ACPAtmospheric Chemistry and PhysicsACPAtmos. Chem. Phys.1680-7324Copernicus GmbHGöttingen, Germany10.5194/acp-15-5827-2015Ocean mediation of tropospheric response to reflecting and absorbing
aerosolsXuY.yangyang@ucar.eduhttps://orcid.org/0000-0001-7173-7761XieS.-P.National Center for Atmospheric Research, Boulder, CO 80303, USAScripps Institution of Oceanography, University of California, San
Diego, La Jolla, CA 92093, USAY. Xu (yangyang@ucar.edu)27May201515105827583324January201525February201523April20155May2015This work is licensed under a Creative Commons Attribution 3.0 Unported License. To view a copy of this license, visit http://creativecommons.org/licenses/by/3.0/This article is available from https://acp.copernicus.org/articles/15/5827/2015/acp-15-5827-2015.htmlThe full text article is available as a PDF file from https://acp.copernicus.org/articles/15/5827/2015/acp-15-5827-2015.pdf
Radiative forcing by reflecting (e.g., sulfate, SO4) and absorbing (e.g.,
black carbon, BC) aerosols is distinct: the former cools the planet by
reducing solar radiation at the top of the atmosphere and the surface,
without largely affecting the atmospheric column, while the latter heats the
atmosphere directly. Despite the fundamental difference in forcing, here we
show that the structure of the tropospheric response is remarkably similar
between the two types of aerosols, featuring a deep vertical structure of
temperature change (of opposite sign) at the Northern Hemisphere (NH)
mid-latitudes. The deep temperature structure is anchored by the slow
response of the ocean, as a large meridional sea surface temperature (SST)
gradient drives an anomalous inter-hemispheric Hadley circulation in the
tropics and induces atmospheric eddy adjustments at the NH mid-latitudes. The
tropospheric warming in response to projected future decline in reflecting
aerosols poses additional threats to the stability of mountain glaciers in
the NH. Additionally, robust tropospheric response is unique to aerosol
forcing and absent in the CO2 response, which can be exploited for
climate change attribution.
Introduction
Greenhouse gas-induced global warming is partially masked (Ramanathan and
Feng, 2008) by the accompanying increase in anthropogenic aerosols (Smith et
al., 2011). The relative contribution of the aerosol masking effect on global
temperature is hard to quantify for the following reasons: (a) some aerosols
(e.g., black carbon (BC) and organics) absorb sunlight and heat the planet
(Bond et al., 2013) and (b) aerosol microphysical effects on clouds are
complex (Rosenfeld et al., 2013). Many ongoing efforts aim to reduce
uncertainties in radiative forcing (Xu et al., 2013) and quantify the surface
temperature response to aerosols (Levy et al., 2013). The atmospheric
circulation response to reflecting aerosols has important effects on the
regional climate (e.g., the Indian monsoon; Bollasina et al., 2011) and
hydrological cycle (Shindell et al., 2012; Hwang et al., 2013). Much
attention has been given to absorbing aerosols for the direct atmospheric
heating effect, including BC (Meehl et al., 2008) and dust (Vinoj et al.,
2014). It is often argued that, by heating directly the atmosphere, absorbing
aerosols can greatly perturb the atmospheric temperature structure, causing
changes in stability and circulation (Lau et al., 2006). The atmospheric
response, especially that of clouds, is hypothesized to be sensitive to the
vertical profile of atmospheric heating (Koch and Del Genio, 2010).
Reflecting aerosols, however, are hinted to be less effective in driving
large-scale circulation changes (Allen et al., 2012).
While previous studies (e.g., Xie et al., 2013; Ocko et al., 2014) focused on
radiative forcing and climate impacts of aerosols on surface temperature and
precipitation (Table S1 in the Supplement), few looked at the tropospheric
response. Using climate model simulations, we show that the atmospheric
responses (temperature and circulation) to reflecting and absorbing aerosols
are surprisingly similar in structure (aside from a sign difference). Both
responses feature a deep vertical temperature structure at the Northern
Hemisphere (NH) mid-latitudes, with a meridional shift in the westerly jet.
Such a strong atmospheric temperature response to absorbing aerosols has been
commonly linked to direct solar absorption in the atmosphere (Lau et al.,
2006). We demonstrate, however, that changes in the sea surface temperature
(SST) gradient and mid-latitude eddies are instrumental in creating a common
deep vertical temperature in response to both types of aerosols, despite the
fundamental difference in their forcing structure.
MethodsThe global climate model
CESM1 (Community Earth System Model 1) is a coupled
ocean–atmosphere–land–sea-ice model. CESM1 climate projections for the
twenty-first century have been documented extensively (Meehl et al., 2013).
The anthropogenic forcings in CESM1 include long-lived greenhouse gases
(GHGs), as well as tropospheric ozone, stratospheric ozone, sulfate aerosols,
and black and primary organic carbon aerosols. The three-mode aerosol scheme
(MAM3) provides internally mixed representations of aerosol number
concentrations and masses (Liu et al., 2012). Aerosol indirect forcing is
included for both liquid and ice phase clouds (Gettelman et al., 2010).
The aerosol emission inventory is from the standard representative
concentration pathway as described in Lamarque et al. (2010). However, the
present-day emission level of BC is adjusted from the standard model emission
inventory to account for the potential model underestimation of BC
atmospheric heating. Our previous analysis (Xu et al., 2013) shows that such
a correction improves simulated radiative forcing compared to the direct
observations. Without the observational constrains, simulated BC forcing (and
associated temperature response) would be lower by about a factor of 2. In
addition to the atmospheric heating, deposition of BC particles onto snow
surface with high albedo would reduce surface albedo and contribute to
surface warming (Huang et al., 2011). The land model of CESM incorporates the
SNICAR (Snow and Ice Aerosol Radiation) module, which represents the effect
of aerosol deposition (BC, organic carbon and dust) on surface albedo
(Flanner et al., 2007).
Note that in this study we used BC, a strong absorber, to characterize
absorbing aerosols that also include dust and organic aerosols. Similarly, we
used SO4 to characterize reflecting aerosols, although dust and organic
aerosols are also partially reflecting. This approach provided a clearer
contrast between these two types of aerosol forcing.
Model experiments
Fully coupled model simulations with instantaneous forcing. We used a
394-year, pre-industrial simulation as the control case. Starting from the
end of the 319th year, we ran the simulations for 75 years, with the last
60 years of output analyzed. This allows the first 15 years for model spin-up
to establish a quasi-equilibration with changes in radiative forcing (Long et
al., 2014). The forcing is imposed by increasing BC emissions (as a proxy for
absorbing aerosols) and SO2 emissions (a precursor of SO4, as a proxy
for reflecting aerosols) instantaneously from pre-industrial levels to the
present-day level. This methodology is similar to the classical CO2
doubling experiment (Manabe and Wetherald, 1975). The long averaging time
(60 years in the perturbed simulation versus 394 years for the pre-industrial
control simulations) enabled us to dampen the influence of decadal natural
variability and to obtain a clear effect due to aerosol perturbation. To
increase the signal-to-noise ratio in the BC case (due to a smaller BC
forcing), five ensembles of perturbed simulations were conducted.
The twentieth century transient simulations using fully coupled model, with
time-evolving sulfate forcing. The details of the simulations can be found in
Meehl et al. (2013). The resolution of both atmosphere and ocean models is
1∘ by 1∘ for the coupled simulations (Experiments a and b) in
this study.
The atmospheric-only simulations with instantaneous forcing. The model
setting and imposed forcing are identical to a., but SST is fixed at a
pre-industrial level, with only seasonal variability. The model was also run
for 75 years.
The SST perturbation experiment. The SST was perturbed according to the
zonal mean of the CESM SO4 Experiment a (Fig. S1 in the Supplement). This
corresponds to a temperature profile that varies from 0 ∘C at
90∘ S to -0.5 ∘C at the Equator, and then to
-1.2 ∘C at 90∘ N. The SST perturbation did not include
any longitudinally varying pattern, as our focus here was to understand the
zonal averaged temperature response. The perturbed model was run for 25 years
(with 10 years of daily output for eddy flux analysis). The resolution of the
atmospheric model is 2∘ by 2∘ for the uncoupled simulations
(Experiments c and d) in this study.
Tropospheric response linked to SST gradient
BC atmospheric radiative forcing is concentrated at 30∘ N and
extends well above the boundary layer to the free atmosphere (Fig. 1), a
structure determined by atmospheric concentration, and indirectly by emission
sources. Intuitively, solar absorption by BC results in atmospheric warming.
Indeed, BC (Fig. 1 upper panels) induces a warming maximum in the NH
mid-latitude troposphere (350 mb, 30 to 40∘ N) in the coupled
ocean–atmosphere model, which dwarfs the upper tropical and Arctic warming.
This simple thermodynamic mechanism seems consistent with the fact that the
magnitude of BC warming is much larger in the boreal summer (JJA) than in the
boreal winter (DJF) (Fig. 2, upper panels), due to solar insolation.
(Left) Heating rate (∘C day-1) due to an increase in
BC, SO4 and CO2 atmospheric concentrations. The heating rate is
diagnosed by contrasting two sets of 5-year atmospheric-only simulations with
pre-industrial and present-day emissions/concentrations, respectively.
(Right) Annually averaged atmospheric temperature in response due to the
forcing of BC, SO4 and CO2. The color scale for SO4 is reversed. The
magnitude of the color scale is chosen considering the difference in
top-of-atmosphere forcing (Table S1).
Interestingly, SO4 also induces a similar enhanced tropospheric cooling at
the mid-latitudes (Figs. 1 and 2). For easy comparison, the response is
reversed in sign to be positive. The deep atmospheric response is unexpected
from the weak, direct atmospheric forcing of reflecting aerosols (Fig. 1,
middle left). Also contradictory to the above thermodynamic argument for BC,
the temperature response to SO4 is of a similar magnitude in DJF and JJA
(Fig. 2). The CO2 response features a structure of amplified upper
tropical troposphere warming (maximum at around 300 mb), which is a robust
feature due to thermodynamical adjustment of the tropical atmosphere to
maintain a moist adiabatic lapse rate there. The lower tropospheric
atmospheric temperature over the Arctic also has a larger response, mostly
due to stronger snow albedo feedback.
Temperature response (∘C) as a function of latitude and
pressure to BC (first row), SO4 (second row), and SO4-induced SST
perturbation (SO4_SST) (third row). The left and right columns are the
DJF and JJA averages, respectively. Note that the color scales for SO4 and
SO4_SST are reversed.
Similar to Fig. 2, but for fast (first and third columns) and slow
components (second and fourth columns) of temperature response (in
∘C). The fast component is calculated by running the atmospheric-only
(fixed SST) simulation with perturbed atmospheric compositions, while the
slow component is the difference between the total (Fig. 2) and fast
components. The color scale for SO4 is reversed.
The climate response may be decomposed into fast and slow components, defined
as the atmospheric response without and due to SST change, respectively
(Ganguly et al., 2012). The BC temperature response results predominately
from the fast component in the summer due to direct atmospheric heating
(Fig. 3), but the slow response dominates in the winter. The SO4 fast
response, due to the lack of atmospheric forcing, is strikingly small (except
in summer polar regions where air temperature above sea ice is free to
change), despite aerosol indirect forcing through fast adjustment of clouds
being allowed. The SO4 slow response in winter features a narrow maximum
around 30∘ N, and the summer mid-latitude response is weaker and
extends into the upper tropics. Therefore, the slow component of the response
due to SST change is entirely responsible for the SO4 deep atmospheric
response and partially responsible for the BC response.
The dominant role of SST in causing the deep atmospheric response is further
confirmed by a set of perturbed-SST experiments, in which the zonal mean SST
change in the full SO4 simulation (Fig. S1) is applied to the
atmospheric-only model, but with no radiative forcing. The model response to
the perturbed SST (third row of Fig. 2) is remarkably similar to the SO4
slow response (Fig. 3), explaining a large fraction of the total response
(second row of Fig. 2). The boundary layer air temperature (below 850 mb) is
closely tied to the underlying SST because of turbulent mixing, while at the
mid-latitudes, the free atmospheric temperature is not tied to the SST
because the atmosphere is stably stratified. However, changes in the SST may
affect the free troposphere through the changes in tropical circulations and
mid-latitude eddy, which we explore next.
Understanding zonal mean circulation changes
Figure 4 shows the circulation responses to aerosols in terms of meridional
overturning stream function (positive values indicate clockwise circulation)
and zonal averaged zonal wind (positive values indicate westerly winds). Note
that the responses of SO4 and BC are similar in space but of opposite
signs. SO4 cooling in the NH induces an anomalous Hadley cell that rises
in the SH and sinks in the NH (also shown in Ocko et al., 2014). The
atmospheric model forced by SO4-induced SST change largely reproduces the
Hadley cell response (Fig. 4, bottom left), highlighting the importance of
the inter-hemispheric SST gradient. Consistent with the Hadley cell response,
the NH jet stream shifts equatorward in response to SO4, and vice versa to
BC. Following the thermal wind relationship (the maximum temperature gradient
sets the maximum zonal wind), the equatorward shift of westerly winds must be
accompanied by a deep cooling structure (Figs. 1 and 2).
(left) Zonal mean meridional stream function change
(109 kg s-1), in response to BC (first row), SO4 (second row),
and SO4-induced SST perturbation (SO4_SST) (third row). The
climatological stream function is shown in contour lines with an interval of
40. The negative values (blue shading and dashed lines) of the stream
function indicate that the meridional flow is counter-clockwise. (right)
Zonal mean zonal wind (U) change under various cases. The climatological
jet stream is around 30 to 60∘ N at 250 mb (line contours). Under
SO4 forcing, the NH jet stream shifts significantly
equatorward.
The color scale for the SO4 response in Fig. 4 is not reversed as in
previous temperature figures, in order to depict the real direction of
circulation change. The magnitude of changes in response to BC is weaker due
to a smaller forcing magnitude (Table S1). In addition, the SO4-induced
Hadley cell change is inter-hemispheric across the Equator, while the
BC-induced Hadley cell change appears more confined to the NH. The same for
the jet stream shift. This is probably because of the geographic difference
in BC and SO4 forcing (amid both are stronger in NH than SH), which may
influence the Pacific and Atlantic branches of jets differently.
Eddy fluxes that transport heat and momentum in meridional directions are
instrumental in maintaining the climatological mid-latitude jets. Here we use
the Eliassen–Palm (EP) flux to diagnose how eddy flux adjustment in response
to aerosols leads to changes of zonal winds. The EP flux vector, with its
vertical component depicting the meridional heat flux and its meridional
component depicting the equatorward meridional momentum flux, is calculated
using 10-year daily data from the control and the perturbed SO4_SST
simulations following Holton (2004).
The NH annual mean EP flux and its divergence (in contour) are shown in
Fig. 5a. Over extratropical atmosphere, EP flux convergence (negative value)
suggests that meridional heat eddy flux (the vertical component of EP flux)
acts to slow the westerly wind aloft (Holton, 2004). However, the strong
equatorward wave propagation in the mid-latitude troposphere (meridional
component of EP flux) is acting to extract momentum from the tropics to the
mid-latitude, therefore maintaining the westerly wind at 40–60∘N.
The Eliassen–Palm (EP) flux (vector) and its divergence (contour).
(a) The climatology. (b) The change due to SO4-induced
SST perturbation (SO4_SST). The convergence (blue) and divergence (red)
of the EP flux correspond to a deceleration and acceleration of the westerly
mean flow, respectively. (c) Contributions of the stationary eddy to
the change shown in (b). This was calculated using a 10-day average,
instead of a daily average. Transient eddies are the difference between the
total and stationary contribution (not shown). (d) NH winter (DJF)
average, not the annual average shown (b). Note that the color scale
and reference vectors are different across the panels.
Under the SO4-induced SST perturbation, the EP flux change is found to be
strongest at the NH mid-latitudes 30–40∘ N, equatorward side of its
climatology (Fig. 5b). Poleward EP flux anomalies reduce the climatological
equatorward wave propagation. In the middle troposphere (400–800 mb), EP
flux convergence (blue) decelerates the vertically average westerly wind at
50–60∘ N, while EP flux divergence (red) tends to accelerate the
westerly wind at 30–40∘ N. Therefore, westerly winds shift
equatorward in response to SO4 (Fig. 4). Of the total eddy flux,
stationary eddies contribute about 60 % (Fig. 5c), with the rest coming
from transient eddies. The EP flux change occurs predominately during the NH
winter, because the background mid-latitude wave activity is stronger. This
is shown by the larger vectors in Fig. 5d and stronger EP flux divergence
(red) at 30–40∘ N.
The change in EP flux is consistent with that in the stationary wave
refractive index as wave propagation is mainly from a high refractive index
region to a low refractive index region (Held and Hou, 1980; Fig. S2). The
quasi-geostrophic refractive index and its change under SST perturbation were
calculated following Limpasuvan and Hartmann (2000). In the climatology
(Fig. S2a), the high refractive index is located at the mid and high
latitudes, and the tropics are mainly occupied by a smaller refractive index,
facilitating the equatorward propagation of mid-latitude wave activities
(Fig. 5a, also seen in Sun et al., 2013). The refractive index negative
anomaly due to perturbed SO4_SST is mainly found in the NH mid-latitude
regions (Fig. S2b), which causes the reduction of wave propagation to the
Equator (Fig. 5b).
The above diagnosis explains the SO4-induced deep tropospheric cooling and
associated equatorward shift of the westerly jet at the NH mid-latitudes.
Firstly, the intensified NH Hadley cell accelerates the upper tropospheric
westerly jets in the subtropics. Secondly, the EP flux divergence accelerates
the westerly jet on the equatorward flank of the mean Hadley cell, while the
jet is decelerated on the poleward flank due to EP flux convergence. Both the
Hadley and eddy adjustments are anchored by the SST change with strong
meridional gradients. Aqua-planet model experiments exploring the response to
an idealized mid-latitude heating (Ceppi et al., 2013) supported our
arguments here about the coupled adjustments of the Hadley circulation and
mid-latitude jets to realistic aerosol forcing.
Conclusions
Our results show that, despite the fundamental difference in forcing
structure, BC and SO4 share common atmospheric response patterns. The
common response is mediated by the ocean through sea-surface temperature
gradient, and is insensitive to microphysical representations of aerosols.
This highlights the importance of ocean–atmosphere interactions in shaping
large-scale patterns of climate response (Xie et al., 2010), a process
overlooked so far in aerosol–climate connection.
The deep mid-latitude warming in response to BC contributes to the retreat of
mountain glaciers in the NH near-anthropogenic BC emissions including the
Alps (Painter et al., 2013) and the Himalayas. Although the cooling effect on
the free troposphere is rarely discussed, SO4 aerosols may have mitigated
glacier retreats elsewhere in the past. Into the future, declining SO4
aerosols may lead to an elevated atmospheric warming and pose a threat to
mountain snow packs. This implies that more stringent controls on BC and GHGs
are needed to mitigate the snow pack retreat.
The tropospheric temperature and circulation response to SO4 is also found
in the twentieth century transient simulation (Fig. S3) and the twenty-first
century multi-model projections (Rotstayn et al., 2014). This suggests that
the deep temperature structure at the mid-latitudes is a robust feature of
aerosol-induced climate change, probably insensitive to model sub-grid
physics. The dynamic response involving the inter-hemispheric Hadley
circulation is weak in the case of CO2 and presumably other
hemispherically symmetrical forcing (such as solar and volcanic activities).
The importance of the SST pattern has been noted previously (Ramanathan et
al., 2005; Xu and Ramanathan, 2010; Friedman et al., 2013; Xie et al., 2013),
and our study reveals a fundamental difference in the mid-latitude
atmospheric responses to CO2 and aerosol forcing. This difference can be
exploited to improve the detection and attribution of climate change (Lu et
al., 2008; Santer et al., 2013). Because aerosol forcing involves stronger
mid-latitude storm track adjustments, our result also has implications for
the attribution and projection of extreme events (e.g., blockings).
The Supplement related to this article is available online at doi:10.5194/acp-15-5827-2015-supplement.
Acknowledgements
The authors wish to thank I. Held, P. Ceppi, and L. Sun for discussions, and
J. Barsugli for sharing codes for EP flux. Y. Xu is supported by the Regional
and Global Climate Modeling Program (RGCM) of the US Department of Energy's
Office of Science (BER, DE-FC02-97ER62402) and the National Center for
Atmospheric Research (NCAR) Advanced Study Programme (ASP) postdoctoral
fellowship; and S. P. Xie by the National Science Foundation (NSF). NCAR is
funded by the NSF. Edited by: J. P. Huang
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