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important consequences for the livelihood of billions of people. At orbital time scales,
the most prominent factor controlling the monsoon is local summer insolation, as
indicated by
fluctuations of monsoon strength coinciding remarkably well with
changes in the Earth
s precessional cycle (Cruz et al.
2005
; Wang et al.
2001
,
2008
).
Moreover, records of monsoon variability derived from paleoclimate proxy data
suggest increased monsoon intensity during warm interglacials and shorter term
interstadials compared to cold glacials and stadials (i.e., Cruz et al.
2005
,
2009
;
Fleitmann et al.
2003
; Wang et al.
2001
,
2008
). The climate of the Eemian (
'
130
*
thousand years (ka) before present (BP)
C
warmer than that of the Holocene (Kukla et al.
2002
; Leduc et al.
2010
), so that an
intensi
—
115 ka BP) was probably by 1
2
°
-
ed hydrological cycle and thus monsoon can be expected due to the tem-
perature dependence of the atmospheric water holding capacity (see e.g., Cruz et al.
2005
). Although climate model results show more intense monsoon systems under
warmer Eemian climate (Braconnot et al.
2008
; Kutzbach et al.
2008
), from
paleoproxy records no clear evidence is available for differences in monsoon intensity
between different interglacials or interstadials. Therefore, in the present study a new
paleoprecipitation record for the Indian monsoon will be compared with results from
a comprehensive climate model simulating the Holocene and the Eemian.
2 Materials and Methods
The model used in the present study is the Kiel Climate Model (KCM; Park et al.
2009
), which is a coupled atmosphere-ocean-sea ice general circulation model. The
atmosphere is represented by ECHAM5 (Roeckner et al.
2003
) using the numerical
resolution T31L19 corresponding to 3.75
on a great circle. ECHAM5 is coupled to
the ocean model NEMO, consisting of the OPA9 ocean circulation (Madec
2008
)
and the LIM2 sea-ice model (Fichefet and Morales Maqueda
1997
) with a horizontal
resolution of approx. 2
°
°×
2
°
and increased meridional resolution (0.5
°
) close to the
13
C of sedimentary leaf wax
n-alkanes from a marine sediment core collected on the Bengal deep-sea fan (core
SO188 17286-1; 19
equator. We measured a
first set of samples for
D and
δ
δ
E) as paleoindicators for precipitation and
vegetation changes, respectively. Analysis of
°
44
′
48
′′
N, 89
°
52
′
76
′′
13
Cofn-alkanes is described
in detail by Wang et al. (
2013
). A preliminary chronology for this core covering the
last 135 ka was established by tuning the
D and
δ
δ
18
O record of planktonic foraminifera to
δ
18
O curve of the Greenland NGRIP ice core (NGRIP-members
2004
) until
126 ka BP and to SPECMAP (Martinson et al.
1987
) and between 115 ka BP and
135 ka BP to fully cover the Eemian.
Simulations with the KCM (Park et al.
2009
) were carried out as quasi steady
state (time slice) simulations forced by changes in orbital parameters for the
respective early and late Holocene and Eemian periods, corresponding to prein-
dustrial, 9.5 ka BP and 126 ka BP (Khon et al.
2010
,
2012
; Schneider et al.
2010
;
Salau et al.
2012
). Modeled precipitation is integrated over the Ganges-Brahmaputra
catchment area in order to be directly comparable with data from the sediment record
the
δ
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