Palaeogeography, Palaeoclimatology, Palaeoecology 310 (2011) 464–470
Contents lists available at ScienceDirect
Palaeogeography, Palaeoclimatology, Palaeoecology
j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / p a l a e o
Comments on “Anti-phase oscillation of Asian monsoons during the Younger Dryas
period: Evidence from peat cellulose δ 13C of Hani, Northeast China” by B. Hong, Y.T.
Hong, Q.H. Lin, Yasuyuki Shibata, Masao Uchida, Y.X. Zhu, X.T. Leng, Y. Wang and C.C.
Cai [Palaeogeography, Palaeoclimatology, Palaeoecology 297 (2010) 214–222]
Martina Stebich a,⁎, Jens Mingram b, Robert Moschen c, Annett Thiele d, Christian Schröder e
a
Senckenberg Research Institute and Natural History Museum, Research Station for Quaternary Palaeontology, Weimar, Germany
Deutsches GeoForschungsZentrum, Dep.5.2, Climate Dynamics and Landscape Evolution, Potsdam, Germany
c
Institute of Bio- and Geosciences: Agrosphere (IBG-3), Research Centre Juelich, Juelich, Germany
d
APB-BirdLife Belarus, Minsk, Belarus
e
Greifswald University Institute of Botany and Landscape Ecology, Greifswald, Germany
b
a r t i c l e
i n f o
Article history:
Received 15 February 2011
Received in revised form 7 June 2011
Accepted 8 June 2011
Available online 16 June 2011
Keywords:
East Asian Monsoon
Lake sediments
Younger Dryas
Late Glacial period
Palynology
Varves
a b s t r a c t
In their recent paper, Hong et al. (2010; Anti-phase oscillation of Asian monsoons during the Younger Dryas
period: Evidence from peat cellulose δ13C of Hani, Northeast China, Palaeogeography, Palaeoclimatology,
Palaeoecology 297, 214–222) discuss bulk peat sample cellulose δ13C data from a fen in northeast China as a
proxy for East Asian summer monsoon intensity during the Late Glacial period. Based on their own results,
cited papers, and an extensive re-interpretation of sedimentological and palynological data from nearby Lake
Sihailongwan, ,Hong et al. (2010) construct a hypothesis of contrasting moisture conditions in northern and
southern China, with wet conditions in the north during the Younger Dryas period and an anti-phase
behaviour of Indian- and East Asian summer monsoon intensity. However, we do not approve of the reinterpretation of our Lake Sihailongwan data by ,Hong et al. (2010) and must strongly reject it. We show here
that neither the ,Hong et al. (2010) fen data, nor the Lake Sihailongwan data or any other cited data allow for
the sound assumption of an intensified East Asian summer monsoon in northeastern China during the
Younger Dryas. The Late Glacial variability of the fen data found by ,Hong et al. (2010) can be easily explained
by changes in the plant assemblage down core and thus by the composition of the peat. Furthermore, the use
of bulk peat cellulose δ13C data as a precipitation proxy remains unproven for that area. Hence, there is no
basis for a model contrasting Indian and East Asian summer monsoons during that period.
© 2011 Elsevier B.V. All rights reserved.
1. Background and motivation
During the last decade, several papers were published based on
stable isotope investigations of extracted cellulose from northeastern
Chinese peat bog and fen material for palaeoclimate reconstructions
and evaluations of solar–terrestrial linkages Hong et al., 2000, 2001,
2005, and the two in large parts identical papers of B. Hong et al.
(2009) and Y.T. Hong et al. (2009).
The most recent publication by Hong et al. (2010) presents new
δ 13C data from peat cellulose of the Hani mire for the Late Glacial
period, with a special focus on the Younger Dryas and the transition to
the Holocene. Using the δ 13 C cellulose data as a proxy for
precipitation, and in particular for the intensity of the East Asian
Summer Monsoon (EASM), Hong et al. (2010) reconstruct a wet
DOI of original article: 10.1016/j.palaeo.2010.08.004.
⁎ Corresponding author.
E-mail address: [email protected] (M. Stebich).
0031-0182/$ – see front matter © 2011 Elsevier B.V. All rights reserved.
doi:10.1016/j.palaeo.2011.06.004
Younger Dryas (YD) and come to the far-reaching conclusion of an
“anti-phase oscillation of Asian monsoons during the Younger Dryas
period”. We admire the independent approach of Hong et al., 2010
which does not obediently follow widely accepted theories. However,
because Hong et al., (2010) extensively discuss data from nearby Lake
Sihailongwan (published in Schettler et al., 2006 and Stebich et al.,
2009) and re-interpret these data in support of their theory, we feel
strongly obliged to comment their paper.
As their new find for the YD climate in northern China, Hong et al.
(2010) state in the abstract that “Both the peat cellulose record and a
pollen record from Lake Sihailongwan sediment indicate an abrupt
increase in precipitation in the region during the Younger Dryas
period”. Furthermore, in their conclusion they deduce that “Hani peat
cellulose isotope and the Sihailongwan lake pollen indicate the same
phenomenon, that the climate of the YD cold period in north-eastern
China is wet, having increased precipitation.” With respect to Lake
Sihailongwan (SHL) pollen and sediment data, the interpretation
given by Hong et al. (2010), although almost entirely based on results
M. Stebich et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 310 (2011) 464–470
published in Stebich et al. (2009), is the opposite of what we
concluded. In this paper, we would like to examine the claims of Hong
et al. (2010), and discuss them in the context of a wider regional
pattern of precipitation during the YD.
2. The supportive evidence selected by Hong et al. 2010
Before presenting and discussing their own data, Hong et al. state
that “available knowledge of the EASM during the YD is still quite
vague, and that research studies examining lake sediments and loess
paleosols have suggested an abrupt strengthening of the EASM during
the YD interval (Kelts et al., 1989; An et al., 1993; Wang et al., 1994a,
1994b; Zhou et al., 1996), which means changes in the EASM and
IOSM had an anti-phase relationship during this period”. However,
available environmental and climate proxy data for the YD period of
the East Asian Summer Monsoon region are far from vague or sparse
(Hulu Cave, Wang et al., 2001; Dongge Cave, Yuan et al., 2004; Lake
Suigetsu, Nakagawa et al., 2006; Kossler et al., 2011, Huguang Maar,
Yancheva et al., 2007; Lake Sihailongwan, Parplies et al., 2008 etc.).
Contrary to the claims of Hong et al. (2010), there are several papers
that provide evidence of a predominantly cool-dry YD period in
eastern and northeastern China. Herzschuh (2006) evaluates 75
palaeoclimate records from monsoonal Central Asia, and out of 24
sites with records of the YD in northeastern and eastern China, only
three localities show “wet” or “moderately wet” conditions, whereas
21 sites show “moderate dry” or “dry” conditions. New speleothem
records from the Qinling Mountains south of the loess plateau show
similar Late Glacial features (Zhang et al., 2010) and an increase in
summer monsoon intensity at the start of the Holocene (Cai et al.,
2010).
Most surprisingly, Hong et al. (2010) do not cite the available highresolution pollen data from the Hani mire itself – obviously generated
from the same core – which reveals a picture more consistent with
pollen data from Lake SHL. Yu et al. (2008) state that “From
12.8 cal ka B.P. to 12.1 cal ka B.P., herbs expanded rapidly and dominated the area, indicating a dryer climate which may be part of the
Younger Dryas event”. Upon closer inspection of the literature, the
choice of the four papers cited by Hong et al. (2010) is very selective
and does not seem to be very supportive of their hypothesis of an
increased summer monsoon during the YD period. A closer look at the
details of the mentioned papers follows:
1. Kelts et al. (1989) postulate from their Lake Qinghai profile a
“seasonal input of melt- or freshwater during a wetter episode
from about 12,000 to 10,500 years” and a “stronger seasonal
meltwater signal which suggests higher inflow rates”, but refer to
their results as preliminary. Their age model is based on dates of 5
occurrences of aquatic Ruppia or algal threads – without discussing
the reservoir effect – for the whole profile. The analytical error of
these datings results in a 2-sigma range of calibrated ages of more
than 1000 years. Later investigations by Ji et al. (2005) estimate the
reservoir effect at 1039 years. Based on a multi-proxy approach
(which included pollen data as well) Ji et al. (2005) depict two
“Colder and more arid phases … possibly corresponding to the
Older Dryas and Younger Dryas events in Europe.”
2. From the Baxie section of the southwest margin of the Chinese
Loess Plateau, An et al. (1993) find evidence for a “strengthened
summer monsoon climate of Younger Dryas age”, but present
alternative interpretations as well, based on uncertainties of the
age model.
Later studies from the same region, including the Baxie section,
come to the conclusion that “from 12,800 to ~ 11,700 cal yr B.P.
(11,000 to 10,000 14C yr B.P.) … the proxy indicators show the
sudden onset of a much colder dry period corresponding to the
European Younger Dryas” (Zhou et al. 1999).
465
3. The paper attributed by Hong et al. to Wang et al. (1994a) presents
no evidence of increased precipitation during the Younger Dryas at
all, but finds instead that “… the record of Younger Dryas event of
Hulun Lake is characterised by cold and dry fluctuation”. However,
the correct citation of the English paper version must be Wang et
al., (1994b) “The Record of Younger Dryas Event in Lake Sediments
From Jalai Nur, Inner Mongolia”, Chinese Science Bulletin 39/10,
831–835. The pages given by Hong et al. (2010) for Wang et al.
(1994a) contain the Chinese version from Chinese Science Bulletin
39/4, 348–351. But also the original Chinese version concludes a
dry and cold Younger Dryas period: “由此可见,呼伦湖的新仙女木事件记
录具有冷干波动的特点” (Wang et al., 1994a). Independently, and
based on sedimentology and diatom analyses, Xue et al. (2003) find
that in the Hulun Lake basin an “An abrupt lake level drop and dry
climatic conditions occurred during 11,200–10,600 (uncal) yr B.P.”
4. Zhou et al. (1996) study two sections in the desert–loess transition
zone of central China: the Midiwan section and the Yangtaomao
profile. Based on their paleomonsoon proxy records from peat and
aeolian sand-paleosol sequences, Zhou et al. (1996) reconstruct a
“rapid oscillation from cold-dry conditions (11,200–10,600 14C yr B.P.)
to cool-humid conditions (10,600–10,200 14C yr B.P.), followed by
a return to cold-dry climate (10,200–10,000 14C yr B.P.)” which
means that the climate “of Younger Dryas time in central China
thus shows an oscillation from cold-dry to cool-humid to cold-dry
conditions”.
3. On the re-interpretation of Lake SHL pollen data by Hong et al.
(2010)
In their discussion of the Hani mire data, Hong et al. (2010)
extensively rely on data from the Lake SHL pollen record published in
Stebich et al. (2009) and interpret them in favour of their theory. Most
unfortunately, Hong et al. (2010) neither do consider the dynamics of
Late Glacial vegetation developments, nor discuss moisture conditions
in the context of the Late Glacial temperature changes. Hence we must
correct the misinterpretations of our data by Hong et al. (2010) in
detail.
In their section on the Bølling–Allerød warming period, Hong et al.
(2010) cite Lake SHL pollen data as having an increase in “some
drought-tolerance taxa, such as Ulmus pumila and Quercus mongolica”
in the period between 14,200 and 13,700 cal yr BP. However, in our
paper we did not give details on Ulmus or Quercus species, and did not
relate Ulmus or Quercus genera to drought tolerance. In fact, the Late
Glacial pollen assemblages (Stebich et al., 2009) reveal a mosaic-like
occurrence of taiga, cool-coniferous forests, cool-mixed forests, and
steppe, indicating a sensitive ecotone situation. Thus, the vegetation
shows a series of changes triggered by rapid climate variations. In
comparison to the glacial vegetation (Stebich et al., 2007, 2009), the
most significant feature of the late-glacial vegetation development is
the substantial spread of the (cool-) temperate Ulmus and Fraxinus.
Ulmus, which is also an element of the interstadial plant associations,
may occur in various biomes (cold-mixed, cool coniferous, cool-mixed
and temperate deciduous forests) indicating a heat accumulation of
900–N1200 GDD5 during the growing season and winter temperatures of at least −19 °C (Prentice et al., 1992). Fraxinus, however, is
definitely a constituent of cool-mixed and temperate deciduous
forests, which requires N1200 GDD5 and coldest-month temperatures
of at least − 15 °C (Prentice et al., 1992). Nevertheless, (cool-)
temperate trees may grow under colder winter temperatures if
sufficient moisture (i.e., snow cover) outweighs the effects of winter
temperatures that are normally too cold for temperate trees (up to
−26 °C; Mokhova et al., 2009). Consequently, Ulmus and Fraxinus are
indicators of warm conditions during the growing season.
For the period between 13,700 and 12,700 cal yr BP, Hong et al.
(2010) conclude from their δ 13C data that “the EASM has strengthened and rainfall has increased”, and from our Lake SHL data
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“vegetation changes that do not correspond to changes in the varve
thickness of Lake SHL and the geochemical index during this period”.
First, neither we nor Schettler et al. (2006) define a geochemical index
for any period of the Lake SHL record. We always discuss multi-proxy
data in context. Unfortunately, Hong et al. apply a rather static view of
vegetation and do not take into account vegetation dynamics in times
of rapid climate change, as during the Late Glacial period. What we
state – which perhaps led Hong et al. (2010) to their assertion – is that
“in contrast to the documented vegetation change around
13,600 cal yr BP, no corresponding feature was found in the
sedimentological, geochemical or isotopic records”. We conclude
that this dynamic spread of mixed broadleaf forests at approximately
13,600 cal yr BP was the result of a critical population density
threshold reached at that time. Contrary to the (mis)interpretation
of our data by Hong et al. (2010), we find a climatic deterioration
during the later Allerød, between 13,100 and 12,900 cal yr BP
(corresponding to the European Gerzensee Oscillation), with clear
expression in the palynological-, sedimentological-, and geochemical
data.
For the YD period, Hong et al., 2010 state an “increase of wetness”
from increased percentages of Larix, Picea, Abies, and Cyperaceae. The
transition into the YD is, however, mainly characterised by significant
declines of Ulmus and Fraxinus, which gives way to a recovery of
boreal woodland. Indeed, Picea and Abies suggest sufficient moisture
supply, but the spread of boreal conifers and boreal summergreen
trees (e.g., Betula) combined with a decrease of (cool-) temperate
elements at the beginning of the YD also indicate lower temperatures
during the winter and/or the growing season. Accordingly, the
increased percentages of Larix, Picea, and Abies during the YD do not
necessarily indicate an intensified summer monsoon. The temperature decrease (i.e., summer temperature) is underlined by higher
densities of shrub alder. Under the assumed permafrost conditions,
however, there would be an additional supply of soil water from the
seasonal snowmelt and soil thawing that could not drain due to the
underlying permafrost. Even in arid regions, permafrost soils provide
sufficient moisture for tree growth (Bonan and Shugart, 1989).
Furthermore, changes in the amount of herb pollen (including
Cyperaceae) appear rather weakly in the Lake SHL during the YD, as
indicated by the percentage and influx diagram (Stebich et al., 2009).
Cyperaceae in particular, which are often used as an indicator of wet
conditions, show higher percentages during the last glacial part of the
Lake SHL record (Stebich et al., 2009), which can be clearly attributed
to the enhanced soil moisture of seasonally thawing ground but not to
intensified summer monsoons. Therefore, the herb pollen of the Lake
SHL record cannot be used as a substantial argument for the
evaluation of the overall environmental change at the beginning of
the YD.
Also the Lake SHL pollen spectra between 11,700 and 11,200 cal yr
BP are far from consistent with those from the YD as stated in Hong
et al. (2010); instead, they represent a transitional period from coldadapted, more open YD forest stands towards a dense, Ulmus- and
Fraxinus-dominated woodland with a thermophilous, species-rich
shrub layer (including Rosaceae). During the second part of this zone,
full interglacial taxa begin to immigrate (including Tilia and Juglans).
4. The use of varve thickness and geochemistry of Lake SHL
sediments as palaeoclimate proxies
For the late-glacial climate amelioration, Hong et al. (2010) relate the
increase of varve thickness from Lake SHL exclusively to the “instability
of the soil horizon” and an “increase of the erosion effect”. Varve
thickness is indeed a valuable environmental proxy, but has to be
interpreted in context of prevailing climate, lake type, and sediment
microfacies (Brauer et al., 2009). Small boreal lakes receive only little
allochthonous supply from runoff, except for spring melt season. Under
conditions of colder climates, allochthonous input from runoff can
increase dramatically due to less vegetated or even barren soils in the
lake's catchment, reduced seepage and increased surface runoff during
warm season (e.g. Anderson and Dean, 1988; Zolitschka, 2006).
As reported in Schettler et al. (2006) and Stebich et al. (2009), our
sediment microfacies investigations are based entirely on petrographic thin sections that cover the complete Lake SHL profile,
including the Late Glacial period. These thin sections show a massive
increase of seasonal diatom blooms from approximately 14,200 cal yr
BP, which cannot be explained solely by unstable soils. Furthermore,
thin sections from the last glacial period show a frequent occurrence
of graded layers with large amounts of terrestrial plant detritus
originating from unstable soils of the lake's catchment — most likely
from the surficial summer melting of permafrost soils; but this type of
graded layering does not occur frequently during the Late Glacial
period, with the exception in some parts of the Younger Dryas.
Concordantly, both sediment density and the ratio of Al2O3 to biogenic
silica (Schettler et al., 2006) show decreased values, particularly
between 14,200 and 13,700 cal yr BP.
In contrast, for the increase in varve thickness during the Younger
Dryas period, Hong et al. (2010) “consider that an abrupt increase in
the rainfall amount would be the first basic dynamic reason” and
assume that for the “YD cold period the Hani peat cellulose δ 13C value
displayed a continuous and obvious decrease, indicating an increase in
the rainfall amount and an increase in intensity of the EASM” which,
in their opinion, is “in general consistent with the climatic condition
indicated by the pollen assembly of cold-tolerance Betula and
hygrophilous Abies, Larix, and Picea, as well as by an obvious increase
in varve thickness at Lake Sihailongwan” (Hong et al., 2010).
They further assert that “the existence of permafrost may have
buffered soil erosion so that a thick varve did not form in the Lake
Sihailongwan sediment”. We are not sure as to what Hong et al.
(2010) mean by “a thick varve”? In varved lake sediment records,
sediment microfacies investigations are usually performed to detect
different varve types and their relation to changing environmental
conditions, not single varves. Outstanding single layers are mostly
event layers, which may indeed provide valuable clues about changes
of the palaeo-environment. As shown in previous investigations (Chu
et al., 2005; Mingram et al., 2004; Schettler et al., 2006) Lake SHL is a
small, closed, groundwater-fed lake with a very small catchment. In
recent sediments derived from freeze and gravity cores and in
sediment trap samples, by far the most pronounced seasonal
allochthonous influx is from ice- and snowmelt in late spring,
comprising both local material and, to a larger extent, more distantly
originated dust. Permafrost conditions at Lake SHL, contrary to the
assumption of Hong et al. (2010), do not hinder soil erosion but
increase it due to less dense vegetation on the slopes, and solifluction.
We have clear evidence from the last glacial as well as from the Late
Glacial period – most relevantly on the YD – of increased local soil
erosion, manifested in numerous brownish graded event layers
composed mainly of reworked local soil material.
After the Younger Dryas period (11,700–11,200 cal yr BP), varves
exhibit regular, fine seasonal lamination with clastic spring melt
layers and mixed organic–minerogenic summer layers. In comparison
to the preceding YD period, varve thickness did not increase for the
reason Hong et al. (2010) provide to make the Lake SHL varve data fit
their model (“the largest difference is the absence of an obvious
increase in varve thickness, which may be related to the low surface
temperature and to permafrost which stabilizes the soil”). On the
contrary, due to stabilised, non-permafrost soils with denser
vegetation cover after the YD period thin, regular varves appear.
Under permafrost conditions at Lake SHL, graded layers with
reworked soil material frequently occur, which are clearly seen as
spikes in the varve thickness data.
Fully concordant with palynological and sedimentological data,
geochemical investigations of Lake SHL sediments reveal peaks in the
dissolved influx of molybdenum and uranium and increased total
M. Stebich et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 310 (2011) 464–470
organic carbon content at the end of the YD period, clearly indicating
climatic amelioration with increased groundwater inflow (Schettler
et al., 2006).
5. The nature of δ 13C signal and the relationship to changes in
regional vegetation
For their interpretation of the δ 13C signal, Hong et al. (2010) imply
stable peat composition over time, but detailed botanical macroremain studies of other cores from the Hani mire prove a strong
variability in the species composition that the peat is derived from,
particularly in the lower part formed during the Late Glacial period,
where brown moss peat occurs (Schröder et al., 2007; Thiele et al.,
2008). Fig. 1 illustrates the local vegetation change at two different
parts of the fen since the Late Glacial, indicating a change in water
origin rather than monsoon driven wetness change.
Moreover, even if one assumes a monospecific Carex composition,
the recent δ 13C Carex data from the Hani mire (Hong et al., 2005)
show a variation of more than 3‰. Consequently, almost all of the
variability of the δ 13C time series shown in Fig. 2B of Hong et al.
(2010) could be explained solely by isotopic differences between
different Carex taxa. If we assume a change in plant composition as
observed by Thiele and Schröder (2007; zone Hani-D, ca. 6.40–7.40 m,
with up to 80% Calliergonella cuspidata, which has δ 13C of 27.55‰,
representing the lowest δ 13C value of the recent Hani mire vegetation
measured by Hong et al., 2005), even the most extreme values of the
Δδ 13C graph can be easily explained without any change in
environmental conditions (which, of course, may have influenced
peat composition change).
From another core taken near the centre of the Hani mire,
Yamamoto et al. (2010) investigate compound-specific carbon
isotope ratios of leaf-wax n-alkanes and revealed vegetational change
as the most likely cause for isotopic variability from 14.9 to 13.2 kyr
and 12.7 to 11.6 kyr. Hence, different lines of evidence show large
changes in plant communities of the Hani mire during the Late Glacial
period, and the assumption of a “relative stability of the floristic
composition” by Hong et al. (2010) must be dismissed.
13
467
temperature or CO2 concentration might have played a major role
in the relative competition of C3 and C4 vegetation …, hence the
carbon isotopic composition in paleosols and lake sediment
sequences.”
Aiming at a palaeovegetation reconstruction, Wang et al. (2008)
find the expected negative correlation between the δ 13C value of C3
plants and annual precipitation for arid and semi-arid regions of
northern China but used bulk organic matter, not cellulose. Moreover,
Wang et al. (2008) state the importance of δ 13C changes in soil organic
matter during decomposition, with continued 13C-enrichment in
deeper horizons as well.
Hong et al. (2003), although they analyse both bulk peat material
and Carex mulieensis from the Hongyuan peat bog core (eastern
Tibetan Plateau), provide no data on a relationship between
precipitation and plant cellulose δ 13C-data of the studied material;
they refer instead to a general, summarised relationship between both
humidity and air temperature and the δ 13C of C3 plants. Finally, they
consider their record as a proxy of the combined effects of humidity
and temperature.
Hong et al. (2005) again provide no data on a relationship between
precipitation and the plant cellulose δ 13C-data of the studied
sequence, but refer to Hong et al. (2001, 2003) with the general
statement that “the stable carbon isotope composition of peat plant
cellulose can be served as a proxy indicator for the Asian summer
monsoon activity. The smaller δ 13C value indicates the stronger
activity of the summer monsoon or the moister-warmer climate
condition in the monsoon region. On the contrary, the larger δ 13C
value indicates the weaker activity of the monsoon or the dryer-colder
climate”. Both papers investigate Holocene peat sections, assume only
minor variability in peat composition and refer to general relationships between the δ 13C of C3 plant cellulose and precipitation without
any regional or local data verification. Furthermore, and contrary to
Hong et al. (2010), they relate plant cellulose δ 13C data to Asian
summer monsoon strength in general without distinguishing between temperature and moisture effects.
7. On δ 18O and temperatures
6. On δ C and rainfall
Hong et al. (2010) state a direct and linear relation between bulk
peat cellulose δ 13C data and rainfall amount, but no proof is given
for the assumed relation between δ 13C bulk peat cellulose data and
EASM precipitation amounts for northeastern China. Interestingly,
F. Oldfield (2001) has already called for such proof for the use of δ 18O
data from the nearby Jinchuan peat bog (Hong et al., 2000, 2001) as a
proxy for air temperature. In his reply to the comment by F. Oldfield
on Hong et al. (2000), Hong (2001) argues that the resolution of peat
samples is insufficient to calibrate stable isotope data with meteorological data and give preference to “well-known climate events with
climatic information derived from archaeological and historical
records”. In their study, Hong et al. (2010) do not provide any kind
of calibration or comparison with local meteorological or historical
data. Instead, they attempt to substantiate the assumed relationship
between the δ 13C bulk peat cellulose data and precipitation based on
four citations: Lee et al. (2005), Wang et al. (2008), and Hong et al.
(2003, 2005) and the above-cited and mis-interpreted Lake SHL data.
Based on their investigations along a semi-arid to arid transect, Lee
et al. (2005) find that “the carbon isotope composition of bulk organic
matter in sediment sequences should be a dependable proxy for
precipitation in the Holocene.” For the Late Glacial period, however,
Lee et al. (2005) state that “There have been dramatic changes in
temperature and atmospheric CO2 concentrations during glacial and
interglacial transitions. The combination of the climatic parameters in
central East Asia may have been quite different from that of presentdays. Therefore, at the glacial–interglacial temporal resolution,
For the transition from the YD to the Holocene, Hong et al. (2010)
state that for the period between 11,700 and 11,200 cal yr BP “the
Hani peat δ 18O record shows an obvious cooling of the temperature.”
We do not wish to discuss the δ 18O data in detail here but refer again
to the comment by F. Oldfield (2001) on the paper by Hong et al.
(2000) on the Jinchuan mire, which we think is entirely valid for the
Hani mire δ 18O data as well. Hong et al. (2010) use the cellulose δ 18O
data on bulk peat samples published by Y.T. Hong et al. (2009) and the
nearly identical paper of B. Hong et al. (2009) as a proxy for air
temperature. They do not give any proof for the relation between Hani
mire δ 18O data and local or regional air temperature variations.
Furthermore, the measured recent δ 18O data on two different modern
Carex species already differ by approximately 4.5‰ (although it is not
specified whether whole plant or plant cellulose data are given by B.
Hong et al. (2009) and Y.T. Hong et al. (2009)).
The evidence cited by Hong et al. (2010) of a link between the
EASM and the North Atlantic ice rafting debris (IRD) events comes
from an earlier publication of data from the Hani mire by Y.T. Hong
et al. (2005) with reference to Bond et al. (2001) for the IRD events.
Later works on the North Atlantic did not support the existence of one
pervasive cycle of IRD events during the Holocene but instead found
different modes of IRD variability (Moros et al., 2006; Andrews et al.,
2009) and a “minimum sea-ice cover from 11,500 to 6000 cal years
BP” (De Vernal and Hillaire-Marcel, 2006). The second reference cited
by Hong et al. (2010), to Schettler et al. (2006), is invalid because
there are no IRD events or IRD-related monsoonal variations
mentioned in that paper.
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M. Stebich et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 310 (2011) 464–470
Fig. 1. Time-dependant compositional variability of the Hani mire. a) Lithology and age–depth profile according to Zhou et al. (2010). b) Lithology and macrofossil analysis according to Thiele et al. (2008). Simplified macrofossil diagram, data
according to percent by volume or (n) absolute occurrences. Explanation of macrofossil zones: A — terrestrialisation of the water body; former water body is indicated by remnants of Nuphar spec. seeds, terrestrialisation by Menyanthes trifoliata seeds
and Potentilla palustris nuts; B — brown moss peat, Calliergonella cuspidata stems/leaflets, indicating the formation of initial percolation mire (shaded); C — zone of disturbance with sand layer allowing connexion to the profile of Zhou et al. 2010;
D — radicel peat of percolation mire, mostly formed by sedges; E — radicel-sphagnum peat, increase of Sphagnum remnants, Eriophorum vaginatum cataphylls and others imply acidification and decrease of groundwater supply.
M. Stebich et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 310 (2011) 464–470
8. Notes on age model and resolution of the Hani peat and Lake
SHL sediment records
Hong et al. (2010) base their age model on distinct AMS 14C dates
with linear interpolation between them. The reported 1σ uncertainty
range of the inferred calibrated ages for the Late Glacial period
(samples HA9 to HA14) is 500 years on average (min. 172 years, max.
673 years), which would allow for larger shifts of the data and
correlations with other zones from neighbouring sections such as
Sihailongwan. Moreover, Hong et al. (2010) use only 1σ age ranges
and seem to use an individual selection of calibrated ages from the
whole uncertainty range for some dates, e.g., for sample HA9 the
youngest age possible (10,745 cal yr BP from a 1σ range of 11,163–
10,745) and for sample HA13 an older age (13,138 cal yr BP from a 1σ
range of 13,171–12,999).
Also problematic is the apparent spacing of their data. Samples
HA10 and HA11 are only 5 cm apart from one another in profile depth
but show an age difference of 693 years. On the contrary, samples
HA11 and HA12 are separated by 35 cm but show nearly identical
ages. Consequently, in their Fig. 2B, Hong et al. show only one data
point between HA11 and HA12, and this is by far the most outstanding
negative excursion of their δ 13C data series.
The Sihailongwan age model is more robust because it is based
entirely on varve counts corrected by a factor of 1.0622, which is
derived from the deviation of varve chronology when compared to
14
C-chronology based on 40 calibrated AMS 14C dates (Schettler et al.,
2006; Mingram et al., 2009; Stebich et al., 2009). Because of the use of
continuously varved sediments as a basis for the age model, only very
small internal age shifts are possible.
9. Conclusions
To summarise, we do not approve of the re-interpretation of our
Lake SHL data by Hong et al. (2010) and must strongly reject it. We
also disagree with the view of Hong et al. (2010) of a regional division
into a wetter YD in northern China and a drier YD in southern China,
which in our opinion is based on re- and misinterpretations of cited
papers and unvalidated (without any modern calibration) use of bulksample cellulose δ 13C as a precipitation proxy. Therefore, we see no
need to discuss their far-reaching conclusions on the evolution of the
Chinese monsoon during the Late Glacial period.
In fact, there is a clear need for further high-resolution multi-proxy
studies and the development of new proxies, which should lead to
improved and/or new models of palaeoclimate change with an
enhanced focus on regional to local developments. Moreover, interpretation and correlation of palaeoclimatic records requires a more critical
consideration of the age models and the climate/proxy relationship and
proxy data interpretation. On the other hand, developing new models
from data of bulk samples of unknown composition, with neither
sufficient time controls nor proven climate/proxy relations, is perhaps
not the required step forward. Instead, there is a need for the
preliminary step of demonstrating the relationship between precipitation changes and the cellulose δ 13C data on widespread moss- and
higher plant genera (namely Carex) from peat bogs. Unfortunately, this
is a laborious task that will probably not lead to easy and fast rewards,
but a first step has been already been taken for Sphagnum mosses. Today,
we know that the separation of bulk Sphagnum plant remains from peat
samples is inadequate for tracking past environmental and climate
change because a significant offset between the stable isotope
compositions of different Sphagnum plant components has been
observed in modern and sub-fossil Sphagnum plant material (Loader
et al., 2007; Moschen et al., 2009). Thus, a physical differentiation of
single, easily identifiable plant components prior to stable isotope
analyses is a strict necessity to avoid systematic misinterpretations of
the resulting isotope time series, and should be combined with site-
469
specific investigations of environmental controls on plant-stable isotope
data to develop a modern calibration technique.
Acknowledgements
For translation of papers published in Chinese, we owe great
thanks to Y.-Q. Hu. We thank an anonymous reviewer for the useful
comments.
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