- About
- Editorial board
- Articles
- Special issues
- Highlight articles
- Manuscript tracking
- Book reviews
- Subscribe to alerts
- Peer review
- For authors
- For reviewers
- EGU publications

Journal cover
Journal topic
**Nonlinear Processes in Geophysics**
An interactive open-access journal of the European Geosciences Union

Journal topic

- About
- Editorial board
- Articles
- Special issues
- Highlight articles
- Manuscript tracking
- Book reviews
- Subscribe to alerts
- Peer review
- For authors
- For reviewers
- EGU publications

- About
- Editorial board
- Articles
- Special issues
- Highlight articles
- Manuscript tracking
- Book reviews
- Subscribe to alerts
- Peer review
- For authors
- For reviewers
- EGU publications

- Abstract
- Introduction
- An energy balance model for climate change
- Anthropocene climate forecasts
- Conclusions and future work
- Appendix A: Determination of parameters in the Anthropocene EBM
- Code and data availability
- Author contributions
- Competing interests
- Acknowledgements
- Financial support
- Review statement
- References

**Research article**
08 Jul 2020

**Research article** | 08 Jul 2020

Anthropocene climate bifurcation

- Department of Mathematics and Statistics, University of Guelph, 50 Stone Road West, Guelph, ON, N1G 2W1, Canada

- Department of Mathematics and Statistics, University of Guelph, 50 Stone Road West, Guelph, ON, N1G 2W1, Canada

**Correspondence**: William Finlay Langford (wlangfor@uoguelph.ca)

**Correspondence**: William Finlay Langford (wlangfor@uoguelph.ca)

Abstract

Back to toptop
This article presents the results of a bifurcation analysis of a simple energy balance model (EBM)
for the future climate of the Earth. The main focus is on the following question: can the nonlinear
processes intrinsic to atmospheric physics, including natural positive feedback mechanisms, cause
a mathematical bifurcation of the climate state, as a consequence of continued anthropogenic
forcing by rising greenhouse gas emissions? Our analysis shows that such a bifurcation could
cause an abrupt change to a drastically different climate state in the EBM, which is warmer and more
equable than any climate existing on Earth since the Pliocene epoch. In previous papers, with
this EBM adapted to paleoclimate conditions, it was shown to exhibit saddle-node and cusp
bifurcations, as well as hysteresis. The EBM was validated by the agreement of its predicted
bifurcations with the abrupt climate changes that are known to have occurred in the paleoclimate
record, in the Antarctic at the Eocene–Oligocene transition (EOT) and in the Arctic at the
Pliocene–Paleocene transition (PPT). In this paper, the EBM is adapted to fit Anthropocene
climate conditions, with emphasis on the Arctic and Antarctic climates. The four Representative
Concentration Pathways (RCP) considered by the IPCC (Intergovernmental Panel on Climate Change) are used to model future CO_{2}
concentrations, corresponding to different scenarios of anthropogenic activity. In addition, the
EBM investigates four naturally occurring nonlinear feedback processes which magnify the warming
that would be caused by anthropogenic CO_{2} emissions alone. These four feedback mechanisms
are ice–albedo feedback, water vapour feedback, ocean heat transport feedback, and atmospheric
heat transport feedback. The EBM predicts that a bifurcation resulting in a catastrophic climate
change, to a pre-Pliocene-like climate state, will occur in coming centuries for an RCP with
unabated anthropogenic forcing, amplified by these positive feedbacks. However, the EBM also
predicts that appropriate reductions in carbon emissions may limit climate change to a more
tolerable continuation of what is observed today. The globally averaged version of this EBM has
an equilibrium climate sensitivity (ECS) of 4.34 K, near the high end of the likely range reported
by the IPCC.

How to cite

Back to top
top
How to cite.

Kypke, K. L., Langford, W. F., and Willms, A. R.: Anthropocene climate bifurcation, Nonlin. Processes Geophys., 27, 391–409, https://doi.org/10.5194/npg-27-391-2020, 2020.

1 Introduction

Back to toptop
Today, there is widespread agreement that the climate of the Earth is changing, but the precise trajectory of future climate change is still a matter of debate. Recently there has been much interest in the possibility of “tipping points” (or bifurcation points) at which abrupt changes in the Earth climate system occur (see Brovkin et al., 1998; Ghil, 2001; Alley et al., 2003; Seager and Battisti, 2007; Lenton et al., 2008; Ditlevsen and Johnsen, 2010; Lenton, 2012; Ashwin et al., 2012; Barnosky et al., 2012; Drijfhout et al., 2015; Bathiany et al., 2016; North and Kim, 2017; Steffen et al., 2018; Dijkstra, 2019; Wallace-Wells, 2019). Section 12.5.5 in IPCC (2013) gives an overview of such potential abrupt changes. At such points, a small change in the forcing parameters (whether anthropogenic or natural forcings) may cause a catastrophic change in the state of the system. In order to prepare for future climate change, it is of great importance to know if such abrupt changes can occur and, if so, when and how they will occur. Some authors have suggested that conventional general circulation models (GCMs) may be “too stable” to provide reliable warning of these sudden catastrophic events (Valdes, 2011), and that the study of paleoclimates may be a better guide to how abrupt changes may occur (Zeebe, 2011). Steffen et al. (2018) asked the fundamental question: “Is there a planetary threshold in the trajectory of the Earth System that, if crossed, could prevent stabilization in a range of intermediate temperature rises?”

In this paper, a simple energy balance model (EBM) is used to investigate the possible occurrence of
such a threshold (tipping point or bifurcation point) in the climate of the Anthropocene. Energy
balance models assume that the climate is in an equilibrium state, for which “energy in equals
energy out” at each point of the Earth's surface and atmosphere. Thus, time dependence is
eliminated from the model, greatly simplifying the analysis. In the literature, EBMs have played an
important role in understanding climate and climate change
(Budyko, 1968; Sellers, 1969; Sagan and Mullen, 1972; North et al., 1981; Thorndike, 2012; Kaper and Engler, 2013; Dijkstra, 2013; Payne et al., 2015; Hartmann, 2016; North and Kim, 2017; Ghil and Lucarini, 2020). Many of those EBMs have
exhibited bifurcations. The present paper presents an EBM with more accurate diagnostic equations
than the early EBMs, built upon the basic laws of geophysics and including nonlinear feedback
processes that amplify anthropogenic CO_{2} forcing. A rigorous mathematical bifurcation
analysis of this EBM has been presented in Kypke and Langford (2020). That analysis gave a mathematical proof of
the existence of a cusp bifurcation in the EBM, complete with a determination of the centre manifold
and of the universal unfolding parameters, as functions of the relevant physical parameters. The
existence of the cusp bifurcation implies the coexistence of two distinct stable equilibrium
climate states (bistability), as well as the existence of hysteresis; that is, two abrupt one-way
transitions between these two states (via fold bifurcations) exist in the EBM. The present paper
extends those results from the paleoclimate model in Kypke and Langford (2020) to the Anthropocene climate model
studied here. This Anthropocene model gives predictions of climate changes in the 21st century and
beyond.

One advantage of an EBM over a more complex GCM is that it facilitates the exploration of specific cause and effect relationships, as particular climate forcing factors are varied or ignored. Another advantage of an EBM is that rigorous mathematical analysis often can prove the existence of certain behaviours, such as bistability and bifurcations, that could only be surmised from numerical evidence, or missed, in more complicated models. Four versions of the EBM are considered here: a globally averaged temperature model and three regional models corresponding to Arctic, Antarctic, and tropical climates.

This EBM was validated in Dortmans et al. (2019), where it was applied to known paleoclimate changes. It successfully “predicted” the abrupt glaciation of Antarctica at the Eocene–Oligocene transition (EOT) and the abrupt glaciation of the Arctic at the Pliocene–Pleistocene transition (PPT), using bifurcation analysis. The transitions in the model were congruent with the abrupt cooling, from warm, equable “hothouse” climate conditions to a cooler state like the climate of today, with ice-capped poles, as indelibly recorded in the geological record at the EOT and PPT.

In adapting the previous paleoclimate EBM to the climate of the Anthropocene, this paper explores the possibility of a “reversal” of those two paleoclimate transitions; that is, a transition from today's climate with ice-capped poles to an equable hothouse climate state, such as existed in the Pliocene and earlier. It provides new mathematical evidence suggesting that catastrophic climate change in polar regions is inevitable in the coming decades and centuries if current anthropogenic forcing continues unabated. The EBM also suggests that if appropriate mitigation strategies are adopted (as recommended by the IPCC), such an outcome can be avoided.

The EBM of this paper has been kept as simple as possible, while incorporating the nonlinear physical processes that are essential to our exploration of bifurcation behaviour. In that sense, it follows in the tradition of simple energy balance models of Budyko (1968), Sellers (1969), North et al. (1981), and others. However, this EBM is but the first step in the authors' study of a hierarchy of nonlinear models of increasing complexity. That hierarchy is outlined in the concluding Sect. 4.

2 An energy balance model for climate change

Back to toptop
The EBM is a simple two-layer model, with layers corresponding to the surface and atmosphere,
respectively; see Fig. 1, which is based on Payne et al. (2015), Trenberth et al. (2009), and Wild et al. (2013). The symbols in
Fig. 1 are defined in the caption of Fig. 1 and in Table 1.
In the EBM of Fig. 1, the surface receives shortwave radiant energy
${F}_{\mathrm{S}}=(\mathrm{1}-{\mathit{\xi}}_{\mathrm{A}}-{\mathit{\xi}}_{\mathrm{R}})Q$ from the sun, where *Q* is the incident
solar radiation and *ξ*_{A} and *ξ*_{R} are the fractions of *Q* absorbed by the
atmosphere and reflected back to space by clouds, respectively. The values of *ξ*_{A}
and *ξ*_{R} are obtained from Trenberth et al. (2009) and Wild et al. (2013) (see the appendix and
Table 1). At the surface, a fraction *α**F*_{S} is reflected back to
space, where *α* is the surface albedo, and the remainder, (1−*α*)*F*_{S}, is
absorbed by the surface. The surface re-emits longwave radiant energy of intensity
${I}_{\mathrm{S}}=\mathit{\sigma}{T}_{\mathrm{S}}^{\mathrm{4}}$ (Stefan–Boltzmann law) upward into the atmosphere. The
atmosphere contains greenhouse gases that absorb a fraction *η* of the radiant energy
*I*_{S} from the surface, and the remainder, (1−*η*)*I*_{S}, escapes to space.
The atmosphere re-emits radiant energy of total intensity *I*_{A}. Of this radiation
*I*_{A}, a fraction *β**I*_{A} is directed downward to the surface, and the remaining fraction
(1−*β*)*I*_{A} goes upward and escapes to space. Balancing the energy flows represented in
Fig. 1 leads directly to the following dynamical system:

$$\begin{array}{ll}{\displaystyle}{c}_{\mathrm{S}}{\displaystyle \frac{\mathrm{d}{T}_{\mathrm{S}}}{\mathrm{d}t}}=& {\displaystyle}\phantom{\rule{0.125em}{0ex}}(\mathrm{1}-\mathit{\alpha})(\mathrm{1}-{\mathit{\xi}}_{\mathrm{A}}-{\mathit{\xi}}_{\mathrm{R}})Q+{F}_{O}+\mathit{\beta}{I}_{\mathrm{A}}\\ \text{(1)}& {\displaystyle}& {\displaystyle}\phantom{\rule{0.125em}{0ex}}-\mathit{\sigma}{T}_{\mathrm{S}}^{\mathrm{4}}-{F}_{\mathrm{C}},\text{(2)}& {\displaystyle}{c}_{\mathrm{A}}{\displaystyle \frac{\mathrm{d}{I}_{\mathrm{A}}}{\mathrm{d}t}}=& {\displaystyle}\phantom{\rule{0.125em}{0ex}}{F}_{\mathrm{A}}+\mathit{\eta}\mathit{\sigma}{T}_{\mathrm{S}}^{\mathrm{4}}-{I}_{\mathrm{A}}+{F}_{\mathrm{C}}+{\mathit{\xi}}_{\mathrm{A}}\phantom{\rule{0.25em}{0ex}}Q,\end{array}$$

where Eq. (1) represents surface energy balance and Eq. (2) represents
atmosphere energy balance. Here *c*_{S} and *c*_{A} are specific heat rate
constants derived in Kypke (2019) and Kypke and Langford (2020) and listed in Table 1. There are three
heat transport terms: *F*_{A} is *atmospheric* heat transport, *F*_{O}
is *ocean* heat transport (both horizontally), and *F*_{C} is
*conductive/convective* heat transport, vertically from the surface to the atmosphere.

In Eqs. (1) and (2) there is an asymmetry, in that temperature
*T*_{S} is used to represent the state of the surface in Eq. (1), but radiant
energy intensity *I*_{A} is used instead of temperature to represent the state of the
atmosphere in Eq. (2). Note that either temperature variables or energy intensity variables could
have been used in either equation if we assume the Stefan–Boltzmann law
(${I}_{\mathrm{S}}=\mathit{\sigma}{T}_{\mathrm{S}}^{\mathrm{4}}$ and ${I}_{\mathrm{A}}=\mathit{\u03f5}\mathit{\sigma}{T}_{\mathrm{A}}^{\mathrm{4}}$,
where *ϵ*=0.9 is the emissivity of dry air). The use of *T*_{S} as state variable
in surface Eq. (1) is the obvious choice, since the surface temperature is the most
important variable in the EBM. However, the choice of *I*_{A} instead of *T*_{A}
in Eq. (2) is less obvious. The atmosphere has thickness. In the actual atmosphere,
temperature decreases with height above the surface at a rate called the *lapse rate*.
Therefore, there is not just one value of temperature *T*_{A} for the atmosphere. However,
we can define a single value of radiant energy intensity *I*_{A}, corresponding to the
total energy emitted vertically by the slab of atmosphere, and use this instead of temperature in
the energy balance Eq. (2). A second reason for the use of *I*_{A} instead of
*T*_{A} in Eq. (2) is that this facilitates the use of the ICAO Standard
Atmosphere lapse rate, as explained in the paragraph following Eqs. (4)
and (5) below, and that allows for a more realistic representation of the greenhouse gas
behaviour of water vapour.

The first step of the analysis of system Eqs. (1) and (2) is a rescaling of
temperature *T* (in kelvin) to a new nondimensional temperature *τ* with *τ*=1
corresponding to the freezing temperature of water (*T*_{R}=273.15 K). Then all
variables and parameters in the system can be made nondimensional by the scalings

$$\begin{array}{ll}{\displaystyle}{\mathit{\tau}}_{\mathrm{A}}& {\displaystyle}={\displaystyle \frac{{T}_{\mathrm{A}}}{{T}_{\mathrm{R}}}},\phantom{\rule{1em}{0ex}}{\mathit{\tau}}_{\mathrm{S}}={\displaystyle \frac{{T}_{\mathrm{S}}}{{T}_{\mathrm{R}}}},\phantom{\rule{1em}{0ex}}q={\displaystyle \frac{Q}{\mathit{\sigma}{T}_{\mathrm{R}}^{\mathrm{4}}}},\phantom{\rule{1em}{0ex}}{f}_{\mathrm{O}}={\displaystyle \frac{{F}_{\mathrm{O}}}{\mathit{\sigma}{T}_{\mathrm{R}}^{\mathrm{4}}}},\\ \text{(3)}& {\displaystyle}{f}_{\mathrm{A}}& {\displaystyle}={\displaystyle \frac{{F}_{\mathrm{A}}}{\mathit{\sigma}{T}_{\mathrm{R}}^{\mathrm{4}}}},\phantom{\rule{1em}{0ex}}{f}_{\mathrm{C}}={\displaystyle \frac{{F}_{\mathrm{C}}}{\mathit{\sigma}{T}_{\mathrm{R}}^{\mathrm{4}}}},\phantom{\rule{1em}{0ex}}{i}_{\mathrm{A}}={\displaystyle \frac{{I}_{\mathrm{A}}}{\mathit{\sigma}{T}_{\mathrm{R}}^{\mathrm{4}}}},\phantom{\rule{1em}{0ex}}\mathit{\omega}={\displaystyle \frac{\mathrm{\Omega}}{{T}_{\mathrm{R}}}},{\displaystyle}s& {\displaystyle}={\displaystyle \frac{\mathit{\sigma}{T}_{\mathrm{R}}^{\mathrm{3}}}{{C}_{\mathrm{S}}}}\cdot t,\phantom{\rule{1em}{0ex}}\mathit{\chi}={\displaystyle \frac{{C}_{\mathrm{S}}}{{C}_{\mathrm{A}}\mathit{\sigma}{T}_{\mathrm{R}}^{\mathrm{3}}}}=\mathrm{54.26},\end{array}$$

where *s* is dimensionless time and *χ* is the dimensionless heat rate constant. The surface and
atmosphere energy balance Eqs. (1) and (2) in nondimensional variables are
then

$$\begin{array}{ll}{\displaystyle \frac{\mathrm{d}{\mathit{\tau}}_{\mathrm{S}}}{\mathrm{d}s}}=& {\displaystyle}(\mathrm{1}-\mathit{\alpha}({\mathit{\tau}}_{\mathrm{S}}\left)\right)(\mathrm{1}-{\mathit{\xi}}_{\mathrm{A}}-{\mathit{\xi}}_{\mathrm{R}})q\\ \text{(4)}& {\displaystyle}& {\displaystyle}+{f}_{O}+\mathit{\beta}{i}_{\mathrm{A}}-{\mathit{\tau}}_{\mathrm{S}}^{\mathrm{4}}-{f}_{\mathrm{C}},\text{(5)}& {\displaystyle \frac{\mathrm{1}}{\mathit{\chi}}}{\displaystyle \frac{\mathrm{d}{i}_{\mathrm{A}}}{\mathrm{d}s}}=& {\displaystyle}{f}_{\mathrm{A}}+\mathit{\eta}({\mathit{\tau}}_{\mathrm{S}};\mathit{\mu},\mathit{\delta}){\mathit{\tau}}_{\mathrm{S}}^{\mathrm{4}}-{i}_{\mathrm{A}}+{f}_{\mathrm{C}}+{\mathit{\xi}}_{\mathrm{A}}\phantom{\rule{0.25em}{0ex}}q.\end{array}$$

In Fig. 1, the atmosphere is shown to be a uniform slab, even though the actual
atmosphere is not a uniform slab. The essential nonlinear processes in the atmosphere, which the
model must capture, are the heating effects of the greenhouse gases CO_{2} and H_{2}O.
According to the Beer–Lambert law, the absorptivity of these gases is determined by their
*optical depths*. Therefore, in the model we substitute for optical depth in the slab, the
values of optical depth that these gases would have in the ICAO Standard Atmosphere
(ICAO, 1993), which is a good approximation to the actual atmosphere. In the ICAO model, the rate
of change of temperature with altitude is assumed to be a negative constant −Γ, called the
ICAO *lapse rate*; see Table 1. The concentration of CO_{2} is
independent of temperature, but the concentration of H_{2}O decreases with altitude as the
temperature decreases, according to the Clausius–Clapeyron relation (Pierrehumbert, 2010). Then the
optical depth of H_{2}O is obtained by integration of the Clausius–Clapeyron relation from the
surface to the tropopause, resulting in Eq. (6). In this way, the simple slab model has
greenhouse effects close to those of these two gases in the actual atmosphere, where the temperature
is not constant but decreases with altitude. This use of the ICAO lapse rate differentiates the
present EBM from all previous EBMs in the literature. The calculation gives the absorptivity *η*
due to greenhouse gases, in the atmosphere Eqs. (2) or (5), as

$$\begin{array}{ll}{\displaystyle}\mathit{\eta}({\mathit{\tau}}_{\mathrm{S}};\mathit{\mu},\mathit{\delta})=& {\displaystyle}\mathrm{1}-\mathrm{exp}[-\mathit{\mu}{G}_{\mathrm{C}}-\mathit{\delta}{G}_{\text{W2}}\cdot \\ \text{(6)}& {\displaystyle}& {\displaystyle}\underset{{\mathit{\tau}}_{\mathrm{S}}-\mathit{\gamma}Z}{\overset{{\mathit{\tau}}_{s}}{\int}}{\displaystyle \frac{\mathrm{1}}{\mathit{\tau}}}\mathrm{exp}\left({G}_{\text{W1}}\right[{\displaystyle \frac{\mathit{\tau}-\mathrm{1}}{\mathit{\tau}}}\left]\right)d\mathit{\tau}],\end{array}$$

where *μ* is the concentration of CO_{2} in the atmosphere, measured in molar parts per
million; *δ* is the relative humidity of water vapour $(\mathrm{0}\le \mathit{\delta}\le \mathrm{1})$;
$\mathit{\gamma}=\mathrm{\Gamma}/{T}_{\mathrm{R}}$ is the nondimensionalized lapse rate (ICAO, 1993); and *Z* is the
tropopause height. (Since methane acts similarly to carbon dioxide as a greenhouse gas, it may be
assumed that *μ* includes also the effects of methane.) Equation (6) is derived using
fundamental laws of atmospheric physics: the Beer–Lambert law, the ideal gas law and the
Clausius–Clapeyron equation; see Dortmans et al. (2019) for more details. In Eq. (6),

$$\begin{array}{ll}{\displaystyle}{G}_{\mathrm{C}}& {\displaystyle}={\displaystyle \frac{\mathrm{1.52}{k}_{\mathrm{c}}{P}_{\mathrm{A}}}{{\mathrm{10}}^{\mathrm{6}}g}},\phantom{\rule{1em}{0ex}}{G}_{\text{W1}}={\displaystyle \frac{{L}_{\mathrm{v}}}{{R}_{\mathrm{W}}{T}_{\mathrm{R}}}},\\ \text{(7)}& {\displaystyle}& {\displaystyle}\phantom{\rule{1em}{0ex}}\text{and}\phantom{\rule{1em}{0ex}}{G}_{\text{W2}}={\displaystyle \frac{{k}_{\mathrm{W}}\cdot {T}_{\mathrm{R}}\cdot {\mathit{\rho}}^{\text{sat}}\left({T}_{\mathrm{R}}\right)}{\mathrm{\Gamma}}}\end{array}$$

are dimensionless physical constants determined in Dortmans et al. (2019), and *k*_{C} and
*k*_{W} are absorption coefficients for CO_{2} and H_{2}O, respectively; see
Table 1.

In the surface Eqs. (1) or (4), *α* is the albedo of the surface
$(\mathrm{0}\le \mathit{\alpha}\le \mathrm{1})$ and *α* depends strongly on temperature *τ*_{S} near the
freezing point (*τ*_{S}=1). Typical values of surface albedo are 0.6–0.9 for snow,
0.4–0.7 for ice, 0.2 for crop land, and 0.1 or less for open ocean. Therefore, in the high Arctic,
as the ice-cover melts, the albedo will transition from a high value such as
*α*_{c}=0.7 for snow/ice to a low value such as *α*_{w}=0.08 for
open ocean. Some authors have assumed this change in albedo to be a discontinuous step function
(Dortmans et al., 2018). However, all variables in this EBM have annually averaged values. As the Arctic
thaws, the annual average albedo will transition gradually, over a number of years, from its high
value for year-round ice-covered surface to its low value for year-round open ocean. Therefore, in
this paper we use a smoothly varying albedo function, which better models this gradual transition
from high to low albedo, as the mean temperature rises through the freezing point
(*τ*_{S}=1). It is modelled by the hyperbolic tangent function:

$$\begin{array}{}\text{(8)}& \mathit{\alpha}({\mathit{\tau}}_{\mathrm{S}},\mathit{\omega})={\displaystyle \frac{\mathrm{1}}{\mathrm{2}}}\left(\right[{\mathit{\alpha}}_{\mathrm{w}}+{\mathit{\alpha}}_{\mathrm{c}}]+[{\mathit{\alpha}}_{\mathrm{w}}-{\mathit{\alpha}}_{\mathrm{c}}\left]\mathrm{tanh}\right({\displaystyle \frac{{\mathit{\tau}}_{\mathrm{S}}-\mathrm{1}}{\mathit{\omega}}}\left)\right).\end{array}$$

Here the parameter *ω* controls the steepness of this switch function. Analysis of polar ice
data in recent years confirms that this function gives a good fit to the decline of ice cover and
albedo in the Arctic with $\mathit{\omega}=\mathrm{\Omega}/{T}_{\mathrm{R}}=\mathrm{0.01}$ (Pistone et al., 2014; Dortmans et al., 2019). The
dependence of Arctic Ocean sea-ice thickness on surface albedo parametrization in models has been
investigated in Björk et al. (2013), where alternative albedo schemes are compared. The nonlinear
dependence of albedo on temperature, as in Eq. (8), has been shown to lead to
hysteresis behaviour (Stranne et al., 2014; Dortmans et al., 2019).

Paleoclimate data are difficult to obtain and in general can only be inferred from proxy data. The situation is different for the Anthropocene. There is now an abundance of land-based and satellite climate data. Therefore, the EBM in this paper can be refined to take advantage of the additional data. The appendix details the changes made in this EBM from that which was presented in Dortmans et al. (2019) and Kypke and Langford (2020), to improve its accuracy for the Anthropocene. These changes do not change the fundamental behaviour of the EBM, including the existence of bifurcations, but they do make the numerical predictions more reliable. Table 1 of this paper may be compared with the corresponding Table 1 of Kypke and Langford (2020) to see how parameter values have been updated.

The total absorptivity *η*, given in Eq. (6) for paleoclimates, is now modified to
reflect the fact that clouds absorb a fraction *η*_{Cl} of the outgoing longwave
radiation. In this paper

$$\begin{array}{ll}{\displaystyle}\mathit{\eta}({\mathit{\tau}}_{\mathrm{S}},\mathit{\mu},\mathit{\delta})=& {\displaystyle}\mathrm{1}-(\mathrm{1}-{\mathit{\eta}}_{\text{Cl}})\cdot \mathrm{exp}[-\mathit{\mu}{G}_{\mathrm{C}}-\mathit{\delta}{G}_{\text{W2}}\phantom{\rule{0.33em}{0ex}}\cdot \\ \text{(9)}& {\displaystyle}& {\displaystyle}\underset{{\mathit{\tau}}_{\mathrm{S}}-\mathit{\gamma}Z}{\overset{{\mathit{\tau}}_{s}}{\int}}{\displaystyle \frac{\mathrm{1}}{\mathit{\tau}}}\mathrm{exp}\left({G}_{\text{W1}}\right[{\displaystyle \frac{\mathit{\tau}-\mathrm{1}}{\mathit{\tau}}}\left]\right)\mathrm{d}\mathit{\tau}];\end{array}$$

see the appendix.

The vertical heat transport term *F*_{C} has been modified to take into account both
sensible and latent heat transport (Kypke, 2019). See the appendix, where the following formula is
obtained (here in nondimensional form).

$$\begin{array}{ll}{\displaystyle}{f}_{\mathrm{C}}\left({\mathit{\tau}}_{\mathrm{S}}\right)=& {\displaystyle \frac{{H}_{\mathrm{1}}\mathit{\gamma}Z}{({\mathit{\tau}}_{\mathrm{S}}-\mathit{\gamma}Z)}}+{\displaystyle \frac{{H}_{\mathrm{2}}}{({\mathit{\tau}}_{\mathrm{S}}-\mathit{\gamma}Z)}}\phantom{\rule{0.33em}{0ex}}\phantom{\rule{0.33em}{0ex}}\cdot \\ \text{(10)}& {\displaystyle}& {\displaystyle}({e}^{\left[{G}_{\text{W1}}\frac{{\mathit{\tau}}_{\mathrm{S}}-\mathrm{1}}{{\mathit{\tau}}_{\mathrm{S}}}\right]}-\mathit{\delta}{e}^{\left[{G}_{\text{W1}}\frac{{\mathit{\tau}}_{\mathrm{S}}-\mathit{\gamma}Z-\mathrm{1}}{{\mathit{\tau}}_{\mathrm{S}}-\mathit{\gamma}Z}\right]}),\end{array}$$

where *H*_{1} and *H*_{2} are nondimensional constants:

$$}{\displaystyle}{H}_{\mathrm{1}}={\displaystyle \frac{{C}_{\mathrm{D}}U{c}_{\mathrm{P}}{P}_{\mathrm{0}}}{\mathit{\sigma}{T}_{\mathrm{R}}^{\mathrm{4}}{R}_{\mathrm{A}}}}\phantom{\rule{1em}{0ex}}\text{and}\phantom{\rule{1em}{0ex}}{H}_{\mathrm{2}}={\displaystyle \frac{{C}_{\mathrm{D}}U{L}_{\mathrm{v}}{P}^{\text{sat}}\left({T}_{\mathrm{R}}\right)}{\mathit{\sigma}{T}_{\mathrm{R}}^{\mathrm{5}}{R}_{\mathrm{W}}}}.$$

In Eqs. (4) and (5),
at equilibrium (i.e. $\frac{\mathrm{d}\cdot}{\mathrm{d}\widehat{s}}=\mathrm{0}$),
the state variable *i*_{A} is easily eliminated,
leaving a single equation with a single state variable *τ*_{S},

$$\begin{array}{ll}{\displaystyle}\mathrm{0}=& {\displaystyle}(\mathrm{1}-\mathit{\alpha}({\mathit{\tau}}_{\mathrm{S}}\left)\right)(\mathrm{1}-{\mathit{\xi}}_{\mathrm{A}}-{\mathit{\xi}}_{\text{Cl}})q+{f}_{\mathrm{O}}+\mathit{\beta}{f}_{\mathrm{A}}\\ \text{(11)}& {\displaystyle}& {\displaystyle}-{f}_{\mathrm{C}}(\mathrm{1}-\mathit{\beta})+\mathit{\beta}q{\mathit{\xi}}_{\mathrm{A}}-{\mathit{\tau}}_{\mathrm{S}}^{\mathrm{4}}(\mathrm{1}-\mathit{\beta}\mathit{\eta}({\mathit{\tau}}_{\mathrm{S}}\left)\right).\end{array}$$

In this subsection, we outline the proof that the cusp bifurcation, which was proven to exist in the Paleoclimate EBM (Kypke and Langford, 2020), in fact persists in this Anthropocene EBM Eqs. (4) and (5). Therefore, the conclusions of that paper carry over to this paper. Readers not interested in these mathematical details may skip this subsection.

The right-hand sides of Eqs. (4) and (5) can be represented by a single vector function $F:{\mathbb{R}}^{\mathrm{2}}\times {\mathbb{R}}^{\mathrm{4}}\to {\mathbb{R}}^{\mathrm{2}}$. Then an equilibrium point $({\overline{\mathit{\tau}}}_{\mathrm{S}},{\overline{i}}_{\mathrm{A}})$ of Eqs. (4) and (5), at which $\frac{\mathrm{d}{\mathit{\tau}}_{\mathrm{S}}}{\mathrm{d}t}=\frac{\mathrm{d}{i}_{\mathrm{A}}}{\mathrm{d}t}=\mathrm{0}$, is a solution of

$$\begin{array}{}\text{(12)}& F({\overline{\mathit{\tau}}}_{\mathrm{S}},{\overline{i}}_{\mathrm{A}};\mathit{\rho})=\mathrm{0},\end{array}$$

where *ρ* represents four physical parameters that may be varied in the model,

$$\begin{array}{}\text{(13)}& \mathit{\rho}\equiv \mathit{\{}\mathit{\mu},\mathit{\delta},{F}_{\mathrm{O}},\mathit{\omega}\mathit{\}}.\end{array}$$

See Table 1 for definitions of these parameters. Since the codimension of the cusp
bifurcation is only two, there is some redundancy in the choice of these four parameters. For
application to future climates, the parameter pair (*F*_{O},*μ*) is of primary interest.
Equilibrium points $({\overline{\mathit{\tau}}}_{\mathrm{S}},{\overline{i}}_{\mathrm{A}})$ satisfying Eq. (12)
have been computed in Kypke (2019). Having computed the equilibrium point
$({\overline{\mathit{\tau}}}_{\mathrm{S}},{\overline{i}}_{\mathrm{A}})$ satisfying Eq. (12), the system may be
translated to the origin $(x,y)=(\mathrm{0},\mathrm{0})$, in new state variables defined by

$$}{\displaystyle}(x,y)\equiv ({\mathit{\tau}}_{\mathrm{S}}-{\overline{\mathit{\tau}}}_{\mathrm{S}},{i}_{\mathrm{A}}-{\overline{i}}_{\mathrm{A}}),$$

$$\begin{array}{ll}{\displaystyle \frac{\mathrm{d}x}{\mathrm{d}s}}=& {\displaystyle}\phantom{\rule{0.125em}{0ex}}(\mathrm{1}-\mathit{\alpha})(\mathrm{1}-{\mathit{\xi}}_{\mathrm{A}}-{\mathit{\xi}}_{\text{Cl}})q+{f}_{\mathrm{O}}-{f}_{\mathrm{C}}\\ {\displaystyle}& {\displaystyle}\phantom{\rule{0.125em}{0ex}}+\mathit{\beta}(y+{\overline{i}}_{\mathrm{A}})-(x+{\overline{\mathit{\tau}}}_{\mathrm{S}}{)}^{\mathrm{4}},\\ \text{(14)}& {\displaystyle \frac{\mathrm{d}y}{\mathrm{d}s}}=& {\displaystyle}\phantom{\rule{0.125em}{0ex}}\mathit{\chi}\left[{f}_{\mathrm{A}}+{f}_{\mathrm{C}}+q{\mathit{\xi}}_{\mathrm{A}}+\mathit{\eta}(x+{\overline{\mathit{\tau}}}_{\mathrm{S}}{)}^{\mathrm{4}}-(y+{\overline{i}}_{\mathrm{A}})\right].\end{array}$$

For Eq. (14) to have a steady-state bifurcation at the equilibrium point
(0,0), the Jacobian **J** of *F* in Eq. (12) must have a zero eigenvalue, *λ*_{1}=0, at
that point. (A Hopf bifurcation is not possible in this system.) For stability, the second
eigenvalue satisfies *λ*_{2}<0. The corresponding eigenvectors, *e*_{1},*e*_{2},
form an *eigenbasis*. A linear transformation takes (*x*,*y*) coordinates to
*eigenbasis coordinates* (*u*,*v*), where $(x,y)=\mathbf{T}(u,v)$, and the columns of **T** are the
normalized eigenvectors *e*_{1},*e*_{2} in Eq. (17), below. Then in
(*u*,*v*) coordinates, the 2-D system Eq. (14) becomes

$$\begin{array}{ll}{\displaystyle \frac{\mathrm{d}u}{\mathrm{d}s}}=& {\displaystyle}\phantom{\rule{0.125em}{0ex}}{\displaystyle \frac{\mathrm{1}}{\mathit{\varphi}}}\left[\right(\mathrm{1}-\mathit{\alpha}\left)\right(\mathrm{1}-{\mathit{\xi}}_{\mathrm{A}}-{\mathit{\xi}}_{\text{Cl}})q+{f}_{\mathrm{O}}+\mathit{\beta}{f}_{\mathrm{A}}\\ \text{(15)}& {\displaystyle}& {\displaystyle}-(\mathrm{1}-\mathit{\beta}){f}_{\mathrm{C}}+\mathit{\beta}{\mathit{\xi}}_{\mathrm{A}}\phantom{\rule{0.25em}{0ex}}q+(\mathrm{1}-\mathit{\beta}\mathit{\eta})(u-kv+{\overline{\mathit{\tau}}}_{\mathrm{S}}{)}^{\mathrm{4}}]{\displaystyle \frac{\mathrm{d}v}{\mathrm{d}s}}=& {\displaystyle \frac{\mathrm{1}}{\mathit{\varphi}}}[-\mathrm{\ell}[(\mathrm{1}-\mathit{\alpha})(\mathrm{1}-{\mathit{\xi}}_{\mathrm{A}}-{\mathit{\xi}}_{\text{Cl}})q+{f}_{\mathrm{O}}]\\ {\displaystyle}& {\displaystyle}+(\mathrm{\ell}+\mathit{\chi}){f}_{\mathrm{C}}+\mathit{\chi}{f}_{\mathrm{A}}-(\mathrm{\ell}\mathit{\beta}+\mathit{\chi})[\mathrm{\ell}u+v+{\overline{i}}_{\mathrm{A}}]\\ {\displaystyle}& {\displaystyle}+(\mathrm{\ell}+\mathit{\chi}\mathit{\eta})(u-kv+{\overline{\mathit{\tau}}}_{\mathrm{S}}{)}^{\mathrm{4}}],\end{array}$$

where

$$\begin{array}{}\text{(16)}& \mathrm{\ell}={f}_{C\mathrm{0}}^{\prime}+\mathrm{4}{\mathit{\eta}}_{\mathrm{0}}{\overline{\mathit{\tau}}}_{\mathrm{S}}^{\mathrm{3}}+{\mathit{\eta}}_{\mathrm{0}}^{\prime}{\overline{\mathit{\tau}}}_{\mathrm{S}}^{\mathrm{4}},\phantom{\rule{1em}{0ex}}k={\displaystyle \frac{\mathit{\beta}}{\mathit{\chi}}},\phantom{\rule{1em}{0ex}}\mathit{\varphi}=\mathrm{1}+k\mathrm{\ell}\end{array}$$

and

$$\begin{array}{}\text{(17)}& {\mathit{e}}_{\mathrm{1}}=\left(\begin{array}{c}\mathrm{1}\\ \mathrm{\ell}\end{array}\right),\phantom{\rule{1em}{0ex}}{\mathit{e}}_{\mathrm{2}}=\left(\begin{array}{c}-k\\ \mathrm{1}\end{array}\right).\end{array}$$

Recall that the parameters *μ* and *δ*, representing CO_{2} concentration and water
vapour relative humidity, respectively, enter into Eq. (15) through the function
*η* defined in Eq. (9). For more details, see Kypke (2019), Kypke and Langford (2020), and Kuznetsov (2004).

If the *Centre Manifold Theorem* applies to Eq. (15), then there exists
a flow-invariant centre manifold, which is tangent to the *u* axis. The applicability of this theorem has
been verified, and the centre manifold has been computed for the present Anthropocene EBM as was
done for the paleoclimate EBM in Kypke (2019) and Kypke and Langford (2020). Details are omitted here. A phase portrait
for Eq. (15) in a neighbourhood of the cusp equilibrium point, together with a portion
of this centre manifold (in red), is shown in Fig. 2 in (*u*,*v*) coordinates. In this
figure, trajectories quickly collapse to the centre manifold around the equilibrium point (0,0),
as predicted by the *Centre Manifold Theorem*. The cusp equilibrium manifold for
Eq. (15) in normal form is shown in Fig. 3. Here *ζ*_{1} and *ζ*_{2} are
the normal form unfolding parameters for the cusp bifurcation. Note the coexistence of three
equilibrium points with different values of *u* (two stable and the middle one unstable) inside the
cusp-shaped region, but only one equilibrium point (stable) outside of that region.

3 Anthropocene climate forecasts

Back to toptop
In this section, the EBM of Sect. 2 is applied to the present and future climates of the
Earth to investigate the possibility of climate bifurcations (or tipping points) in the
Anthropocene. The principal parameters chosen to be explored are carbon dioxide concentration
*μ*, ocean heat transport *F*_{O}, and relative humidity *δ*. The EBM is adapted
locally to three separate regions, namely the Arctic, Antarctic, and the tropics, in
Sect. 3.2, 3.3, and 3.4, respectively.

Carbon dioxide production due to human activities has been well documented as a driver of climate
change in the Anthropocene. Projections of future atmospheric CO_{2} levels are available
under various future scenarios; we follow the Representative Concentration Pathways of
IPCC (2013), which is described in Sect. 3.1. Ocean heat transport is a difficult quantity to
predict, as many different factors influence the transport of heat to various regions of the world
via the oceans. Changes in temperature can change ocean heat transport which in turn affects local
temperatures. This is ocean heat transport feedback, which is explored in
Sect. 3.2.2. Similarly, the role of atmospheric heat transport feedback is
investigated briefly in Sect. 3.2.2. In addition, water vapour is a powerful
greenhouse gas, with a positive feedback effect that is investigated in
Sect. 3.2.3.

In Sect. 3.5, a globally averaged model is considered, mainly for the purpose of determining the global equilibrium climate sensitivity (ECS) of this EBM, for comparison with the ECS of other models as reported in IPCC (2013) and Priostosescu and Huybers (2017); see Sect. 3.6.

The IPCC has standardized four Representative Concentration Pathways (RCPs), which are used for
projections of future carbon dioxide concentrations; see van Vuuren et al. (2011) and Box TS.6 in
IPCC (2013). These RCPs are scenarios for differing levels of anthropogenic forcings on the
climate of the Earth and represent differing global societal and political “storylines”. The
scenarios are named RCP 8.5, RCP 6.0, RCP 4.5, and RCP 2.6, after their respective peak radiative
forcing increases in the 21st century. That is, in the year 2100, RCP 8.5 will reach its maximum
radiative forcing due to anthropogenic emissions of +8.5 W m^{−2} relative to the year
1750. This scenario is understood as one where emissions continue to rise and are not mitigated in
any way. RCP 6.0 corresponds to +6.0 W m^{−2} and RCP 4.5 corresponds to
+4.5 W m^{−2} relative to 1750. These are stabilization scenarios, where greenhouse gas
emissions level off to a target amount by the end of the century. Finally, RCP 2.6 corresponds to
+2.6 W m^{−2} in 2100, relative to 1750. This is a mitigation scenario, where strong
steps are taken to eliminate the increase and eventually reduce anthropogenic greenhouse gas
emissions. Figure 4 shows the carbon dioxide concentrations projected to the year 2500
in the IPCC scenarios for the four RCPs. The carbon dioxide increase is relatively moderate for
RCPs 2.6 and 4.5 and even decreasing eventually for RCP 2.6. The increase for RCP 6.0 is larger, and
it is drastic for RCP 8.5.

These RCPs represent hypothetical forcings due to human activity up to the year 2100. Beyond 2100,
they assume a “constant composition commitment”, where emissions are kept constant, which serves
to stabilize the scenarios beyond 2100 (IPCC, 2013). Of course, emissions could continue to
increase or be greatly reduced (“zero emissions commitment”) at some point in the future.
However, the constant emission commitment dataset provided in IPCC (2013), and shown in
Fig. 4, is what is utilized in this work. In the following sections we enter the
CO_{2} concentrations *μ* shown in Fig. 4, one at a time along each of the four
IPCC RCPs, into the versions of our EBM for the Arctic. Antarctic, etc., and we then let the EBM
climate evolve quasi-statically along each CO_{2} pathway.

Figure 5a is the manifold of equilibrium states for the EBM parameterized by
(*F*_{O},*μ*) for the case of an Arctic climate under Anthropocene conditions.
Figure 5b is the projection of this manifold onto the parameter plane, showing the fold
bifurcation lines as boundaries between coloured regions. Parameter values are as in
Table 1, with *F*_{A}=104 W m^{−2} and *δ*=0.6 taken as
the nominal values for the modern Arctic throughout this section, except in Fig. 8.
These figures were computed as in Kypke and Langford (2020). The cusp point, seen in Fig. 3, still
exists but is not visible in Fig. 5, because it is outside of the range of parameters
included in the figure.

In Fig. 5a, today's climate state lies on the lower (blue) portion of the equilibrium manifold, as shown by the red dot. The upper (yellow) portion represents a coexisting warm equable climate state, similar to the climate of Earth in the Pliocene and earlier. The intermediate (green) portion represents an unstable (and unobservable) climate state.

Similarly, in Fig. 5b, the blue area represents unique cold stable states, yellow
represents unique warm stable states, and the green region is the overlap region between the two
fold bifurcations, where both warm and cold states coexist. Hence, on moving in the
(*F*_{O},*μ*) parameter space starting from the blue region, crossing the green region,
and into the yellow region, there would be no observable change in climate on crossing from blue to
green; however, crossing the boundary from green to yellow would cause a catastrophic jump from cold
to warm stable climate states. Alternatively, if the (*F*_{O},*μ*) parameter values moved
from the yellow, through the green, into the blue region in Fig. 5b, there would be no
abrupt change in climate state on crossing the yellow–green boundary, but a sudden transition to
a cold state would occur at the green–blue boundary. This behaviour is called *hysteresis*.

The paramount question considered in this paper can now be phrased as follows. Can a bifurcation
leading to a warmer and more equable climate state be expected in the EBM if it is allowed to
evolve along one of the four RCPs in Fig. 4? In Fig. 5b, this bifurcation
would correspond to crossing the line of fold bifurcations separating the green and yellow regions
on increasing *μ* and possibly *F*_{O}.

Figure 6 shows the increase in surface temperature in the Arctic region, using
historical data from the year 1900 to the present, and then the EBM forecasts up to the years 2100
and 2300, holding *μ* to each of the four RCPs in Fig. 4, and with constants
${F}_{\mathrm{O}}=\mathrm{10}\phantom{\rule{0.125em}{0ex}}\mathrm{W}\phantom{\rule{0.125em}{0ex}}{\mathrm{m}}^{-\mathrm{2}}$, *F*_{A} = 104 W m^{−2}, and
*δ*=0.6. The temperature change shown is relative to the Arctic temperature of the EBM in
the year 2000, which was −28.90 ^{∘}C (*τ*_{S}=0.8942).

Figure 6a may be compared to the results in Figs. AI.8 and AI.9 in IPCC (2013).
Those figures use the RCPs of Fig. 4 and an ensemble of climate models to forecast
surface temperature changes for the Arctic up to the year 2100 for land and sea separately. IPCC
Fig. AI.8 displays the winter months of December–February and AI.9 is for June–August.
Figure 6 of this paper does not distinguish surface covering, so it is given as an annual average
value. Figure 6a shows predictions of the EBM to the year 2100, which is the same
time frame as for the IPCC GCM projections. Both use the RCP scenarios to determine hypothetical
CO_{2} concentrations (*μ*) by year, up to year 2100. Then both use these CO_{2}
concentrations as input to the respective models (EBM or GCM) to determine predicted climate changes
for the same period. They are in good agreement if a weighted average of the sea and land
temperature changes are considered and if the winter months are more representative of an annual
value for the Arctic climate.

Supported by the agreement until year 2100 between Fig. 6a and the IPCC reported values
of temperature change on the RCP paths, the EBM forecast was then extended to the year 2300 (see
Fig. 6b), which uses the same parameter values as Fig. 6a. It exhibits
a saddle-node bifurcation for RCP 8.5 near the year 2170. Following the disappearance of the cooler
equilibrium state in this saddle-node bifurcation, bifurcation theory tells us that there exists
a neighbourhood in the state space, where the saddle node once existed, inside of which trajectories
move slowly through a so-called “ghost equilibrium” that is a remnant of the saddle node. (The
transit time has inverse square-root dependence on the unfolding parameter in that neighbourhood.)
Outside of that neighbourhood, trajectories evolve with velocity determined by *c*_{S} and
*c*_{A} in Eqs. (1) and (2) to the upper stable equilibrium point.
(The vertical arrow in Fig. 6b represents that transition but is not an actual solution
of the dynamical system, and similarly for the vertical arrows in Figs. 7 and
10.) This bifurcation on RCP 8.5 results in a drastic increase in temperature: the mean
Arctic temperature jumps by $+\mathrm{26.5}\phantom{\rule{0.125em}{0ex}}{}^{\circ}\mathrm{C}$ to the new value of
$+\mathrm{22.5}\phantom{\rule{0.125em}{0ex}}{}^{\circ}\mathrm{C}$. This is warmer than the mean annual Arctic temperature in the Pliocene
but is consistent with what existed in the Eocene (Willard et al., 2019). Because of simplifications made in
this EBM, these numbers should not be taken literally as quantitatively accurate forecasts; however,
the qualitative existence of a dramatic increase in temperature due to a bifurcation must be taken
seriously. The topological methods employed in this work ensure that bifurcation in this model is
a mathematically persistent feature that will be manifest in all “nearby” models; see
Kypke and Langford (2020).

The other three RCPs show no such jump in Fig. 6, and indeed all three stay well below freezing. However, it must be borne in mind that the IPCC assumptions (used here) have all four RCPs levelling off to a target value by the end of this century; see Fig. 4. That may be overly optimistic.

In Fig. 6, the only forcing parameter that was changing was the CO_{2}
concentration (assumed due to anthropogenic forcing). Now we incorporate changes to ocean and
atmosphere heat transport, which may amplify the effects of increasing CO_{2} alone.

There is evidence that ocean heat transport, *F*_{O}, into the Arctic is increasing. For
example, Koenigk and Brodeau (2014) project ocean heat transport above 70 ^{∘}N to increase from
0.15 PW in 1860 to 0.2 and 0.3 PW in 2100, for RCP 2.6 and 8.5, respectively. At
the same time, they find in their model that atmospheric heat transport decreases slightly, from
1.65 PW in 1850 to 1.6 PW (1.5 PW) for RCP 2.6 (RCP 8.5). These authors
found that, in a stable climate state that ensures global energy conservation, *F*_{O} and
*F*_{A} tend to be out of phase; see, for example, the coupled climate model in Koenigk and Brodeau (2014).
However, Yang et al. (2016) show that in a more realistic situation when the climate is perturbed by both
heat and freshwater fluxes, the changes in *F*_{O} and *F*_{A} may be in phase.
We assume the latter situation in this paper; see Fig. 7b.

In our model, climate forcings *F*_{O} and *F*_{A} are expressed as power per unit
area (W m^{−2}) determined as follows (see Table 2). First, the surface
area of the Arctic region is estimated. The Arctic is taken to be the surface of the Earth above
the 70th parallel; as such the surface area is

$$\begin{array}{}\text{(18)}& \text{Arctic Surface Area}=\mathrm{2}\mathit{\pi}{R}^{\mathrm{2}}(\mathrm{1}-\mathrm{cos}\mathit{\theta}),\end{array}$$

where *R* is the radius of the Earth, 6371 km, and *θ* is 90^{∘} minus the
latitude. Hence, the surface area of the Earth above 70 ^{∘}N is approximately
1.538×10^{13} m^{2}. This leads to atmospheric and ocean heat fluxes into the Arctic
as summarized in Table 2. Because the change in *F*_{A} is small relative
to the changes in *μ* and *F*_{O}, *F*_{A} is kept constant at an intermediate
value of 104 W m^{−2} in Figs. 5 to 7a.

Figure 7a shows the change in Arctic surface temperature (relative to the year 2000
temperature of $-\mathrm{28.90}\phantom{\rule{0.125em}{0ex}}{}^{\circ}\mathrm{C}$) for the four RCPs with the ocean heat flux
*F*_{O} increasing linearly on each RCP until the year 2100, as projected in Koenigk and Brodeau (2014),
using their data for *F*_{O} in Table 2 but with constant
*F*_{A}=104. Beyond the year 2100, until 2300, the ocean heat transport, *F*_{O},
, is held constant at its 2100 value. In this scenario, the onset of the jump (via a fold
bifurcation) to a warm equable climate is advanced dramatically. The bifurcation for RCP 8.5 occurs
in the year 2118, more than 40 years earlier than was the case with a constant *F*_{O} in
Fig. 6. The temperatures before and after the jump in 2118 between two stable states
are −4.6 and +24.7 ^{∘}C, respectively, which is a sudden increase of
29.3 ^{∘}C. The other three RCPs remain below the freezing temperature.

Figure 7b shows the same scenario as in 7a, with increasing *μ* and *F*_{O}
but with the atmospheric heat transport, *F*_{A}, also increasing linearly from
104 W m^{−2} in the year 2000 to 129 W m^{−2} in 2100, being constant thereafter.
In this case, the saddle-node bifurcation occurs even earlier for RCP 8.5, and a new bifurcation
appears for RCP 6.0. Both of these changes make mitigation more challenging.

Overall, water vapour is known to be the most powerful greenhouse gas in the atmosphere
(Dai, 2006; Pierrehumbert, 2010; IPCC, 2013). Warming of the surface causes evaporation of more water vapour, which
causes further greenhouse warming and a further rise in surface temperature. Thus, water vapour
amplifies the warming due to other causes. This is called water vapour feedback. Empirical studies
such as Dai (2006) show that the increase in surface specific humidity *H* with surface
temperature *T* is close to that predicted by the Clausius–Clapeyron equation as in
Eq. (6) (outside of desert regions). The relative humidity, RH or *δ*, changes
little with surface temperature, even as the specific humidity *H* increases (Serreze et al., 2012). For
paleoclimates, Jahren and Sternberg (2003) have described an Eocene Arctic rain forest with RH estimated to be
*δ*=0.67. Modern data, from a variety of sources, suggest similar values of Arctic RH. For
example, Shupe et al. (2011) report Arctic RH at the surface over 0.7 and atmospheric RH at 2.5 km
altitude near 0.6, with relatively small seasonal and spatial variations.

Therefore, in the EBM Eqs. (1) and (2), it is assumed that the greenhouse
warming effect of water vapour is determined by the Clausius–Clapeyron relation as in
Eq. (6) and the RH *δ* remains constant. We assumed a value of *δ*=0.60 for
the Arctic atmosphere in the previous section.

The Clausius–Clapeyron equation tells us that below the freezing temperature
(*τ*_{S}=1) the concentration of water vapour is very low and therefore it has very
little greenhouse effect. However, above freezing, if a source of water is available (e.g. oceans),
then the concentration of water vapour and its greenhouse warming effect increase rapidly. This is
shown clearly in Fig. 8, where the three curves show different levels of relative
humidity, *δ*, but all assume that CO_{2} follows RCP 8.5. Here, the reference
temperatures in year 2000 are as follows: red curve $-\mathrm{28.899}\phantom{\rule{0.125em}{0ex}}{}^{\circ}\mathrm{C}$, green curve
$-\mathrm{29.418}\phantom{\rule{0.125em}{0ex}}{}^{\circ}\mathrm{C}$, and blue curve $-\mathrm{30.320}\phantom{\rule{0.125em}{0ex}}{}^{\circ}\mathrm{C}$. On each of the curves of
Fig. 8, the dashed portions with negative slope are unstable, while portions with
positive slope are asymptotically stable (in the sense of Liapunov). Bistability (the coexistence
of two stable solutions) occurs sooner when water vapour is present. The lower continuous blue
curve with *δ*=0 shows no thawing (*τ*_{S}<1) in this range.

This EBM for the Anthropocene Arctic predicts that a bifurcation will occur, leading to catastrophic
warming of the Arctic if CO_{2} emissions continue to increase along RCP 8.5 without
mitigation. This is true in the model even if ocean and atmosphere heat transport remain unchanged.
The amplifying effects of ocean and atmosphere heat transport can make this abrupt climate change
become even more dramatic and occur even earlier. Water vapour feedback further intensifies global
warming. However, the EBM predicts that RCPs with reduced CO_{2} emissions (due, for example,
to effective mitigation strategies if introduced soon enough) may avert the drastic consequences of
such a bifurcation.

Further work on Anthropocene Arctic climate modelling will include the effects of other positive
feedback mechanisms, e.g. the greenhouse effects of methane and CO_{2} that will be
released as the permafrost thaws in the Arctic and the Hadley cell feedback that may occur as
global circulation patterns change. With such additional amplification in the Arctic taken into
account, and no mitigation strategies in place, a saddle-node bifurcation resulting in a transition
to a warmer Arctic climate state may occur even earlier than predicted by the present model.

It has long been recognized that the climate of the Southern Hemisphere is generally colder than
that of the Northern Hemisphere for a number of reasons (Feulner et al., 2013). In particular the Antarctic
is colder than the Arctic. The Antarctic climate is affected by the Antarctic Circumpolar Current
(ACC), which flows freely west to east around Antarctica in the Southern Ocean, unimpeded by
continental barriers. The ACC blocks the poleward heat transport from the warm oceans of the
Southern Hemisphere (Hartmann, 2016). Hence, *F*_{O} is restricted to be below
20 W m^{−2} in the Antarctic EBM. Additionally, the cold surface albedo
*α*_{C}=0.8 is greater for the Antarctic than the Arctic, because the snow and ice
that covers the continent is more pure than that in the Arctic. Cloud albedo is reduced, with
a value of 7 % as opposed to the value of 12.12 % for the Arctic (Pirazzini, 2004). Atmospheric
heat transport is *F*_{A}=97 W m^{−2}, as determined in Zhang and Rossow (1997). Finally,
the Antarctic region is much drier than the Arctic; hence, a relative humidity of *δ*=0.4 is
used.

For the Antarctic, Fig. 9a is the equilibrium manifold for the energy balance model
parameterized by (*F*_{O},*μ*), and Fig. 9b is the projection of the fold
bifurcations onto the parameter plane. In Fig. 9b, the yellow area represents a warm
stable climate, the blue area a cold stable climate, and the green area represents the overlap
region (between the two fold bifurcations) where bistability exists. It can be seen in
Fig. 9 that a bifurcation from a cold state to a warm state in the EBM cannot occur for
an ocean heat transport value of less than *F*_{O}=12 W m^{−2} and a carbon
dioxide concentration less than *μ*=3000 ppm.

Figure 10a shows the temperature change, following each of the RCP curves in the
Antarctic, extended to the year 2300. The reference temperature is $-\mathrm{32.78}\phantom{\rule{0.125em}{0ex}}{}^{\circ}\mathrm{C}$ for
the year 2000, and the value of ocean heat transport into the Antarctic is assumed to be
*F*_{O}=14 W m^{−2}, an annual mean for the sea-ice zone (approximately
70^{∘}S) of the Antarctic, as determined in Wu et al. (2001). This scenario does not exhibit
a bifurcation from a cold climate state to a warm state. This suggests that for the Antarctic,
a change in *μ* alone is not sufficient to cause a hysteresis loop to exist, between two
coexisting stable states, in the context of modern and near-future carbon dioxide concentrations.

Figure 10b presents the Antarctic model for values of *F*_{O} that increase with
time as *μ* increases. The value of *F*_{O} is kept constant at 5 W m^{−2}
until the year 2000, after which time it increases linearly up to the year 2100, where it has
a value of *F*_{O}=20 W m^{−2}, after which it is held constant again. This
increase might represent an increase in sea levels, caused by thawing of the Arctic, loss of the
Antarctic ice shelves, and subsequent increase in ocean heat transport (Koenigk and Brodeau, 2014). The first
thing to notice is that a hysteresis loop now exists. With increasing CO_{2} on RCP 8.5, there
is a jump between stable states, from $\mathrm{\Delta}T=+\mathrm{29.1}\phantom{\rule{0.125em}{0ex}}{}^{\circ}\mathrm{C}$ to
$\mathrm{\Delta}T=+\mathrm{56.5}\phantom{\rule{0.125em}{0ex}}{}^{\circ}\mathrm{C}$, occurring in year 2225, where Δ*T* is the change in
surface temperature relative to the year 2000 temperature of $-\mathrm{32.78}\phantom{\rule{0.125em}{0ex}}{}^{\circ}\mathrm{C}$. The
above-freezing average temperatures on the upper warm branch of equilibrium states are consistent
with estimates of Antarctic temperatures in the Eocene (Passchier et al., 2013). Such a transition would imply
melting of the Antarctic ice cap and a drastic rise in sea levels around the world. The return
bifurcation from warm to cold (with time reversed) is visible in Fig. 10b. The
“cold-to-warm” transition occurs at a later time in the Antarctic than in the Arctic and at
a higher temperature. The differences between the Antarctic and Arctic bifurcations in the EBM are
due to the differences in ocean heat transport and ice albedos. The difference could be larger if
other factors are taken into account, e.g. the Antarctic has thicker ice and hence more
heat is required to melt enough ice to cause a change in albedo. This could be represented with
a larger value of *ω* in the tanh switch function; then, a greater temperature change is
required for the function to switch from a cold albedo value to a warm one.

Next, the EBM is adapted to model the climate of the tropics by choosing parameter values that are
annual mean, zonally averaged values at the Equator. This gives insolation
*Q*=418.8 W m^{−2} and relative humidity *δ*=0.8. Heat transports
${F}_{\mathrm{A}}=-\mathrm{38}$ W m^{−2} and ${F}_{\mathrm{O}}=-\mathrm{39}$ W m^{−2} are both negative,
because heat is transported away from the Equator towards the poles (Hartmann, 2016). The shortwave
cloud cooling *ξ*_{R} (the albedo of the clouds) is also greater in the tropics, and the
surface has a lower albedo. The value *ξ*_{R}=22.35 *%* is determined in the appendix
from the global energy budgets of Trenberth et al. (2009) and Wild et al. (2013).

As the tropics have annual average temperatures well above the freezing point of water, ice–albedo
feedback is absent in the tropics and a bifurcation from a cold stable state to a warm state can not
occur under Anthropocene conditions. However, if forced to low *F*_{O}, *F*_{A}
values and very low carbon dioxide levels, the climate state known as “snowball Earth”
(Kaper and Engler, 2013; Pierrehumbert, 2010) is a possibility. That scenario is not explored in this paper, as it is not
relevant to the Anthropocene.

The large relative humidity of *δ*=0.8 in the tropics serves to mitigate the radiative
forcing of increasing CO_{2}. Water vapour is a much more effective greenhouse gas than carbon
dioxide (Pierrehumbert, 2010). The atmosphere of the tropics contains more water vapour (the product of
greater relative humidity and warmer temperatures). The total atmospheric longwave absorption
*η* is given in Eq. (9). In the EBM of the tropics, the water vapour content is so
high that *η*_{W} (and thus total absorptivity *η*) is almost “maxed out” at
*η*≈1. Hence the total absorptivity is dominated by the contribution due to water
vapour, and an increase in CO_{2} concentration has little additional greenhouse warming
effect.

Figure 11 reveals relatively low increases in temperature for the tropics compared to the
poles, as CO_{2} concentration increases along the RCPs. Both the absence of a bifurcation
and the mitigation due to a large existing water vapour greenhouse forcing taken together cause the
temperature increase relative to the year 2000 to be less than 2^{∘}C, in all four RCP
scenarios.

Changes in globally averaged temperature can be modelled more easily than changes in regional
temperatures due to the fact that, in a globally averaged equilibrium model, overall net horizontal
transport of energy, by the oceans and the atmosphere, are both zero. Thus, the two-layer EBM
Eqs. (4) and (5), globally averaged with *F*_{O}=0 and
*F*_{A}=0, are simplified as follows.

$$\begin{array}{}\text{(19a)}& {\displaystyle \frac{\mathrm{d}{\mathit{\tau}}_{\mathrm{S}}}{\mathrm{d}\widehat{s}}}& {\displaystyle}=(\mathrm{1}-\mathit{\alpha})(\mathrm{1}-{\mathit{\xi}}_{\mathrm{A}}-{\mathit{\xi}}_{\mathrm{R}})q-{f}_{\mathrm{C}}+\mathit{\beta}{i}_{\mathrm{A}}-{\mathit{\tau}}_{\mathrm{S}}^{\mathrm{4}},\text{(19b)}& {\displaystyle \frac{\mathrm{d}{i}_{\mathrm{A}}}{\mathrm{d}\widehat{s}}}& {\displaystyle}=\mathit{\chi}[{f}_{\mathrm{C}}+q{\mathit{\xi}}_{\mathrm{A}}+\mathit{\eta}{\mathit{\tau}}_{s}^{\mathrm{4}}-{i}_{\mathrm{A}}].\end{array}$$

Here *α* is as defined in Eq. (8), *η* is as in
Eq. (9), and *f*_{C} is as in Eq. (10).
Parameters *ξ*_{A} and *ξ*_{R} are as in Table 1 and
Sect. 2.1.

Figure 12 shows the change in globally averaged equilibrium surface temperature relative
to the year 2000 global average (*τ*_{S}=1.064, 17.59 ^{∘}C), as
determined by the EBM Eq. (19) to the year 2300. It is assumed that CO_{2}
evolves with time along each of the four RCPs defined in IPCC (2013) and displayed in
Fig. 4. The other parameters, assumed constant, are as follows. The global relative
humidity *δ* is fixed at a value of 0.74. This is determined from Dai (2006), where it lies
at the lower end of a range of averages. Surface albedo is highly variable regionally, so a global
average was calculated from Wild et al. (2013), much like the atmospheric shortwave absorption and the
cloud albedo. From Fig. 1 of Wild et al. (2013), of the global average solar radiation of
185 W m^{−2} that reaches the surface, a portion of 24 W m^{−2} is reflected. Thus,
the global average surface albedo is $\frac{\mathrm{24}}{\mathrm{185}}=\mathrm{0.13}=\mathrm{13}\phantom{\rule{0.125em}{0ex}}\mathit{\%}$. The values for cloud albedo
and atmospheric shortwave radiation are calculated as follows. The global average incident solar
radiation *Q* at the top of the atmosphere is 340 W m^{−2}, of which
$\frac{\mathrm{100}-\mathrm{24}}{\mathrm{340}}=\mathrm{0.2235}=\mathrm{22.35}\phantom{\rule{0.125em}{0ex}}\mathit{\%}$ is reflected by clouds and
$\frac{\mathrm{79}}{\mathrm{340}}=\mathrm{0.2324}=\mathrm{23.24}\phantom{\rule{0.125em}{0ex}}\mathit{\%}$ is absorbed by the atmosphere. The planetary boundary layer
(PBL) altitude is 600 m (Ganeshan and Wu, 2016). Finally, for the purposes of sensible and latent heat
transport (see appendix), the wind speed *U* is 5 m s^{−1} (Nugent et al., 2014) and the drag
coefficient is ${C}_{\mathrm{D}}=\mathrm{1.5}\times {\mathrm{10}}^{-\mathrm{3}}$.

*Equilibrium climate sensitivity* (ECS) is a useful and widely adopted tool used to estimate
the effects of anthropogenic forcing in a given climate model. The ECS of a model is defined as the
change in the globally averaged surface temperature, after equilibrium is obtained, in response to
a doubling of atmospheric CO_{2} levels (IPCC, 2013; Knutti et al., 2017; Priostosescu and Huybers, 2017). The starting carbon dioxide
concentration is that of the preindustrial climate, taken to be *μ*=270 ppm. The doubled value
is then *μ*=540 ppm. Since the Earth has not yet experienced a doubling of CO_{2}
concentration since the industrial revolution, these numbers cannot be verified. Calculation of the
global ECS for the EBM of this paper facilitates comparisons with other climate models as reported
in IPCC (2013).

Table 3 gives both the nondimensional *τ*_{S} and the degree Celsius
temperature values for the *μ*=270 climate, the *μ*=540 climate, and the temperature
difference. This difference is the ECS of the global EBM of this paper.

For the models used in IPCC (2013), ECS values lie within a *likely* range of 1.5 to
4.5 ^{∘}C. Values of less than 1 ^{∘}C are deemed to be *extremely unlikely*, and values greater than 6 ^{∘}C are *very unlikely* (IPCC, 2013). The
value of 4.343 ^{∘}C calculated for this global EBM lies just inside the
*likely* range, at the high end. Recent work gives evidence that statistical climate models
based on historical data tend to lie on the lower end of *likely* ECS values, with a range of
1.5 to 3 ^{∘}C, whereas nonlinear GCMs tend to have larger ECS values (Priostosescu and Huybers, 2017).
Therefore, as the global EBM presented in this paper is nonlinear and is based on physical rather
than statistical modelling, it may be expected to fall on the side of larger ECS values.

Local ECS values may be determined for each of the three regional models, the Arctic, Antarctic, and the tropics, as defined in Sects. 3.2 to 3.4. These values are given in
Table 4. In all cases, *F*_{O} values are kept constant at their minimal
values: 10 W m^{−2} for the Arctic, 14 W m^{−2} for the Antarctic, and
−39 W m^{−2}, for the tropics. The regional ECS values are high, 7.95 and 7.54
^{∘}C, respectively, for the Arctic and Antarctic, and low, 1.27
^{∘}C for the tropics. Although the Earth has not yet experienced a doubling of
CO_{2} concentrations since the industrial revolution, these ECS values are consistent with
observations to date.

4 Conclusions and future work

Back to toptop
The analysis of this paper shows that a cusp bifurcation can occur in an energy balance model (EBM),
which could lead to hysteresis behaviour and an abrupt warming of the Anthropocene climate, to
a climate state that is like nothing that has existed on Earth since the Pliocene. The model has
been constructed from fundamental nonlinear processes of atmospheric physics. This bifurcation is
most likely to occur in the Arctic climate. It would lead to catastrophic warming, if increases in
atmospheric CO_{2} continue on their current pathway. However, if the increase in atmospheric
CO_{2} is mitigated sufficiently, this bifurcation can still be avoided. Climate changes in
the Arctic, Antarctic, and tropics are compared. The globally averaged equilibrium climate
sensitivity (ECS) of the EBM is 4.34 ^{∘}C, which is at the high end of the likely
range reported in IPCC (2013).

Future work will strengthen the conclusions of this paper by extending this simple EBM to more comprehensive models, which still allow rigorous bifurcation analysis to be performed. In the first generalization, the two-layer model will be replaced with a column model, with the atmosphere extending continuously from the surface to the tropopause. The ICAO Standard Atmosphere assumption will be replaced with a Schwarzschild radiation model of the atmosphere (Pierrehumbert, 2010), which will determine the lapse rate from the solution of a two-point boundary value problem. This Schwarzschild column model will be used to study, in addition to the positive feedback processes of this paper, the amplifying effects of methane from permafrost feedback in the Arctic.

The next generalization will combine the ideas of these energy balance models with a 3-D Navier–Stokes–Boussinesq partial differential equation (PDE) model, representing the convectively driven atmosphere as a fluid in a rotating spherical shell, presented in Lewis and Langford (2008) and Langford and Lewis (2009). In that model, surface temperatures were prescribed as boundary conditions; however, with the assumption of an energy balance constraint, the meridional surface temperature gradient will be determined implicitly. Meridional heat transport also will be determined in this model and may be compared with the poleward heat transports in the EBM of this paper. The code to solve this PDE model for a cusp bifurcation has been written. Later, guided by these results, the climate bifurcations found in these analytical models will be sought in an open-source general circulation model. This hierarchy of models is expected to add credibility to the prediction, presented here, that the Earth's climate system is capable of exhibiting dramatic topological changes (bifurcations) in the Anthropocene, leading to a climate state that resembles the pre-Pliocene climate of Earth.

Appendix A: Determination of parameters in the Anthropocene EBM

Back to toptop
The determination of the physical parameters appearing in the EBM (Eqs. 4 and
5) is summarized here. In most cases, these parameters were determined in earlier
papers of the authors dealing with paleoclimates (Dortmans et al., 2019; Kypke and Langford, 2020). The focal point is the scaled,
two-dimensional Eqs. (4) and (5) in Sect. 2.1, with *α*
given by Eq. (8), *η* given by Eq. (9), and *f*_{C} given by
Eq. (10).

First consider the incoming solar radiation Q. A fraction *ξ*_{A} is absorbed by the
atmosphere and a fraction *ξ*_{R} is reflected by the clouds, as seen in
Fig. 1 and Eqs. (4) and (5). These fractions were
determined in Kypke (2019) and appendix A1 of Dortmans et al. (2019), using data for the global average
energy budget described in Trenberth et al. (2009), Wild et al. (2013); and Kim and Ramanathan (2012). The globally averaged values were
*ξ*_{A}=0.2324 and *ξ*_{R}=0.2235 as listed in Table 1.
For the polar regions, the albedo of clouds is less than elsewhere. Using data collected in the
Surface Heat Budget of the Arctic Ocean (SHEBA) program (Intrieri et al,, 2002; Shupe and Intrieri, 2003), the revised polar value of
*ξ*_{R}=0.1212 was determined in Kypke (2019).

Clouds in the atmosphere have a second main effect, the absorption of a fraction, *η*_{Cl},
of the longwave radiation outgoing from the surface of the Earth (Hartmann, 2016). This effect warms
the Earth's atmosphere. Through data in the SHEBA program, the longwave cloud forcing in the Arctic
was found to be 51 W m^{−2} in the paper of Shupe and Intrieri (2003). Using this SHEBA data,
Kypke (2019) determined the fraction *η*_{Cl}=0.255, as used in Eq. (9);
see Table 1.

The absorption coefficients *k*_{C} for carbon dioxide and *k*_{W} for water
vapour in the atmosphere (see Table 1) were calculated using an empirical
approach, based on the modern-day global energy budget (Trenberth et al., 2009; Wild et al., 2013). Figure 1 in
Trenberth et al. (2009) provides the global mean surface radiation as 396 W m^{−2}, along with an
atmospheric window of 40 W m^{−2}. This atmospheric window, $\frac{\mathrm{40}}{\mathrm{396}}\approx \mathrm{0.1}$,
is then equal to 1−*η*. Schmidt et al. (2010) provide percentage contributions of carbon dioxide and
water vapour and clouds in an all-sky scenario, based on simulations using modern climate conditions
from the year 1980. The calculated values for *η*_{C} and *η*_{W} are then
used to calculate the corresponding optical depths, *λ*_{C}, and *λ*_{W}
for the case of the modern atmosphere, and these are used to solve for *k*_{C} and
*k*_{W}, which then appear in the *G*_{C} and *G*_{W2} terms, respectively,
in Table 1 and Eq. (7).

The vertical transport of sensible and latent heat is a difficult and complicated process to model,
so many approximations are made to keep it within the scope of this work. For more details, see
Kypke (2019). The heat transports are modelled via *bulk aerodynamic exchange* formulae
describing fluxes between the surface and the atmosphere (Pierrehumbert, 2010; Hartmann, 2016). For the sensible heat
flux, *c*_{p} is the specific heat of the air being heated, and *T* is its temperature.
The bulk aerodynamic formula for sensible heat (SH) is

$$\begin{array}{}\text{(A1)}& \text{SH}={c}_{\mathrm{p}}\phantom{\rule{0.33em}{0ex}}\mathit{\rho}\phantom{\rule{0.33em}{0ex}}{C}_{\text{DS}}\phantom{\rule{0.33em}{0ex}}U({T}_{\mathrm{S}}-{T}_{\mathrm{A}}),\end{array}$$

where *C*_{DS} is the drag coefficient for temperature and *U* is the mean horizontal wind
velocity. The density of the atmosphere *ρ* is determined as a function of both surface
temperature *T*_{S} and altitude *Z* using the barometric formula, and a constant lapse
rate Γ (ICAO, 1993) is used to determine the temperature gradient.

In the case of latent heat (LH), the moisture content is represented by *L*_{v}*r*, where
*L*_{v} is the latent heat of vaporization of water and *r* is the *mass mixing ratio* of condensable air to dry air (Pierrehumbert, 2010). Then

$$\begin{array}{}\text{(A2)}& \text{LH}={L}_{\mathrm{v}}\mathit{\rho}{C}_{\text{DL}}U({r}_{\mathrm{S}}-{r}_{\mathrm{A}}),\end{array}$$

where *C*_{DL} is the drag coefficient for moisture. In the following, for simplicity, we
set

$$}{\displaystyle}{C}_{\text{DS}}={C}_{\text{DL}}\equiv {C}_{\mathrm{D}}.$$

The mass mixing ratio is equal to the saturation mixing ratio times the relative humidity. The
saturation mixing ratio depends on the saturation vapour pressure, which is a function of
temperature as given by the Clausius–Clapeyron Eq. (6). The sensible and latent heat
transports are combined into a single term, *F*_{C}, which replaces the *F*_{C}
term that was introduced in Dortmans et al. (2019). This term is defined here as a function of surface
temperature *T*_{S}.

$$\begin{array}{ll}{\displaystyle}{F}_{\mathrm{C}}\left({T}_{\mathrm{S}}\right)=& {\displaystyle \frac{{C}_{D}U}{({T}_{\mathrm{S}}-\mathrm{\Gamma}Z)}}[{\displaystyle \frac{{c}_{\mathrm{p}}{P}_{\mathrm{0}}\mathrm{\Gamma}Z}{{R}_{\mathrm{A}}}}+{\displaystyle \frac{{L}_{\mathrm{v}}{P}^{\text{sat}}\left({T}_{\mathrm{R}}\right)}{{R}_{\mathrm{w}}}}\\ \text{(A3)}& {\displaystyle}& {\displaystyle}({e}^{\left[{G}_{\text{W1}}\frac{{T}_{\mathrm{S}}-{T}_{\mathrm{R}}}{{T}_{\mathrm{S}}}\right]}-\mathit{\delta}{e}^{\left[{G}_{\text{W1}}\frac{{T}_{\mathrm{S}}-\mathrm{\Gamma}Z-{T}_{\mathrm{R}}}{{T}_{\mathrm{S}}-\mathrm{\Gamma}Z}\right]})]\end{array}$$

Here, *P*_{0} is the atmospheric pressure at the surface (*Z*=0) and *R*_{A} is the ideal gas
constant specific to dry air. This equation is scaled by $\frac{\mathrm{1}}{\mathit{\sigma}{T}_{\mathrm{R}}^{\mathrm{4}}}$ to
nondimensionalize it, creating ${f}_{\mathrm{C}}=\frac{{F}_{\mathrm{C}}}{\mathit{\sigma}{T}_{\mathrm{R}}^{\mathrm{4}}}$ in
Eq. (10), where *T*_{R}=273.15 K is the reference temperature. As
this *f*_{C} represents energy moving from the surface to the atmosphere, it is subtracted
from the surface Eq. (4) and added to the atmosphere Eq. (5).
A different model of *F*_{C}, used by the authors in Dortmans et al. (2019), was a simple functional
form calibrated to empirical data. The result was a relationship between *F*_{C} and
*T*_{S} that is quantitatively very similar to that given by Eq. (A3). In
Kypke and Langford (2020), *F*_{C} was set equal to zero for the paleoclimate Arctic and Antarctic
models for simplicity, since both SH and LH are very small for temperatures that are below
freezing.

Code and data availability

Back to toptop
Code and data availability.

The MATLAB files, used to solve the model equations and to produce the figures in this article, have been posted online at https://doi.org/10.5281/zenodo.3834344 (Kypke et al., 2020).

Author contributions

Back to toptop
Author contributions.

The original conception of this research project was due to WFL. Most of the mathematical analysis and computations were performed by KLK, starting with his MSc Thesis (Kypke, 2019). ARW co-supervised the thesis and provided valuable contributions and insights throughout the course of the work. All authors contributed to the writing of the article.

Competing interests

Back to toptop
Competing interests.

The authors declare that they have no conflict of interest.

Acknowledgements

Back to toptop
Acknowledgements.

The authors thank Marcus Garvie for his support of this work and thank the two NPG reviewers for suggestions and constructive criticisms that have improved the quality of this paper.

Financial support

Back to toptop
Financial support.

This research has been supported by the Natural Sciences and Engineering Research Council of Canada (grant no. RGPIN-2015-05090).

Review statement

Back to toptop
Review statement.

This paper was edited by Stefano Pierini and reviewed by Michel Crucifix and Peter Ditlevsen.

References

Back to toptop
Alley, R. B., Marotzke, J., Nordhaus, W. D., Overpeck, J. T., Peteet, D. M., Pielke Jr., A., Pierrehumbert, R. T., Rhines, P. B., Stocker, T. F., Talley, L. D., and Wallace, J. M.: Abrupt climate change, Science, 299, 2005–2010, 2003. a

Ashwin, P., Wieczorek, S., Vitolo, R., and Cox, P.: Tipping points in open systems: bifurcation, noise-induced and rate-dependent examples in the climate system, Philos. T. Roy. Soc. A, 370, 1166–1184, 2012. a

Barnosky, A.D., Hadly, E. A., Bascompte, J., Berlow, E. L., Brown, J. H., Fortelius, M., Getz, W. M., Harte, J., Hastings, A., Marquet, P. A., Martinez, N. D., Mooers, A., Roopnarine, P., Vermeij, G., Williams, J. W., Gillespie, R., Kitzes, J., Marshall, C., Matzke, N., Mindell, D. P., Revilla, E., and Smith, A. B.: Approaching a state shift in Earth's biosphere, Nature, 486, 52–58, 2012. a

Barron, E. J., Thompson, S. L., and Schneider, S. H.: An ice-free Cretaceous? Results from climate model simulations, Science, 212, 10–13, 1981.

Bathiany, S., Dijkstra, H., Crucifix, M., Dakos, V., Brovkin, V., Williamson, M. S., Lenton, T. M., and Scheffer, M.: Beyond bifurcation: using complex models to understand and predict abrupt climate change, Dyn. Stat. Clim. Syst., 1, 1–31, https://doi.org/10.1093/climsys/dzw004, 2016. a

Björk, G., Stranne, C., and Borenäs, K.: The sensitivity of the Arctic Ocean sea ice thickness and its dependence on the surface albedo parameterization, J. Climate, 26, 1355–1370, 2013. a

Brovkin, V., Claussen, M., Petoukhov, V., and Ganopolski, A.: On the stability of the atmosphere-vegetation system in the Sahara/Sahel region, J. Geophys. Res., 103, 31613–31624, 1998. a

Budyko, M. I.: The effects of solar radiation variations on the climate of the Earth, Tellus, 21, 611–619, 1968. a, b

Dai, A.: Recent climatology, variability and trends in global surface humidity, J. Climate, 19, 3589–3606, 2006. a, b, c

Dijkstra, H. A.: Nonlinear Climate Dynamics, Cambridge University Press, New York, 2013. a

Dijkstra, H. A.: Numerical bifurcation methods applied to climate models: analysis beyond simulation, Nonlin. Process. Geophys., 26, 359–369, 2019. a

Ditlevsen, P. D. and Johnsen, S. J.: Tipping points: Early warning and wishful thinking, Geophys. Res. Lett., 37, L19703, https://doi.org/10.1029/2010GL044486, 2010. a

Dortmans, B., Langford, W. F., and Willms, A. R.: A conceptual model for the Pliocene paradox, in: Recent Advances in Mathematical and Statistical Methods, Springer Proceedings in Mathematics and Statistics vol. 259, edited by: Kilgour, D. M., Kunze, H., Makarov, R., Melnik, R., Wang, X., IV AMMCS International Conference, Waterloo, Canada, 20–25 August 2017, Springer Nature Switzerland AG, Gewerbestrasse 11, 6330 Cham, Switzerland, pp. 339–349, 2019. a

Dortmans, B., Langford, W. F., and Willms, A. R.: An energy balance model for paleoclimate transitions, Clim. Past, 15, 493–520, https://doi.org/10.5194/cp-15-493-2019, 2019. a, b, c, d, e, f, g, h, i, j, k

Drijfhout, S., Bathiany, S., Beaulieu, C., Brovkin, V., Claussend, M., Huntingford, C., Scheffer, M., Giovanni Sgubin, G., and Swingedouw, D.: Catalogue of abrupt shifts in Intergovernmental Panel on Climate Change climate models, P. Natl. Acad. Sci. USA, 112, E5777–E5786, https://doi.org/10.1073/pnas.1511451112, 2015. a

Feulner, G., Rahmstorf, S., Levermann, A., and Volkwardt, S.: On the origin of the surface air temperature difference between the hemispheres in Earth's present-day climate, J. Climate, 26, 7136–7150, 2013. a

Ganeshan, M. and Wu, D. L.: The open-ocean sensible heat flux and its significance for Arctic boundary layer mixing during early fall, Atmos. Chem. Phys., 16, 13173–13184, https://doi.org/10.5194/acp-16-13173-2016, 2016. a

Ghil, M.: Hilbert problems for the geosciences in the 21st century, Nonlin. Process. Geophys., 8, 211–222, 2001. a

Ghil, M. and Lucarini, V.: The Physics of Climate Variability and Climate Change, arXiv 1910.00583v2, 26 February 2020. a

Hartmann, D. L.: Global Physical Climatology, 2nd edition, Elsevier, Amsterdam, 2016. a, b, c, d, e

International Institute for Applied Systems Analysis (IIASA): The Representative Concentration Pathways data for Fig. 4, available at: http://www.iiasa.ac.at/web-apps/tnt/RcpDb/dsd?, last access: 2019. a

International Civil Aviation Organization: Manual of the ICAO Standard Atmosphere: extended to 80 kilometres, Third Edition, ICAO, 1000 Sherbrooke Street West Suite 400, Montreal, Quebec, Canada H3A 2R2, 1993. a, b, c

Intrieri, J. M., Fairall, C. W., Shupe, M. D., Persson, P. O. G., Andreas, E. L., Guest, P. S., and R. E. Moritz, R. E.: An annual cycle of Arctic surface cloud forcing at SHEBA, J. Geophys. Res., 107, 13–14, 2002. a

Intergovernmental Panel on Climate Change: Climate Change 2013: The Physical Science Basis, in: Contribution of Working Group I to the Fifth Assessment Report of the Intergovernmental Panel on Climate Change, edited by: Stocker, T. F., Qin, D., Plattner, G. K., Tignor, M., Allen, S. K., Boschung, J., Nauels, A., Xia, Y., Bex, V., Midgley, P. M., Cambridge Univ. Press, Cambridge, UK and New York, NY, USA, available at: http://www.ipcc.ch/ (last access: 11 March 2019), 2013. a, b, c, d, e, f, g, h, i, j, k, l, m, n

Jahren, A. H. and Sternberg, L. S. L.: Humidity estimate for the middle Eocene Arctic rain forest, Geology, 31, 463–466, 2003. a

Kaper, H. and Engler, H.: Mathematics and Climate, Society for Industrial and Applied Mathematics, Philadelphia, 2013. a, b

Kim, D. and Ramanathan, V.: Improved estimates and understanding of global albedo and atmospheric solar absorption, Geophys. Res. Lett., 39, L24704, https://doi.org/10.1029/2012GL053757, 2012. a

Knutti, R., Rugenstein, M. A. A., and Hegerl, G. C.: Beyond equilibrium climate sensitivity, Nat. Geosci., 10, 727–744, 2017. a

Koenigk, T. and Brodeau, L.: Ocean heat transport into the Arctic in the twentieth and twenty-first century in EC-Earth, Clim. Dynam., 42, 3101–3120, 2014. a, b, c, d, e, f

Kuznetsov, Y. A.: Elements of Applied Bifurcation Theory, 3rd edition, Applied Mathematical Sciences vol. 112, Springer-Verlag, New York, 2004. a

Kypke, K. L.: Topological Climate Change and Anthropogenic Forcing, M.Sc. Thesis, Department of Mathematics and Statistics, University of Guelph, Guelph, Ontario, Canada, 137 pp., 2019. a, b, c, d, e, f, g, h, i, j, k

Kypke, K. L. and Langford, W. F.: Topological climate change, Int. J. Bifurc. Chaos, 30, 2030005, https://doi.org/10.1142/S0218127420300050, 2020. a, b, c, d, e, f, g, h, i, j, k, l

Kypke, K. L., Langford, W. F., and Willms, A. R.: MATLAB files for figures for “Anthropocene Climate Bifurcation” in Nonlinear Processes in Geophysics, https://doi.org/10.5281/zenodo.3834344, 2020. a

Langford, W. F. and Lewis, G. M.: Poleward expansion of Hadley cells, Can. Appl. Math. Quart., 17, 105–119, 2009. a

Lenton, T. M.: Arctic climate tipping points, AMBIO, 41, 10–22, 2012. a

Lenton, T. M., Held, H., Kriegler, E., Hall, J. W., Lucht, W., Rahmstorf, S., and Schellnhuber, H. J.: Tipping elements in the Earth's climate system, P. Natl. Acad. Sci. USA, 105, 1786–1793, 2008. a

Lewis, G. M. and Langford, W. F.: Hysteresis in a rotating differentially heated spherical shell of Boussinesq fluid, SIAM J. Appl. Dyn. Syst., 7, 1421–1444, 2008. a

North, G. R. and Kim, K.-Y.: Energy Balance Climate Models, Wiley–VCH, Weinheim, Germany, 2017. a, b

North, G. R., Cahalan, R. F., and Coakley, J. A.: Energy balance climate models, Rev. Geophys. Space Phys., 19, 91–121, 1981. a, b

Nugent, A. D., Smith, R. B., and Minder, J. R.: Wind speed control of tropical orographic convection, J. Atmos. Sci., 71, 2695–2712, 2014. a

Passchier, S., Bohaty, S. M., Jiménez-Espejo, F., Pross, J., Röhl, U., van de Flierdt, T., Escutia, C., and Brinkhuis, H.: Early Eocene to middle Miocene cooling and aridification of East Antarctica, Geochem. Geophy. Geosy., 14, 1399–1410, 2013. a

Payne, A. E., Jansen, M. F., and Cronin, T. W.: Conceptual model analysis of the influence of temperature feedbacks on polar amplification, Geophys. Res. Lett., 42, 9561–9570, 2015. a, b

Pierrehumbert, R. T.: Principles of Planetary Climate, Cambridge University Press, Cambridge, UK, 2010. a, b, c, d, e, f, g

Pirazzini, R.: Surface albedo measurements over antarctic sites in summer, J. Geophys. Res., 109, D20118, https://doi.org/10.1029/2004JD004617, 2004. a

Pistone, K., Eisenman, I., and Ramanathan, V.: Observational determination of albedo decrease caused by vanishing Arctic sea ice, P. Natl. Acad. Sci. USA, 111, 3322–3326, 2014. a

Proistosescu, C. and Huybers, P. J.: Slow climate mode reconciles historical and model-based estimates of climate sensitivity, Sci. Adv., 3, e1602821, https://doi.org/10.1126/sciadv.1602821, 2017. a, b, c

Sagan, C. and Mullen, G.: Earth and Mars: Evolution of atmospheres and surface temperatures, Science, 177, 52–56, 1972. a

Schmidt, G. A., Ruedy, R. A., Miller, R. L., and Lacis, A. A.: Attribution of the present day total greenhouse effect, J. Geophys. Res., 115, D20106, https://doi.org/10.1029/2010JD014287, 2010. a

Seager, R. and Battisti, D. S.: Challenges to our understanding of the general circulation: abrupt climate change, in: The Global Circulation of the Atmosphere: Phenomena, Theory, Challenges, edited by: Schneider, T. and Sobel, A. S., Princeton University Press, Princeton, NJ, pp. 331–371, 2007. a

Sellers, W. D.: A global climate model based on the energy balance of the Earth–atmosphere system, J. Appl. Meteorol., 8, 392–400, 1969. a, b

Serreze, M. C., Barrett, A. P., and Stroeve, J.: Recent changes in tropospheric water vapor over the Arctic as assessed from radiosondes and atmospheric reanalyses, J. Geophys. Res., 117, D10104, https://doi.org/10.1029/2011JD017421, 2012. a

Shupe, M. D. and Intrieri, J. M.: Cloud radiative forcing of the Arctic surface: The influence of cloud properties, surface albedo, and solar zenith angle, J. Climate, 17, 616–628, 2003. a, b

Shupe, M. D., Walden, Von P., Eloranta, E., Uttal, T., Campbell, J. R., Starkweather, S. M., and Shiobara, M.: Clouds at Arctic atmospheric observatories. Part I: occurrence and macrophysical properties, J. Appl. Meteorol. Clim., 50, 626–644, 2011. a

Steffen, W., Rockström, J., Richardson, K., Lenton, T. M., Folke, C., Liverman, D., Summerhayes, C. P., Barnosky, A. D., Cornell, S. E., Crucifix, M., Donges, J. F., Fetzer, I., Lade, S. J., Scheffer, M., Winkelmann, R., and Schellnhuber, H. J.: Trajectories of the Earth system in the Anthropocene, P. Natl. Acad. Sci. USA, 115, 8252–8259, https://doi.org/10.1073/pnas.1810141115, 2018. a, b

Stranne, C., Jakobsson, M., and Björk, G.: Arctic Ocean perennial sea ice breakdown during the Early Holocene Insolation Maximum, Quart. Sci. Rev., 92, 123–132, 2014. a

Thorndike, A.: Multiple equilibria in a minimal climate model, Cold Reg. Sci. Technol., 76, 3–7, https://doi.org/10.1016/j.coldregions.2011.03.002, 2012. a

Trenberth, K. E., Fasullo, J. T., and Kiehl, J.: Earth's global energy budget, B. Am. Meteorol. Soc., 90, 311–323, 2009. a, b, c, d, e, f

Valdes, P.: Built for stability, Nat. Geosci., 4, 414–416, 2011. a

van Vuuren, D. P., Edmonds, J., Kainuma,M., Riahi, K., Thomson, A., Hibbard, K., Hurtt, G. C., Kram, T., Krey, V., Lamarque, J.-F., Masui, T., Meinshausen, M., Nakicenovic, N., Smith, S. J. and Rose, S. K.: The representative concentration pathways: an overview, Clim. Change, 109, 5–31, 2011. a

Wallace-Wells, D.: The Uninhabitable Earth: Life After Warming, Penguin Random House LLC, New York, 2019. a

Wild, M., Folini, D., Schär, C., Loeb, N., Dutton, E. G. and König-Langlo, G.: The global energy balance from a surface perspective, Clim. Dynam., 40, 3107–3134, 2013. a, b, c, d, e, f, g

Willard, D. A., Donders, T. H., Reichgelt, T., Greenwood, D. R., Sangiorgi, F., Peterse, F., Nierop, K. G. J., Frieling, J., Schouten, S., and Sluijs, A.: Arctic vegetation, temperature, and hydrology during Early Eocene transient global warming events, Global Planet. Change, 178, 139–152, 2019. a

Wu, X., Budd, W. F., Worby, A. P., and Allison, I.: Sensitivity of the Antarctic sea-ice distribution to oceanic heat flux in a coupled atmosphere-sea-ice model, Ann. Glaciol., 33, 577–584, 2001. a

Yang, H., Zhao, Y.-Y., and Liu, Z.-Y.: Understanding Bjerknes compensation in atmosphere and ocean transports using a coupled box model, J. Climate, 29, 2145–2160, 2016. a, b

Zeebe, R. E.: Where are you heading Earth?, Nat. Geosci. 4, 416–417, 2011. a

Zhang, Y.-C. and Rossow, W. B.: Estimating meridional energy transports by the atmospheric and oceanic circulations using boundary fluxes, J. Climate, 10, 2358–2373, 1997. a

Short summary

The climate of Earth is governed by nonlinear processes of geophysics. This paper presents energy balance models (EBMs) embracing these nonlinear processes which lead to positive feedback, amplifying the effects of anthropogenic forcing and leading to bifurcations. We define bifurcation as a change in the topological equivalence class of the system. We initiate a bifurcation analysis of EBMs of Anthropocene climate, which shows that a catastrophic climate change may occur in the next century.

The climate of Earth is governed by nonlinear processes of geophysics. This paper presents...

Nonlinear Processes in Geophysics

An interactive open-access journal of the European Geosciences Union