Palaeogeography, Palaeoclimatology, Palaeoecology 265 (2008) 148–158

Contents lists available at ScienceDirect

Palaeogeography, Palaeoclimatology, Palaeoecology

j ourna l homepage: www.e lsev ie r.com/ locate /pa laeo

Cooling-driven climate change at 12–11 Ma: Multiproxy records from a longfluviolacustrine sequence at Guyuan, Ningxia, China

Hanchao Jiang a,⁎, Junliang Ji b, Ling Gao c, Zihua Tang a, Zhongli Ding a

a Key Laboratory of Cenozoic Geology and Environment, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing, 100029, Chinab Key Laboratory of Biogeology and Environmental Geology, Ministry of Education, Wuhan, 430074, Chinac Department of Geological Sciences, Indiana University, 1001 East 10th Street, Bloomington, IN 47405-1405, USA

⁎ Corresponding author. Tel.: +86 10 82998379; fax: +E-mail addresses: hanchaojiang@mail.igcas.ac.cn,

hanchaojiang@hotmail.com (H. Jiang).

0031-0182/$ – see front matter © 2008 Elsevier B.V. Aldoi:10.1016/j.palaeo.2008.05.006

A B S T R A C T

A R T I C L E I N F O

Article history:

To explore the evolution o Received 26 January 2008Received in revised form 2 May 2008Accepted 17 May 2008

Keywords:Mid-Miocene coolingClimate changeMultiproxy continental recordsSikouziNingxia

f climate through the Miocene, especially during the Middle Miocene climatetransition, multiproxy continental records were analyzed by susceptibility (SUS), lightness (L⁎), totalinorganic carbon (TIC) and total organic carbon (TOC) content from the 2900-m-thick fluviolacustrinesediment sequence at Guyuan, Ningxia, China and compared with the redness (a⁎) and the pollen humidityindices from the same section. The published Miocene records of sedimentation rate, sediment grain size,biota turnover, and palaeoproductivity from other regions were also compiled and compared. Results showthat multiproxies of the Sikouzi section have changed substantially since 12–11 Ma ago. SUS, L⁎ and TIC haveincreased while a⁎, TOC and the pollen humidity index decreased, all implying that the palaeoclimate in thestudy area has got cooler and drier since that time.This climate change also left imprints in many other regions, probably resulting from global cooling and thedevelopment of the East Antarctic Ice Sheet since about 14 Ma. Several positive feedback mechanisms areproposed to have effectively modulated and magnified the Mid-Miocene global cooling, including vegetationchange, greenhouse gas (atmospheric CO2 and water vapor) fluctuations. They led to the strengthing of oceanand atmospheric circulation, significant lowering of sea level and increasing aridification over the globe,especially in middle-high latitudes. Since 12–11 Ma ago, a continuously cooler and drier climate may havecaused the substantial intensification of physical weathering and exhumation on the earth's surface, theincreased sediment accumulation rates on the land and in the sea, and the inception of productivity increasein the oceans. Nevertheless, when and where these mechanisms operated and how they were tied to eachother remain ambiguous, deserving further investigation.

© 2008 Elsevier B.V. All rights reserved.

1. Introduction

It is well-established that the global climate was warm during theEarly to early Middle Miocene and began to cool from the late MiddleMiocene following the expansion of the East Antarctic Ice Sheet (EAIS)(e.g., Shackleton and Kennett, 1975a; Miller et al., 1987; Zachos et al.,2001). However, the course of climatic evolution during the Miocene,especially during the Middle Miocene climate transition (MMCT)remains uncertain. Magnesium/calcium data from Southern Oceanplanktonic foraminifera demonstrate that high-latitude (~55°S)Southwest Pacific sea surface temperatures (SSTs) cooled 6° to 7 °Cfrom 14.2 to 13.8 Ma (Shevenell et al., 2004). Based on the synthesizedoxygen isotope information of benthic foraminifera, Miller et al.

86 10 62010846.

l rights reserved.

(1987) and Zachos et al. (2001) concluded that themajor growth of theEAIS took place between ~15 and 13Ma, concurrent with a deep oceancooling of ~3 °C (Holbourn et al., 2005). This is coincident with theformation of a large Antarctic ice sheet by 13 Ma ago inferred from theδ18O record of benthic foraminifera of DSDP Site 289 in the westernequatorial Pacific (Woodruff et al., 1981). Deep-sea benthic foramini-fera at Hole 757B showed a decline in diversity following themajor icesheet formation in Antarctica ~13.5 Ma (Singh and Gupta, 2005).Nevertheless, δ18Obenthic values from ODP Sites 1192 and 1195 drilledon the Marion Plateau indicate that a conservative estimate for thesea-level fall of 50.0±5.0 m occurred between 13.6 and 11.4 Mawith amaximumdeepwater cooling of 2.4 °C and at least 73% of the EAISwasformed in this period (John et al., 2004). Interpreting the covariancebetween benthic (Cibicidoides spp.), surface-dwelling (Globigerinoidesquadrilobatus), and deep-dwelling planktonic (Globoquadrina dehis-cens) foraminiferal δ18O, Flower and Kennett (1993) conclude that theMiddle Miocene δ18O positive shift occurred in several steps: a +0.2‰shift between ~14.6 and 14.0 Ma, a +0.5‰ shift from ~13.6 to13.55Ma, and a +0.3‰ shift from 12.5 to 12.1Ma, probably responding

Fig. 1. Map showing the mean annual precipitation in the Chinese Loess Plateau and the locality of the studied section, compiled from Xiao et al. (2006) and Feng et al. (2004).

Fig. 2. Monthly changes of mean annual temperature, mean annual precipitation andmean annual latent evaporation in the Sikouzi area. Data from the Database of ChineseNatural Resources are averaged by observations of the years 1959–1988.

149H. Jiang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 265 (2008) 148–158

to different phases of growth of the EAIS. It seems to agree with theidea that the volume of the EAIS reached its peak by 12–10Ma inferredfrom the oxygen isotope records of DSDP Sites 277, 279, 281 and 284(Shackleton and Kennett, 1975a,b).

During the late Middle Miocene, high-latitude temperaturesdropped dramatically but low-latitude temperatures remained rela-tively constant, leading to the significant increase in the polar–equatortemperature gradient (Savin et al., 1975). Such a change in themeridional temperature gradient brought increased aridity to theglobe, especially to the middle to high latitude continental regionsincluding Australia (Robert et al., 1986; Stein and Robert, 1986), Africa(Retallack, 1992), North America (Webb, 1977; Wolfe, 1985) and SouthAmerica (Pascual and Janreguizar, 1990), and may have fostered thedevelopment of grasslands and the consequent evolution of grass-land-adapted biota (Flower and Kennett, 1994). However, fewcontinental records during the MMCT are reported from the Asianregion (Wang et al., 2001; Chung and Koh, 2005). In northern China,there are numerous Tertiary sedimentary basins where thick lacus-trine deposits should have the potential to provide us with importantinformation on long-term climate history including the evolutioncourse of MMCT (e.g. Li et al., 1997; Fang et al., 2003; Pares et al., 2003;Horton et al., 2004; Wang and Deng, 2005). Recently, a well-exposed,2880-m thick fluviolacustrine sequence was discovered and paleo-magnetically dated (Jiang et al., 2007) in the Sikouzi area, Guyuan,Ningxia, China. This late Cenozoic sequence permits an exploration ofthe course of Miocene climate evolution, especially during the MMCT.

In this study, Lightness (L⁎), Susceptibility (SUS), total inorganiccarbon (TIC) and total organic carbon (TOC) content were measuredand analyzed, and compared with redness (a⁎, Jiang et al., 2007) andpollen data (Jiang and Ding, 2008) from the Sikouzi sequence.Concurrently, a variety of other Miocene records of sedimentationrate, grain-size, palaeoproductivity etc. were compiled. Togetherthey provide a firm basis for understanding the influence of Mid-Miocene climate change on East Asia and even on the globalsedimentary environment, and provide better boundary conditions

for the establishment of late Cenozoic regional and global climatemodels.

2. Geographical and geological settings

The Sikouzi area is situated on the east side of the LiupanMountainsand influenced by the East-Asian summer monsoon (Fig. 1). Themodern vegetation in the Sikouzi area is classified as temperate steppe(Wu, 1980; Hou, 1983). At present, the climate of the Sikouzi area istemperate and characterized by relatively hot-humid summers andcold-drywinters. Over the past 30 years, themean annual temperature(MAT) has been 6.2 °C, with a July average of 18.9 °C and a Januaryaverage of −8.3 °C (Fig. 2).Mean annual precipitation (MAP) is 478mm,and over 60% of the precipitation falls during July–September with apeak mean rainfall of 109.1 mm in August (Fig. 2). Mean annual latent

Fig. 3. Lithology and magnetostratigraphy of the Sikouzi sequence adapted from Jianget al. (2007).

150 H. Jiang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 265 (2008) 148–158

evaporation reaches 1772 mm, which is 3.7 times the annualprecipitation (Fig. 2).

In the study area, reddish and brownish fluviolacustrine sediments,with a thickness of 2880m, arewell exposed along a 5265-m stretch ofthe Qingshui River. The cross-section map and several photographs ofthe typical outcrops are presented in Jiang et al. (2007).

This fluviolacustrine succession, resting pseudo-conformably onPaleogene sandstones, can be roughly divided into two parts basedon the lithology (Fig. 3). The lower part (2880–488 m depth) iscomposed of lacustrine shore and deltaic sediments (Zhang et al.,2005), which mainly consist of alternating layers of grayish andreddish mudstone, muddy siltstone, siltstone, and sandstone. Thingypsum layers occur occasionally within the sediments. The upperpart of the succession (488–0 m depth) is made up of fluviolacustrineor alluvial deposits, being composed mainly of grayish and brownishmudstone, siltstone, sandstone, and conglomerate layers. Conglom-erate layers are present throughout the portion, and their thicknessand frequency show a dramatic increase upwards (Fig. 3).

3. Field and laboratory methods

In the field, sampling according to lithologywas carried out for SUSmeasurement at intervals of 0.03–12.8 m, for L⁎ measurement atintervals of 0.1–7.8m, and for TIC and TOCmeasurement at intervals of5–20 m along the river valley, where the Sikouzi section is exposed.Sampling sites were selected where possible in mudstone, siltymudstone, muddy siltstone and siltstone. At each site the surfaceweathered material was removed and a fresh sample was taken. Fromtop to bottom, 3494 SUS samples, 2207 L⁎ samples, and 400 TIC andTOC samples were collected. Age control for all the samples is derivedfrom a detailed palaeomagnetic record and biostratigraphic data(Jiang et al., 2007) (Fig. 3).

Bulk magnetic susceptibility wasmeasured for 3494 samples usinga Bartington MS2 susceptibility meter. For L⁎ measurement, a sample(~1.5 g) was first dried at 40 °C for 24 h and then was crushed to dustwithout destroying its grain size. Finally, color reflectance wasmeasured using a handheld Minolta-CM2002 spectrophotometer atthe Institute of Geology and Geophysics, Chinese Academy of Sciences(CAS). TIC and TOC concentration for 199 selected samples at aninterval of about 0.1 Ma was measured at the Carbon-ConcentrationAnalysis Lab, Institute of Geology and Geophysics, CAS, using theprocedure described by Xiao et al. (2006).

There is no common time scale for all records mentioned in thisstudy, but the Miocene characteristic Chron C5n/C5n.2n is alwayspresent in different records. The base line of C5n/C5n.2n spanning~1.0Ma is selected as the reference point and has the age of 10.949 Maaccording to the Geomagnetic Polarity Timescale (Cande and Kent,1995; Berggren et al., 1995). It is noteworthy that the new Neogenetime scale of ATNTS 2004 (Astronomically Tuned Neogene Time Scale,Lourens et al., 2004; Ogg and Smith, 2004) puts the basis of C5n.2n at11.040 Ma. But we believe that this minor change in dating does notinfluence our discussion and conclusion. Because we center on longsequences of climate and environment changes on the tectonic timescale and time resolution for various records is relatively low, theuncertainy caused by different age models is relatively small.

4. Results

The SUS, L⁎, TOC and TIC records of the Sikouzi section are shownin Fig. 4. On the whole, the SUS curve is visibly separated into an earlyperiod and a late period about 12.0 Ma (Fig. 4(a)). Before 12.0 Ma ago,the SUS value generally fluctuated between zero and 24.4 with anaverage of 6.9 apart from an ephemeral peak about 17.5–17.0Ma. Since11.0 Ma ago, the SUS fluctuating amplitude increased markedly withan increased mean value of 11.7. From 12.0 to 8.2 Ma, the SUSincreased distinctly. During the period 4.1–0.08 Ma it reached thelargest fluctuating amplitude of 1.0–31.6 and the highest mean valueof 14.1. Noticeably, the SUS shows two small peaks of 14.8 and 17.5 at11.39 Ma and 11.08 Ma ago, respectively (Fig. 4(a)).

Similar to variation of the SUS curve,11.4Ma divides the L⁎ curve ofthe Sikouzi section into 2 sections, an early period and a late period(Fig. 4(b)). The period before 12.0 Ma ago is marked by L⁎ values of

Fig. 4. The SUS, L⁎, TOC, TIC, a⁎ and pollen humidity records of the Sikouzi section plotted against palaeomagnetic ages.

151H. Jiang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 265 (2008) 148–158

generally less than 58.0, ranging from 46.6 to 64.2 with an average of53.6. Since 12.0 Ma ago, L⁎ values are usually higher than 58.0, varyingbetween 51.3 and 70.3 and averaging to 60.5. During the period 4.3–0.08 Ma, L⁎ values increased further and often exceeded 61.0,oscillating between 57.1 and 67.6 with a mean value of 62.5 (Fig. 4(b)).

Commonly lower than 1.2% before 14.0 Ma ago, the TIC content ofthe Sikouzi sequence undulated between 0.21 and 5.47% and averaged1.02% (Fig. 4(c)). Since 11.0 Ma ago, the TIC generally had a highercontent than 1.2%, fluctuating between 0.51 and 7.53% with a meanvalue of 1.84%. From 14.0 to 11.0 Ma, the TIC curve shows a weakly

increasing trend. On the whole, the TIC curve seems to be separatedinto an early period and a late period by two peaks of 5.39% and 7.53%occurring at 11.0 Ma and 10.6 Ma, respectively (Fig. 4(c)).

As shown in Fig. 4(d), the TOC values of the Neogene Sikouzisection are low and the highest of 0.73% occurred at 11.3 Ma ago. Thistime point separates the TOC curve into an early period and a lateperiod. Before 12.5 Ma ago, almost 60% of the samples bearing TOChave the highest values when compared to the rest of the section,though individual TOC concentrations in this period are not high andgenerally less than 0.2%. Since 11.3 Ma ago, TOC content decreased so

152 H. Jiang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 265 (2008) 148–158

notably that more than 75% of the samples had TOC values of zeroeven if a few peaks also showed a clear decreasing trend (Fig. 4(d)).

According to Jiang et al. (2007), 11.5 Ma is a dividing line whichseparates the redness (a⁎) record of the Sikouzi section into an earlyperiod and a late period (Fig. 4(e)) in spite of a decrease in rednessbeginning 12.7 Ma ago. The period before 11.5 Ma ago is generallycharacterized by higher redness values than 12.0, ranging between17.7 and 9.6 with a mean value of 13.2. Since 11.5 Ma ago, the rednessvalue commonly decreased to below 12.0, varying from 8.0 to 13.2 andaveraging to 10.1. In contrast to variation in the SUS and L⁎ records, theredness shows a rapid declining trend during the period 12.7–11.5 Ma(Fig. 4(e)).

In the palynologic literature, the A/C ratio (Artemisia/Chenopodia-caeae, El-Moslimany, 1990) has usually been applied to the pollenspectra from semi-arid to arid regions, particularly in central Asia (e.g.,Van Campo, 1996; Rhodes et al., 1996) in order to distinguish steppefrom desert and to reconstruct changes in precipitation. According toJiang and Ding (2008), the 20-Ma pollen record of the Sikouzi sectionis dominated by herb and shrub taxa and four kinds of halophyteinclude Chenopodiaceae (C), Ephedra (E), Nitraria (N), and Tamarica-ceae (T), mainly growing in desert environments (Bews, 1927; Hou,1983). Then the ratio of (100-C-E-N-T)/(C+E+N+T) was used as anindicator of changes in relative humidity (Fig. 4(f)).

Similarly, the early period for the humidity curve inferred from thepollen record can be distinguished from the late one if the periodbetween 20 and 0.08Ma is divided at 11.5Ma (Fig. 4(f)). Before 11.5Maago, the humidity index usually exceeded 4.0, fluctuating between0.28 and 22.09 with an average of 8.32. Since 11.5 Ma ago, thehumidity index was often less than 4.0, varying from 0.02 to 28.57 andaveraging to 3.01 in spite of brief peaks around 8.02Ma ago and duringthe period 1.02–0.74 Ma. From 14.5 to 11.5 Ma, the humidity indexdecreased in a fluctuating stepwise manner (Fig. 4(f)).

5. Discussion

The above six proxies of the Sikouzi section share a commonfeature, that is to say, they all shift irrreversibly at 12–11 Ma. Thenwhat do these Neogene shifts mean? It is well known that the SUShas been widely applied in palaeoclimate reconstruction to eolian(e.g., Kukla et al., 1988; An et al., 1991; Ding et al., 1999; Hao and Guo,2005), lacustrine (e.g., Thouveny et al., 1994; Schwab et al., 2004) andmarine sediments (e.g., Bloemendal and deMenocal, 1989; Schmiederet al., 2000). Related to humidity fluctuation, L⁎ values and TICconcentrations within arid to semi-arid lakes or basins are considered

Table 1Comparison of different proxies before and after 11/10 Ma ago for 9 collected sites

No. Site Latitude Longitude Duration(Ma)

P

1 Chiang Muan Formation, northern Thailand(Suganuma et al., 2006)

18°–19°N

98.5°–101°E

9.8–13.0 M

2 DSDP598, Southeast Pacific(Rea and Bloomstine, 1986)

19.1°S 125°W 0.17–15.61

Eg

3 Khutia Khola section, far western Nepal(Ojha et al., 2000)

28°54′N 80°40′E 13.3–7.65

M

4 Kotal Kund, Potwar Plateau, Pakistan(Johnson et al., 1982)

32°45′–33°15′N

72°32′–73°25′E

13.46–2.47

M

5 LL44-GPC3, central North Pacific(Janecek and Rea, 1983)

30°20′N 157°49′W 0.3–20.08

M

6 ODP1085, Cape Basin, South Atlantic (Twichellet al., 2002; Diester-Haass et al., 2004)

29°22.4′S 13°59.4′E 3.0–14.6 TL

7 ODP757, Ninetyeast Ridge(Hermoyian and Owen, 2001)

17°1.5′S 88°10.9′E 2.1–16.4 P

8 ODP758, northern Ninetyeast Ridge(Klootwijk et al., 1992)

5.4°N 90.4°E 0–18.7 M

9 Surai Khola section, West Nepal(Appel et al., 1991)

27°45′N 82°50′E 1.8–13.65

M

to reflect variation in the carbonate content and higher L⁎ values andTIC concentrations denotemore carbonate content and thusmore aridconditions (e.g., Abdul Aziz et al., 2003, 2004; Xiao et al., 2006). Theredness is probably related to temperature significantly influencingoxidation of sediments in arid to semi-arid regions (Jiang et al., 2007).As indicated by the humidity index, SUS, L⁎, TIC, TOC concentrationsand redness of the Sikouzi fluviolacustrine sediments probably reflectthe control of the climate evolution on environment changes. Before12 Ma, the higher humidity index and TOC concentrations and thelower TIC concentrations and L⁎ values indicate that the East-Asiansummer monsoon was strong during the Early to Middle Miocene(Jiang and Ding, 2008) and brought much more moisture to the studyarea, favoring vegetation and soil development. This would not onlyincrease the TOC content in sediments in the Sikouzi area but alsomitigate aridity resulting from intense evaporation and thus decreasethe TIC contents and L⁎ values. During the Early to Middle Miocene,the higher redness values and the lower SUS imply that temperatureof that time was relatively high and favored the further oxidation ofmagnetite, leading to the redness increase and the SUS decrease.

Since 11 Ma ago, the humidity index has basically remained lowvalues, implying significant aridification in the study area and theweakening of the East-Asian summermonsoon so that vegetation andsoil developmentwas suppressed. This interpretation is supported byadistinct decline in the TOC contents during this period. The increase ofSUS, L⁎ and TIC contents and the decrease of redness suggest that,under the cooler and more arid climate, physical weatheringintensified and increased the abundance of magnetite in sediments.Because of relatively rapid accumulation, abundant magnetite had noenough time to oxidize and thus enhanced the SUS value anddecreased the redness of sediments.

In short, the above six proxies of the Sikouzi section indicate thatthe substantial climate change occurred at 12–11Ma in the study area.This climate change has also been revealed by various proxies fromprevious studies in different regions. As shown inTable 1, these proxiesinclude mean sedimentation rate (MSR), eolian grainsize, erosion rate,total organic carbon content (TOC), mean accumulation rate of organiccarbon (MARTOC), linear sedimentation rate (LSR), benthic foraminif-eral accumulation rate (BFAR), and so on. Here we take several repre-sentative regions as examples.

5.1. Qaidam, Hexi Corridor and Linxia basins

According to Sun et al. (1999), the ostracod Cyprideis occurs widelyin the shallow brackish-water lake environments and Cyprideis-

roxy Before 11/10 Ma After 11/10 Ma

SR 3.69 cm/kyr 5.50 cm/kyr

olianrainsize

8.04φ 7.27φ

SR ~0.32 mm/yr 0.47 mm/yr

SR 23 cm/kyr 52 cm/kyr

AR 15.7 mg cm−2 kyr−1 97.9 mg cm−2 kyr−1

OC, MARTOC,SR, BFAR

0.18%, 4.46 mg cm−2 kyr−1, 2.22 cm/kyr, 0.25⁎103 cm−2 kyr−1

0.80%, 56.42 mg cm−2 kyr−1,5.32 cm/kyr, 1.70⁎103 cm−2 kyr−1

MAR 0.34 µmol cm−2 kyr−1 0.74 µmol cm−2 kyr−1

SR 0.35 cm/kyr 1.14 cm/kyr

SR 24.83 cm/kyr 43.77 cm/kyr

153H. Jiang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 265 (2008) 148–158

bearing sediments from the Qaidam basin usually contain gypsumwith salinity reaching up to 10.23‰ (Institute of Exploration andDevelopment, Qinghai Petroleum Administration and Nanjing Insti-tute of Geology and Palaeontology, 1988). The thriving of Cyprideissince ~12 Ma ago in the Youshashan Formation of the Qaidam basinsuggests onset of an intense arid climate (Yang et al., 2000). At thebeginning of the Youshashan Formation in the Qaidam basin, biotitecontents increased distinctly upwards and remained higher thanbefore (Rieser et al., 2005), probably implying that a drier climateweakened chemical weathering and facilitated preservation of theunstable mineral of biotite. The volumetric accumulation ratesdoubled from around 11 Ma ago and showed a major rise after thebeginning of the Pliocene (5.3 Ma ago) in the Qaidam and the HexiCorridar basins (Metivier et al., 1998). Consistent with the δ18Orecords of lacustrine carbonate (Dettman et al., 2003), fossil toothenamel and paleosols from the Linxia basin (Wang and Deng, 2005),all the above indicate that the regional climate along the northeastmargin of the Tibetan Plateau became drier since 12–11 Ma ago.

Fig. 5. Comparison between the eolian grainsize from DSDP 598, the mean sedimentation raheavy minerals from Pakistan.

5.2. Pakistan, Nepal and their neighbouring regions

As shown in Fig. 5(a), the hornblende content within heavyminerals increased since 11 Ma ago in the Siwalik sediments exposednear the Chinji Village, Potwar Plateau, Pakistan (Johnson et al., 1985).From ~11 Ma to the present, K-feldspar sand increased clearly in theNeogene Dumri Formation and the Siwalik Group in western Nepal,suggesting that physical erosion of granitic rocks strengthened since~11 Ma ago (DeCelles et al., 1998). Inferred from fossil plant leaves, atropical wet evergreen forest was replaced by a tropical deciduousforest since the beginning of the early Late Miocene in the Surai Kholaarea and the annual dry season became prolonged to 4–7 months(Walter, 1962; Awasthi and Prasad, 1990; Awasthi et al., 1994).Moreover, all of the magnetostratigraphic investigations indicatethat MSR increased abruptly since 10 Ma (about equal to 11 Ma forCK95, same hereafter) ago both for the Siwalik sediments at KotalKund (Johnson et al., 1982) and near the Chinji Village (Johnson et al.,1985), Potwar Plateau, Pakistan (Fig. 5(b)) and for the Surai Khola

tes (MSR) from ODP758, West Nepal and Pakistan, and the hornblende content within

154 H. Jiang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 265 (2008) 148–158

section (Appel et al., 1991, Fig. 5(c)) and the Khutia Khola section (Ojhaet al., 2000), western Nepal (Table 1). In addition, sedimentation rateand detrital heavy minerals of Leg 116 ODP Sites 717, 718 and 719 fromthe distal Bengal fan in the central Indian Ocean indicate that calcicamphibole eroded from the Higher Himalayas and sedimentation rateincreased abruptly at 10.9 Ma ago (Amano and Taira, 1992). As for thelacustrine and fluvial sediments from the Miocene Chiang MuanFormation in northern Thailand (Table 1), the MSR was 3.69 cm/kyrbefore 11 Ma ago and increased to 5.50 cm/kyr since 11 Ma ago(Suganuma et al., 2006). More pulsed faunal turnover in the NeogeneSiwaliks of Northern Pakistan occurred between 14 and 12 Ma andmore constant turnover is evident after 11 Ma ago (Barry et al., 1995).

5.3. Northern Indian Ocean

ODP Leg 121 drilled three cores, Sites 756, 757 and 758, on theNinetyeast Ridge, among which the age of site 758 was achieved bythe combination of biostratigraphic and magnetostratigraphic inves-tigations. As shown in Fig. 5(d), MSR was calculated at an interval of0.1 Ma based on the Neogene age-depth record of Site 758 provided byKlootwijk et al. (1992). Before ~11 Ma ago MSR was low with anaverage of 0.35 cm/kyr, but from ~11 Ma ago it increased abruptly andremained high with a mean value of 1.14 cm/kyr (Table 1). Thephosphorus mass accumulation rate (PMAR) in Site 757 showed lowvalues between ~17 and 11 Ma, and they started to increase since11 Ma ago (Table 1). It reached a peak that lasted until ~9 Ma ago,

Fig. 6. Correlation of the mean sedimentation rate (MSR) of the Sikouzi section, Guyuan, Ninvariation at ODP Site (Miller et al., 1996) (c) and Neogene sea level record (Haq et al., 1987) (dare obtained using geomagnetic boundaries as age controls and then linear interpolating. N

reflecting an increase in productivity during the Late Miocene,probably caused by an overall increase in nutrient flux into theworld oceans (Hermoyian and Owen, 2001; Gupta et al., 2004). Guptaet al. (2004) related a widespread increase in productivity in thePacific, Indian and Atlantic oceans during the Late Miocene to a globalcooling event caused by increased glaciation in the Antarctic region.The mass accumulation rate of the terrigenous component of deep seasediments was quantified by Rea (1992) for eleven DSDP and ODP sitesin the northern Indian Ocean. Depositional patterns at these sites allshow very low input prior to 11 Ma, a five-fold increase in terrigenousflux starting at 11 Ma ago, and distinct flux maxima in the LateMiocene and the Middle Pliocene.

5.4. Central North Pacific Ocean

The mass accumulation rate and grain size distribution of the totaleolian component isolated from a large-diameter piston core, LL44-GPC3, recovered from the central North Pacific, have been used toevaluate changes in eolian sedimentation and in the intensity ofatmospheric circulation that have occurred during the past 70 Ma(Janecek and Rea, 1983). During the period 20.08–11.36 Ma, theaccumulation rate varied from11.4 to 21.7mg cm−2 kyr−1 and averagedto 15.7 mg cm−2 kyr−1 (Table 1). From 9.93 to 3.86 Ma it increasedsignificantly, ranging from 15.5 to 28.2 mg cm−2 kyr−1 with a meanvalue of 23.4 mg cm−2 kyr−1 (Janecek and Rea, 1983). This is consistentwith the global synchronous increase of average sedimentation rates

gxia, China (a), the normalized mean sediment flux (MSF) to the Indian Ocean (b), δ18O). Calculation for the Sikouzi sectionwas made at 100,000-year intervals. The timescalesote that the largest sea-level fall of about 78 m occurred at about 11.5 Ma ago.

155H. Jiang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 265 (2008) 148–158

for the Atlantic, Indian, and Pacific oceans during the period 11.2–5.2 Ma (Davies et al., 1977, this interval recalculated by Diester-Haasset al. (2004)). As shown in Fig. 6 (A) and (B), both theMSRof the Sikouzisection and the normalized mean sediment flux to the Indian Oceanshowed a significant increase since 12–11 Ma ago. A global synchro-nous increase in total sedimentation rates since the early LateMiocenepoints to significantly different modes of continental weathering,which may be due to alternate states of atmospheric circulationmarked by a reduction in global precipitation (Davies et al., 1977).

5.5. Southern Ocean

Grainsize variations of eolian dust extracted from Southern Oceanpelagic sediments provide a good record of southern hemisphereatmospheric processes. As shown in Fig. 5(e) and Table 1, mean eoliangrainsize data in DSDP 598 from Southeast Pacific Ocean show anincrease from 8.04φ before 10 Ma ago to 7.27φ after 10 Ma ago,indicating a significant increase in the intensity of atmosphericcirculation (Rea and Bloomstine, 1986). An increase in iron supplyassociated with enhanced dust inputs could be responsible for highermarine phytoplankton production (Calvo et al., 2004). Multiproxiesfrom the Cape Basin reveal the detailed evolution of paleoproductivityand nutrient supply in the Southeast Atlantic Ocean during the LateCenozoic, including TOC content (TOC%), mean accumulation ratesof TOC (MARTOC), linear sedimentation rates (LSR), and benthicforaminifera accumulation rates (BFAR) (e.g., Twichell et al., 2002;Diester-Haass et al., 2004). As shown in Fig. 7 and Table 1, TOC%,

Fig. 7. Comparison between the benthic foraminiferal accumulation rates (BFAR), the linear sand the total organic carbon content (TOC%) from ODP Site 1085 of the Cape basin in South

MARTOC, LSR and BFAR maintained low fluctuating values before 11.5–11.0 Ma ago, and then showed an increase with intense fluctuationsince 11.5–11.0 Ma ago. Moreover, variations of abundance anddiversity of radiolarians from ODP 1143 in the northern South ChinaSea indicate that paleoproductivity began to increase ~12–11 Ma agoand reached a maximum at ~8.24 Ma ago (Chen et al., 2003). Also at12–11Ma from the northern South China Sea, the grain size coarsenedin ODP Site 1146 (Wan et al., 2007) and the δ13C benthic positivelyexcursed at ODP Site 1148 (Zhao et al., 2001). The black carbon isotopevalues displayed a spike around 11 Ma at ODP Sites 1148 and 1147,implying an expansion of C4 plants (Jia et al., 2003). Middle throughUpper Miocene (17–5 Ma) BFARs from a depth transect in the westernequatorial Indian Ocean (DSDP Site 237, ODP Sites 707, 709 and 710)began to increase significantly at shallower sites and slowly at deepersites from 11 Ma ago (Smart et al., 2007).

5.6. South and North America

The Washington Cascade Range is a complex, polygeneticmountain range that dominates the topographic, climatic, and culturalconfigurations ofWashington State. Samples on thewest flank suggestonset of a rapid exhumation since the early Late Miocene (Reinerset al., 2002). Accelerated exhumation rates since the early LateMiocene are also observed in the Coastal Mountains of BritishColumbia and southeast Alaska, ~1500 km to the north (Reinerset al., 2002). Magnetostratigraphic study of Neogene nonmarine strataof the Mendoza Precordillera suggests the initiation of exhumation of

edimentation rates (LSR), the mean accumulation rate of total organic carbon (MARTOC)east Atlantic Ocean.

156 H. Jiang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 265 (2008) 148–158

the Frontal Cordillera to the northwest of the study region since~11.7 Ma ago (Irigoyen et al., 2000). A multiple data set that includestiming of facies changes, paleocurrent data, accumulation rates, andprovenance allows a temporal correlation to be made betweendeposition of the La Pilona Formation and exhumation of the FrontalCordillera to the northwest of the study site (Irigoyen, 1997), implyingthe beginning of exhumation in the region at ~11.7 Ma ago.

5.7. Global synthesis

To sum up, a variety of proxies from different regions including thestudy area of Guyuan, Ningxia, China, changed significantly at 12–11 Ma ago. Apparently, this is more than a coincidence in time andmust have been modulated by some mechanisms in the world. Suchmechanisms involve tectonic movement and climate change withoutquestion, but the former can not explain continuous increases since11 Ma ago for various late Cenozoic proxies from different regionsdistant from each other and with such large amplitudes for differentproxies in the world. It is only climate change that can leavesimultaneous imprints in different regions on the globe. However,since the Mid-Miocene global cooling began at ~14 Ma ago (Shevenellet al., 2004) as mentioned above, why had the nearly simultaneoustransformation of values for various proxies in different regionsoccurred around 12–11 Ma? Here, we propose that several well-knownmechanismsmay have contributed to the transfer of the globalcooling signatures since 14 Ma into different regions by 12–11Ma ago.

With the development of EAIS and the lowering of sea level duringthe late Middle Miocene, the main source regions for water vapor, thesubtropical and tropical oceans, have significantly decreased in area(Barron, 1985). Atmospheric cooling reduces the amount of watervapor held in the atmosphere and produces additional coolingbecause of the reduced greenhouse effect (Ruddiman, 2002). Thisfeedback associated with a decrease in atmospheric water vapor maybe as important as decreasing the sources (Wright and Miller, 1996).As mentioned above, there was a large-scale increase in oceanicprimary productivity during the Late Miocene–Early Pliocene, whichwould have induced an increase in CO2 drawdown out of theatmosphere, altering the global carbon cycle and potentially leadingto atmospheric cooling (Hermoyian and Owen, 2001), though pCO2

variability probably played a secondary role in forcing climate changeduring the Miocene (Pagani et al., 1999). On the other hand, the globalcooling during the late Middle Miocene caused the replacement oftwo-storey (forest) by single-storey vegetation types (tundra ordesert, Wolfe, 1980; Mudie and Helgason, 1983; Head et al., 1989) inmiddle-high latitudes, leading to increased albedo of snow cover andinducing additional cooling (Dutton and Barron, 1997).

The persistent global cooling led to increased meridional tem-perature gradients which caused a northward shift of the AntarcticPolar Front and thus probably strengthened the oceanic circulation.Intensified oceanic circulation induced stronger upwelling or verticalmixing in the surface layer for the Atlantic Ocean (Siesser, 1980;Twichell et al., 2002), the Pacific Ocean (Chen et al., 2003), and theIndian Ocean (Smart et al., 2007) and elevated productivity.

Following global cooling and the development of EAIS during themiddle–late Miocene transition, the glacioeustatic lowstand wasrecorded in the coastal onlap–offlap charts (Haq et al., 1987), whichcan be biostratigraphically correlated with the cooling events in thestable isotope stratigraphy (Miller et al., 1991). Times of falling,rapidly fluctuating, or low sea level generally correspond with timesof enhanced input of sediment to the deep sea (Haq et al., 1987;Greenlee and Moore, 1988). According to Haq et al. (1987), the sea-level fall at 11.5Ma ago, corresponding to the cooling eventMi5 (Fig. 6(C)), is the most significant in the Neogene and sea-level was about78m lower than today (Fig. 6(D)). It supports the idea that the volumeof EAIS reached its peak by 12–10Ma (Shackleton and Kennett,1975a,b). Sea-level fall during this time period presumably resulted in

exposure of the Bering land bridge, allowing the Hipparion fauna toimmigrate from North America to Eurasia at ~11.1 Ma ago (Garces etal., 1997).

Although these mechanisms mentioned abovemay have played animportant role on the transfer of the cooling signatures since 14 Maago into different regions by 12–11 Ma, when and where theyoperated and how they were tied to each other are ambiguous. In thisrespect, the influence of the Antarctic glaciation history on the globalclimate change during the MMCT deserves further investigation in thefuture.

6. Conclusions

In this paper, SUS, L⁎, TIC and TOC content were analyzed for theSikouzi fluviolacustrine sediments at Guyuan, Ningxia, China, andcompared with a⁎ and the pollen humidity index from the samesection. Before 12–11 Ma ago, SUS, L⁎ and TIC values were low whileTOC, a⁎ and the humidity index values were high, which completelyreversed from 12–11 Ma. Changes for these proxies indicate that thepalaeoclimate in the study area became colder and drier since 12–11 Ma. This climate change also had similar imprints in other regionsof the world, being explained by global cooling during the Middle–Late Miocene transition and the significant development of EAIS. Thiscourse not only strengthened the meridional temperature gradientsand global aridity in themiddle-high latitudes but also intensified theoceanic and atmospheric circulation and the major falling of globalsea level. Several well-known positive feedback mechanisms pre-sumably suppressed the rise of global temperature and the thaw ofpolar ice sheet during the Middle–Late Miocene transition, mainlyincluding vegetation changes in the middle–high latitudes andgreenhouse gas variations (atmospheric CO2 and water vapor). Allof the above are considered to have caused the substantialintensification of physical weathering and exhumation in the world,the increased sediment accumulation rates on land and in the sea, andthe inception of productivity increase in the oceans since 12–11 Maago. Nevertheless, when and where these mechanisms operated andhow they were tied to each other remain ambiguous, and more high-resolution records are needed to clarify these mechanisms in thefuture.

Acknowledgements

We are grateful to Profs. A. Peter Kershaw (editor), A.K. Gupta andone anonymous reviewer for their valuable comments and sugges-tions. We thank Profs. A. Peter Kershaw (editor) and A.K. Gupta fortheir help with English improvement. This work was financiallysupported by the National Natural Science Foundation of China(Grants 40730208, 40502017, 40730210).

References

Abdul Aziz, H., Krijgsman, W., Hilgen, F.J., Wilson, D.S., Calvo, J.P., 2003. An astronomicalpolarity timescale for the late middle Miocene based on cyclic continentalsequences. J. Geophys. Res. 108 (B3), 2159. doi:10.1029/2002JB001818.

Abdul Aziz, H., van Dam, J., Hilgen, F.J., Krijgsman, W., 2004. Astronomical forcing inUpper Miocene continental sequences: implications for the Geomagnetic PolarityTime Scale. Earth Planet. Sci. Lett. 222, 243–258.

Amano, K., Taira, A., 1992. Two-phase uplift of Higher Himalayas since 17 Ma. Geology20, 391–394.

An, Z.S., Kukla, G., Porter, S.C., Xiao, J.L., 1991. Magnetic susceptibility evidence of monsoonvariation on the Loess Plateau of central China during the last 130,000 years. Quat. Res.36, 29–36.

Appel, E., Rosler, W., Corvinus, G., 1991. Magnetostratigraphy of the Miocene–Pleistocene Surai Khola Siwaliks in West Nepal. Geophys. J. Int. 105, 191–198.

Awasthi, N., Prasad, M., 1990. Siwalik plant fossils from Surai Khola area,Western Nepal.Paleobotanist 38, 298–318.

Awasthi, N., Sarkar, S., Prasad, M., 1994. Vegetation and palaeoenvironment of Siwaliksuccession in Surai Khola area, Nepal. Himalayan Geol. 15, 107–119.

Barron, E.J., 1985. Explanation of the Tertiary global cooling trend. Palaeogeogr.Palaeoclimatol. Palaeoecol. 50, 45–61.

157H. Jiang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 265 (2008) 148–158

Barry, J.C., Morgan, M.E., Flynn, L.J., Pilbeam, D., Jacobs, L.L., Lindsay, E.H., Raza, S.M.,Solounias, N.,1995. Patterns of faunal turnover anddiversity in theNeogene Siwaliksof Northern Pakistan. Palaeogeogr. Palaeoclimatol. Palaeoecol. 115, 209–226.

Berggren, W.A., Kent, D.V., Swisher III, C.C., Aubry, M.-P., 1995. A revised Cenozoicgeochronology and chronostratigraphy. In: Berggren, W.A., Kent, D.V., Aubry, M.-P.,Hardenbol, J. (Eds.), Geochronology, Time Scales and Global Stratigraphic Correla-tion. SEPM Spec. Publ., vol. 54, pp. 129–212.

Bews, J.W., 1927. Studies in the ecological evolution of the angiosperms. The NewPhytologist 16, 65–84.

Bloemendal, J., deMenocal, P., 1989. Evidence for a change in the periodicity of tropicalclimate cycles at 2.4 Myr from whole-core magnetic susceptibility measurements.Nature 342, 897–900.

Calvo, E., Pelejero, C., Logan, G.A., De Deckker, P., 2004. Dust-induced changes inphytoplankton composition in the Tasman Sea during the last four glacial cycles.Paleoceanography 19, PA2020. doi:10.1029/2003PA000992.

Cande, S.C., Kent, D.V., 1995. Revised calibration of the geomagnetic polarity timescalefor the Late Cretaceous and Cenozoic. J. Geophys. Res. 100, 6093–6095.

Chen, M.H., Wang, R.J., Yang, L.H., Han, J.X., Lu, J., 2003. Development of east Asiansummer monsoon environments in the late Miocene: radiolarian evidence fromSite 1143 of ODP Leg 184. Mar. Geol. 201, 169–177.

Chung, C.-H., Koh, Y.-K., 2005. Palynostratigraphic andpalaeoclimatic investigations on theMiocene deposits in the Pohang area, South Korea. Rev. Palaeobot. Palynol. 135, 1–11.

Davies, T.A., Hay, W.W., Southam, J.R., Worsley, T.R., 1977. Estimates of Cenozoic oceanicsedimentation rates. Science 197, 53–55.

DeCelles, P.G., Gehrels, G.E., Quade, J., Ojha, T.P., Kapp, P.A., Upreti, B.N., 1998. Neogeneforeland basin deposits, erosional unroofing, and the kinematic history of theHimalayan fold-thrust belt, western Nepal. Geol. Soc. Amer. Bull. 110 (1), 2–21.

Dettman, D.L., Fang, X.M., Garzione, C.N., Li, J.J., 2003. Uplift-driven climate change at12 Ma: a long δ18O record from the NE margin of the Tibetan plateau. Earth Planet.Sci. Lett. 214, 267–277.

Diester-Haass, L., Meyers, P.A., Bickert, T., 2004. Carbonate crash and biogenic bloom in thelate Miocene: evidence from ODP Sites 1085, 1086, and 1087 in the Cape Basin,southeast Atlantic Ocean. Paleoceanography 19, PA 1007. doi:10.1029/2003PA000933.

Ding, Z.L., Xiong, S.F., Sun, J.M., Yang, S.L., Gu, Z.Y., Liu, T.S., 1999. Pedostratigraphy andpaleomagnetism of a ~7.0 Ma eolian loess-red clay sequence at Lingtai, LoessPlateau, north-central China and the implications for paleomonsoon evolution.Palaeogeogr. Palaeoclimatol. Palaeoecol. 152, 49–66.

Dutton, J.F., Barron, E.J., 1997. Miocene to present vegetation changes: a possible piece ofthe Cenozoic cooling puzzle. Geology 25 (1), 39–41.

El-Moslimany, A.P., 1990. Ecological significance of common nonarboreal pollen:examples from drylands of the Middle East. Rev. Palaeobot. Palynol. 64, 343–350.

Fang, X.M., Garzione, C., Van der Voo, R., Li, J.J., Fan, M.J., 2003. Flexural subsidence by29 Ma on the NE edge of Tibet from the magnetostratigraphy of Linxia Basin, China.Earth Planet. Sci. Lett. 210, 545–560.

Feng, Z.-D., Wang, H.B., Olson, C., Pope, G.A., Chen, F.H., Zhang, J.W., An, C.B., 2004.Chronological discord between the last interglacial paleosol (S1) and its parentmaterial in the Chinese Loess Plateau. Quat. Int. 117, 17–26.

Flower, B.P., Kennett, J.P., 1993. Middle Miocene ocean/climate transition: High-resolution stable isotopic records from DSDP Site 588A, southwest Pacific.Paleoceanography 8, 811–843.

Flower, B.P., Kennett, J.P., 1994. The middle Miocene climatic transition: East Antarcticice sheet development, deep ocean circulation and global carbon cycling.Palaeogeogr. Palaeoclimatol. Palaeoecol. 108, 537–555.

Garces, M., Cabrera, L., Agusti, J., Maria Pares, J., 1997. Old World first appearance datumof “Hipparion” horses: Late Miocene large-mammal dispersal and global events.Geology 25 (1), 19–22.

Greenlee, S.M., Moore Jr., T.C., 1988. Recognition and interpretation of depositionalsequences and calculation of sea-level changes from stratigraphic data-offshoreNew Jersey and Alabama Tertiary. In: Wilgus, C.K., Hastings, B.S., Kendall, C.G.St.C.,Posamentier, H.W., Ross, C.A., Van Wagoner, J.C. (Eds.), Sea-Level Changes—AnIntegrated Approach. Society of Economic Paleontologists andMineralogists SpecialPublication, vol. 42, pp. 329–353. Tulsa, OK.

Gupta, A.K., Singh, R.K., Joseph, S., Thomas, E., 2004. Indian Ocean high-productivityevent (10–8 Ma): linked to global cooling or to the initiation of the Indianmonsoons? Geology 32 (9), 753–756.

Hao, Q.Z., Guo, Z.T., 2005. Spatial variations of magnetic susceptibility of Chinese loessfor the last 600 kyr: implications for monsoon evolution. J. Geophys. Res. 110,B12101. doi:10.1029/2005JB003765.

Haq, B.U., Hardenbol, J., Vail, P.R., 1987. Chronology of fluctuating sea levels since theTriassic. Science 235, 1156–1167.

Head, M.J., Norris, G., Mudie, P.J., 1989. Palynology and dinocyst stratigraphy of theMiocene in ODP Leg 105, Hole 645E, Baffin Bay. In: Srivastava, S.P. (Ed.), Proceedingsof the Ocean Drilling Program. scientific results, vol. 105. Ocean Drilling Program,College Station, Texas, pp. 467–514.

Hermoyian, C.S., Owen, R.M., 2001. Late Miocene–early Pliocene biogenic bloom:evidence from low-productivity regions of the Indian and Atlantic Oceans.Paleoceanography 16 (1), 95–100.

Holbourn, A., Kuhnt, W., Schulz, M., Erlenkeuser, H., 2005. Impacts of orbital forcingand atmospheric carbon dioxide on Miocene ice-sheet expansion. Nature 438,483–487.

Horton, B.K., Dupont-Nivet, G., Zhou, J., Waanders, G.L., Butler, R.F., Wang, J., 2004.Mesozoic–Cenozoic evolution of the Xining-Minhe and Dangchang basins, north-eastern Tibetan Plateau: magnetostratigraphic and biostratigraphic results.J. Geophys. Res. 109, B04402. doi:10.1029/2003JB002913.

Hou, X.Y., 1983. Vegetation of China with reference to its geographical distribution. Ann.Mo. Bot. Gard. 70, 509–549.

Institute of Exploration and Development, Qinghai Petroleum Administration, NanjingInstitute of Geology and Palaeontology, CAS, 1988. A Research on Tertiary OstrocodeFauna from the Qaidam Basin, Qinghai Province. Nanjing University Press, Nanjing,pp. 29–30 (in Chinese).

Irigoyen,M.V.,1997. Magnetic polarity stratigraphy and geochronological constraints onthe sequence of thrusting in the Principal and Frontal cordilleras and thePrecordillera of the Argentine Central Andes (33° S latitude) [Ph.D. thesis]: Ottawa,Canada, Carleton University, 392 p.

Irigoyen, M.V., Buchan, K.L., Brown, R.L., 2000. Magnetostratigraphy of Neogene Andeanforeland-basin strata, lat33°S, Mendoza Province, Argentina. Geol. Soc. Amer. Bull.112 (6), 803–816.

Janecek, T.R., Rea, D.K., 1983. Eolian deposition in the northeast Pacific Ocean: Cenozoichistory of atmospheric circulation. Geol. Soc. Amer. Bull. 94, 730–738.

Jia, G.D., Peng, P.A., Zhao, Q.H., Jian, Z.M., 2003. Changes in terrestrial ecosystem since30 Ma in East Asia: Stable isotope evidence from black carbon in the South ChinaSea. Geology 31 (12), 1093–1096.

Jia, H.C., Ding, Z.L., 2008. A 20 Ma pollen record of East-Asian summer monsoonevolution from Guyuan, Ningxia, China. Palaeogeogr. Palaeoclimatol. Palaeoecol.265, 30–38. doi:10.1016/j.palaeo.2008.04.016.

Jiang, H.C., Ding, Z.L., Xiong, S.F., 2007. Magnetostratigraphy of the Neogene Sikouzisection at Guyuan, Ningxia, China. Palaeogeogr. Palaeoclimatol. Palaeoecol. 243,223–234.

John, C.M., Karner, G.D., Mutti, M., 2004. δ18O and Marion Plateau backstripping:combining two approaches to constrain late middle Miocene eustatic amplitude.Geology 32, 829–832.

Johnson, N.M., Opdyke, N.D., Johnson, G.D., Lindsay, E.H., Tahirkheli, R.A.K., 1982.Magnetic polarity stratigraphy and ages of Siwalik Group rocks of the PotwarPlateau, Pakistan. Palaeogeogr. Palaeoclimatol. Palaeoecol. 37, 17–42.

Johnson, N.M., Stix, J., Tauxe, L., Cerveny, P.F., Tahirkheli, R.A.K., 1985. Paleomagneticchronology, fluvial processes, and tectonic implications of the Siwalik deposits nearChinji village, Pakistan. J. Geol. 93, 27–40.

Klootwijk, C.T., Gee, J.S., Peirce, J.W., Smith, G.M., 1992. Neogene evolution of theHimalayan–Tibetan region: constraints from ODP Site 758, northern NinetyeastRidge; bearing on climatic change. Palaeogeogr. Palaeoclimatol. Palaeoecol. 95,95–110.

Kukla, G., Heller, F., Liu, X.M., Xu, T.C., Liu, T.S., An, Z.S., 1988. Pleistocene climates inChina dated by magnetic susceptibility. Geology 16, 811–814.

Li, J.J., Fang, X.M., Van der Voo, R., Zhu, J.J., Macniocaill, C., Cao, J.X., Zhong, W., Chen, H.L.,Wang, J.L., Wang, J.M., Zhang, Y.C., 1997. Late Cenozoic magnetostratigraphy (11–0 Ma) of the Dongshanding and Wangjiashan sections in the Longzhong Basin,western China. Geol. Mijnb. 76, 121–134.

Lourens, L., Hilgen, F., Shackleton, N.J., Laskar, J., Wilson, D., 2004. The Neogene period.In: Gradstein, F., et al. (Ed.), A Geological Time Scale 2004. Cambridge UniversityPress, pp. 409–440.

Metivier, F., Gaudemer, Y., Tapponnier, P., Meyer, B., 1998. Northeastward growth of theTibet Plateau deduced from balanced reconstruction of two depositional areas: theQaidam and Hexi Corridor basins, China. Tectonics 17 (6), 823–842.

Miller, K.G., Fairbanks, R.G., Mountain, G.S., 1987. Tertiary oxygen isotope synthesis, sealevel history, and continental margin. Paleoceanography 2 (1), 1–19.

Miller, K.G., Wright, J.D., Fairbanks, R.G., 1991. Unlocking the icehouse: Oligocene–Miocene oxygen isotopes, eustasy and margin erosion. J. Geophys. Res. 96,6829–6848.

Miller, K.G., Mountain, G.S., the Leg 150 Shipboard Party, Members of the New JerseyCoastal Plain Drilling Project, 1996. Drilling and Dating New Jersey Oligocene–Miocene sequences: ice volume, global sea level, and Exxon records. Science 271,1092–1095.

Mudie, P.J., Helgason, J., 1983. Palynological evidence for Miocene climatic cooling ineastern Iceland about 9.8 Myr ago. Nature 303, 689–692.

Ogg, J.G., Smith, A.G., 2004. The geomagnetic polarity time scale. In: Gradstein, F., et al.(Ed.), A Geological Time Scale 2004. Cambridge University Press, pp. 63–86.

Ojha, T.P., Butler, R.F., Quade, J., DeCelles, P.G., Richards, D., Upreti, B.N., 2000. Magneticpolarity stratigraphy of the Neogene Siwalik Group at Khutia Khola, far westernNepal. Geol. Soc. Amer. Bull. 112 (3), 424–434.

Pagani, M., Arthur, M.A., Freeman, K.H., 1999. Miocene evolution of atmospheric carbondioxide. Paleoceanography 14 (3), 273–292.

Pares, J.M., Van der Voo, R., Downs, W.R., Yan, M.D., Fang, X.M., 2003. Northeastwardgrowth and uplift of the Tibetan Plateau: magnetostratigraphic insights from theGuide Basin. J. Geophys. Res. 108 (B1), 2017. doi:10.1029/2001JB001349.

Pascual, R., Janreguizar, E.O., 1990. Evolving climates and mammal faunas in CenozoicSouth America. J. Human Evol. 19, 23–60.

Rea, D.K., 1992. Delivery of Himalayan sediment to the northern Indian Ocean and itsrelation to global climate, sea level, uplift, and seawater strontium,. In: Duncan, R.,et al. (Ed.), Synthesis of results from scientific drilling in the Indian Ocean. AmericanGeophysical Union Geophysical Monograph, vol. 70, pp. 387–402.

Rea, D.K., Bloomstine, M.K., 1986. Neogene history of the south Pacific tradewinds:evidence for hemispherical asymmetry of atmospheric circulation. Palaeogeogr.Palaeoclimatol. Palaeoecol. 55, 55–64.

Reiners, P.W., Ehlers, T.A., Garver, J.I., Mitchell, S.G., Montgomery, D.R., Vance, J.A.,Nicolescu, S., 2002. Late Miocene exhumation and uplift of theWashington CascadeRange. Geology 30 (9), 767–770.

Retallack, G.J., 1992. Middle Miocene fossil plants from Fort Ternan (Kenya) andevolution of African grasslands. Paleobiology 18, 383–400.

Rhodes, T.E., Gasse, F., Ruifen, L., Fontes, J.-C., Wei, K.Q., Bertrand, P., Gibert, E., Melieres,F., Tucholka, P., Wang, Z.X., Cheng, Z.Y., 1996. A late Pleistocene–Holocene lacustrinerecord from Lake Manas, Zunggar (northern Xinjiang, western China). Palaeogeogr.Palaeoclimatol. Palaeoecol. 120, 105–121.

158 H. Jiang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 265 (2008) 148–158

Rieser, A.B., Neubauer, F., Liu, Y.J., Ge, X.H., 2005. Sandstone provenance of north-western sectors of the intracontinental Cenozoic Qaidam basin, western China:tectonic vs. climatic control. Sediment. Geol. 177, 1–18.

Robert, C., Stein, R., Acquaviva, R., 1986. Cenozoic evolution and significance of clayassociations in the New Zealand region of the southwest Pacific, Leg 90. In: Kennett,J.P., von der Borch, C.C., et al. (Eds.), Init. Rep. DSDP, vol. 90, pp. 1225–1238.

Ruddiman, W.F., 2002. Earth's climate: Past and future. W.H. Freeman and Company,New York, pp. 1–465.

Savin, S.M., Douglas, R.G., Stehli, F.G., 1975. Tertiary marine paleotemperatures. Geol.Soc. Amer. Bull. 86, 1499–1510.

Schmieder, F., von Dobeneck, T., Bleil, U., 2000. The Mid-Pleistocene climate transitionas documented in the deep South Atlantic Ocean: initiation, interim state andterminal event. Earth Planet. Sci. Lett. 179, 539–549.

Schwab, M.J., Neumann, F., Litt, T., Negendank, J.F.W., Stein, M., 2004. Holocenepalaeoecology of the Golan Heights (Near East): investigation of lacustrinesediments from Birkat Ram crater lake. Quat. Sci. Rev. 23, 1723–1731.

Shackleton, N.J., Kennett, J.P., 1975a. Paleotemperature history of the Cenozoic andinitiation of Antarctic glaciation: oxygen and carbon isotopic analyses in DSDP Sites277, 279 and 281. Initial Rep. Deep Sea Drill. Proj. 29, 743–755.

Shackleton, N.J., Kennett, J.P., 1975b. Late Cenozoic oxygen and carbon isotopic changesat DSDP site 284: implications for glacial history of the northern hemisphere andAntarctica. Initial Rep. Deep Sea Drill. Proj. 29, 801–807.

Shevenell, A.E., Kennett, J.P., Lea, D.W., 2004. Middle Miocene southern ocean coolingand Antarctic cryosphere expansion. Science 305, 1766–1770.

Siesser, W.G., 1980. Late Miocene origin of the Benguela upwelling system off NorthernNamibia. Science 208, 283–285.

Singh, R.K., Gupta, A.K., 2005. Systematic decline in benthic foraminiferal speciesdiversity linked to productivity increases over the last 26 Ma in the Indian Ocean.J. Foraminiferal Res. 35 (3), 219–227.

Smart, C.W., Thomas, E., Ramsay, A.T.S., 2007. Middle–late Miocene benthic foraminiferain a western equatorial Indian Ocean depth transect: paleoceanographic implica-tions. Palaeogeogr. Palaeoclimatol. Palaeoecol. 247, 402–420.

Stein, R., Robert, C., 1986. Siliciclastic sediments at Sites 588, 590 and 591: Neogene andPaleogene evolution in the southwest Pacific and Australian climate. In: Kennett, J.P.,von der Borch, C.C. (Eds.), Init. Rep. DSDP, 90, pp. 1437–1455.

Suganuma, Y., Hamada, T., Tanaka, S., Okada, M., Nakaya, H., Kunimatsu, Y., Saegusa, H.,Nagaoka, S., Ratanasthien, B., 2006. Magnetostratigraphy of the Miocene ChiangMuan Formation, northern Thailand: implication for revised chronology of theearliest Miocene hominoid in Southeast Asia. Palaeogeogr. Palaeoclimatol.Palaeoecol. 239, 75–86.

Sun, Z.C., Feng, X.J., Li, D.M., Yang, F., Qu, Y.H., Wang, H.J., 1999. Cenozoic Ostracoda andpalaeoenvironments of the northeastern Tarim Basin, western China. Palaeogeogr.Palaeoclimatol. Palaeoecol. 148, 37–50.

Thouveny, N., de Beaulieu, J.-L., Bonifay, E., Creer, K.M., Gulot, J., Icole, M., Johnsen, S.,Jouzel, J., Reille, M., Williams, T., Williamson, D., 1994. Climate variation in Europeover the past 140 kyr deduced from rock magnetism. Nature 371, 503–506.

Twichell, S.C., Meyers, P.A., Diester-Haass, L., 2002. Significance of high C/N ratios inorganic-carbon-rich Neogene sediments under the Benguela Current upwellingsystem. Org. Geochem. 33, 715–722.

Van Campo, E., 1996. Holocene environmental changes in Bangong Co basin (WesternTibet). Part 2: the pollen record. Palaeogeogr. Palaeoclimatol. Palaeoecol. 120,49–63.

Walter, H., 1962. Plant associations in the humid tropics of India as affected by climatewith special reference to periods of drought in the monsoon region. Humid TropicsResearch Program, Final Rep. UNESCO NS/HT. 6.

Wan, S.M., Li, A.C., Stuut, J.W., Xu, F.J., 2007. Grain-size records at ODP Site 1146 from thenorthern South China Sea: implications on the East Asian monsoon evolution since20 Ma. Sci. China Series D 50 (10), 1536–1547.

Wang, W.-M., Saito, T., Nakagawa, T., 2001. Palynostratigraphy and climatic implicationsof Neogene deposits in the Himi area of Toyama Prefecture, Central Japan. Rev.Palaeobot. Palynol. 117, 281–295.

Wang, Y., Deng, T., 2005. A 25 m.y. isotopic record of paleodiet and environmentalchange from fossil mammals and paleosols from the NE margin of the TibetanPlateau. Earth Planet. Sci. Lett. 236, 322–338.

Webb, S.D., 1977. A history of savanna vertebrates in the New World. Part I: NorthAmerica. Annu. Rev. Ecol. Syst. 8, 355–380.

Wolfe, J.A., 1980. Tertiary climates and floristic relationships at high latitudes in theNorthern Hemisphere. Palaeogeogr. Palaeoclimatol. Palaeoecol. 30, 313–323.

Wolfe, J.A.,1985. Distribution ofmajor vegetational types during theTertiary. In: Sundquist,E.T., Broecker, W.S. (Eds.), The Carbon Cycle and Atmospheric CO2: Natural VariationsArchean to Present (Geophys. Monogr. 32). AGU, Washington, D.C., pp. 357–375.

Woodruff, F., Savin, S.M., Douglas, R.G., 1981. Miocene stable isotope record: a de-tailed deep Pacific Ocean study and its paleoclimatic implications. Science 212,665–668.

Wright, J.D., Miller, K.G., 1996. Control of North Atlantic Deep Water circulation by theGreenland-Scotland Ridge. Paleoceanography 11, 157–170.

Wu, Z.Y., 1980. Vegetation of China. Science Press, Beijing. 1144pp. (in Chinese).Xiao, J.L., Wu, J.T., Si, B., Liang, W.D., Nakamura, T., Liu, B.L., Inouchi, Y., 2006. Holocene

climate changes in the monsoon/arid transition reflected by carbon concentrationin Daihai Lake of Inner Mongolia. The Holocene 16, 551–560.

Yang, P., Sun, Z.C., Li, D.M., Jing, M.C., Xu, F.T., Liu, H.M., 2000. Ostracoda extinction andexplosion events of the Mesozoic–Cenozoic in Qaidam Basin, Northwest China.J. Palaeogeogr. 2 (3), 69–74 (in Chinese with English abstract).

Zachos, J., Pagani, M., Sloan, L., Thomas, E., Billups, K., 2001. Trends, rhythms, andaberrations in global climate 65 Ma to present. Science 292, 686–693.

Zhang, J., Ma, Z.J., Ren, W.J., 2005. The sedimentary characteristics of Cenozoic strata incentral and southern Ningxia and their relationships with the development of theQinghai–Tibetan Plateau. Acta Geol. Sin. 79, 757–773 (in Chinese with Englishabstract).

Zhao, Q.H., Wang, P.X., Cheng, X.R., Wang, J.L., Haung, B.Q., Xu, J., Zhou, Z., Jian, Z.M.,2001. A record of Miocene carbon excursions in the South China Sea. Sci. ChinaSeries D 44 (10), 943–951.