The cause of the seasonal variation in the oxygen isotopic composition of
precipitation along the western U.S. coast
Nikolaus Buenning1, Lowell Stott1, Max Berkelhammer2, Kei Yoshimura3
1Department
of Earth Sciences, University of Southern California; 2Department of Atmospheric and Oceanic Sciences, University of Colorado; 3Department of Natural Environmental Studies, University of Tokyo
For further information contact at: [email protected]
Introduction and Motivation
Model Results
Climate simulations of the 21st century project a drier climate throughout most of California and the American
southwest. This poses a major problem for the western U.S., where there is an already inadequate supply of
water and a growing demand. The western U.S. is also a region that has experienced protracted periods of
drought. Figure 1 shows two 20th century records from northern California: rainfall rate and the 18O/16O
composition (hereafter 18O) of tree cellulose, which is believed to record precipitation 18O values (18Op).
The curves in Figure 1 shows that drought intervals in California coincide with periods with anomalously low
18Op. Because isotopic fractionation takes place during phase changes (Figure 2), water isotopes are often
used to trace moist processes in the atmosphere [Friedman et al., 1993; Berkelhammer et al., 2012], and can
thus provide information on drought mechanisms. Therefore, understanding the cause of temporal variations
in 18Op is of clear importance. This study aims to understand the controls on the seasonal variations in 18Op
along the western U.S. coast.
To validate the model, we compared simulated
seasonal 18Op cycles with measured cycles from 16
different locations in western coastal states. Both
modeled and measured values reveal a drop in 18Op
during the winter months (Figures 4 and 5). This
drop is primarily driven by equilibrium fractionation
during condensation, as can be seen from the model
experiments (Figure 6). Figure 7 shows that the
model experiment (NOFRAC2) removes the
equator-to-pole gradient in vapor 18O values; thus,
the seasonal 18Op cycle could be related tropical vs
mid-latitude moisture transport. Figure 8 reveals that
the NOFRAC2 simulation also removes the vertical
surface-to-tropopause gradient in vapor 18O values,
so the the seasonal 18Op cycle might also be due to
seasonal changes in condensation height. Vapor
tagging simulations were used to resolve this issue.
Figure 1. Sacramento rainfall anomaly (blue line) relative to
20th century mean (low-pass filtered, 10-year cutoff). Lower
(black/red) is the 18O of Bristlecone Pine cellulose. Intervals of
California drought are orange boxes.
Model
The primary tool for investigating
isotopic variability is the isotopeincorporated Global Spectral Model
(IsoGSM) [Yoshimura et al., 2008]. This
atmospheric general circulation model is
forced with prescribed sea surface
temperatures and sea-ice conditions,
though the dynamical wind and
temperature fields were spectrally nudged
to the NCEP/NCAR reanalysis version 1.
The global simulations were performed
from 1953 through 2010 with a horizontal
resolution of approximately 1.85° x 1.85°.
In addition to a control simulation, we
performed a suite of model experiments
that “turn off” certain fractionation
processes to see how the seasonal cycle in
18Op responds (Table 1).
Figure 2. Simplified schematic of
preferential evaporation of lighter
isotope and preferential
condensation of heavier isotope.
Model Experiments
Tagging Simulations
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Figure 4 Observed and simulated season
18Op cycles, showing the wintertime drop in
18Op.
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Figure 5. Long-term mean
(1953-2010) simulated season
cycles of 18Op for grid cells
located along the western North
American coast. The southern
most grid cell (red) is located in
southern California, and the
northern most grid cell (violet) is
located in British Columbia).
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To decipher whether the seasonal 18Op cycle was due
to atmospheric circulation variations, a vapor tagging
simulation (TAGLAT) was performed where vapor
was tagged within a boxed region in the tropics and a
boxed region of the mid-latitudes (Figure 7). To
determine the role of condensation height, a vertical
tagging simulation was conducted (TAGLEV) where
two tags are applied to vapor: one below and one
above the =0.85 level. The TAGLAT simulation
reveals a larger (smaller) proportion of tropical (midlatitude) moisture rains out on the west coast during
winter (Figures 9a-b), which is inconsistent with the
expected result. During winter there is also a drop
(rise) in the fraction of precipitation with lower
(upper)
level tags
(Figure suggest
9c-d).
These
results
that the
seasonal 18Op cycle is due to
condensation height.
Figure 9. Seasonal cycles of fraction of precipitation
with (a) tropical (b) mid-latitude, (c) low level, and (d)
upper level tag. Results are for same locations as in
Figure 5.
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Table 1. Name and description of IsoGSM simulations
Sim. name
Simulation Description
CTRL
Unperturbed control simulation
NOFEQ1
Equilibrium fractionation during ocean water evaporation is turned off.
NOFEQ2
Equilibrium fractionation during vapor condensation is turned off.
CONFEQ2
Equilibrium fractionation during ocean evaporation is set to a constant.
CONFEQ2
Equilibrium fractionation during vapor condensation is set to a constant.
NORNEV
No isotope effects associated with raindrop evaporation and exchanges.
NOFKI1
Kinetic fractionation during ocean water evaporation is turned off.
NOFKI2
Kinetic fractionation during vapor deposition onto ice is turned off.
TAGLAT
Tropical vs mid-latitude vapor tagging simulation.
TAGLEV
Lower vs Upper troposphere vapor tagging simulation.
Figure 3. IsoGSM simulations are spectrally nudged to the
NCEP/NCAR Reanalysis 1.
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Figure 6. Simulated long-term mean seasonal 18Op cycles
(relative to an annual mean) for a grid-cell located at coastal
Oregon. The control simulation is represented by the solid blue
curves in all panels. Top panels compare the control simulation
with the experiments that turn off or set constant equilibrium
fractionation during ocean evaporation (a) and vapor
condensation (b). Bottom panels compare the control
simulation with the experiment that turns off post-condensation
exchanges (c) and the experiments that turn off kinetic
fractionation (d).
Figure 8. Simulated annual mean
vertical profile of vapor 18O
values for the control simulation
(CTRL, blue) and the simulation
that turned off equilibrium
fractionation during vapor
condensation (NOFEQ2, yellow).
The profiles are taken from a grid
cell near Santa Cruz, California.
Figure 7. Annual mean 18O value of
vapor for (a) the control simulation and
(b) the simulation that turns off
equilibrium fractionation during vapor
condensation. Boxed regions in (a)
indicate the tagging regions for the
TAGLAT simulation
References
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Figure 10. Vertical profiles of vertical pressure velocity
(a) and horizontal divergence for two locations during
January (blue) and July (red). The solid curves represent
the west coast at 37° and the dashed curves show the west
coast at 41°. All curves are weighted by daily
precipitation. The seasonal change in condensation height
is caused by seasonal changes to vertical winds, which is a
result of divergence in the upper troposphere.
Implications
Figure 11. Mean 200 mb wind speeds (closed
contours) and geopotential heights (open
contours) weighted by precipitation from one
location (red asterisk). Top panels are July;
bottom panels are January. Winds converge
towards the location in July and slightly diverge
during January, causing the seasonal change in
vertical winds and condensation height.
• The main conclusion of this work is condensation height controls the simulated seasonal cycle of the 18O/
16O composition of precipitation along the western U.S. coast.
• The effect from condensation height adds another layer of complexity when interpreting climate proxies
(Figure 1) based on isotopes in precipitation (i.e., 18O or D of tree cellulose, speleothems, or leaf wax nalkanes). There is a strong possibility that these interannual variations are influenced by condensation
height.
Acknowledgements: Acknowledgments. We acknowledge with gratitude Christopher Lehmann and the National Atmospheric Deposition Program (NADP) for providing precipitation
samples. We recognize Miguel Rincon and Jack Zang for their help with the isotope analysis of NADP samples. Funding for this work was provided by the NOAA/CPO Climate Change &
Detection Program: Paleoclimate Studies (grant NA10OAR4310129). This work was also supported by a grant from the National Science Foundation (award AGS-1049238).