.27	4.52	0.55	8.21	Sandy silt
42	84–86	8372.50	18.41	64.90	16.69	3.19	28.69	3.58	0.40	8.89	Sandy silt
43	86–88	8477.14	14.97	69.91	15.12	2.79	29.74	3.86	0.42	9.10	Clayey silt
44	88–90	8581.79	20.99	63.20	15.81	2.40	30.80	4.15	0.45	9.30	Sandy silt
45	90–92	8686.43	24.33	61.45	14.22	2.46	32.17	4.91	0.56	8.77	Sandy silt
46	92–94	8791.07	23.57	61.10	15.33	3.16	31.62	4.90	0.56	8.76	Sandy silt
47	94–96	8895.71	19.65	64.20	16.15	4.08	30.10	4.90	0.56	8.75	Sandy silt
48	98–100	9105.00	27.50	57.93	14.57	0.92	30.28	4.42	0.53	8.31	Sandy silt
Average			43.87	48.77	7.36	1.37	34.87	8.07	0.77	10.44
Minimum			12.25	7.56	0.52	0.19	22.48	2.76	0.40	4.45
Maximum			91.92	74.56	16.69	4.08	47.02	15.79	1.25	14.07

OM: organic matter; TOC: total organic carbon; TN: total nitrogen.

Ages: measured interpolated. Compositions: maximum in bold, minimum in bold italics.

Table 4.

Major oxides (%) and weathering index of the sediments from the Kukkal sediment core.

S.no.	Depth (cm)	Age	SiO₂	Al₂O₃	Fe₂O₃	TiO₂	MgO	K₂O	P₂O₅	CaO	MnO	Na₂O	CIA	CIW	PIA	CWI
1	2–4	17.58	41.01	14.06	6.29	0.94	0.57	0.55	0.59	0.68	0.08	0.01	91.88	91.96	95.13	8.99
2	4–6	29.30	40.06	14.02	6.71	0.94	0.58	0.54	0.57	0.68	0.08	0.01	91.88	91.96	95.08	9.14
3	8–10	52.74	40.52	14.73	6.46	0.93	0.58	0.55	0.58	0.69	0.07	0.01	92.17	92.25	95.27	8.73
4	14–16	87.90	41.50	13.39	6.13	0.95	0.56	0.56	0.60	0.67	0.08	0.01	91.56	91.65	94.97	9.26
5	18–20	339.34	39.59	13.31	6.95	0.95	0.59	0.54	0.56	0.68	0.09	0.01	91.56	91.65	94.87	9.59
6	24–26	1058.25	40.18	14.14	6.47	1.01	0.71	0.55	0.57	0.68	0.07	0.01	91.96	92.05	95.17	9.90
7	28–30	1537.53	39.35	13.77	5.38	0.94	0.60	0.58	0.61	0.65	0.06	0.01	91.67	91.75	95.19	9.21
8	34–36	2256.45	46.65	16.77	7.75	1.13	0.63	0.57	0.50	0.43	0.05	0.01	94.29	94.37	97.34	6.42
9	38–40	2735.72	36.65	18.73	7.38	1.25	0.66	0.61	0.46	0.30	0.04	0.01	95.31	95.38	98.31	5.19
10	44–46	3531.07	31.66	17.73	6.87	1.13	0.71	0.52	0.47	0.27	0.03	0.02	95.65	95.73	98.34	5.66
11	48–50	4163.21	32.30	19.01	5.09	1.12	0.60	0.53	0.47	0.25	0.02	0.01	96.03	96.10	98.61	4.50
12	54–56	5111.43	32.16	19.16	6.01	1.09	0.64	0.53	0.42	0.23	0.02	0.01	96.11	96.18	98.70	4.64
13	58–60	5743.57	32.01	19.31	6.93	1.06	0.69	0.53	0.37	0.22	0.01	0.01	96.20	96.26	98.79	4.77
14	64–66	6691.79	31.68	19.30	8.03	1.07	0.79	0.52	0.38	0.21	0.01	0.01	96.26	96.33	98.80	5.26
15	68–70	7323.93	31.35	19.28	9.14	1.08	0.89	0.51	0.38	0.21	0.01	0.01	96.33	96.39	98.82	5.75
16	74–76	7849.29	34.11	21.00	9.30	1.13	0.73	0.58	0.39	0.15	0.01	0.01	96.58	96.64	99.22	4.22
17	78–80	8058.57	37.63	22.21	7.50	1.19	0.88	0.64	0.46	0.18	0.02	0.01	96.38	96.44	99.10	4.85
18	84–86	8372.50	37.65	22.05	7.63	1.18	0.90	0.65	0.47	0.18	0.02	0.01	96.31	96.37	99.09	4.99
19	88–90	8581.79	37.67	21.89	7.76	1.18	0.92	0.65	0.47	0.19	0.02	0.01	96.25	96.31	99.07	5.13
20	92–94	8791.07	37.49	21.48	7.37	1.12	0.89	0.64	0.45	0.18	0.02	0.02	96.22	96.29	99.05	5.06
21	98–100	9105.00	37.31	21.06	6.98	1.07	0.85	0.63	0.43	0.18	0.01	0.02	96.20	96.27	99.04	5.00
Average			37.07	17.92	7.05	1.07	0.71	0.57	0.49	0.38	0.04	0.01	94.61	94.68	97.52	6.49
Minimum			31.35	13.31	5.09	0.93	0.56	0.51	0.37	0.15	0.01	0.01	91.56	91.65	94.87	4.22
Maximum			46.65	22.21	9.30	1.25	0.92	0.65	0.61	0.69	0.09	0.02	96.58	96.64	99.22	9.90

CIA: Chemical Index of Alteration; CIW: Chemical Index of Weathering; PIA: Plagioclase Index of Alteration; CWI: Chemical Weathering Intensity.

Maximum in bold and minimum in italic.

Table 5.

Trace element content (ppm) of the sediments from the Kukkal sediment core.

S.no.	Depth (cm)	Age	Sr	Y	Zr	V	Cr	Ni	Cu	Zn	Rb	Nb	Pb	Th	U	Ba
1	2–4	17.58	78.75	15.75	222.95	66.50	69.65	54.95	33.35	78.75	38.80	8.90	40.75	20.25	0.60	498.50
2	4–6	29.30	79.35	15.50	236.25	72.00	70.25	55.80	33.40	79.95	41.30	9.40	37.95	20.10	0.65	516.50
3	8–10	52.74	76.70	16.10	209.20	76.00	72.40	54.50	33.40	75.50	37.80	8.70	40.00	20.80	0.30	468.00
4	14–16	87.90	80.80	15.40	236.70	57.00	66.90	55.40	33.30	82.00	39.80	9.10	41.50	19.70	0.90	529.00
5	18–20	339.34	82.00	14.90	263.30	68.00	68.10	57.10	33.40	84.40	44.80	10.10	35.90	19.40	1.00	565.00
6	24–26	1058.25	74.30	15.20	257.50	79.00	69.00	54.70	31.90	80.50	46.30	8.40	36.10	19.90	0.90	464.00
7	28–30	1537.53	79.30	15.30	277.60	92.00	61.40	59.60	33.20	84.30	46.90	8.90	42.70	20.90	0.80	483.00
8	34–36	2256.45	59.70	16.20	297.10	100.00	70.60	62.60	35.10	74.70	41.70	11.70	27.30	22.10	1.00	350.00
9	38–40	2735.72	61.40	17.90	325.90	95.00	97.80	67.70	50.50	90.40	45.90	12.70	42.80	29.10	1.40	331.00
10	44–46	3531.07	53.40	20.10	226.40	106.00	103.20	60.80	48.20	74.60	36.80	13.30	36.40	30.00	1.50	283.00
11	48–50	4163.21	51.70	24.30	277.20	92.00	96.00	62.70	54.60	72.90	36.00	12.90	38.60	34.00	1.60	277.00
12	54–56	5111.43	50.85	26.35	246.65	94.00	92.15	65.80	51.90	70.80	37.80	11.85	34.10	35.35	1.90	272.00
13	58–60	5743.57	50.00	28.40	216.10	96.00	88.30	68.90	49.20	68.70	39.60	10.80	29.60	36.70	2.20	267.00
14	64–66	6691.79	48.60	28.65	223.20	121.00	94.00	71.55	47.25	71.90	42.90	12.00	29.40	33.80	1.60	283.50
15	68–70	7323.93	47.20	28.90	230.30	146.00	99.70	74.20	45.30	75.10	46.20	13.20	29.20	30.90	1.00	300.00
16	74–76	7849.29	51.10	30.20	257.60	174.00	103.90	83.20	44.40	78.20	51.50	13.00	37.70	33.10	1.00	303.00
17	78–80	8058.57	52.40	31.80	234.60	93.00	105.10	89.50	50.00	93.20	55.10	14.30	36.20	35.00	1.20	314.00
18	84–86	8372.50	52.00	31.10	249.10	120.50	107.80	91.55	50.40	98.95	59.50	14.45	37.55	32.45	1.10	319.50
19	88–90	8581.79	51.60	30.40	263.60	148.00	110.50	93.60	50.80	104.70	63.90	14.60	38.90	29.90	1.00	325.00
20	92–94	8791.07	50.15	31.40	235.55	119.50	108.95	89.95	50.10	103.05	64.55	13.70	35.25	27.05	1.35	322.50
21	98–100	9105.00	48.70	32.40	207.50	91.00	107.40	86.30	49.40	101.40	65.20	12.80	31.60	24.20	1.70	320.00
Average			60.95	23.15	247.35	100.31	88.72	69.54	43.29	83.05	46.78	11.66	36.17	27.37	1.18	371.02
Minimum			47.20	14.90	207.50	57.00	61.40	54.50	31.90	68.70	36.00	8.40	27.30	19.40	0.30	267.00
Maximum			82.00	32.40	325.90	174.00	110.50	93.60	54.60	104.70	65.20	14.60	42.80	36.70	2.20	565.00

Maximum in bold and minimum in bold italic.

Figure 4.

Down core variations of major elements normalized by Al and CWI values showing distinct variation during short extreme events such as the ‘RWP’ and ‘MWP’.

Calculation of chemical indices

The degree of rock alteration and weathering can be estimated quantitatively using several geochemical indices (e.g. Duzgoren-Aydin et al., 2002; Maslov et al., 2003; Price and Velbel, 2003; Yudovich and Ketris, 2011). However, for the charnockite and gneisses exposed around the study area, weathering indices to date have not been calculated. Most of these indices are expressed in terms of molecular or weight percentages of various oxides or groups of oxides. During chemical weathering, more labile minerals, such as feldspar and plagioclase, are depleted in Ca²⁺, K+ and Na+ and transformed into minerals that are more stable under surface conditions. As the intensity of weathering increases, rocks and sediments become enriched in Al, Ti, Fe and Mn (Minyuk et al., 2014).

Chemical Index of Alteration (CIA): The CIA was defined as follows by Nesbitt and Young (1982) using molecular proportions:

CIA = [\frac{A l_{2} O_{3}}{(A l_{2} O_{3} + CaO + NaO + K_{2} O)}] \times 100

with CaO being the amount of CaO incorporated in the silicate fraction of the rock. The CIA measures the proportion of Al₂O₃ versus more labile oxides, and reflects the relative amount of feldspars and clay minerals in a sample. A CIA value of 50 corresponds to unaltered albite, anorthite and potassic feldspar. Typical values of the CIA are 30–40 for basalt, 45–55 for granites and granodiorites, 75–85 for illite, 75 for muscovite and close to 100 for kaolinite and chlorite (Nesbitt and Young, 1982).

Chemical Index of Weathering (CIW). This index was proposed by Harnois (1988) and is calculated using the equation:

CIW = [\frac{A l_{2} O_{3}}{(A l_{2} O_{3} + CaO + N a_{2} O)}] \times 100

with Al₂O₃ treated as an immobile component and CaO and Na₂O as mobile elements. K, in contrast, is not included in this index because it may be leached and/or accumulated in the residual weathering products. Fedo et al. (1995) suggested that Al in this index be used without any correction for its occurrence in K-feldspar, and therefore, generally, K-feldspar-rich rocks yield very high CIW values. Potassic granite, for example, has a CIW value of 80, and fresh K-feldspar has an index of 100, values that are similar to those of the residual products of chemical weathering, for example, 80 for smectite and 100 for kaolinite, illite and gibbsite.

Plagioclase Index of Alteration (PIA). This index (Fedo et al., 1995) estimates the weathering intensity of plagioclase feldspar and is calculated as follows:

PIA = [\frac{(A l_{2} O_{3} - K_{2} O)}{(A l_{2} O_{3} + CaO + N a_{2} O + K_{2} O)}] \times 100

Chemical Weathering Intensity (CWI). Oxides such as CaO, MgO and Na₂O are considered more soluble and mobile, while Al₂O₃, SiO₂ and TiO₂ are considered to be more insoluble and resistant (Engstrom and Wright, 1984; Mackereth, 1966). Sun et al. (2010) proposed the CWI as an alternative to the CIA in order to understand the intensity of weathering and apparently wet or dry conditions. The CWI was calculated as follows, using molecular proportions:

CWI = [\frac{(CaO + MgO + N a_{2} O)}{A l_{2} O_{3}}] \times 100

Palynology

Forty eight sediment samples (taken at 2-cm intervals) were analysed. Ten grams of each sample was treated with warm 10% potassium hydroxide for removing the OM coagulation from the sediments and later sieved through 150 ASTM mesh (105 µm) in order to remove roots/rootlets or any other macro organic/inorganic parts in the sediment. The sieved filtrate fraction was allowed to settle overnight and the supernatant was drained. The residue was treated with 40% hydrofluoric acid (HF) for 2–3 days depending on the sand content in sediments. The HF was decanted after centrifuging the samples at 2000 r/min for 10 min.

The sample was then acetolysed, following Fægri and Iversen (1989), using glacial acetic acid, and later after centrifuging and decanting, the samples were treated with warm acetolysis mixture (anhydrous acetic acid and concentrated H₂SO₄ in 9:1 ratio). Decanting of all the acid was done after centrifuging. The residue was washed with distilled water and decanting was done by centrifuging every time at 2000 r/min for 10 min. After decantation, the samples were passed through 650 mesh size (~10 µm). The residue was then made up to 5 mL in 50% glycerine and water. Whole of the 5-mL sample was studied, and the total count of each ecological group presented in Figures 6 and 7 includes pollen sum. Arboreals and non-arboreals along with aquatic pollen, monocots, pteridophytic spores and the ubiquitous taxa were included in pollen sum because these perhaps grew in the near vicinity of the studied site and their relative percentages provide insight into its equilibrium with the prevailing climatic conditions in the past (Figures 6 and 7). The 650 mesh (10 µm) fraction was collected and mounted on a glass slide in glycerine jelly for palynological study under a light microscope (Olympus BX-52). Pollen atlases of Guinet (1962), Vasanthy (1976), Nayar (1990), Tissot et al. (1994), Farooqui et al. (2010) and reference slides from the Herbarium collection of the Birbal Sahni Institute of Palaeobotany were consulted for identification of the palynomorphs. The phytolith cells were observed in the palynological slides and only qualitative report is presented in the present work. The quantitative analysis through standard methodology is in progress and will be published later.

Results

Lithology

Based on the colour and sediment texture, three lithological units were recognized in the core. Lithological unit I (100–47 cm) from the lower end of the sediment core is represented by black coloured (10 Y 2/1) (Munsell colour scale codes), fine sandy silt, clayey silt and silty sand, with peaty layers, overlain by lithological unit II (47–28 cm) which is brownish black (10 Y 2/3) in colour, moderately sorted, sandy silt and silty sand, with peat laminae within sub-angular to angular grains that are poorly sorted. Lithological unit III (28–0 cm) is brownish black (10 Y 3/1) in colour and consists of sandy silt alternating with silty sand laminae, with minute specks of peat with well rounded sand grains. This lithological unit contains abundant roots in varying stages of decay and root pores (Table 1, Figures 2 and 3).

Chronology

Radiocarbon dates (Figure 2, Table 2) range from 9105 yr BP (99 cm depth, ~9255 BC) to near modern (99.7 pMC at 14 cm depth, ~AD 1650–1955 or 300 yr BP). The plot of depth versus age indicates that the sedimentation rate from the early-Holocene to recent fluctuated, with greater sedimentation rate during 9100–7600 and less after 7600. Sedimentation rate from 7640 to 300 yr BP was fairly constant. The small variations are as follows: 0.19 mm/a from 9105 to 7640 yr BP; 0.06 mm/a from 7640 to 3215 yr BP (71–43 cm depth), 0.08 mm/a from 3215 to about 300 yr BP and about 0.4 mm/a since the last 300 years. Above 14 cm, the deposition rate is higher, with a minimum value, calculated by attributing an age of 300 yr BP at 14 cm, of 0.4 mm/a.

Stable carbon isotopes

Values of δ¹³C of organic carbon in the sediment show a general upward decrease from near −16‰ at 98–100 cm (9105 yr BP) to near −23‰ at present (Table 2). The transition can best be explained as the result of a decreasing proportion of OM derived from C₄ plants over time. Both C₃ and C₄ sedges grow at present in a peat bog in a similar tropical montane environment in the Nilgiri Hills, Tamil Nadu State, India (Rajagopalan et al., 1999).

Sediment texture

Textural analysis of clastic fraction of the Kukkal sediment core is shown in Figure 3. The sand percentage ranges from 12.5% (depth 70 cm) to 92% (depth 50 cm). The silt percentage ranges from 8% (depth 50 cm) to 75% (depth 70 cm), while the clay ranges from 0.5% (depth 50 cm) to 17% (depth 85 cm). Three main zones are apparent: below about 65 cm, with silt > sand and clay > 10%; from about 60 to 40 cm, sand predominant, clay < 5%; above 40 cm, sand and silt about equal, and clay < 5% (Figure 3). Most of the sediment core contains >25% OM (peaty in nature) consistent with a lacustrine or wetland environment. Recurring peaty sediment suggests that a wetland environment may have prevailed over much or whole of the interval represented by the sediment core, prior to construction of the dam.

Sediment chemistry

TOC content increases steadily up the section from about 3.5% at 85–100 cm to 15.8% at 1 cm unit I to unit III (Figure 3, Table 3). Similarly, TN content increases gradually from unit I 0.4% (depth 85–100 cm) to unit III 1.0–1.3% (depth 0–30 cm), and the C/N ratio increases, mimicking the TN content (unit I to unit III) gradually from about 8% (depth 85–100 cm) to 14% (depth 1 cm). A C/N ratio 13–14 suggests a sub-equal mixture of algal and vascular plant contributions (Meyers and Lallier-Vergès, 1999). Intervals of low C/N ratio (<9) occur at 12–14, 22–24, 30–34, 40–44 and 76–100 cm core depth (Figure 3). The sample at 2–4 cm probably represents the present artificial lacustrine environment and has a C/N ratio of 14, indicating that vascular and algal plant material is deposited in shallow standing water in this environment. At the depth of 40–35 cm, a concurrent decrease in OM, TOC and TN immediately follows a sharp decrease in sand deposition (Figure 3).

The vertical profile of the major elements normalized by Al versus depth (Figure 4) shows a distinct variation with depth, and the major oxide content (Table 4) shows that the lake sediments have SiO₂ varying from 31% (depth 70 cm) to 47% (depth 35 cm), with Al₂O₃ from 13% (depth 20 cm) to 22% (depth 80 cm) and Fe₂O₃ from 5% (depth 50 cm) to 9% (depth 75 cm). TiO₂ varies from 0.9% (depth 10 cm) to 1.3% (depth 40 cm). K₂O, CaO, MgO, MnO and Na₂O are present in small (<2%) amounts. The relative concentrations of the major oxides in the sediment core can be summarized as SiO₂ > Al₂O₃ > Fe₂O₃ > TiO₂ > MgO > CaO > K₂O > P₂O₅ > MnO > Na₂O. Most of the oxide trends show two prominent features: decreases in SiO₂, Al₂O₃, Fe₂O₃ and TiO₂ at 30–35 cm, accompanied by increases in P₂O₅, MnO and CaO, and at 70–80 cm, decreases in SiO₂, Al₂O₃, P₂O₅ and K₂O (Tables 3 and 4). There is an overall significant decrease of Al₂O₃ (about 10%) since 9105 yr BP (Figure 4).

In the sediment sequence of the Kukkal Lake core, the values of the CWI increase from 4–6 (characteristic of depths below 40 cm) to 8–10 (characteristic of depths above 30 cm) (Table 4). Generally, higher values of CWI indicate stronger weathering and presumably wetter conditions and a strong monsoon (Sun et al., 2010). Concurrent decreases occur in the values of the CIA, PIA and CIW between 40 and 30 cm (Table 4). The shifts in all four indices reflect an increase in CaO concurrent with a decrease in Al₂O₃ between 40 and 30 cm. Scatter plots of the CIA versus CIW and PIA values reflect extreme silicate weathering of the bedrock exposed around the lake catchment and the deposition of extremely altered sediments into the lake (Figure 5).

Figure 5.

Scatter plots of CIA versus CIW and CIA versus PIA values of the core.

Palynology

Three palynological Phases I–III were identified on the basis of relative percentages of ecological groups identified during the study. A cluster analysis of the results was carried out and the display of palynological spectrum was prepared following Grimm (1987).

Phase I (90–98 cm)

This phase shows low arboreal pollen (7.1%) and high non-arboreal pollen (16.8%). The average relative percentage of Poaceae and Cyperaceae accounts to 35.1%. Highest percentage (~40%) of pteridophytic spores (Monolete and Triletes) was recorded. The percentage of Arecaceae remained low. The arboreals constitute Acacia (1.7%), Bignoniaceae (3.6%), Elaeocarpus (2.4%), Eugenia (1.2%), Eurya (1.2%), Shorea/Hopea (2.4%); and non-arboreals comprise Malvaceae (2.1%), Apiaceae (2.2%), Artemisia (3.3%), Asteraceae (9.3%), Chenopodiaceae (0.9%). Cyperaceae (7.0%), Poaceae (28.1%), Arecaceae (8.1%), Monolete spore (16.5%) and Trilete spore (16.2%) were recorded.

Phase II (30–90 cm)

Poor count was observed in this phase with no arboreal pollen. The average relative percentages of these are Asteraceae (15.8%), Cyperaceae (4.9%), Poaceae (49.2%), Arecaceae (26.7%), Monolete spore (7.2%) and Trilete spore (5.4%).

Phase III (30–2 cm)

This phase is comparatively good in pollen/spore preservation. Total 29.1% of arboreal and 18.3% of non-arboreal pollen were recorded. Aquatic pollen account to about 11.6%. The ubiquitous taxa such as Poaceae and Cyperaceae account to 20.9%, along with pteridophytic spore showing 20.1% of the total count. Arecaceae accounts to about 4.9% of the total count. The arboreals constitute Acacia (1.7%), Bignoniaceae (3.6%), Elaeocarpus (2.3%), Eugenia (3.4%), Euphorbiaceae (4.3%), Eurya (0.5%), Shorea/Hopea (4.2%), Myrtaceae (5.8%), Syzygium (4.3%) and Terminalia (3.3%). The non-arboreals constitute Lamiaceae (2.3%), Malvaceae (1.2%), Apiaceae (2.9%), Artemisia (0.8%), Asteraceae (6.0%) and Chenopodiaceae (0.9%). The aquatic pollen present were Myriophyllum (0.7%) and Lemna (10.9%). Cyperaceae (11.9%) and Poaceae (9.0%) along with Arecaceae (4.9%), Monolete spore (8.1%) and Trilete spore (7.1%) were recorded (Figures 6 and 7).

Figure 6.

Pollen spectra in sediments of Kukkal Lake.

Figure 7.

A comparison of various palaeoclimate proxies: (a) Globigerina bulloides percent from the ODP Hole 723A, Arabian Sea 39; (b) haematite % in core MC52 in the North Atlantic; (c) Oxygen isotope record from Dongee Cave, southern China; (d) Hongyuan peat deposit; and (e) Qunf cave δ¹⁸O and (f to h) Sand, C/N and CWI of Kukkal Lake (present study).

Discussion

Palynological environment

Increased non-arboreal pollen along with highest percentage of Monolete/Trilete fern spores with low arboreal pollen in Phase I indicate moist, fluctuating and non-static climatic conditions with sparse vegetation, suggesting more open land with herbaceous taxa. The Phase II exhibits a complete absence of arboreal pollen followed by scarce non-arboreals with occurrence of high pollen of Poaceae and Arecaceae, suggesting increased open land and drier climatic conditions. Fern spores (Monoletes and Triletes) too decrease significantly, suggesting low moisture conditions. The beginning of Phase II shows total absence of Cyperaceae and pteridophytic spores, pointing towards cooler and dry 8200 yr BP global event which is also evident from different parts of Indian subcontinent (Dixit et al., 2014). The Phase III shows good percentage of arboreals and comparatively quite low non-arboreals. Poaceae increased along with Arecaceae and pteridophytic spores, suggesting the shift in climatic conditions from comparatively drier to moist conditions (strengthened precipitation) favouring the vegetation in the catchment vicinity. Overall palynological results in Phase III indicate the lacustrine ecosystem as the silt percentage increased along with high OM, TN and carbon/nitrogen ratio. Major elements too show high percentage in this phase. Phases I and II show fluctuating values with respect to all the studied proxies, suggesting fluctuating climatic conditions, but Phase III indicates static environmental conditions and the existence of lake in the studied area perhaps since ~3000 yr BP (Figure 6).

The increase in fern spores and their ecology, however, remains poorly understood because of the sustained lack of information on the gametophyte generation (Limm and Dawson, 2010). It is unsurprising that we know little about the factors influencing the decline/abundance of fern species (Bucharová et al., 2010; Kelsall et al., 2004; Testo and Watkins, 2013). By studying the development, reproductive biology and stress ecophysiology of fern gametophyte generation, it is noted that at both 20°C and 25°C temperatures, the sporophyte formation was affected (Limm and Dawson, 2010). It generally adapts itself from well to cooler (temperate) moist climate and also has the capacity to transfer moisture from the fog within the forest ecosystem even when the precipitation is comparatively low. In this study, it is observed that high percentage of fern spores (both Trilete and Monolete spores) correspond to the time period when relatively high arboreal/non-arboreals are recorded. Thus, it is inferred that high percentage of fern spores recorded here indicate an increase in moisture level with relatively cooler climate enhancing the stabilization of forest cover around the lake. An integration of all the proxies studied is presented below.

Sedimentary environment

The chronology of the core, and inferred measurements of sedimentation rate, indicates a sedimentary environment that has been continuous and close to stable in this small basin for the past 9000 years. However, high proportion of sand deposited between 9100 and 7600 (60–45 cm) interval indicates a sediment flux from the nearby catchment area and/or faster accumulation rates and less after 7600 yr BP, and a decrease in sand flux after 1700 yr BP probably because of drier conditions. Peat-like deposits of OM are consistent with a wetland environment like those present elsewhere in montane valleys of the tropical Western Ghats. Variations in the C/N ratio of OM may indicate changes in the abundance of algae in the basin and, by implication, changes in the amount of water depth in the lake (Figure 3). The basin may have been lacustrine for relatively brief intervals corresponding to 12–14, 22–24, 30–34, and 40–44 cm, in addition to a longer interval corresponding to 76–100 cm in the core (Figure 4). The present-day lacustrine environment is because of damming. Superimposed on this sedimentary system are changes revealed by our climate proxy measurements – both unidirectional steady changes and abrupt events.

Unidirectional changes

The chemistry of the OM in the core has evolved steadily over the past 9000 years. TOC, TN and C/N all show an upward increase in the core (Figure 3). The values of δ¹³C also suggest an increase. A progression from mainly C₄ to mainly C₃ species in the basin is indicated, and this is in support with an increasingly wet climate. Increasing rainfall may also be responsible for the increasing amount of OM over time, preservation being more likely when the basin sediments remain wet for long periods. The CaO/MgO ratio in the silicate fraction also increases towards the top of the core section, and Al₂O₃ decreases.

Stepwise changes

Several proxies show stepwise changes in the range of 30–40 cm. Major oxide measurements change in such a way that all four chemical weathering indices calculated for this study shift to new values corresponding to more intense weathering over that depth interval. Fe₂O₃, MnO and P₂O₅ also show shifts. The changes in chemical indices are concurrent with increases in Ca/Al and Si/Al (Figure 4), but there is no change in the relative abundance of clay (Figure 3). Proportions of Asteraceae, Poaceae and Palmae or Arecaceae decrease markedly in the record over this depth interval (Figure 6). These changes immediately follow a change in the character of clastic sedimentation from sand-dominant to sand and silt in approximately equal proportions. They are succeeded rapidly by a further change in the pollen record, namely, enhanced preservation of arboreal pollen. Ferns increase in abundance earlier, at depths of 46–40 cm in the core. This change is at ca. 3000 yr BP (at 40 cm) to 1800 yr BP (at 30 cm). A stepwise change similar to that reflected in enhanced preservation of pollen and the appearance of diatoms just above 30 cm in the Kukkal core and occurred at 1500 yr BP in Pookode Lake, Kerala, at an altitude of 770 m.a.s.l (Veena et al., 2014).

As noted before, the shifts in the chemical indices appear to relate to increased deposition of CaO as silicate minerals, and concurrent increase in Si/Al. A change in clay mineralogy, with deposition of Ca-smectite above 30 cm, is a possible way of accounting for the change. The increase in CWI indicates a greater intensity of weathering, but deposition of sediment bearing this signature is likely to have lagged behind the climate stimulus responsible for the change. Together, most of the observations suggest a period of increasing rainfall. The changes may even have begun as early as 6000 yr BP, when sand began to dominate the clastic sediment, and have led to more abundant fern vegetation by 3500 yr BP, and this forest vegetation largely replacing grasslands with palms by 1700 yr BP. This change might account for the decrease in sand deposition at about 1700 yr BP, if the change to forest vegetation led to less energetic runoff.

The increase in sand deposition near 6000 yr BP could be a separate stepwise event, independent of the events beginning around 3500 yr BP. A further stepwise event is indicated at 90 cm in the core (about 8700 yr BP), prior to which some arboreal pollen and fern spores were preserved. Above this horizon, the relative abundances of sand, silt and clay change abruptly towards the top in the sediment core. Major oxide trends show inflections at about 80 cm, possibly resulting from the same event, with a lag reflecting the time required for weathering chemistry to respond.

Human effects

The higher sediment deposition rate corresponds to events of the last 350 years (17–18 cm in depth to the core top) and is the only evidence consistent with human-caused changes in the lake catchment. As noted above, abandoned paddy terraces are present along the east side of the catchment. Agriculture may have become established in this montane site since the early 17th century. Cutting of trees in the forested part of the catchment may also have contributed to the increased sedimentation rate.

Regional and global correlations

The Kukkal lacustrine/wetland basin has responded to hydrological changes, principally changes in the amount of precipitation from the SWM, over the 9000-year record presented here. Several other proxy records of similar length are available for comparison in the region of the SWM and the EAM. Oxygen isotope records from radiometrically dated speleothems (Cheng et al., 2009; Fleitmann et al., 2003, 2004, 2007; Liu et al., 2013; Morrill et al., 2013a, 2013b) show a general unidirectional decrease in δ¹⁸O since about 8000 yr BP (Figure 7). The trend is interpreted by the authors, on the basis of the isotope amount effect, as evidence for a gradual increase in intensity of the SWM and EAM, and is consistent with the unidirectional changes at Kukkal, also interpreted as evidence of an increase in monsoon intensity. Similar unidirectional changes occur in another proxy, for example, the abundance of Globigerina bulloides in seafloor sediment in the Arabian Gulf near Oman (Gupta et al., 2003). The general increase in SWM rainfall is consistent with the results of Yadava and Ramesh (2005), who interpreted a speleothem record as indicating the highest rainfall around AD 1666 (‘LIA’).

The speleothem records also show a trough in the δ¹⁸O between 8500 and 8000 yr BP, interpreted in part as a decrease in monsoon intensity beginning at 8200 yr BP and lasting 100–150 years. An event close to this time is present in the Kukkal sediments at an interpolated age of 8700 yr BP, where it is marked by the replacement of savanna grassland and inflections in the curves of major oxide abundances. Given the uncertainties of the age interpolation, the proxy records may all be recording the same event.

At Kukkal, the more pronounced stepwise event between 3000 and 1800 yr BP has no clear counterpart in the speleothem or Globigerina bulloides records (Figure 7). Possible counterpart events in climate records from the Indian subcontinent include the warmer and humid event at 1500 yr BP at Pookode Lake, Kerala (Veena et al., 2014) and speleothem δ¹⁸O record from Orissa, India (Yadava and Ramesh, 2005). In the case of the latter, the authors interpret the record as indicating increasing rainfall up until 1500 yr BP, followed by a period of decreasing rainfall that is not evident at Kukkal. A high-rainfall phase at 3000–2300 yr BP is recorded in the discharge of the Indus River (Von Rad et al., 1999a, 1999b). However, this is at variance with a long drought in the Thar Desert, India, between 3600 and 2000 yr BP (Bryson and Bryson, 1996; Enzel et al., 1999; Swain et al., 1983) or aridity and weakening of the summer monsoon between 4500 and 2000 yr BP in the continental records (Swain et al., 1983). The lake records show a maximum Holocene wet period from 6000 to 4000 yr BP and a change to a drier climate after ~4000 yr BP. It is possible that some of the moisture at these sites in the Thar desert received the winter precipitation, which might explain as to why they are not in synch with other monsoon sites that respond to summer precipitation.

Climatic events of global scale during the last 3000 years include the ‘Roman Warm Period’ (‘RWP’) (2500–1600 yr BP (Wang et al., 2012, 2013), the ‘MWP’ (790–620 yr BP) and the ‘LIA’ (330–80 yr BP) (Bianchi and McCave, 1999). These phenomena have been proposed on the basis of multi-proxy studies in records from higher latitudes (e.g. Denniston et al., 2000; Lehmkuhl, 1997; Liu et al., 2013; Sukumar et al., 1993; Von Rad et al., 1999a, 1999b; Zhou et al., 1991, 2011). Crowley and Lowery (2000) described the ‘LIA’ as a mild cooling event in the Northern Hemisphere. Hu et al. (2001, 2008) determined that the ‘LIA’ led to ~1.7°C of cooling around Alaska (Keigwin, 1996). The ‘LIA’, therefore, has a large temporal and spatial variability across the globe (Intergovenrmental Panel on Climate Change (IPCC), 2001). The stepwise event at 3000–1800 yr BP in the Kukkal proxies appear to anticipate the ‘RWP’ by a few hundred years, but the interpolated dates are only approximate; the inflections in the curves may in fact correspond with the initiation of the ‘RWP’. There are no events at Kukkal that correspond convincingly to the ‘MWP’ and the ‘LIA’. Even though Zonneveld et al. (1997) suggested a rapid teleconnection between the climates of high latitude and the SWM, the only evidence of such relationships in the Kukkal core is the stepwise event that may correspond to the beginning of the ‘RWP’.

A possible scenario for Kukkal

When comparing and contrasting climatic proxy records from different environments, it must be borne in mind that each environment will have its own set of responses to a particular climatic forcing. Terrestrial records involving plant associations, in particular, may have much longer response times than the cave speleothems and marine sediments. In the case of the Kukkal record, it is possible to conceive of the evolution of the basin as the protracted response to a century or two of dry climate at about 8000 yr BP, followed by increasing rainfall. Savanna is initially replaced by grassland with palms, sedges and ferns appear later, and all are eventually succeeded by forest in a process that lasts several thousand years. Evolving vegetation and the amount of rainfall control the weathering and the character and availability of clastic sediments. Eventual establishment of mature Shola forest limits the energy of runoff and the availability of coarser clastics. Pollen preservation is poor following the period of dry climate, and recovers only when the SWM intensity increases more rapidly, augmenting the amount of standing water in the wetland/lake basin.

Conclusion

In summary, the following points emerge from this study:

The Kukkal sediment core reflects lacustrine or wetland depositional environments of early-Holocene age (9000 yr BP) to modern age. Four radiocarbon ages were used for producing an age model showing a slow average sedimentation rate of 0.11 mm/yr, except in the upper 14 cm of core, where the rate was 0.4 mm/yr or greater. The increased rate may reflect human activity because of deforestation, urbanization and agricultural practices around the lake.

Rapid climate change is indicated by the degree of preservation of OM and changes in its composition. Superimposed on this trend are two stepwise changes, at about 8000 yr BP and between about 3000 and 1800 yr BP. The former change appears to correspond to a dry spell of 100–150 years recorded elsewhere in India. The latter corresponds to a rapid intensification of the SWM, and may correlate with the initiation of the ‘RWP’, the two phenomena coinciding in time. The ‘MWP’ and the ‘LIA’ are not clearly represented in the record.

The C/N ratio of Kukkal Lake sediments ranges from 14.02 to 8.31, indicating that the organic carbon originates from a mixture of lacustrine algae, vascular plants and terrestrial plants. Stable carbon isotope data suggest early dominance of C₄ plants such as sedges, and a progressive increase in the ratio of C₃ vegetation.

Chemical weathering indices indicate extreme silicate weathering over the past 9000 years, with a stepwise increase in intensity beginning about 3000 yr BP. Palynological results reveal a succession of vegetation types, beginning with savanna prior to about 8000 yr BP, succeeded by grassland with palms and Asteraceae species. A vegetational change is observed just prior to 3200 yr BP, with ferns being abundant and the Shola forest established between 3200 and 1800 yr BP.

Footnotes

Funding

The author(s) received no financial support for the research, authorship and/or publication of this article.

References

Achyuthan

Kar

Eastoe

(2007a) Late Quaternary environmental changes and evolution of the Tal Chappar salt lakes, Rajasthan. Journal of Paleolimnology 38: 493–507.

Achyuthan

Nagasundaram

Gourlan

. (2014) Mid-Holocene Indian summer monsoon variability off the Andaman Islands, Bay of Bengal. Quaternary International 349: 232–244.

Achyuthan

Quade

Roe

. (2007b) Stable isotopes of pedogenic carbonates from the eastern margins of the Thar Desert, Rajasthan. Quaternary International 162: 50–60.

Allchin

(1997) Origins of a Civilization: The Prehistory and Early Archaeology of South Asia. New York: Viking.

Kutzbach

. (2000) Asynchronous Holocene optimum of the East Asian monsoon. Quaternary Science Reviews 19: 743–762.

Anderson

Overpeck

Gupta

(2002) Increase in the Asian, Southwest Monsoon during the past four centuries. Science 297(5581): 596–599.

Bagyaraj

Gurugnanam

(2011) Significance of morphometry studies, soil characteristics, erosion phenomena and landform processes using remote sensing and GIS for Kodaikanal Hills, a global biodiversity hotpot in Western Ghats, Dindigul District, Tamil Nadu, South India. Research Journal of Environmental and Earth Sciences 3(3): 221–233.

Bera

Farooqui

(2000) Mid-Holocene vegetation and climate of South Indian Montane. Journal of the Palaeontological Society of India 45: 49–56.

Bianchi

McCave

(1999) Holocene periodicity in North Atlantic climate and deep-ocean flow south of Iceland. Nature 397: 515–517.

10.

Bryson

(1996) Application of a global volcanicity time-series on high-resolution paleoclimatic modeling of the eastern Mediterranean. Water Science and Technology Library 31: 1–19.

11.

Bucharová

Münzbergová

Tájek

(2010) Population biology of two rare fern species: Long life and long-lasting stability. American Journal of Botany 97: 1260–1271.

12.

Carver

(1971) Procedure in Sedimentary Petrology. Chicago, IL: Wiley Interscience.

13.

Chamyal

Maurya

Bhandari

. (2002) Late Quaternary geomorphic evolution of the lower Narmada Valley, western India: Implications for neotectonic activity along the Narmada-Son Fault. Geomorphology 46: 177–202.

14.

Chauhan

Vogelsang

(2006) Climate induced changes in the circulation and dispersal patterns of fluvial sources during late Quaternary in the middle Bengal fan. Journal of Earth System Science 115: 379–386.

15.

Chauhan

Shiva

Suneethi

(2004) Fluvial influx and weathering history of the Himalayas since Last Glacial Maximum – Isotopic, sedimentological and magnetic records from the Bay of Bengal. Current Science 87: 509–515.

16.

Chauhan

Vogelsang

Basavaiah

. (2010) Reconstruction of the variability of SW monsoon during the past 3 ka from the continental margin of the south eastern Arabian Sea. Journal of Quaternary Science 25(5): 798–807.

17.

Cheng

Fleitmann

Edwards

. (2009) Timing and structure of the 8.2 kyr BP event inferred from δ¹⁸O records of stalagmites from China, Oman and Brazil. Geology 37: 1007–1010.

18.

Clift

Plumb

(2008) The Asian Monsoon: Causes, History and Effects. London: Cambridge University Press.

19.

Cronin

Dwyer

Kamiya

. (2003) ‘Medieval Warm Period’, ‘Little Ice Age’ and 20th century temperature variability from Chesapeake Bay. Global and Planetary Change 36: 17–29.

20.

Crowley

(2000) Causes of climatic change over the past 1000 years. Science 289: 270–277.

21.

Crowley

Lowery

(2000) How warm was the ‘Medieval Warm Period?’ Ambio 29: 51–54.

22.

Denniston

Gonzalez

Asmerom

. (2000) Speleothem evidence for changes in Indian summer monsoon precipitation over the last similar to 2300 yr. Quaternary Research 53: 196–202.

23.

Dixit

David

Hodell

. (2014) Abrupt weakening of the Indian summer monsoon at 8.2 kyr B.P. Earth and Planetary Science Letters 391: 16–23.

24.

Duplessy

(1982) Glacial to interglacial contrast in the northern Indian Ocean. Nature 295: 494–498.

25.

Duzgoren-Aydin

Aydin

Malpas

(2002) Re-assessment of chemical weathering indices: Case study on pyroclastic rocks of Hong Kong. Engineering Geology 63: 99–119.

26.

Engstrom

Wright

Jr (1984) Chemical stratigraphy of lake sediments as a record of environmental change. In: Hawroth

Lund

JWG

(eds) Lake Sediments and Environmental History. Leicester: Leicester University Press, pp. 11–69.

27.

Enzel

Ely

Mishra

. (1999) High resolution Holocene environmental changes in the Thar Desert, northwestern India. Science 284: 125–127.

28.

Esper

Cook

Schweingruber

(2002a) Low-frequency signals in long tree-ring chronologies and the reconstruction of past temperature variability. Science 295: 2250–2253.

29.

Esper

Schweingruber

Winiger

(2002b) 1300 yr of climatic history for western Central Asia inferred from tree-rings. The Holocene 12: 267–277.

30.

Esper

Shiyatov

Mazepa

. (2003) Temperature-sensitive Tien Shan tree ring chronologies show multi-centennial growth trends. Climate Dynamics 21: 699–706.

31.

Fægri

Iversen

(1989) Textbook of Pollen Analysis. Chichester: Wiley.

32.

Farooqui

Sekhar

(2011) Climate change and vegetation succession in Lalitpur area, Uttar Pradesh (India) during late Holocene. Tropical Ecology 52(1): 69–77.

33.

Farooqui

Ray

Farooqui

. (2010) Palynodiversity, sea level and climate prior to 40 ka in Indian Peninsula, Kerala. Quaternary International 213: 2–11.

34.

Fedo

Nesbitt

Young

(1995) Unravelling the effects of potassium metasomatism in sedimentary rocks and paleosols, with implications for paleoweathering conditions and provenance. Geology 23: 921–924.

35.

Fleitmann

Burns

Mangini

. (2007) Holocene ITCZ and Indian monsoon dynamics recorded in stalagmites from Oman and Yemen (Socotra). Quaternary Science Reviews 26: 170–188.

36.

Fleitmann

Burns

Mudelsee

. (2003) Holocene forcing of the Indian monsoon recorded in a stalagmite from southern Oman. Science 300: 1737–1739.

37.

Fleitmann

Burns

Neff

. (2004) Palaeoclimatic interpretation of high-resolution oxygen isotope profiles derived from annually laminated speleothems from southern Oman. Quaternary Science Reviews 23: 935–945.

38.

Gaudette

Flight

Toner

. (1974) An inexpensive titration method for the determination of organic carbon in recent sediments. Journal of Sedimentary Petrology 44: 249–253.

39.

Gonzalez-Carranza

Hooghiemstra

Vélez

(2012) Major altitudinal shifts in Andean vegetation on the Amazonian flank show temporary loss of biota in the Holocene. The Holocene 22(11): 1227–1241.

40.

Grimm

(1987) CONISS: A FORTRAN 77 program for stratigraphically constrained cluster analysis by the method of incremental sum of squares. Computers & Geosciences 13: 13–35.

41.

Guinet

(1962) Pollens d’Asie tropicale. Travaux de la Section Scientifique et Technique, Rapport Institut Francais de Pondichéry, Insitut Français de Pondichéry 5(1): 1–8.

42.

Gupta

(2004) Origin of agriculture and domestication of plants and animals linked to early Holocene climate amelioration. Current Science 87(1): 54–59.

43.

Gupta

Anderson

Overpeck

(2003) Abrupt changes in Asian southwest monsoon during the Holocene and their links to the North Atlantic Ocean. Nature 421: 354–356.

44.

Gupta

Das

Anderson

(2005) Solar influence on the Indian summer monsoon during the Holocene. Geophysical Research Letters 32: 2–5.

45.

Gupta

Anderson

Pandey

. (2006) Adaptation and human migration, and evidence of agriculture coincident with changes in the Indian summer monsoon during the Holocene. Current Science 90: 1082–1090.

46.

Harnois

(1988) The CIW index: A new chemical index of weathering. Sedimentary Geology 55: 319–322.

47.

Henderson

Huang

. (2008) Quantification of Holocene Asian monsoon rainfall from spatially separated cave records. Earth and Planetary Science Letters 266(3): 221–232.

48.

Ito

Brown

. (2001) Pronounced climatic variations in Alaska during the last two millennia. Proceedings of the National Academy of Sciences of the United States of America 98(19): 10552–10556.

49.

Intergovenrmental Panel on Climate Change (IPCC) (2001) Climate change 2001: The scientific basis. In: Houghton

Ding

Griggs

. (eds) Contribution of Working Group I to the Third Assessment Report of the Intergovernmental Panel on Climate Change. Cambridge: Cambridge University Press, pp. 1–373.

50.

Kaufman

Ager

Anderson

. (2004) Holocene thermal maximum in the western Arctic (0–180 W). Quaternary Science Reviews 23(5–6): 529–560.

51.

Keigwin

(1996) The ‘Little Ice Age’ and ‘Medieval Warm Period’ in the Sargasso Sea. Science 274: 1504–1508.

52.

Kelsall

Hazard

Leopold

(2004) Influence of climate factors on demographic changes in the New York populations of the federally-listed Phyllitis scolopendrium (L.) Newm. var. americana. Journal of the Torrey Botanical Society 131: 161–168.

53.

Kudrass

Hofman

Doose

. (2001) Modulation and amplification of climatic changes in the Northern Hemisphere by the Indian summer monsoon during the past 80 ky. Geology 29: 63–66.

54.

Kumaran

KPN

Limaye

Nair

. (2008) Palaeoecological and palaeoclimate potential of subsurface palynological data from the late Quaternary sediments of south Kerala Sedimentary Basin, southwest India. Current Science 95(4): 515–526.

55.

Lehmkuhl

(1997) Late Pleistocene, late glacial and Holocene glacier advances on the Tibetan Plateau. Quaternary International 38–39: 77–83.

56.

Limaye

Kumaran

KPN

Nair

. (2010) Cyanobacteria as potential biomarkers of hydrological changes in the late Quaternary sediments of south Kerala Sedimentary Basin, India. Quaternary International 213: 79–90.

57.

Limm

Dawson

(2010) Polystichum munitum (dryopteridaceae) varies geographically in its capacity to absorb fog water by foliar uptake within the redwood forest ecosystem. American Journal of Botany 97(7): 1121–1128.

58.

Liu

Henderson

. (2013) Links between the East Asian monsoon and North Atlantic climate during the 8200 year event. Nature Geoscience 6: 117–120.

59.

Loring

Rantala

RTT

(1992) Manual for the geochemical analyses of marine sediments and suspended particulate matter. Earth-Science Reviews 32: 235–283.

60.

Lückge

Doose-Rolinski

Khan

. (2001) Monsoonal variability in the northeastern Arabian Sea during the past 5000 yr: Geochemical evidence from laminated sediments. Palaeogeography, Palaeoclimatology, Palaeoecology 167: 273–286.

61.

Mackereth

FJH

(1966) Some chemical observations on post glacial lake sediments. Philosophical Transactions of the Royal Society of London, Series B: Biological Sciences 250: 165–213.

62.

Maslov

Krupenin

Gareev

(2003) Lithological, lithochemical, and geochemical indicators of paleoclimate: Evidence from riphean of the southern Urals. Lithology and Mineral Resources 38: 427–446.

63.

Mayewski

Rohling

Stager

. (2004) Holocene climate variability. Quaternary Research 62: 243–255.

64.

Meher-Homji

(1967) Phytogeography of a south Indian Hill station. Bulletin of the Torrey Botanical Club 94: 230–242.

65.

Meher-Homji

Gupta

(1999) A critical appraisal of vegetation and climate changes during Quaternary in the Indian region. Proceedings of the Indian National Science Academy, Part B: Biological Sciences 65(5): 205–244.

66.

Meyers

Lallier-Vergès

(1999) Lacustrine sedimentary organic matter records of late Quaternary paleoclimate. Journal of Paleolimnology 21: 345–372.

67.

Minyuk

Borkhodoevl

Wennrich

(2014) Inorganic geochemistry data from Lake El’gygytgyn sediments: Marine isotope stages 6–11. Climate of the past 10: 467–485.

68.

Morrill

Anderson

Bauer

. (2013a) Proxy benchmarks for intercomparison of 8.2 ka simulations. Climate of the past 9: 423–432.

69.

Morrill

LeGrande

Renssen

. (2013b) Model sensitivity to North Atlantic freshwater forcing at 8.2 ka. Climate of the past 9: 955–968.

70.

Nagasundaram

(2014) Middle to late-Holocene paleoenvironmental changes inferred from a marine sediment core off Landfall Island, Bay of Bengal. Unpublished PhD Thesis, Anna University, 136 pp.

71.

Nagasundaram

Achyuthan

Ahmad

(2014) Monsoonal changes inferred from the Middle to late-Holocene sediments off Landfall Island, North Andaman. Arabian Journal of Geosciences 7: 3513–3523.

72.

Nayar

(1990) Pollen Flora of Maharashtra State, India. New Delhi: Today and Tomorrow’s Printers and Publishers, 150 pp.

73.

Nesbitt

Young

(1982) Early Proterozoic climates and plate motions inferred from major element chemistry of lutites. Nature 199: 715–717.

74.

Overpeck

Anderson

Trumbore

. (1996) The southwest Indian monsoon over the last 18,000 yr. Climate Dynamics 12: 213–225.

75.

Piper

(1947) Soil and Plant Analysis. Adelaide: University of Adelaide Press, 368 pp.

76.

Ponton

Giosan

Eglinton

. (2012) Holocene aridification of India. Geophysical Research Letters 39(3): L3407. Available at: http://dx.doi.org/10.1029/2011GL050722.

77.

Possehl

(1996) Mehrgarh. In: Fagan

(ed.) Oxford Companion to Archaeology. Oxford: Oxford University Press, pp. 436–438.

78.

Prell

Kutzbach

(1987) Monsoon variability over the past 150,000 yr. Journal of Geophysical Research 82: 8411–8425.

79.

Prell

Van Campo

(1986) Coherent response of Arabian Sea upwelling and pollen transport to late Quaternary monsoonal winds. Nature 323: 526–528.

80.

Price

Velbel

(2003) Chemical weathering indices applied to weathering profiles developed on heterogeneous felsic metamorphic parent rocks. Chemical Geology 202: 397–416.

81.

Rajagopalan

Ramesh

Sukumar

(1999) Climatic implications of δ¹³C and δ¹⁸O ratios from C₃ and C₄ plants growing in a tropical montane habitat in southern India. Journal of Bioscience 24: 491–498.

82.

Rawat

Gupta

Sangode

. (2015) Late Pleistocene–Holocene vegetation and Indian summer monsoon record from the Lahaul, northwest Himalaya, India. Quaternary Science Reviews 114: 167–181.

83.

Sant

Wadhavan

Ganjoo

. (2011) Linkage of Proglacial processes from Last glacial to recent inferred form spity sequence, Ley valley, Ladakh Himalaya. Journal of the Geological Society of India 74: 147–156.

84.

Shah

Bhattacharya

(2009) Tree-ring analysis of sub-fossil woods of Pinus wallichiana from Ziro Valley, Arunachal Pradesh, northeast Himalaya. Journal of the Geological Society of India 74: 503–508.

85.

Sharma

Joachimski

Tobschall

. (2006) Correlative evidence of monsoon variability, vegetation change and human habitation in Senai Lake deposit, Ganga plain. Current Science 90: 973–978.

86.

Shi

Kong

Wang

(1993) The important climatic fluctuations and events in China during the Holocene megathermal. Science in China, Series B 23: 865–873.

87.

Singhvi

Kale

(2008) Palaeoclimate Studies in India: Last Ice Age to the Present (IGBP-WCRP-SCOPE Report Series: 4). New Delhi: Indian National Science Academy.

88.

Singhvi

Bhatt

Glennie

. (2012) India, Arabia and adjacent regions. In: Metcalfe

Nash

(eds) Quaternary Environmental Change in the Tropics. New York: Wiley, pp. 151–206.

89.

Singhvi

Rupakumar

Thamban

. (2010) Instrumental, terrestrial and marine records of the climate of South Asia during the Holocene: Present status, unresolved problem and societal aspects. In: Mitra

Sharma

(eds) Global Environmental Change in South Asia: A Regional Perspective. New Delhi: Capital Publications, pp. 54–124.

90.

Sinha

Cannariato

Stott

. (2007) A 900-year (600 to 1500 A.D.) record of the Indian summer monsoon precipitation from the core monsoon zone of India. Geophysical Research Letters 34(16): 1–5.

91.

Sirocko

Garbe-Schönberg

Devey

(2000) Processes controlling trace element geochemistry of Arabian Sea sediments during the last 25,000 yr. Global and Planetary Change 26: 217–303.

92.

Sirocko

Garbe-Schösnberg

Mclntyre

. (1996) Teleconnections between the subtropical monsoons and high-latitude climates during the last deglaciation. Science 272: 526–529.

93.

Sirocko

Sarnthein

Erlenkeuser

. (1993) Century scale events in monsoon climate over the past 24,000 yr. Nature 364: 322–324.

94.

Staubwasser

Sirocko

Grootes

. (2002) South Asian monsoon climate change and radiocarbon in the Arabian Sea during early and middle Holocene. Paleoceanography 17(4): 1063.

95.

Staubwasser

Sirocko

Grootes

. (2003) Climate change at the 4.2 ka BP termination of the Indus valley civilization and Holocene south Asian monsoon variability. Geophysical Research Letters 30: 1425.

96.

Stuiver

Reimer

(1993) Extended ¹⁴C data base and revised CALIB 3.0 ¹⁴C age calibration program. Radiocarbon 35: 215–230.

97.

Stuiver

Reimer

(2010) CALIB 5.0.2, WWW program and documentation. Available at: www.calib.qua.ac.uk.

98.

Sukumar

Ramesh

Pant

. (1993) A δ¹³C record of late Quaternary climate change from tropical peats in southern India. Nature 364: 703–706.

99.

Sun

(2012) The quantitative reconstruction of the paleoclimate between 5200 and 4300 cal BP in the Tianshui Basin, NW China. Climate of the past 8: 625–636.

100.

Sun

Ding

Dai

(2010) Changing links between South Asian summer monsoon circulation and tropospheric land-sea thermal contrasts under a warming scenario. Geophysical Research Letters 37: L02704.

101.

Swain

Kutzbach

Hastenrath

(1983) Estimates of Holocene precipitation for Rajasthan, India, based on pollen and lake-level data. Quaternary Research 19: 1–17.

102.

Testo

Watkins

Jr (2013) Understanding mechanisms of rarity in pteridophytes: Competition and climate change threaten the rare fern Asplenium scolopendrium var. americanum (Aspleniaceae). American Journal of Botany 100(11): 2261–2270.

103.

Thamban

Rao

Schneider

(2002) Reconstruction of late Quaternary monsoon oscillations based on clay mineral proxies using sediment cores from the western margin of India. Marine Geology 186: 527–539.

104.

Tissot

Chikhi

Nayar

(1994) Pollen of Wet Evergreen Forests of the Western Ghats, India. Pondichéry: Institut Français; All India Press, 132 pp.

105.

Tiwari

Ramesh

Bhushan

. (2006) Paleoproductivity variations in the equatorial Arabian Sea: Implications for east African and Indian summer rainfalls and the El Nino frequency. Radiocarbon 48: 17–29.

106.

Tiwari

Ramesh

Somayajulu

BLK

. (2005) Solar control of southwest monsoon on centennial time scales. Current Science 89(9): 1583–1589.

107.

Van Campo

Cour

Hang

(1996) Holocene environmental changes in Bangong Co basin (western Tibet), part 2: The pollen records. Palaeogeography, Palaeoclimatology, Palaeoecology 120: 49–63.

108.

Vasanthy

(1976) Pollens des montagnes du Sud de l’Indie. Travaux de la Section Scientifique et Technique. Rapport Insitut Français de Pondichéry 15: 1–74.

109.

Veena

Achyuthan

Eastoe

. (2014) A multi-proxy reconstruction of monsoon variability in the late-Holocene, south India. Quaternary International 325: 63–73.

110.

Verschuren

(2003) Lake-based climate reconstruction in Africa: Progress and challenges. Hydrobiologia 500: 315–330.

111.

Verschuren

Laird

Cumming

(2000) Rainfall and drought in equatorial east Africa during the past 1100 yr. Nature 403: 410–414.

112.

Von Rad

Schaaf

Michels

. (1999a) A 5000-yr record of climate change in varved sediments from the oxygen minimum zone off Pakistan, northeastern Arabian Sea. Quaternary Research 51: 39–53.

113.

Von Rad

Schulz

Riech

. (1999b) Multiple monsoon controlled breakdown of oxygen-minimum conditions during the past 30,000 yr documented in laminated sediments off Pakistan. Palaeogeography, Palaeoclimatology, Palaeoecology 152: 129–161.

114.

Wang

Surge

Mithen

(2012) Seasonal temperature variability of the Neoglacial (3300–2500 BP) and Roman Warm Period (2500–1600 BP) reconstructed from oxygen isotope ratios of limpet shells (Patella vulgata), northwest Scotland. Palaeogeography, Palaeoclimatology, Palaeoecology 317–318: 104–113.

115.

Wang

Surge

Walker

(2013) Seasonal climate change across the Roman Warm period/Vandal Minimum transition using isotope sclerochronology in archaeological shells and otoliths, southwest Florida, USA. Quaternary International 308–309: 230–241.

116.

Wang

Cheng

Edwards

. (2005) The Holocene Asian monsoon: Links to solar changes and North Atlantic climate. Science 308: 854–857.

117.

Xia

Y-M

Wang

P-F

(2001) Peat record of climate change for the last 3000 yr in Yangmu, Mishan region of Sanjiang plain. Journal of Geographical Sciences 11(4): 454–460.

118.

Yadav

Singh

(2002) Tree-ring-based spring temperature patterns over the past four centuries in western Himalaya. Quaternary Research 57: 299–305.

119.

Yadava

Ramesh

(2005) Monsoon reconstruction from radiocarbon dated tropical Indian speleothems. The Holocene 15: 48–59.

120.

Yudovich

Ketris

(2011) Geochemical Indicators of Lithogenesis. Syktyvkar: Geoprint, 742 pp.

121.

Zhou

Man

. (2011) A comparison of the ‘Medieval Warm Period’, ‘Little Ice Age’ and 20th century warming simulated by the FGOALS climate system model. Chinese Science Bulletin 56(28): 3028–3041.

122.

Zhou

Head

. (1991) ¹⁴C dating and measurements of climatic proxy indices of loess sequence to record paleomonsoon variation on the Loess Plateau of China during the last 16000 yr (abstr). Radiocarbon 33: 259.

123.

Zonneveld

KAF

Ganssen

Troelstra

. (1997) Mechanisms forcing abrupt fluctuations of the Indian Ocean summer monsoon during the last deglaciation. Quaternary Science Reviews 16: 187–201.