Palaeogeography, Palaeoclimatology, Palaeoecology 160 (2000) 291–299
www.elsevier.nl/locate/palaeo
C /C vegetation evolution over the last 7.0 Myr in
3 4
the Chinese Loess Plateau: evidence from
pedogenic carbonate d13C
Z.L. Ding *, S.L. Yang
Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, People’s Republic of China
Received 16 September 1999; received in revised form 3 February 2000; accepted for publication 10 February 2000
Abstract
The stable carbon and oxygen isotopic compositions of soil carbonate were measured on an eolian loess and red
clay sequence at Lingtai, the Chinese Loess Plateau. This sequence is composed of 130 m of Tertiary red clay deposits
with a basal age of 7.05 Ma overlain by 175 m of Pleistocene loess. In the field we identified ca. 110 carbonate nodule
horizons in the red clay and 27 nodule horizons in the loess. These carbonate nodule horizons are formed by leaching
and re-precipitation of carbonate from the eolian material. The d13C record of soil carbonate indicates a major
expansion of C plants at ca. 4.0 Myr in the Loess Plateau. This event is comparable in timing with the expansion of
4
C plants in northern North America (Cerling et al., 1997. Nature 389, 153–158) but is ca. 3 million years later than
4
the C biomass expansion in Pakistan (Quade et al., 1989. Nature 342, 163–166). The pedogenic characteristics of
4
the soils and the d18O record in the red clay suggest that the C plant expansion in the Loess Plateau was not driven
4
by local climatic changes, which may support Cerling et al.’s (1997) assertion that the decline of atmospheric CO
2
levels in the Neogene is responsible for this global vegetation change. Our record also implies that the Tibetan Plateau
could have been uplifted to a critical height in the late Miocene, thus resulting in the formation of the atmospheric
Great East-Asia Trough. © 2000 Elsevier Science B.V. All rights reserved.
Keywords: atmospheric CO levels; carbon isotope record; eolian deposits; paleovegetation
2
1. Introduction
The relative proportion of C and C plants in
3
4
local biomass can be inferred from the d13C compositions of soil organic matter and pedogenic
carbonate (Cerling, 1984; Cerling et al., 1989).
Pure C and C plants have d13C values ranging
3
4
from ca. −22 to −30‰ and −10 to −14‰,
* Corresponding author. Tel.: +86-10-6200-8111;
fax: +86-10-6491-9140.
E-mail address: [email protected] (Z.L. Ding)
respectively (Bender, 1971; Winter et al., 1976;
Brown, 1977; Vogel et al., 1978; Farquhar et al.,
1989). Soil organic matter preserves this isotopic
distinction with little or no isotopic fractionation,
but pedogenic carbonate and carbonate in the
dental enamel of fossil herbivores are enriched in
13C by ca. 14 (25°C ) to 17‰ (0°C ) with respect
to source carbon (Lee-Thorp and van der Merwe,
1987; Cerling et al., 1989; Wang et al., 1994).
Therefore, under conditions of moisture stress,
d13C values of −8 and +4‰ in pedogenic carbonate indicate nearly pure C and C ecosystems,
3
4
0031-0182/00/$ - see front matter © 2000 Elsevier Science B.V. All rights reserved.
PII: S0 0 3 1 -0 1 8 2 ( 0 0 ) 0 0 07 6 - 6
292
Z.L. Ding, S.L. Yang / Palaeogeography, Palaeoclimatology, Palaeoecology 160 (2000) 291–299
respectively ( Ehleringer et al., 1986; Ehleringer
and Cooper, 1988; Farquhar et al., 1989).
Recently, Cerling et al. (1997) have reported
changes in the carbon isotope ratios of fossil tooth
enamel from Pakistan, Africa, North America and
South America since the late Miocene. The authors
determined that there was a global expansion of
C biomass between 8 and 4 million years (Myr)
4
ago, with the C plant expansion occurring first in
4
low latitudes and later in higher latitudes. In
western Europe, there is no evidence suggesting a
significant C component in the diet of large
4
mammals at any time. To explain this variable
spatial–temporal pattern, Cerling et al. (1997)
developed a model in which both atmospheric
CO level and temperature are thought to have
2
controlled the global spatial and temporal evolution of the C /C plants. As the continental weath3 4
ering increased, due to the quick uplifting of the
Tibetan Plateau in the Neogene (Raymo and
Ruddiman, 1992), atmospheric CO concentration
2
has been lowered to a critical level that favors C
4
over C plants in the low latitudes with an elevated
3
temperature. With the further decrease in atmospheric CO levels, the crossover point favoring
2
C over C plants will be reached later in higher
4
3
latitudes with relatively low temperatures (Cerling
et al., 1997). This means that d13C records of
geological sediments are ideal in reconstructing the
history of mutual interactions between the lithosphere, biosphere and atmosphere of the Earth
system.
However, this atmospheric CO -climate2
ecosystem linkage is challenged by two new pieces
of evidence from marine sediments. Based on
boron isotope composition measurements of
planktonic foraminifera from tropical Pacific
Ocean sediments, Pearson and Palmer (1999)
established a pH profile for ancient ocean surface
water, which suggests that atmospheric CO partial
2
pressure during the middle Eocene was probably
similar to modern concentrations or only slightly
higher. This means that the extreme warmth of
middle Eocene environment might not be caused
mainly by high atmospheric CO level. A pCO
2
2
record derived from alkenone d13C measurements
(Pagani et al., 1999) shows that atmospheric
pCO was surprisingly low (180–300 ppmv) over
2
the interval of 15 to 9 Myr and stabilized at
preindustrial values by 9 million years ago.
Absence of evidence for major changes in pCO
2
during the late Miocene led Pagani et al. (1999)
to suggest that the sudden expansion of C vegeta4
tion at ca. 7 Myr could be initiated by the development of low-latitude seasonal aridity and changes
in growing conditions on a global scale, rather
than a decrease in pCO .
2
As plant-type distributions are determined by
several factors, more records of paleovegetation
changes are needed to understand the mechanisms
for the late Miocene C plant expansion. The
4
Chinese Loess Plateau is located to the northeast
of the Tibetan Plateau, where the climate is controlled essentially by the Asian monsoon system.
Many studies have shown that the wind-blown
loess–soil sequence in the Loess Plateau is one of
the most complete records of late Cenozoic climatic
changes in the world (Liu, 1985; Kukla and An,
1989; Rutter et al., 1991; Ding et al., 1999). The
Asian monsoon system has experienced a longterm evolution with the uplift of the Tibetan
Plateau during the late Cenozoic, which may in
turn result in fundamental changes in the ecosystem of the Loess Plateau. In this context, paleovegetation change records from the eolian
successions and comparison with other records in
the world would help to clarify the factors causing
the global expansion of C ecosystems 4.0–
4
7.0 Myr ago.
In this paper we report a 7.0 Myr d13C record
from soil carbonate nodules formed in an eolian
loess and red clay sequence at Lingtai, the Chinese
Loess Plateau.
2. Setting and stratigraphy
Lingtai (Fig. 1) is situated in the middle of the
Loess Plateau at an elevation of ca. 1340 m above
sea level ( latitude 35°00∞33◊N, longitude
107°30∞33◊E ). At present, the mean annual temperature and precipitation at Lingtai are ca. 9.1°C
and 600 mm, respectively. The averaged temperature of July, the warmest month, is ca. 22.3°C.
The modern climate of this region is essentially
controlled by the Asia monsoon system. In the
Z.L. Ding, S.L. Yang / Palaeogeography, Palaeoclimatology, Palaeoecology 160 (2000) 291–299
293
Fig. 1. Schematic map showing the locations where the d13C records used in the paper are derived. $, Lingtai section of the Chinese
Loess Plateau; +, the Siwalik sediments in Pakistan. The dashed line in the North American continent indicates 37°N latitude.
summer season, the East-Asia summer monsoon
brings warm, moisture-laden air masses from tropical oceans to the Loess Plateau, resulting in heavy
rainfall in the region. Over 50% of the annual
precipitation occurs from July to September.
During the winter season, the winter monsoon
winds flowing from the Siberian region prevail in
the Loess Plateau, leading to a dry and cold
climate.
Recently, a loess and red clay sequence, ca.
305 m thick, was discovered and studied at Lingtai
(Ding et al., 1998a, 1999). The loess deposit is ca.
175 m thick, and contains ca. 40 paleosols. Field
observation and magnetic susceptibility and
median grain size (Md) records (Fig. 2) show that
the Lingtai loess and soil sequence can be well
correlated with the standard sections of the Loess
Plateau (Liu, 1985; Kukla and An, 1989; Rutter
et al., 1991). Paleomagnetic studies indicate a basal
age of ca. 2.6 Myr for the Lingtai loess ( Fig. 2).
The paleosols within the loess are generally characterized by a brownish or reddish color, and substantial clay skins. Horizontal carbonate nodule
horizons are generally not observed immediately
below the pedogenic B horizons for the paleosols
in the upper part of the loess. However, scattered
nodules are common at the base of the soils. Most
of the paleosols in the lower part of the loess
sequence have ca. 20–80 cm thick nodule horizons
(Ding et al., 1999).
The red clay underlying the loess is ca. 130 m
thick. Paleomagnetic measurements (Ding et al.,
1998a) suggest that it accumulated from 2.6 to
7.05 Myr (Fig. 2). The magnetic polarity stratigraphy shown in Fig. 2 was established by analyses
of 625 orientated samples at an interval of 15–
25 cm from the red clay sequence. The magnetic
remanence was measured at the Paleomagnetism
Laboratory in the Institute of Geology and
Geophysics, Chinese Academy of Sciences, with
a 2G three-axis cryogenic magnetometer.
Remanence data are obtained after alternative
demagnetization, usually at 20 or 25 mT. In most
cases, an alternative demagnetization at 15–30 mT
leads to an unambiguous polarity assignment.
The Matuyama/Gauss (M/G) reversal occurred
in the upper part of the oldest loess of L33 ( Fig. 2).
Within the Gauss Chron, there are two short
reversed zones occurring respectively at depths of
190 and 196 m. They are interpreted as subchrons
2An.1r and 2An.2r (i.e. the Kaena and Mammoth
Events). The 2An/2Ar boundary (Gauss/Gilbert
boundary) falls near a depth of 206 m, and the
bottom of 2Ar can be defined at the depth of ca.
219.5 m. Below ca. 220 m, the polarity zonation is
readily obtained from the remanence data, which
can be easily correlated to the standard polarity
time scale (Cande and Kent, 1995).
Sedimentological and geochemical studies have
demonstrated that this Tertiary red clay is windblown in origin as is the overlying Pleistocene loess
(Ding et al., 1998b). One of the most striking
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Z.L. Ding, S.L. Yang / Palaeogeography, Palaeoclimatology, Palaeoecology 160 (2000) 291–299
Fig. 2. Magnetic susceptibility and median grain size (Md ) records of the Lingtai red clay–loess sequence with the magnetic reversal
polarity [modified from Ding et al. (1998a)]. Most of the soil units (Si) and some thick loess beds (Li) are indicated.
features of the red clay sequence is the existence
of many horizontal carbonate nodule horizons.
The thickness of these horizons ranges from 10 cm
to >100 cm, with most of the nodules <10 cm in
diameter. The groundmass within most of the
nodule horizons is reddish, weathered soil material.
In addition to the nodule horizons, individual
nodules are commonly scattered throughout the
red clay. The red clay between the nodule horizons
shows a redder color than that of the paleosols in
the overlying Pleistocene loess. The horizons have
been subjected to relatively strong pedogenesis, as
indicated by the relative abundance of clay and
Fe–Mn skins, and are identified as pedological B
horizons. A remarkable difference between the soil
horizons of the red clay and loess records is that
pedogenic A horizons are generally lacking in the
red clay sequence, whereas a complete A–B–C
Z.L. Ding, S.L. Yang / Palaeogeography, Palaeoclimatology, Palaeoecology 160 (2000) 291–299
sequence is readily recognized in the loess. A total
of ca. 120 soil B horizons and ca. 110 intervening
carbonate nodule horizons are identified in the
Lingtai red clay sequence (Ding et al., 1999).
3. Sampling and analysis
In the Pleistocene loess sequence, we took one
carbonate nodule from each of the carbonate
nodule horizons if nodules existed at the base of
the paleosols; 27 nodule samples were collected in
the loess–soil record. In the red clay sequence, one
sample was taken from each thin nodule layer,
while in every thick nodule horizon, two samples
from the upper and lower parts were collected. A
total of 178 samples were taken in the red clay
record.
Nodules were ground to a powder, and converted to CO with 100% phosphoric acid. Carbon
2
and oxygen isotopic ratios were both measured in
a DeltaS gas-source mass spectrometer. Isotopic
results are presented in the usual d notation as the
permil (‰) deviation of the sample CO from the
2
PDB standard
A
R
B
sample −1 ×1000
R
standard
where R=13C/12C or 18O/16O. Replicate analyses
(n=13) of a single, homogenized sample show that
this procedure yields a standard deviation of
±0.05 and ±0.06‰, respectively, for d13C and
d18O measurements.
d=
4. Results
The carbon and oxygen isotopic records of soil
carbonate nodules and magnetic susceptibility are
shown in Fig. 3 for the Lingtai loess and red clay
sequence. The time scale was established by linear
interpolation between paleomagnetic reversal
boundaries (Ding et al., 1998a). The carbon isotope record can be grouped roughly into three
phases. From 7.05 to ca. 4.0 Myr, the d13C values
of soil carbonate fall between −9.6 and −7.8‰,
indicating a C -dominated ecosystem in the Lingtai
3
area. At ca. 4.0 Myr, the stable carbon isotopic
295
compositions of soil carbonate begin to shift
toward more positive values. Most of the d13C
values lie between −7.8 and −5.3‰ from 4.0 to
2.0 Myr (Fig. 3). A single sample at the 3.4 Myr
level yielded the d13C value of −4.7‰, representing the highest d13C value in the entire red clay
sequence. Over this interval, a mixed ecosystem
of C and C plants can be interpreted from the
3
4
stable carbon isotopic compositions of soil carbonate. Therefore, the expansion of C plants on the
4
Loess Plateau occurred at ca. 4.0 Myr, but at no
time did C plants dominate the vegetation on the
4
plateau. From 2.0 Myr to the Holocene, most of
the d13C values in soil carbonate range from −6.5
to −9.2‰ with an averaged value of ca. −8.0‰
( Fig. 3). Compared to the interval of 4.0 to
2.0 Myr, C plants in the Loess Plateau ecosystem
3
increased significantly in the Pleistocene. However,
in the youngest two paleosols (S1 and S2), soil
carbonate is enriched in 13C (Fig. 3), suggesting
another expansion of C plants on the plateau.
4
In the Pakistan stable isotopic records (Quade
et al., 1989; Quade and Cerling, 1995), the shift
of d13C compositions toward more positive values
at ca. 7.0 Myr was accompanied by a clear increase
in the d18O record. This feature is not found in
our d18O record. Changes in soil carbonate d18O
compositions range scatteringly from −11.8 to
−8.0‰ (Fig. 3). There is an overall trend of
d18O compositions toward more positive values
from bottom to top of the loess and red clay
sequence. Changes in d18O compositions of soil
carbonate may be controlled by complicated
factors, including temperature, regional moisture
sources, rainfall amount, and seasonality of precipitation (Quade and Cerling, 1995). The causes for
the trend of d18O changes of soil carbonate in the
Loess Plateau are unclear. We preliminarily think
that this may be partly related with the increase
in seasonality during the past 7.0 Ma due to the
long-term intensification of both the summer and
winter monsoon systems over East Asia (Liu and
Ding, 1993; Ding et al., 1999).
5. Discussion and conclusions
Our carbon isotopic record indicates a significant expansion of C plants at ca. 4.0 Myr on the
4
296
Z.L. Ding, S.L. Yang / Palaeogeography, Palaeoclimatology, Palaeoecology 160 (2000) 291–299
Fig. 3. Magnetic susceptibility and carbon and oxygen isotope records plotted against time. The timescale is constructed by linear
interpolation between magnetic reversal boundaries.
Chinese Loess Plateau. A major question concerning this event is whether or not it was driven by
local climatic changes. The pedogenic characteristics of the soils within the red clay sequence have
recorded the history of the East-Asia summer
monsoon evolution during the late Miocene and
Pliocene in the Loess Plateau. Recent observations
(Ding et al., 1999) show that the Lingtai red clay
sequence can be subdivided into five units according to soil features. The first ( lowest) unit accumulated from 7.05 to 6.2 Myr. Soils in this unit are
characterized by a reddish brown color and a weak
subangular blocky structure with few clay skins,
suggesting a relatively weak summer monsoon.
The second unit, formed during 6.2 to 5.5 Myr,
contains soils showing a moderate subangular
blocky structure and some or common clay skins,
implying a strengthened summer monsoon relative
to the earlier time. The soils of the third unit (5.5
to 3.85 Myr) display many thick clay skins and
many dark Fe–Mn films, interpreted as having
developed under strong summer monsoon conditions. The summer monsoon was significantly
weakened over the interval of 3.85 to 3.15 Ma (the
fourth unit), as suggested by greatly reduced clay
skins on the pedogenic structure surfaces. In the
fifth unit (3.15 to 2.6 Myr), translocated clay skins
in the soils again significantly increase, implying
an intensified summer monsoon climate. This interpreted monsoon evolution suggests that there is
no significant local climatic event at ca. 4.0 Myr,
and thus it seems unlikely that the ecological
change registered in the d13C record was be forced
solely by changes in local climatic conditions. This
idea may be supported also by the soil carbonate
d18O record ( Fig. 3) which shows an almost linear
Z.L. Ding, S.L. Yang / Palaeogeography, Palaeoclimatology, Palaeoecology 160 (2000) 291–299
enrichment in 18O from bottom to top of the
sequence. No peculiar change ca. 4.0 Myr is found
in the record. Therefore, if the d18O trend indeed
represents the increase in seasonality of precipitation, this factor is not the most important in
driving the middle Pliocene C plant expansion on
4
the Loess Plateau.
The modern spatial distribution of C and C
4
3
grasses shows that C grasses dominate in the
4
tropical and subtropical regions, that the transition
to C grasses takes place between ca. 30 and 45°
3
latitudes, and that C grasses dominate at high
3
latitudes. According to the model developed by
Cerling et al. (1997), the photosynthetic efficiency
of C grasses relative to C grasses varies with
3
4
297
both atmospheric CO levels and temperature.
2
This model predicts that with the continuous
decrease in atmospheric CO levels during the late
2
Cenozoic, as deduced from geological records,
expansion of C plants will occur first in the
4
tropical and subtropical regions, and later in the
transitional latitudes because of lower temperatures. Fig. 4 shows the comparison of d13C
records between Pakistan, the Chinese Loess
Plateau and northern North America (>37°N ). It
is evident that the Loess Plateau and northern
North America experienced a synchronous expansion of C plants 4.0 Myr ago, ca. 3 million years
4
later than the C expansion in Pakistan. This may
4
support the model of Cerling et al. (1997),
Fig. 4. Comparison of the C -plant expansion event between the Loess Plateau (pedogenic carbonate d13C ), Pakistan (pedogenic
4
carbonate d13C ) and northern North America ( Equid d13C ). The arrows indicate the positions of the event. The Pakistan data are
from Quade et al. (1989) and the northern North America data from Cerling et al. (1997).
298
Z.L. Ding, S.L. Yang / Palaeogeography, Palaeoclimatology, Palaeoecology 160 (2000) 291–299
although direct evidence for high concentration of
atmospheric CO in the Neogene is lacking (Pagani
2
et al., 1999; Pearson and Palmer, 1999).
The Pakistan carbon isotopic records are
derived from between 32 and 33°N, only 2–3° of
latitude lower than the Lingtai section (35°N ).
The difference in the timing of the C -plant expan4
sion between the two regions is, however, as great
as ca. 3 million years. On the other hand, northern
North America has substantially higher latitudes
than Lingtai, but the two regions witnessed an
approximately synchronous C -plant expansion.
4
This evidence suggests an anomalously low temperature of the Loess Plateau ca. 7.0 Myr ago, according to the Cerling et al. (1997) model. At present,
cold air masses propagated onto the Plateau are
mainly delivered by the northerly winter monsoon
and the westerlies. Changes in grain size values of
four red clay–loess sections demonstrate that the
southward gradient of decreasing grain size in the
Tertiary red clay on the Plateau is negligible (Ding
et al., 1998b), whereas there is a strong negative
gradient in particle size from north to south in the
Loess Plateau for sediments younger than 2.6 Ma.
This decrease in loess particle size is consistent
with the direction of the winter monsoonal winds
flowing from Siberia. It is suggested that the winter
monsoon may not have formed until ca. 2.6 Myr
ago (Ding et al., 1998b). In this case, the lowered
temperature background over the Loess Plateau
during the late Miocene may be caused by
increased propagation of cold air masses by the
westerlies. Meteorological observations ( Wallace,
1983) show that at 500 mb, there is a broad trough,
called the Great East-Asia Trough (GEAT ), centered about the longitudes of Japan and forming
a broad ridge to the west, which produces northwesterly winds in the upper troposphere over
China. The GEAT is the strongest of the quasistationary waves on the globe, which is partly
responsible for the relatively low temperature over
eastern and central China. GCM modeling experiments (Manabe and Terpstra, 1974; Kutzbach
et al., 1989) show that this stationary wave over
Asia is caused by the uplifted Tibetan Plateau.
The greatly-delayed shift of C plants in the Loess
4
Plateau suggests that the Tibetan Plateau may
have reached a critical height in the late Miocene,
thus leading to the formation of the GEAT. This
is consistent with the observations of several geological studies (Mercier et al., 1987; Copeland and
Harrison, 1990; Harrison et al., 1993).
Our d13C record shows a subsequent diminishment of the C -plant component in the period of
4
2 Myr to the late Pleistocene ( Fig. 3). This phenomenon is also observed in the dental enamel
carbon isotope record of fossil herbivores from
northern North America (Cerling et al., 1997). To
date, the cause for this is unknown. As no evidence
is obtained for a subsequent increase in atmospheric CO levels during the Pleistocene, we pro2
pose that decrease in temperature could be the
main factor for this event. Dramatic expansion of
polar ice sheets in the Northern Hemisphere during
late Pliocene and early Pleistocene has been indicated by many geological records (Ruddiman
et al., 1989; Raymo et al., 1989; Shackleton et al.,
1990). Large-scale ice sheet expansion in the polar
area will lead to a decrease in temperature background and an increase in north–south temperature gradient, thereby causing a southward shift
of C taxa in the Northern Hemisphere.
3
Acknowledgements
This research is supported by the NNSF of
China (49525203) and CAS ( KZ951-A1-402). The
authors are greatly indebted to Dr Eve Arnold for
critical comments on the manuscript.
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