True to Milankovitch: Glacial Inception in the
new Community Climate System Model.
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Markus Jochum, Alexandra Jahn, Synte Peacock, David A. Bailey, John T. Fasullo,
Jennifer Kay, Samuel Levis, and Bette Otto-Bliesner
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submitted to Journal of Climate, 1/19/11
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Corresponding author’s address:
National Center for Atmospheric Research
PO Box 3000
Boulder, CO, 80307
1-303-4971743
[email protected]
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Abstract: The equilibrium solution of a fully coupled general circulation model
with present day orbital forcing is compared to the solution of the same model with
the orbital forcing from 115.000 years ago. The difference in snow accumulation
between these two simulations has a pattern and a magnitude comparable to the
ones infered from reconstructions for the last glacial inception. As postulated in
Milankovitch’s hypothesis the only necessary positive feedback is the snow albedo
feedback, which is initiated by reduced melting of snow and sea-ice in the summer.
The ocean meridional heat transport in the two simulations is almost identical, and
the atmospheric meridional heat transport and the sea-ice/low-cloud connection
constitute negative feedbacks that almost fully compensate for the snow albedo
feedback. The fact that the realistic difference in snow accumulation is achieved
with the same model that is used for the fifth assessment report builds trust in the
ability of climate models to anticipate the evolution of climate in the future.
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Introduction
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Over the last 400.000 years Earth’s climate went through several glacial and inter-
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glacial cycles, which are correlated with changes in its orbit and associated changes
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in insolation (Milankovitch 1941; Hays et al. 1976; Huybers and Wunsch 2004).
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Over this time changes in global ice volume, temperature and atmospheric CO2
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have been strongly correlated with each other (Petit et al. 1999). The connection
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between temperature and ice volume is not surprising, and a strong snow/ice -
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albedo feedback makes it plausible to connect both to changes in the Earth’s orbit
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(Huybers and Tziperman 2008, and references therein).
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Models of intermediate complexity (e.g., Crucifix and Loutre 2001; Khodri et
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al. 2001; Wang and Mysak 2002; Khodri et al. 2003; Calov et al. 2005) and flux-
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corrected GCMs (Groeger et al. 2007) have typically been able to simulate a connec-
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tion between orbital forcing and temperature or snow volume. So far, however, free
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running, fully coupled primitive equation General Circulation Models (GCMs) have
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failed to reproduce glacial inception, the cooling and increase in snow and ice cover
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that leads from the warm interglacials to the cold glacial periods. Milankovitch
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(1941) postulated that the driver for this cooling is the orbitally induced reduction
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in Northern Hemisphere summer time insolation and the subsequent increase of
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perennial snow cover. The increased perennial snow cover and its positive albedo
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feedback are, of course, only precursors to ice sheet growth. The GCMs failure to
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recreate glacial inception (see Otieno and Bromwich 2009 for a summary) indicates
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a failure of either the GCMs or of Milankovitch’s hypothesis. Of course, if the hy-
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pothesis would be the culprit, one would have to wonder if climate is sufficiently
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understood to assemble a GCM in the first place. Either way, it appears that re-
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producing the observed glacial/interglacial changes in ice volume and temperature
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represents a good testbed for evaluating the fidelity of some key model feedbacks
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relevant to climate projections. Note that from the GCM perspective glacial incep-
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tion and glacial termination are not symmetric, because for the latter one needs
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an ice sheet model, the development of which is still in its infancy (see Vizcaino et
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al. 2008 for the description of a state of the art model), whereas the current GCMs
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should contain all the necessary components of the climate system to reproduce the
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former.
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The potential causes for GCMs failing to reproduce inception are plentiful, rang-
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ing from numerics (Vettoretti and Peltier 2003) on the GCMs side to neglected feed-
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backs of land, atmosphere, or ocean processes (e.g.; Gallimore and Kutzbach 1996,
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Hall et al. 2005, Jochum et al. 2010, respectively) on the theory side. It is encour-
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aging, though, that for some GCMs it takes only small modifications to produce
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an increase in perennial snowcover (e.g.; Dong and Valdes 1995). Nevertheless, the
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goal for the GCM community has to be the recreation of increased perennial snow-
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cover with a GCM that has been tuned to present day climate, and is subjected to
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changes in orbital forcing only.
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The present study is motivated by the hope that the new class of GCMs that
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has been developed over the last 5 years (driven to some extent by the forthcoming
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fifth assessment report, AR5), and incorporates the best of the climate community’s
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ideas, will finally allow us to reconcile theory and GCM results. It turns out that
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at least one of the new GCMs is true to the Milankovitch hypothesis. The next
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section will describe this GCM and illustrate the impact of changing present day
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insolation to the insolation of 115.000 years ago (115 kya). At least to the present
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authors it was surprising that increased perennial snow cover was achieved with
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only the ice/snow albedo feedbacks postulated by Milankovitch, no other feedbacks
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seem to be necessary. In particular the stability of the Atlantic Meridional Over-
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turning Circulation (AMOC) under changes to the insolation is surprising, and will
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be discussed in Section 3. Section 4 analyzes in detail the Arctic heat budget with
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its multitude of positive and negative feedbacks. Section 5 concludes the study with
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a summary and implications for future model development.
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Model Description and Results
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The numerical experiments are performed using the National Center for Atmo-
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spheric Research (NCAR) latest version of the Community Climate System Model
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(CCSM) (version 4), which consists of the fully coupled atmosphere, ocean, land
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and sea ice models. A description of this version can be found in Gent et al. (2011).
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The ocean component has a horizontal resolution that is constant at 1.125◦ in lon-
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gitude and varies from 0.27◦ at the equator to approximately 0.7◦ in high latitudes.
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In the vertical there are 60 depth levels; the uppermost layer has a thickness of 10
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m, the deepest layer has a thickness of 250 m. The atmospheric component uses a
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horizontal resolution of 0.9◦ × 1.25◦ with 26 levels in the vertical. The sea ice model
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shares the same horizontal grid as the ocean model and the land model is on the
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same horizontal grid as the atmospheric model. The details of the different model
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components will be described in a series of forthcoming papers; for the present
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purpose it is sufficient to know that CCSM4 is a state-of-the-art climate model
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that has improved in many aspects from its predecessor CCSM3 (Gent et al. 2011).
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For the present context the most important improvement is probably the increased
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atmospheric resolution and the switch from a spectral to a finite volume dynamical
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core, because both of these improvements allow for a more accurate representation
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of altitude and therefore land snow cover (e.g.; Vettoretti and Peltier 2003).
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The subsequent sections will analyze and compare two different simulations: an
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1850 control (CONT), in which earth’s orbital parameters are set to the 1990 values
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and the atmospheric composition are fixed at its 1850 values; and a simulation
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identical to CONT, with the exception of the orbital parameters, which are set to
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the values of 115 kya (OP115). CONT was integrated for 700 years, and OP115
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was branched off from CONT after year 500 and continued for 300 years. Unless
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noted otherwise the comparisons will be done between the means of years 651-
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700 of CONT and the years 751-800 of OP115. Although CONT is considered fully
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spun-up the comparison is not quite clean, because there is a small drift in ocean
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temperature and salinity. The two different time intervals were chosen because at
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year 700 the AMOC in OP115 did not reach equilibrium yet, and we did not have
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the computational resources to extend CONT beyond year 700. However, in the
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present context the drift is negligible because it is mostly confined to the abyssal
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ocean (Danabasoglu et al. 2011).
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The particular time period for the orbital forcing of OP115 is chosen because
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the solar forcing at 65◦ N during June differs from its present value by approxi-
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mately 7%, more than during most other epochs (Berger and Loutre 1991). More-
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over, observations suggest that the onset of the last ice age dates around this time
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(Petit et al. 1999; Capron et al. 2010). The shift in orbital parameters means that
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in particular at northern high latitudes there is less insolation during spring and
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early summer months, and more radiation later in the fall (Figure 1). This shift
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in the seasonal distribution of insolation is at the heart of the inception hypoth-
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esis of Milankovitch (1941), because it leads to increased snow and sea-ice cover
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during early summer, and a positive snow/ice albedo feedback. The two main goals
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of the present work are to demonstrate that the 115 kya orbital changes produce
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a realistic increase in snow cover in CCSM4, and to quantify the relevant climate
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feedbacks.
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The difference in orbital parameters between OP115 and CONT lead to a
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warmer tropical band, and cooler northern high latitudes (Figure 2a). There is only
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little response in the southern high latitudes, and these are mostly confined to the
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ocean. To isolate the impact of the ocean on glacial/interglacial dynamics, we re-
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placed the full ocean model with a slab-ocean model (for the detailed methodology,
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please see Danabasoglu and Gent 2009). The results are quite similar (Figure 2b;
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snow and sea-ice differences are similar too, but not shown), with the full ocean
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leading to a slightly weaker response. The exception is the Labrador Sea, which
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shows approximately 2◦ C more cooling with a full ocean model. While tropical and
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southern hemisphere differences in both sets of experiments are similar to the
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coarse resolution results of Jochum et al. (2010, JPLM from here on), the northern
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high-latitude response is significantly weaker. The relative weakness at high lati-
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tudes can be attributed to the different response of the AMOC, and will be focus of
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the next section.
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The colder northern high-latitudes in OP115 are associated with significant in-
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crease in snow depth in the Canadian Archipelgo, and northern Siberia (Figure 2c),
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regions cited as the origination points for the glacial ice sheets in ice sheet recon-
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structions (Svendsen et al. 2004; Kleman et al. 2010). The increase in snow depth
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will be analyzed in section 4, but it is worth noting here that it has been achieved
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without modifying the underlying GCM. The CONT simulation is identical to the
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1850 control simulation of Gent et al. (2011), which has been optimized to produce
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the observed 20th century climate. The fact that CONT can produce a realistic
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snow distribution for a very different mean climate is a major achievement for the
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CCSM, and instills confidence in the results of its future projections.
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In accordance with the Milankovitch (1941) hypothesis (see also Huybers and
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Tziperman 2008), it can be seen that the reduced spring/summer insolation in
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OP115 leads to more perennial snow, and larger summer sea-ice concentration
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(Figure 2d), both of which contribute to the positive snow/ice albedo feedback. Sec-
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tion 4 will provide a quantification of the positive and negative feedback processes
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leading to the northern high latitude cooling.
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The meridional heat transport of the AMOC is a major source of heat for the north-
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ern North Atlantic ocean (e.g.; Ganachaud and Wunsch 2001), but it is also be-
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lieved to be suceptible to small perturbations (e.g.; Marotzke 1990). This raises the
The role of the AMOC and the Labrador Sea gyre
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possibility that the AMOC amplifies the orbital forcing, or even that this amplifi-
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cation is necessary for the northern hemisphere glaciations and terminations (e.g.;
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Broecker 1998). In fact, JPLM demonstrates that at least in one GCM changes
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in orbital forcing can lead to a weakening of the MOC and a subsequent large
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northern hemisphere cooling. Here, we revisit the connection between orbital forc-
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ing and AMOC strength with the CCSM4, which features improved physics and
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higher spatial resolution compared to JPLM.
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It turns out that for CCSM4 the OP115 scenario does not lead to a reduced
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AMOC in contrast to the JPLM results, which show a 30% reduction of the AMOC
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strength under OP115 forcing. After branching off at year 500 of CONT, the AMOC
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in OP115 immediately weakens, and continues to do so until it reaches a minimum
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after approximately 40 years (Figure 3a). It stays weaker for some 50 years, starts
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to recover around year 600, and reaches a mean value of 15.3 Sv (yrs 751-800),
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which is not significantly different from the mean of CONT. The depth of the AMOC
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has not changed significantly either (not shown). The meridional heat transport
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shows a similar evolution (Figure 3b), with final values of 0.74 PW. The strength
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of the subpolar gyre, too, mirrors this decline and subsequent recovery (to 52.7 Sv,
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Figure 3c). By the end of this 300 year development, the surface of the Irminger
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Sea and most of the subsurface subpolar Atlantic is warmer and saltier than in
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CONT, and the surface of the Labrador Sea is colder and fresher (Figure 3d).
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Much of this reorganization of the subpolar Atlantic can be explained by the
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analysis of sea-ice extent and maximum ocean boundary layer depth, the latter
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being by construction the depth of a year’s deepest convective event (Large et al.
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1994). In CONT convective activity can be seen all along the edge of the sea-ice and
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south of Iceland (Figure 4a). In OP115 the different orbital forcing leads initially
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to more extensive sea-ice in the Labrador Sea (Figure 4b), because the melting
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season is shortened (Figure 1). The Irminger Sea, however, is under the influence
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of the warm North Atlantic Drift, which leads to only minor changes in the sea
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ice. The increase in sea-ice concentration insulates the Labrador Sea from atmo-
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spheric forcing, and therefore inhibits convective activity and reduces the strength
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of the subpolar gyre (Figure 3c). South of Iceland, without the effects sea-ice, the
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orbital forcing leads merely to a cooling of the surface water, and therefore a desta-
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bilization of the water column (Figure 4c) and deeper convection (Figure 4b). The
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intermediate depth warming see in Figures 4c and 3d is the direct result of the
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weaker Labrador Sea gyre: the weaker gyre leads to a poleward migration of the
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Gulf Stream and the North Atlantic Drift (Figure 4d), with its associated influx of
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more spicy (warmer and saltier) water. For water of equal density, the atmosphere
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removes buoyancy more efficiently from spicier water (Jochum 2009), so that af-
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ter 300 years the convective activity in the subpolar gyre is quite similar between
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CONT and OP115 (not shown), albeit in a different background state (Figure 3d).
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Thus, the connection between subpolar gyre strength and subtropical spiciness ad-
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vection acts as a negative feedback that stabilizes the gyre.
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Another source of North Atlantic Deep Water is convection in the Norwegian
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Sea, which enters the subpolar gyre through the Denmark Strait and Faroe Bank
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Channel overflows. In CONT, this is a 5.2 (± 0.4) Sv contribution to the AMOC.
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This value is determined by a parameterization and depends largely on the density
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differences between the water masses on either sides of the ridges (Danabasoglu
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et al., 2010). North of the ridges, the changes in salinity and temperature largely
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compensate each other so that the maximum boundary depth and density changes
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are minor (not shown). Thus, the transient response is determined by the behaviour
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of the subpolar gyre, which leads to an initial cooling of the waters on the Atlantic
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side of the ridge, and then a recovery as the gyre and the AMOC strengthen (Figure
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4c). The cooling leads to minimum overflow of 4.1 Sv after 120 years (about 40
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years after the minimum in AMOC and gyre strength) and then recovers towards
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a transport of 5.4 Sv in years 751-800 (not shown). Thus, the variability of the
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overflow strength, too, is controlled by the subpolar gyre.
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In principle, the two mechanisms described above should have come into play
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in JPLM as well. There, however, an increased freshwater export out of the Arc-
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tic led to a halocline catastrophe (see Bryan 1986 for a general discussion of this
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process). It turns out that in CCSM4 with its finer spatial resolution there is a cru-
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cial third process that allows the AMOC to recover: the freshwater flux through the
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Baffin Bay (Figure 5). An analysis of the subpolar freshwater budget shows that its
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largest anomalies are the liquid freshwater transport through the Nares Strait and
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Northwest Passage (via the Baffin Bay), and the sea-ice import through the Fram
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Strait (Figure 6): Immediately after changing the orbital forcing, the import of sea-
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ice through the Fram Strait increases, and continues to increase for another 100
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years. This is also happening in the coarse resolution version and leads to the halo-
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cline catastrophe (JPLM). In the present experiment, however, a good part of the
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increased ice import is compensated for by reduced inflow of liquid freshwater into
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the Baffin Bay. The strength of the flow into Baffin Bay is determined by the sea
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surface height difference between the Arctic and the Labrador Sea (Prinsenberg
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and Bennett 1987; Kliem and Greenberg 2003; Jahn et al. 2010). This difference
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is reduced from 70 cm in CONT to 50 cm during the years 551-600 in OP115 (not
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shown).
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Thus, there are two negative feedbacks by which the effect of orbital forcing on
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the AMOC is minimized. Both work through the subpolar gyre: Firstly, increased
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sea-ice cover reduces its strength; this brings in spicier subtropical water, which is
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more susceptible to convection; and secondly, the reduced gyre strength leads to a
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reduced pressure difference between the Arctic and the Labrador Sea, thereby re-
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ducing the import of freshwater through the Nares Strait and Northwest Passage.
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To what extent the present feedbacks are relevant in the real world will depend
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on an accurate representation of the Arctic and subpolar physics. A detailed dis-
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cussion of model performance and biases is provided in Danabasoglu et al. (2011);
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here we will limit ourselves to a quantification of some of the key processes. As
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in the observations (Dickson and Brown 1994; Vage et al., 2009), deep water for-
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mation occurs in the Irminger, Labrador and the Greenland Seas with maximum
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boundary layer depths of approximately 1000 m in the latter two, and 800 m in
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the former. However, deep convection is not observed south of Iceland and this is
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a major bias in the model. The product of the Denmark Strait and Faroe Bank
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Channel overflow water is well reproduced with 5.2 Sv when compared to the re-
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cent observational estimates of 6.5 Sv (Girton and Sandford 2003; Mauritzen et al.
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2005). The shutdown of the Labrador Sea convection in OP115 leads to 1 Sv reduc-
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tion in the AMOC, which is consistent with recent evidence that the contribution of
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Labrador Sea convection to the AMOC is less than 2 Sv (Pickart and Spall 2007).
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With 19.5 Sv the strength of the AMOC at 26.5 N is close to the observed value
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of 18.7 Sv (Kanzow et al. 2009). The freshwater budget and the sea-ice properties
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of the Artic Ocean are also well represented in CCSM4 (Jahn et al., 2011, to be
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submitted). This general assessment of the model preformance is encouraging, but
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it should be kept in mind that the response of the AMOC to disturbances can be
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critically dependent on poorly constrained mixing processes (e.g.; Schmittner and
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Weaver 2001; Saenko et al. 2003).
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Observational evidence for the past AMOC strength comes from proxy tracers
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obtained from deep ocean cores for the Last Glacial Maximum (LGM, 21 kya). The
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most widely used tracers to infer ocean circulation patterns during the LGM are
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carbon isotopes (δ 13 C), Cadmium-Calcium ratios (Cd/Ca) in benthic foraminifera,
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Protactinium-Thorium ratios (P a/T h) in benthic foraminifera, and estimates of the
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density gradient across the Atlantic basin based on oxygen isotope profiles (δ 18 O).
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Measurements of sortable silts have been used to infer rates of past deep-ocean
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flows. While for each of these tracer studies exist that indicate a weaker or shal-
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lower AMOC during the LGM (e.g.; Lynch-Stieglitz et al. 2007; Gherardi et al.
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2009; Lynch-Stieglitz et al. 2006; McCave et al. 1995), the more recent application
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of rigorous statistical tools suggests that the available database is currently not
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sufficient to reject the hypothesis of an unchanged AMOC during the LGM (e.g.;
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Gebbie and Huybers 2006; Marchal and Curry 2008; Peacock 2010). The present
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authors are aware of only one observation based analysis of the AMOC strength
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that stretches past 115 kya, and it suggests that the AMOC started weakening
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only several thousand years after the beginning of the last glacial inception (Gui-
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hou et al. 2010). Thus, the present model result of 115 kya orbital forcing having no
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impact on the strength of the AMOC is consistent with the available observations.
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The pattern and amplitude of wind stress and precipitation response is similar
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to the one in the coarse resolution study of JPLM, involving minor changes with
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the exception of a stronger Indian Summer Monsoon and stronger westerlies over
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the North Pacific (not shown). In particular the zonally averaged wind stress over
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the Southern Ocean is identical to within 0.5 %, and its maximum is at the same
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latitude. As already illustrated in Figure 2 the main differences between OP115
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and CONT are in the Arctic and will be analyzed here.
The Atmospheric Response
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The difference in orbital forcing between OP115 and CONT leads to larger in-
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coming radiation at the top of the atmosphere (TOA) in the tropics, and smaller
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radiation at high latitudes in the former compared to the latter (Figure 7a), with
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the total incoming solar radiation being 0.3 W/m2 larger in OP115 than in CONT.
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Our focus will be on the northern hemisphere north of 60◦ N, which covers the areas
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of large cooling and increased snow cover (Figure 2). Compared to CONT, the an-
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nual average of the incoming radiation over this Arctic domain is smaller in OP115
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by 4.3 W/m2 (black line), but the large albedo reduces this difference at the TOA
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to only 1.9 W/m2 (blue line, see also Table 1). This is only a minor change, and the
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core piece of the Milankovitch hypothesis is how this signal is spread across the
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seasonal cycle and amplified (e.g.; Huybers and Tziperman 2008): reduced summer
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insolation (Figure 1) prolongs the time during which land is covered with snow,
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and ocean covered with ice (Figure 2d). In CCSM4 this larger albedo in OP115
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leads to a TOA clearsky shortwave radiation that is 8.6 W/m2 smaller than in
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CONT (red line; clearsky radiation is diagnosed during the simulation by omitting
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the clouds from the radiation calculation) - more than four times the original sig-
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nal. The snow/ice albedo feedback is then calculated as 6.7 W/m2 (8.6 W/m2 - 1.9
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W/m2 ). Interestingly, the low cloud cover is smaller in OP115 than in CONT, re-
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ducing the difference in total TOA shortwave radiation by 3.1 W/m2 to 5.5 W/m2
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(green line). Summing up, an initial forcing of 1.9 W/m2 north of 60◦ N, is amplified
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through the snow/ice albedo feedback by 6.7 W/m2 , and damped through a negative
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cloud feedback by 3.1 W/m2 .
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The regions of TOA shortwave difference are confined to areas of increased
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snow and sea-ice (compare Figure 7b with Figure 2d). The negative feedback of
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the clouds is mostly due to a reduction of low-cloud cover over the ocean during
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summer (Figure 7c). High latitude stratus clouds have a similar effect as snow or
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ice: they reflect sunlight. Unlike at lower latitudes, Arctic low cloud formation via
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coupling with the ocean is frequently inhibited by sea ice and surface tempera-
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ture inversions. While the presence of open water in the Arctic does not guarantee
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atmosphere-ocean coupling and low cloud formation, open water has been shown
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to increase Arctic cloud cover when atmosphere-ocean coupling is strong (Kay and
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Gettelman 2009). Since the Arctic Ocean in OP115 is colder and has more sea ice
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cover than in CONT, especially in summer, reductions in low cloud amount are not
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surprising, and serve to counteract the positive ice-albedo feedback from increased
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sea ice cover.
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Because of the larger meridional temperature (Figure 2a) and moisture (not
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shown) gradient, the lateral atmospheric heatflux into the Arctic is increased from
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2.88 to 3.00 PW. This 0.12 PW difference translates into an Arctic average of 3.1
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W/m2 , a negative feedback as large as the cloud feedback. and ten times as large as
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the changes to the ocean meridional heat transport (Figure 3b). Thus, the negative
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feedback of the clouds and the meridional heat transport almost compensate for
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the positive albedo feedback, leading to a total feedback of only 0.5 W/m2 . One way
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to look at these feedbacks is that the climate system is quite stable, with clouds and
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meridional transports limiting the impact of albedo changes. This may explain why
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some numerical models have difficulties creating the observed cooling associated
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with the orbital forcing (e.g.; Jackson and Broccoli 2003).
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Ultimately, of course, a successful simulation of the inception does not necessar-
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ily need cooling, but an increased snow and ice cover to build ice sheets. In principle
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the increased snow depth seen in Figure 2b could be due to increased snow fall or
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reduced snow melt. The global moisture budgets reveal that OP115 has a larger
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poleward moisture transport than CONT (not shown), which is consistent with the
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increased heat transport. This does lead to increased snow fall, but in contrast to
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the results of Vettoretti and Peltier (2003) this is negligible compared to the reduc-
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tion in snow melt (Figure 7d). The global net difference in melting and snowfall
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between OP115 and CONT leads to a snow built-up, the volume increase of which
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is equivalent to a sea-level drop of 20 m in 10.000 years. This is less than the 50 m
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estimate based on sea-level reconstructions between present day and 115 kya (e.g.;
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Lambeck and Chappell 2001), but nontheless it suggests that the model response
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is of the right magnitude.
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It is shown here that the same CCSM4 version that realistically reproduces the
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present day climate, also shows, when subjected to orbital forcing from 115 kya,
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increased areas of perennial snow in exactly the areas where glacial reconstruc-
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tions put the origin of the glacial ice sheets. Furthermore, the difference in snow
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deposition between OP115 and CONT is of the same order of magnitude as the
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reconstructions based on sea level data. CCSM4 achieves this by the most basic
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mechanism, which was already postulated by Milankovitch (1941): reduced north-
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ern hemisphere summer insolation reduces snow and sea-ice melt, which leads to
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larger areas with perennial snow and sea-ice, and increased albedo. This, in turn,
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futher reduces melting, leading to the positive snow/ice-albedo feedback.
Summary and Discussion
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This positive feedback is opposed and almost compensated for by increased
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meridional heat transport in the atmosphere, and by reduced low cloud cover
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mainly over the Arctic ocean. The former is on solid theoretical grounds (e.g.; Stone
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1978) and has support from other GCM studies (e.g.; Shaffrey and Sutton 2006),
380
but the physics of low Arctic clouds is one of the weak points of current GCMs (e.g.;
381
Kay et al. 2011). This does not mean that the CCSM4 response of Arctic clouds
382
to insolation is necessarily wrong. The paucity of observations relevant for sea-
383
ice/cloud feedbacks means, however, that there have been only few opportunities to
17
384
test this aspect of the model.
385
An equally large source of uncertainty is the stability of the AMOC. To our
386
knowledge the only other inception study with a full, freely evolving GCM is JPLM,
387
and they find a 30% reduction of the AMOC strength under 115 kya solar forcing.
388
The present GCM has better physics and higher resolution, so it is tempting to
389
claim that its results are more realistic. However, the fact that CCSM4 is able to
390
recreate a reasonable inception scenario without an active ocean removes the 115
391
kya scenario as a testbed for the ocean model. As discussed in Section 3, the ob-
392
servations are not conclusive either, nor is there a comprehensive theory for the
393
AMOC strength (see Kuhlbrodt et al. 2007 for a recent review). Thus, from the
394
present set of experiments we can only conclude that ocean feedbacks are not nec-
395
essary for glacial inception, but we cannot rule out that the ocean did play a role.
396
While the increase in perennial snow cover and snow deposition is encouraging,
397
the present study is only a first step toward reproducing the glacial inception in a
398
GCM. It still remains to be shown that the OP115 climate allows for the growth of
399
the Scandinavian, Siberian and Laurentide ice sheets. More importantly, though,
400
the lack of any significant southern hemisphere polar response needs explaining
401
(Figure 2). While Petit et al. (1999) suggests that Antarctica cooled by about 10◦ C
402
during the last inception, some studies point to the substantial uncertainties of
403
these ice-core based temperature measurements (e.g.; Stenni et al. 2010). There is,
404
however, clear evidence that during the LGM the Southern Ocean sea-ice cover was
405
much more extensive (Gersonde et al. 2005), so that the GCMs should at least show
406
some increased sea-ice concentrations during glacial inception. Unfortunately, the
18
407
present day simulation of CCSM4 does already have a significant excess of sea ice
408
(Landrum et al., to be submitted to J. Climate), due at least in part to its exces-
409
sively strong Southern Ocean winds (Holland and Raphael 2006). Arguably this
410
leads to a reduced sensitivity to reduced springtime insolation, so that one focus of
411
future model development will have to be improvement of the overly strong South-
412
ern Ocean winds - a problem which, strangely enough, has been unresolved since
413
Boville (1991).
414
Thus, while the present study finally provides numerical support for the Mi-
415
lankovitch hypothesis of glacial inception, it also identifies two foci of future re-
416
search: Firstly, the sensitivity of Arctic clouds to climate fluctation needs to be
417
validated and possibly improved upon; and secondly, the Southern Ocean sea-ice
418
concentration and the strength of the zonal winds need to become more realistic.
419
420
421
Acknowledgements:
422
The authors are supported by NSF through NCAR. The participation of JTF
423
is made possible through funding under NASA Contract # NNG06GB91G. The
424
computations were done on the NSF supercomputing facilities at NCAR. The
425
authors thank Dave Lawrence and Keith Oleson for explaining the intricacies of
426
the snow budget, and Trout fishing in America for inspiration.
427
428
429
19
430
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27
Table 1: Summary of the annual mean heat flux feedbacks north of 60◦ N.
Process
insolation forcing
× present day albedo
snow/ice albedo feedback
low-cloud/sea-ice feedback
merid. heat transp. feedback
total forcing
584
585
Strength [W/m2 ]
+ 4.3
- 2.4
+ 6.7
- 3.1
- 3.1
+ 2.4
586
Figure 1: Seasonal cycle of insolation at the top of the atmosphere for CONT
587
(black) and OP115 (red) at 70◦ N (maximum in June) and 70◦ S. In OP115 the June
588
insolation is 38 W/m2 weaker than in CONT.
589
Figure 2: Difference in annual mean surface temperature between OP115 and
590
CONT (a), and between their respective slab-ocean counterparts (b). c: Difference
591
in annual mean snow depth between OP115 and CONT, d: area where the summer
592
(minimum) snow depth is larger than 1 cm and the summer sea-ice concentration
593
is larger than 40% for CONT (red), and its increase in OP115 (blue),
594
Figure 3: a: Strength of the AMOC (in Sv) at 50◦ N for CONT (black) and OP115
595
(red). The straight black lines denote the mean and mean ± one standard deviation
596
of the annual mean of CONT. Here, and in the next 2 panels, only years 501-700
597
are shown, because CONT ends at year 700; OP115 slowly approaches the val-
598
ues cited in the text. b: Like (a), but for the meridional heat transport (in PW).
599
c: Maximum strength of the subpolar gyre (Sv, typically the maximum is located
600
between Greenland and Labrador). Again, the black line is based on CONT, the red
601
on OP115. Note that all the timeseries have been smoothed with an 11-year run-
602
ning mean. d: Mean temperature (color) and salinity (contour interval: 0.02 psu,
603
minimum: -0.4 psu, maximum: 0.1 psu) difference between Labrador and Norway.
604
Figure 4: a: Mean maximum annual boundary layer depth (m, in color) and
605
annual mean sea ice concentration (contour lines: 10%) for CONT. b: Like (a), but
606
for years 551-600 of OP115. c: Mean temperature profile south of Iceland. d: Depth
607
integrated transport in CONT (color), and difference between OP115 and CONT
608
(OP115-CONT, contour interval: 1 Sv).
609
610
Figure 5: Sea Surface Salinity north of 50◦ N. Also shown are the locations of
several seas and passages that are mentioned in the text.
611
Figure 6: Change in the total freshwater (FW) import from the Arctic Ocean to
612
the subpolar Atlantic (thick black line) in OP115 compared to the mean of years
613
501-700 from CONT. This anomaly is also shown split up into the contribution of
614
the FW import east (red line) and west (blue line) of Greenland, as well as split up
615
into the solid FW import (mainly east of Greenland; solid thin black line) and the
616
liquid FW import (dashed line). The fluxes are computed relative to a salinity of
617
34.8 (Aagaard and Carmack 1989).
618
Figure 7: a: Difference of zonally averaged flux at the top of the atmosphere
619
(OP115-CONT). Black: insolation. Blue: insolation times albedo of CONT. Red:
620
clear-sky net shortwave radiation. Green: net shortwave radiation. b: Spatial
621
pattern of the difference in net shortwave radiation [in W/m2 ]. c: Difference in
622
maximum (summer) low level cloud concentration (color, in %) and minimum
623
(summer) sea-ice concentration (contour interval 5 %). For the sake of clarity
624
only the cloud cover over the ocean is shown; with the exception of the Canadian
625
Archipelago, the cloud cover changes over land are less than 5%. d: Difference in
626
snow melt (color, in mm/day) and snow fall (contour lines: 0.1 mm/day).
627
628
Figure 1: Seasonal cycle of insolation at the top of the atmosphere for CONT (black)
and OP115 (red) at 70◦ N (maximum in June) and 70◦ S. In OP115 the June insolation is 38 W/m2 weaker than in CONT.
629
Figure 2: Difference in annual mean surface temperature between OP115 and
CONT (a), and between their respective slab-ocean counterparts (b). c: Difference
in annual mean snow depth between OP115 and CONT, d: area where the summer
(minimum) snow depth is larger than 1 cm and the summer sea-ice concentration
is larger than 40% for CONT (red), and its increase in OP115 (blue),
Figure 3: a: Strength of the AMOC (in Sv) at 50◦ N for CONT (black) and OP115
(red). The straight black lines denote the mean and mean ± one standard deviation
of the annual mean of CONT. Here, and in the next 2 panels, only years 501-700
are shown, because CONT ends at year 700; OP115 slowly approaches the values cited in the text. b: Like (a), but for the meridional heat transport (in PW).
c: Maximum strength of the subpolar gyre (Sv, typically the maximum is located
between Greenland and Labrador). Again, the black line is based on CONT, the red
on OP115. Note that all the timeseries have been smoothed with an 11-year running mean. d: Mean temperature (color) and salinity (contour interval: 0.02 psu,
minimum: -0.4 psu, maximum: 0.1 psu) difference between Labrador and Norway.
Figure 4: a: Mean maximum annual boundary layer depth (m, in color) and annual mean sea ice concentration (contour lines: 10%) for CONT. b: Like (a), but for
years 551-600 of OP115. c: Mean temperature profile south of Iceland. d: Depth
integrated transport in CONT (color), and difference between OP115 and CONT
(OP115-CONT, contour interval: 1 Sv).
Figure 5: Sea Surface Salinity north of 50◦ N. Also shown are the locations of several
seas and passages that are mentioned in the text.
more FW
import
2500
1500
3
FW import anomaly [km /yr]
2000
1000
500
0
−500
−1000
−1500
less FW
import
−2000
500
550
600
650
simulation year
700
750
800
Figure 6: Change in the total freshwater (FW) import from the Arctic Ocean to
the subpolar Atlantic (thick black line) in OP115 compared to the mean of years
501-700 from CONT. This anomaly is also shown split up into the contribution of
the FW import east (red line) and west (blue line) of Greenland, as well as split up
into the solid FW import (mainly east of Greenland; solid thin black line) and the
liquid FW import (dashed line). The fluxes are computed relative to a salinity of
34.8 (Aagaard and Carmack 1989).
Figure 7: a: Difference of zonally averaged flux at the top of the atmosphere
(OP115-CONT). Black: insolation. Blue: insolation times albedo of CONT. Red:
clear-sky net shortwave radiation. Green: net shortwave radiation. b: Spatial pattern of the difference in net shortwave radiation [in W/m2 ]. c: Difference in maximum (summer) low level cloud concentration (color, in %) and minimum (summer)
sea-ice concentration (contour interval 5 %). For the sake of clarity only the cloud
cover over the ocean is shown; with the exception of the Canadian Archipelago, the
cloud cover changes over land are less than 5%. d: Difference in snow melt (color,
in mm/day) and snow fall (contour lines: 0.1 mm/day).