Nonlinear Earth System Dynamics Determine Biospheric Structure and Function: I—A Primer on How the Climate System Functions as a Heat Engine and Structures the Biosphere

Kittel, Timothy G. F.; Ferron, Kelly

doi:10.3390/cli14020038

Open AccessReview

Nonlinear Earth System Dynamics Determine Biospheric Structure and Function: I—A Primer on How the Climate System Functions as a Heat Engine and Structures the Biosphere

by

Timothy G. F. Kittel

^1,*

and

Kelly Ferron

²

¹

Institute of Arctic and Alpine Research, University of Colorado, Boulder, CO 80303, USA

²

National Geographic-Lindblad Expeditions, 96 Morton Street, 9th Floor, New York, NY 10014, USA

^*

Author to whom correspondence should be addressed.

Climate 2026, 14(2), 38; https://doi.org/10.3390/cli14020038 (registering DOI)

Submission received: 16 July 2024 / Revised: 13 November 2025 / Accepted: 23 November 2025 / Published: 1 February 2026

(This article belongs to the Special Issue Climate System Uncertainty and Biodiversity Conservation)

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Abstract

The Earth’s climate system exhibits nonlinear behavior driven by interactions among the atmosphere, oceans, cryosphere, land, and biosphere. These dynamics have given rise to relatively stable environments that shape the structure and function of the modern biosphere. This review is a primer for conservation practitioners and natural resource managers to develop a deep understanding of how the Earth System works. The key is to recognize that shifts in Earth System dynamics due to global climate change can destabilize the biosphere in unforeseen ways. The potential emergence of novel ecoregions must be a critical factor in adaptation planning for conservation and resource management. We review how thermodynamic constraints and global circulation dynamics determine the distribution of terrestrial and marine biomes. These dynamics stem from the Earth System functioning as a heat engine, transporting excess heat from low to high latitudes. We illustrate how biome climates are organized into climate regimes, with spatial and temporal characteristics linked to complex features of atmospheric and oceanic circulation. At centennial to millennial scales, these dynamics have created a stable envelope of natural variability in climate that has established a long-standing operating space for biota. However, this stability is becoming increasingly uncertain due to the growing positive energy imbalance in the Earth System primarily driven by anthropogenic greenhouse gas emissions. This forcing is leading to disruptive climatic change, putting the biosphere on a trajectory toward new transient states. Such global to regional climatic instability and biospheric restructuring introduce a high level of uncertainty in ecological futures, with major implications for natural resource management, biodiversity conservation strategies, and societal adaptation. We conclude by discussing frameworks for impact assessments and decision making under climate uncertainty.

Keywords:

climate change; biogeography; nonlinear dynamics; atmospheric circulation; ocean circulation; envelope of normal variability; historical range of variability; safe operating space; biodiversity conservation; natural resource management

1. Introduction

1.1. Key Questions

A fundamental question in biodiversity conservation and natural resource management is what magnitude of biogeographic changes under rapid climate change is in store for biota from local to global scales. At the planetary scale, the distribution of biomes is primarily linked to surface climate, such as air temperature and moisture regimes for terrestrial biomes and surface-layer temperature and wind regimes for marine biomes. Surface climate conditions are, in turn, controlled by global atmospheric and ocean circulation. These dynamics are ultimately driven by how the Earth System functions as a heat engine, transferring surplus heat energy from low latitudes to high latitudes.

To envision possible futures, we need to understand these dynamics. This leads to key questions about the climate system:

How do global climate dynamics shape regional climates and the distribution of biomes?
How might global atmospheric and ocean circulation changes induced by increasing greenhouse gases alter terrestrial and marine regional climates?
What do these changes mean for future biospheric structure and function and socioecological stability?

In the face of rapid climate change, these are questions that must be considered in conservation and resource management. Our goal in this paper is to address these questions by imparting a deep understanding of Earth System dynamics. We highlight key points in the Executive Summary (Box 1). This is a contribution to the Special Issue “Climate System Uncertainty and Biodiversity Conservation” in Climate (https://www.mdpi.com/journal/climate/special_issues/bio) and Earth (https://www.mdpi.com/journal/earth/special_issues/climate_biodiversity) (accessed on 12 December 2024).

Box 1. Executive Summary.

Understanding Earth System dynamics and how they determine the structure and function of the biosphere can guide conservation planning and resource management aimed at developing and implementing adaptation and mitigation strategies for climate change. In this summary, we highlight key points presented in this paper that contribute to this understanding. For each point, we identify corresponding sections that explain the hows and whys of these dynamics. A graphical summary is presented below (Box 1, Figure 1).

While highly nonlinear dynamics characterize the climate system, features of these dynamics have remained relatively stable at centennial to millennial timescales. This stability results from system feedbacks that constrain climate variability (Section 2.1). This constraint creates a stable ecological “safe operating space” within which the biosphere has evolved to operate. The stability of this space is threatened by the foreseeable shift of climate regimes to outside the range of natural climate variability. This, along with the poor prospect for climate stabilization, indicates that we may expect a broad-scale reorganization of ecosystems by the end of this century (Section 2.2 and Section 2.3).
Global circulation dynamics are driven by the Earth System functioning as a heat engine to transport excess heat from low to high latitudes (Section 3.1). This transport is accomplished by horizontal and vertical energy transfers within and among the atmosphere, oceans, and other climate system components (Section 3.2). Globally, the resulting atmospheric and ocean circulations are organized by latitude, longitude, altitude, ocean depth, and season.
The distribution of terrestrial, open ocean, and coastal biomes is tied to regional climates, whose spatial and temporal attributes are linked to complex features of global atmospheric and oceanic circulation (Section 4 and Section 5). These biome-organizing features include, for example, the Intertropical Convergence Zone and mid-latitude Jet Streams in the atmosphere (Section 4.3) and subtropical and subpolar gyres in the oceans (Section 5.1). Other key ocean circulation features include upwelling zones and global thermohaline (density-driven) circulation (Section 5.1 and Section 5.2).
The variety of terrestrial biomes in warm regions is greater than in colder ones. This variation is due to a thermodynamic relationship between air temperature and water vapor, where warm air can hold more moisture than cold air. Consequently, tropical, subtropical, and temperate zones have a wider range of arid to humid environments compared to colder regions (Section 4.2).
The four points above mean that complex climate dynamics determine the distribution and structure of terrestrial and marine biomes. This is through both global atmosphere and ocean circulation dynamics and thermodynamic constraints.
Driven primarily by elevated greenhouse gases, the Earth’s global energy budget is out of balance, retaining more energy than is released into space. Since the industrial era began, the accumulated effect of this imbalance has increased the heat energy content of climate system components and energy transfers between them (Section 3.1). Consistent with this is an observed intensification and shift in the location of atmospheric and ocean circulation patterns, which have biospheric consequences (Section 6). Such changes are currently or anticipated to be of a magnitude and character to push regional climates outside their historical ecologically safe operating space, leading to a reorganization of the biosphere with the possible emergence of novel systems.
Observed relationships between climatic means (e.g., monthly or annual mean temperatures and precipitation) and ecological processes are not expected to hold under a changing climate. This is to say, such empirical relationships are nonstationary over time (Section 4.1.1). This arises because climatic controls over key biotic processes operate at timescales other than captured by long-term monthly or annual means. Instead, determining factors are often event-level climate dynamics (i.e., weather events) or legacy dependent (e.g., constrained by previous years’ climate). These dynamics respond nonlinearly to changes in global circulation patterns. Consequently, ecological models based on mean climate–ecology relationships, such as climate envelope distribution models, have limited utility in assessing climate change impacts (Section 4.1.2).
Rapid disruption of climate and the biosphere threatens the provision of ecosystem services, with severe consequences for human socioeconomic systems (Section 7.3). Environmental instability additionally puts biodiversity at risk at genetic, population, community, and seascape/landscape scales. Uncertainty associated with these threats poses a challenge for biodiversity conservation, resource management, and adaptation of socioecological systems. Resource managers and conservation practitioners are well situated to implement vulnerability-based strategies due to their focus on site to regional scales, their detailed knowledge of species and ecosystems of concern, their involvement with local communities, and the need to act without complete information (Section 7.4).

Box 1. Figure 1. Graphical summary illustrating key linkages between Earth System dynamics and biospheric structure and function and disruption pathways under anthropogenic-altered forcing. [Credits: Left—image from CERES Radiation Balance, NASA Scientific Visualization Studio, https://svs.gsfc.nasa.gov/4794. Accessed 9 June 2023. Center—ocean circulation graphic from https://commons.wikimedia.org/w/index.php?curid=20706629. Accessed 3 November 2025. Right—image from NASA SeaWiFS Global Biosphere Composite 1997–1998. https://oceancolor.gsfc.nasa.gov/SeaWiFS/BACKGROUND/Gallery/biosphere.jpg. Accessed 3 November 2025. All public domain].

Box 1. Figure 1. Graphical summary illustrating key linkages between Earth System dynamics and biospheric structure and function and disruption pathways under anthropogenic-altered forcing. [Credits: Left—image from CERES Radiation Balance, NASA Scientific Visualization Studio, https://svs.gsfc.nasa.gov/4794. Accessed 9 June 2023. Center—ocean circulation graphic from https://commons.wikimedia.org/w/index.php?curid=20706629. Accessed 3 November 2025. Right—image from NASA SeaWiFS Global Biosphere Composite 1997–1998. https://oceancolor.gsfc.nasa.gov/SeaWiFS/BACKGROUND/Gallery/biosphere.jpg. Accessed 3 November 2025. All public domain].

1.2. Review Objectives

Earth System dynamics function to shape the biosphere across ecological, spatial, and temporal scales. This review endeavors to present an integrated understanding of these dynamics and how they are shifting as part of global climate change. Shifts in these dynamics act to destabilize biospheric structure and function.

Our review is presented here and in two companion papers. Each subsequent paper expands on details from the previous ones to further emphasize the biosphere’s connection to nonlinear climate dynamics. This first paper illustrates how these nonlinearities create a physical environment that determines the organization of the biosphere into biomes. The second paper explores (1) the Earth’s positive energy imbalance arising primarily from increasing levels of greenhouse gases [1], (2) how feedbacks in the Earth System result in complex climate dynamics, and (3) how novel “no-analog” ecosystems arise. The third paper examines complex linkages between climate and biospheric dynamics at the regional scale. In that paper, we tie changing climatic extremes and climate-forced disturbance regimes to regional ecological vulnerability.

While this review focuses on the dependence of biospheric structure on climate system dynamics, the biosphere also plays a strong role in shaping the physical and chemical climates of terrestrial and ocean systems through biogeophysical and biogeochemical feedbacks at geological and modern era scales [2,3,4,5,6,7,8,9]. We cover key biospheric feedbacks here and in companion papers, particularly where they play a vital role in climate system sensitivity to altered forcing [2,10,11].

This review serves as a primer on climate–biosphere dynamics for conservation practitioners, natural resource managers, and climate change impact policymakers, providing a window into linkages between climate and ecological domains. The objective, however, is not to review our understanding of the ecological consequences of rapid climate change. This has been addressed and continues to be explored by others (e.g., [12,13,14,15,16,17,18]). While parts of this review will be basic knowledge for climate scientists, it uncovers underlying climate dynamics for ecologists and resource managers. Conversely, other sections that are intuitive for ecologists may offer climate scientists new insights into the complexity of how biospheric dynamics are tied to climate.

1.3. Terms Defined

We refer to “climate change” here in a strict sense of current and future anthropogenic climate change, ongoing since the beginning of the industrial era, unless otherwise indicated. A broader sense of the term would encompass changes over longer timescales, including those in a paleoclimatic context, which we reference at times here and in companion papers. Anthropogenic forcings include not only emissions of greenhouse gases (GHGs: CO₂, CH₄, N₂O, etc.) but also altered biogeophysical and biogeochemical processes resulting from land use/land cover change (including altered continental runoff) and aerosol emissions (including aircraft and ship emissions affecting cloud condensation nuclei concentrations) [19,20].

By “historical” or “modern” conditions, we refer to the recent centuries leading up to the point when strong anthropogenic forcing began disrupting climate and the biosphere. Generally, this extends to the mid- or late 20th century, when such disruption is clear [13,18,19,21], although this marker is arbitrary or context dependent.

We use the term “nonlinear” in a systems context to describe behaviors and structures that emerge from complex interactions among system elements [22]. Interactions among Earth System components—the atmosphere, oceans, land, cryosphere, and biosphere—involve climatic, biogeophysical, and biogeochemical processes. These interactions generate positive and negative feedbacks that operate across multiple spatial and temporal scales. In the Earth System, these create complex structures and nonlinear dynamics. See the Glossary (Appendix A) to further explore nonlinearities and Earth System examples.

Other key terms are also defined in Appendix A. The first use of a term in the text from here on is given in italics.

2. Biosphere Stability and Earth System Dynamics

In this section, we examine how climatic and biospheric stability are tied to Earth System dynamics. We explore this in three ways:

Regional climates operate within a limited range of variability maintained by Earth System dynamics.
Strong forcing results in this range being exceeded, causing climatic and ecological reorganization and potentially giving rise to novel states.
Once this forcing abates, climatic and biospheric stabilization to these states will be slow and potentially irreversible on human timescales due to system feedbacks.

2.1. Stationary Envelope of Normal Variability

In its current state (or at least before strong anthropogenic forcing), the climate system is structurally highly nonlinear but relatively stable. Climate system dynamics involve a balance between stabilizing and amplifying feedbacks that, in the net, have maintained terrestrial and marine climates within certain limits [22,23]. Consequently, climate is variable but has been constrained in that variability at centennial to millennial timescales [24,25]. This has led to ecological and hydrologic stability tied to a stationary envelope of normal variability [24,26,27,28,29]. This variability defines a stable ecological “safe operating space” within which the biosphere has evolved [24,30]. Such adaptation has occurred through biota evolving with other dynamic elements of the Earth System via biogeochemical, biogeophysical, and biogeographical processes over multi-centennial to millennial timescales.

The past 2000-year record of extratropical Northern Hemisphere summer land temperatures illustrates a steady envelope of historical variability through the end of the 19th century (Figure 1a). This stability is relative to Late Pleistocene climates, which were highly variable, with decadal-scale sudden warming followed by millennial-scale cooling over northern latitudes [31]. The transition from these climates to the relatively stable climate of the Holocene started around 11,700 years ago [22,32]. This shift saw a reorganization of plant and faunal assemblages and, more generally, of terrestrial biome structures across North America and Europe to what we see today [33,34,35].

The 2000-year stability in extratropical Northern Hemisphere summer land temperatures into the 19th century is also relative to the recent record (1850–2023). Most recently, land temperatures have increased strongly, exceeding the previous two millennia’s envelope of natural variability (Figure 1a) [39]. Such exceedance in temperatures is also seen in global land and sea records [40].

Figure 1b also reveals the developing nonstationarity in climate in terms of a shift in the frequency distribution of summer northern land temperatures from that for the past 2000 years to that since 1850. This includes a shift in the long-term mean from pre-1850 to 1850–2023 (blue-to-red arrow in Figure 1b). Other recent shifts in climate system behavior include changes in the means and extremes of other surface climate variables (e.g., precipitation and wind speed) and global climate dynamics [21,41,42,43]. Also unprecedented over the last millennia are changes in key atmospheric and ocean circulation dynamics since the early 20th century. These include the expansion of subtropical atmosphere high-pressure systems and shifts in subtropical and subpolar currents and deep oceanic circulation [42,44,45,46]. Nonstationarity is also seen in multiyear climate oscillations, such as in El Niño–Southern Oscillation (ENSO), whose multidecadal variability has increased since the end of the 19th century relative to the previous three centuries [47].

This evolving climatic nonstationarity is threatening biospheric integrity. Such ecological disruption is occurring via widespread geographic, phenological, and phenotypic shifts in species [48,49,50]. Shifts in community structure and function has ensued [15,51]. We face the question of whether rates of anthropogenically forced environmental change exceed or will exceed the capacity for species and communities to adapt [52,53]. This presents a challenge for the conservation of not only at-risk species and ecosystems but also for conservation at all levels of biodiversity (i.e., genetic through compositional, structural, and functional biodiversity) [54]. Such pressures on biospheric function lead to the degradation of ecosystem services and jeopardize human socioeconomic stability [51,55,56,57,58,59]. Additionally, the loss of biospheric integrity pushes the Earth System towards biosphere-linked tipping points [17,25,60,61,62].

There have been, nonetheless, multi-century warm and cold periods over the past two millennia that have been ecologically and socially disruptive at the biome level [63,64]. These include, for example, the Medieval Warm Period and the Little Ice Age (Figure 1a). An important distinction between these anomaly periods and the current warming period is that recent decadal temperatures exceed those of the last 2000 years (Figure 1a,b) [39,63]. In the coming centuries, this exceedance is projected to continue to increase or stabilize at yet higher temperatures (discussed more in Section 2.3). Also unprecedented is global synchrony in multidecadal-scale temperature anomalies since the early 20th century, whereas the last 2000 years’ cold and (prior) warm periods lacked such global coherence [65].

The current transgression of climate outside the envelope of natural variability is one of seven global safe operating space boundaries that have been breached [61]. Others include the loss of integrity in land cover, freshwater hydrology, biogeochemical cycles, ocean pH levels (acidification), biodiversity, and biospheric productivity, along with the introduction of disruptive “novel entities” (e.g., synthetic chemicals and genetically modified organisms) [61,66]. Climate change interacts synergistically with these other stressors, amplifying their aggregate impact. Together, rapid climate change and high levels of other human disturbance are anticipated to result in future rates of environmental change unrealized over the last 65–66 M years of evolutionary history [67].

2.2. System Reorganization

We understand from the paleorecord that the climate system and biosphere have switched between different steady states in the past, such as during the aforementioned Late Pleistocene–Holocene transition [22,33,68,69,70,71]. The transition was specifically from the Younger Dryas cold period (12,900 to 11,700 years ago) to the early Holocene. The Younger Dryas ended abruptly, with temperature increases on the order of 5–10 °C over 10–90 years for the North Atlantic region [22,72,73]. The transformation was global, with magnitude and timing varying with latitude and between Northern and Southern Hemispheres. It substantially altered atmospheric and ocean circulation patterns and, therefore, the very nature of surface climates [22,72,74,75,76]. The lesson is that the Earth System can change states suddenly for regions or the globe at timescales as short as decades [22]. Given ongoing, increasing anthropogenic forcing, global and regional changes in state are considered plausible within this century ([21] (Box TS-9)), [25,60].

Whether by gradual shifts or abrupt state changes, anticipated climate change will likely be of a magnitude and manner that will reorganize the biosphere along with other Earth System components [32,77,78,79]. The ecological reassembly would arise from the movement of climate regimes outside the range of normal variability, putting their corresponding ecosystems in an environmental space to which their structure and function are poorly suited [24,32,80,81,82]. Such reorganization would result in communities and ecosystems unrecognizable to us and may be sudden and essentially irreversible [35,83,84]. “Essentially” means that such a change in state is likely irreversible on human timescales.

The unpredictable disassembly and reorganization of biospheric structure and function would lead to new ecological equilibrium states or, more likely, to a trajectory of novel transient states, whether of communities, ecosystems, or ecoregions. The nature of this reorganization would develop from the collective responses of individual species as a function of their adaptive capacities, dispersal capabilities, species interactions, and speciation [12,35,78,84,85,86,87,88,89].

Novel trophic, competitive, and symbiotic species interactions will also emerge as geographic shifts bring together species that have limited or no common evolutionary past [48,90]. New interactions may also arise between coexisting species whose phenologies are not currently synchronized but for whom phenologies become synchronized under a new climate regime. Their climate change responses may lead to key life history stages being coordinated in a way that leads to new tight trophic or mutualistic interactions [91].

Ecological reorganization is already underway globally at genetic through ecosystem levels in marine and terrestrial biomes and, in some systems, critically so [12,15,49,92,93,94,95]. Global climate model (GCM) simulations under a range of GHG emission scenarios suggest that the development of novel climates, ecoregions, and biomes could occur worldwide by the end of this century [32,79,81].

The reorganization into stable novel systems will take a long time due to lagged responses and only when there is some level of the following:

Stabilization of regional climates;
Stabilization of species distributions and species interactions across a region;
Stabilization of species microevolutionary processes and plastic adjustments, the latter including changing morphology (“shape shifting”) and prey switching [12,15,88,96,97,98,99].

2.3. Climate Stabilization

Climate stabilization will not be approached as long as the Earth’s energy budget is out of balance due to strong anthropogenic forcing from elevated GHG concentrations and other human activities [1,100,101,102]. This continued destabilizing forcing is driven by socioeconomic inertia due to factors such as human behavior and established energy infrastructure [102,103].

In addition to socioeconomic inertia, a key question remains: what is the Earth System’s physical inertia given emissions and climate system changes to date? That is, in what timeframe and at what state will climate stabilize if net-zero emissions of GHGs are reached [104]? A “net-zero” scenario refers to when anthropogenic GHG emissions are stabilized, achieved through both reduction in emissions and anthropogenic-facilitated CO₂ removal (such as by CO₂ capture and nature-based solutions) [105]. Once that occurs, Earth System Model (ESM) simulations project atmospheric CO₂ to have a downward trajectory due to ocean and terrestrial carbon uptake [106].

In these simulations, decreasing CO₂ levels result in cooling due to reduced radiative forcing. This cooling is countered by the release of heat from the oceans to the atmosphere from heat previously stored under elevated global temperatures [1,106,107]. As these energy terms come into balance, net-zero emission simulations project that mean global surface temperatures would level off on timescales of decades from and at temperatures close to when net-zero emissions are achieved [101,106,108]. These results suggest that once net zero is reached, global surface temperatures would stabilize fairly rapidly, albeit at temperatures substantially higher than pre-industrial levels. In these projections, the new global surface temperature regime would persist for centuries. The span of ESM results suggests some cooling or additional warming could occur at decadal and century timescales (e.g., ±0.3 °C/50 years) [105,106]. To give a possible timeframe for net-zero climate stabilization, the most ambitious mitigation pathways considered in the IPCC’s 6th assessment [109] would not result in net-zero GHG emissions being reached until the latter half of this century [108].

There are large uncertainties in a net-zero scenario due to competing positive and negative physical and biogeochemical feedbacks involved in determining the trajectory of global surface temperatures [11,105,106,110]. Many of these are poorly understood or not included in the models [105,111]. Additionally, stabilization of global surface temperatures does not necessarily translate into stabilization of regional or biome climates nor of other climate processes, such as precipitation, frequency of extreme temperatures, and sea-level rise [112,113]. Critically, an elevated equilibrium mean temperature places the Earth System at greater risk for passing many climate–biosphere tipping points [114,115,116].

A return to conditions resembling historical climates, while socioecologically ideal, is unlikely, even with a massive draw down of anthropogenic GHGs [117]. This is due to the following:

Long response times of key stabilizing processes (such as deep-ocean carbon sequestration) [101,102,105,117,118];
State changes that are essentially irreversible at centennial or longer timescales of some Earth System components, such as ocean heat content, continental ice sheet mass, sea level, and permafrost carbon [101,102,105,117,118,119,120].

The potential shift in climate regimes outside the envelope of normal climate variability, coupled with the low likelihood of climate stabilization under even the most ambitious socioeconomic pathways, indicates that we can anticipate a broad-scale reorganization of biomes and their constituent ecosystems by the end of this century. To plan for such changes, understanding how the climate system functions to determine biospheric structure and its sensitivity to strong altered forcing is crucial.

3. The Planetary Heat Engine and Global Atmospheric and Oceanic Circulation

3.1. The Climate System as a Heat Engine

In its broadest sense, the Earth’s climate system is a heat engine driven by the equator-to-pole imbalance in radiative inputs vs. outputs (net gain in low latitudes and net loss in high latitudes) (Figure 2) [121]. Atmospheric and oceanic planetary circulation function to equilibrate this imbalance by transporting low-latitude heat poleward [121].

Strong seasonal differences in net radiation balance (Figure 2b) lead to marked changes in atmospheric and ocean circulation from winter to summer in Northern and Southern Hemispheres. The latitudinal band where net radiation balance is near zero shifts seasonally towards the winter hemisphere (light yellow, Figure 2b). This, along with strong wintertime radiative losses at high latitudes, concentrates the latitudinal gradient in net radiation in that hemisphere (June in the Southern Hemisphere and January for the Northern Hemisphere, Figure 2b). This intensifies wintertime mid- and high-latitude atmospheric circulation, such as during winter storms. In contrast, atmospheric circulation in the subtropics and tropics is more vigorous in the summer hemisphere, where radiation input is strongest (e.g., January for the Southern Hemisphere, Figure 2b).

While absorbed solar radiation is a roughly sinusoidal forcing function with latitude (Figure 2a), the latitude-dependent response of global atmosphere and ocean circulation to this forcing is far more complex. The resulting heat energy transfers within and among Earth System components are strongly nonlinearly dependent on latitude, longitude, and season. We explore these global patterns in energy transfers in the next section (Section 3.2). We link the distribution of terrestrial and marine biomes to these patterns in Section 4 and Section 5.

While atmospheric and oceanic circulation function to equilibrate the equator-to-pole imbalance, anthropogenic GHG emissions and other altered forcing are increasing the heat content of the Earth System as a whole [1]. The resulting planetary energy disequilibrium (due to heat retained) is changing global patterns in energy transfers through changes in the location and intensity of atmospheric and oceanic circulation [122,123]. We consider the general sensitivity of these patterns to altered forcing arising from the Earth’s increasing energy content later in the paper (Section 6).

Figure 2. Earth’s radiative balance. (a) Latitudinal distribution of annual absorbed incoming solar radiation (blue line) and outgoing thermal (infrared) radiation (red line), showing net gain in the tropics and subtropics and net loss at higher latitudes. The resulting surplus heat energy in lower latitudes is transferred towards the poles by atmospheric and oceanic circulation (indicated by broad horizontal arrows). (b) Seasonal comparison of net radiation for the Western Hemisphere for boreal winter (January; left) and boreal summer (July; right) for 2009 from satellite observations. Net radiation = incoming solar radiation minus reflected solar and emitted thermal radiation. In (a,b), blue shading indicates net loss to space, with red for net gain [(a) from Webb [124], used under a Creative Commons CC BY 4.0 license. Original source: National Oceanography Centre, UK (NOC) under Creative Commons CC BY 3.0 Unported license. (b) Captured from NASA Scientific Visualization Studio: CERES Radiation Balance. https://svs.gsfc.nasa.gov/4794. Accessed on 9 June 2023. Public domain.].

3.2. Global Atmospheric and Ocean Circulation Driven by the Planetary Heat Engine

In this section, we first describe major atmosphere and oceanic circulation systems. We then relate these dynamics to the planetary heat engine’s energy transfers. From here on, we use “atmosphere” and “troposphere” interchangeably, unless otherwise indicated, where the troposphere is the lower layer of the atmosphere where most weather occurs.

3.2.1. Planetary Circulation Systems

Key atmospheric circulation features, from the equator to the poles, are as follows:

The Intertropical Convergence Zone (ITCZ);
Subtropical High-Pressure zones;
Mid-latitude westerly Jet Streams and embedded cyclonic storms;
Polar High-Pressure zones.

The mean global distribution of these features for January and July is mapped in Figure 3. Their vertical structure in the troposphere is illustrated in Figure 4a. The ITCZ is a pantropical low-pressure region. The Subtropical Highs are semi-permanent high-pressure centers, predominately located over eastern regions of subtropical oceans. Also shown are the winter positions of the Mid-latitude Jet Streams (“Polar Jet Streams” in Figure 3). Wintertime semi-permanent, mid-latitude low-pressure centers are found in the North Pacific and North Atlantic Oceans, known as the Aleutian and Icelandic Lows, respectively (Figure 3a).

Air masses are an attribute of the troposphere arising from these dynamics. Seasonal air masses from continental or marine sources and of high- or low-latitude origins develop because of and are transported by these circulation systems (Figure 4b). Characteristic air masses tend to prevail over a region and, therefore, distinguish terrestrial climate regimes [128,129,130,131,132,133,134,135]. In addition to thermal properties, air masses also differ in moisture content. Figure 5 reflects the instantaneous spatial complexity of air masses in terms of atmospheric water vapor.

Major surface ocean circulation features are, from the equator to the poles, as follows:

Westward equatorial currents and eastward equatorial countercurrents;
Mid-ocean subtropical gyres (with anticyclonic circulation);
Eastern boundary currents;
Western boundary currents;
Mid-latitude and subpolar eastward currents;
Subpolar gyres (with cyclonic circulation);
Subpolar and polar westward currents.

The distribution of these surface currents is mapped in Figure 6. They are forced largely by prevailing winds at the surface (Figure 3) and constraining continental geometry, notably at eastern and western ocean basin boundaries.

3.2.2. Global Heat Energy Transfers Accomplished by Planetary Circulation Systems

The vertical and horizontal structures of the atmosphere and oceans arise from (1) the Earth System’s heat engine transporting excess heat energy from low to high latitudes and (2) geophysical fluid dynamics. This transport occurs via interconnected atmospheric and ocean circulation.

For low latitudes, these transport mechanisms can be described generally as follows:

In the tropics and subtropics, the primary atmospheric circulation is the Hadley Circulation, a thermally direct, vertically overturning circulation (Figure 4a). The Hadley Circulation consists of two Hadley Cells roughly on either side of the equator rotating in opposite directions vertically (the Southern Hemisphere cell is roughly a mirror image of the Northern Hemisphere cell in Figure 4a).
The Hadley Circulation is driven by high solar heating in the tropics and the convergence of the Trade Winds in the ITCZ (Figure 3 and Figure 4a). Solar radiative energy powers heating at the surface and high rates of evaporation. The latter is most important over the oceans. The Trade Wind convergence forces a lifting of the Trades’ warm moist air. This air cools on lifting, resulting in water vapor condensation, cloud formation, and high precipitation. The release of latent heat from condensation additionally powers the ITCZ’s deep convection (to altitudes of 15–20 km). In this way, the Hadley Circulation is fueled largely by the conversion of latent heat energy to kinetic energy and then, with lifting, to potential energy [136]. The return flow aloft transports energy poleward. This transport is strongest between 10 and 30° N/S latitude [137].
As the air aloft cools, it sinks into the subtropics until roughly 30° N and S (Figure 4a). The subsiding air suppresses cloud formation and precipitation and creates the arid Subtropical High-Pressure zones. This is the descending branch of the Hadley Circulation (Figure 4a). The Trade Winds are the return flow at the surface, completing the Hadley Circulation.
From the tropics to the mid-latitudes, poleward heat transport is also accomplished by the oceans, primarily in subtropical gyre western boundary currents (Figure 6). These currents include the Gulf Stream and Kuroshio, as well as those along the eastern coasts of Brazil, South Africa, and Australia. This transport is strongest between 5 and 40° N/S latitude [138,139]. The subtropical gyre eastern boundary currents return cold water equatorward off the mid-latitude and subtropical west coasts of North and South America, Europe, Africa, and Australia (Figure 6).

In the mid- and high latitudes, energy transport is as follows:

From the subtropics, the continuation of thermally direct atmospheric flow to the poles is constrained by the rotation of the Earth (the Coriolis effect). This dynamical constraint turns the winds to the east, forming the Mid-latitude Westerlies and, where strongest, the Mid-latitude Jet Stream (Figure 3 and Figure 4a). This turning results in a mid-latitude buildup of low-latitude warm air against higher-latitude cold air, forming a strong meridional temperature gradient.
Fronts form where these warm and cold air masses meet. For example, in the mid-latitudes, the convergence of warm subtropical air and mid-latitude or subpolar cold air forms the Polar Front (Figure 4b).
Similarly, fronts form in the upper ocean where warm and cold surface currents meet. In the mid-latitudes, poleward-moving subtropical gyre warm water intersects equatorward-flowing colder subpolar gyre water (e.g., the Gulf Stream and Labrador Current in the western North Atlantic; Figure 6). The convergence forms a strong oceanic mid-latitude temperature gradient. The ocean’s temperature gradient supports the development of the troposphere’s mid-latitude meridional temperature gradient [140].
From the subtropics and through the mid-latitudes, atmospheric energy transport is largely by synoptic scale, mid-latitude cyclonic storms [141]. These storms dissipate the mid-latitude temperature gradient by mixing warm air poleward and cold air equatorward. They act primarily in the horizontal, with warm air circling around on the east side of the storm, and cold air circling on the west. Where they meet, either warm air overruns cold or cold air runs under warm air. In both cases, warm air is lifted, leading to condensation, cloud formation, and precipitation. The release of latent heat from condensation further powers these storms, boosting poleward energy transport. This energy transport is strongest between 30 and 60° N/S latitude [137,138]. These storms are the main source of wintertime precipitation in the mid-latitudes. In the Northern Hemisphere, favored locations for storm initiation are Aleutian and Icelandic Low-Pressure centers (Figure 3a).
At higher latitudes, mid-latitude and subpolar air masses meet yet colder polar air, forming a shallow, high-latitude meridional temperature gradient and corresponding Arctic (Antarctic) Front in the Northern (Southern) Hemisphere (Figure 4a,b) [142]. In the Northern Hemisphere summer, Arctic Front development is supported by the sea–land contrast between the open Arctic Ocean and adjacent continents and land-cover feedback arising from the contrast between the boreal woodland and adjacent tundra [130,143]. Associated with the high-latitude meridional temperature gradient are summer Arctic cyclonic storms which transport heat to the polar cap [143,144].
In the Northern Hemisphere oceans, subpolar gyre eastern boundary currents continue heat transport poleward (Figure 6). This is particularly important in the North Atlantic, where the North Atlantic and Norwegian Currents transport Gulf Stream warm water into the Arctic Ocean. Cold water is returned in the subpolar gyre western boundary currents, for example, off the coasts of Greenland, Labrador, and the Kamchatka Peninsula (Figure 6).

Through these processes, the poleward transport of heat energy is traded off between the atmosphere and oceans. This exchange is conducted principally via air–sea surface fluxes of sensible heat and latent heat.

These transport and exchange processes respond nonlinearly to internal and external forcing. Responses include shifts in the strength and location of circulation features at daily and seasonal scales to interannual, decadal, and longer timescales. Multiyear and longer-term dynamics are tied to (1) internal modes of climate variability (e.g., ENSO) and (2) altered forcing arising from, for example, volcanic activity, solar radiation variability, and anthropogenic sources [21,145,146,147,148]. Change in one part of the system results in adjustments in others, with balancing feedbacks.

These dynamics define climate regimes and set the stage for the world’s biomes. This gives us an overall picture of the climate system as follows:

Highly structured with distinct processes governing each region of the atmosphere and oceans, including characteristic atmosphere–ocean exchanges of heat, moisture, and momentum;
Highly nonlinear in how these processes operate and interact to produce key features of atmospheric and ocean circulation.

4. Earth System Dynamics and Terrestrial Biomes

4.1. Terrestrial Biomes and Mean Climate

The distribution of terrestrial biomes (Figure 7) and their structure and function are generally ascribed to regional climate (Box 2) [149]. The simplest delineation of their bioclimates is by mean annual temperature and mean annual precipitation [150,151,152]. Other simple formulations include seasonality in temperature, precipitation, and evapotranspiration as controls over the seasonality of major plant functional types that characterize individual biomes [149,153]. For example, terrestrial biomes are commonly distinguished by the seasonal dynamics of dominant vegetation types, such as evergreen versus cold-deciduous or dry-deciduous forests (Box 2).

4.1.1. Nonstationarity of Biome–Mean Climate Relationships

While biomes can be generally prescribed by annual or seasonal means of surface climate variables, relationships between climate and the structure and function of biomes are more complex. This complexity reflects climate nonlinearities in three ways. First, observed relationships among surface climate variables—such as among minimum and maximum temperatures, precipitation, solar and infrared radiation, and cloud cover—are expected to be restructured as climate dynamics shift. In addition, relationships between long-term mean climate and the frequency distribution of weather events and multiyear variability modes are not physically constrained. Consequently, temporal autocorrelation and covariance structures of surface climate variables are not expected to be stable over time.

Second, the observed correspondence between annual or monthly climatic means and biological processes is not fixed but is inherently process dependent. Critical bioclimatic controls over biotic processes often operate at the level of weather events (i.e., at sub-daily to weekly timescales). Examples are plant tolerances to extreme cold or heat events and mammal survivorship tied to winter storm events [154,155]. Bioclimatic controls can also have legacy or hysteretic effects. An example of legacy effects is herbaceous community structure dependence on the year-to-year sequence of prior precipitation in addition to the current year’s mean and events [156]. These two kinds of complex bioclimatic relationships are tied to (1) event-level climate dynamics or (2) interannual/interdecadal climate variability modes and their teleconnections, respectively. Both respond nonlinearly to changes in global heat energy transfers under climate change [137,157]. Consequently, empirical relationships between mean climate and the structure and function of biomes are not expected to be stable under a changing climate [81]. Nonstationarity between biome composition and mean climate is also observed in the paleorecord [33,35].

Third is the nature of climate regimes. In multivariate climate space, surface climates are naturally clustered into regular recognizable patterns or “regimes” rather than arranged along steady gradients. The result is a level of uniformity in surface climate within and marked contrasts between climate regimes. These regimes define biome climates. They are the direct result of the spatial structure of global atmospheric dynamics (Figure 3, Figure 4 and Figure 5). This clustering is not fixed. Under climate change, shifts in atmospheric circulation patterns will reorganize surface climates into novel regimes.

4.1.2. Climate Nonstationarity and Empirical Ecological Models

The nonstationarity of climate in the three ways just described limits the utility of ecological models that are based on empirical relationships between ecology and mean climate. This is the case for climate envelope species and ecosystem distribution models [158]. Mechanistic models share the same limitation insofar as they rely mean climate relationships. The adaptive capacity of species (e.g., genetic variability, phenotypic plasticity, and behavioral adaptation) and shifting ecological interactions (e.g., competitive exclusion, trophic interactions, and disease ecology) are among other factors that contribute to the nonstationarity of mean climate–ecology relationships [86,158,159,160,161,162,163,164,165,166]. In addition, microhabitat climates important for species persistence [167,168] are not represented by spatially coarse surface climate datasets commonly used in climate change impact assessments. Nonstationarity of climate–ecology relationships, unknown species attributes, complex disequilibrium dynamics, and novel community and trophic dynamics limit the ability of models to predict species and ecosystem distribution changes under novel environmental conditions [15,158,159,160,169].

Box 2. On the nature of terrestrial biomes.

Biomes are the coarsest unit of ecological organization on the globe, both terrestrial and oceanic. For terrestrial biomes, they represent units that, at continental to global scales, have similar ecological structure but not necessarily similar taxa nor evolutionary origins. They are defined primarily by dominant vegetation type and key climatic features (e.g., tropical rain forest and cold temperate desert). They are often described in terms of the following:

Similar structure, e.g., vegetation physiognomy (physical structure), biodiversity, and trophic complexity reflected in plant, animal, and microbial functional types [149,170,171];
Similar regional climate, with similar surface climate (e.g., in terms of thermal, moisture, and solar radiation regimes) [149,172] and regional climate dynamics (such as of summer monsoons and winter cyclonic storms) [173,174];
Similar ecosystem function, e.g., net primary production, biogeochemical processes (including soil and plant nutrient dynamics), and disturbance regime (such as fire and severe wind regimes) [175,176,177,178].

Biomes are a complex set of ecosystems differentiated by local climatic factors (e.g., tied to topography and proximity to water bodies), soil type, and biotic interactions (e.g., dominant species, trophic interactions, and keystone species) [179]. Boundaries between biomes can be abrupt (<100 km wide) or broad with overlapping elements. Regardless, biomes are more similar both internally and from continent to continent than in comparison to neighboring biomes on the same continent.

Relationships between climate and biome structure and function go beyond the simple formulations noted in the main text (Section 4.1). They are instead multivariate in climate space and vary across timescales. Capturing these relationships is the goal of dynamic global vegetation models (DGVMs) (e.g., [180]). DGVMs model the distribution of biomes defined by plant functional groups by simulating vegetation dynamics (e.g., establishment, competition, and phenology) coupled to biogeochemical and ecophysiological processes [181].

Biomes have characteristic rates of net primary production (NPP) [175]. NPP rates are climatically limited by solar radiation, water availability, and warmth (temperature) [172]. The relative importance of these factors varies globally by biome, reflecting corresponding climate regimes (Box 2, Figure 1). The shifting of their importance among biomes is reflected in nonlinear relationships between surface climate parameters and vegetation cover. For example, vegetation cover exhibits threshold or optimal responses to precipitation and temperature [182,183]. Changes in regional climate can shift which limiting factor dominates in a biome. Under a warming climate from the mid- to late 20th century, boreal forests appear to have shifted from being predominately temperature limited to being moisture limited [184].

Box 2. Figure 1. Global distribution of the relative importance of water, temperature, and solar radiation controls (color triangle insert) over terrestrial net primary production (NPP). Note that the wet tropics are primarily light limited (green), warm deserts are water limited (red), cold deserts and grasslands are both water and temperature limited (magenta), and boreal and Arctic systems are variously limited by temperature and light (cyan) and regionally also by water (blend of cyan and magenta). The map is based on satellite estimates of terrestrial NPP (little or no productivity is indicated in gray), climate data (1931–1960), and modeled plant physiological limits set by water balance, temperature, and absorbed photosynthetically active radiation [172,185,186,187]. [Adapted from Running et al. [172]; used with permission from the American Institute of Biological Sciences/Oxford University Press].

4.2. Phase-Change Thermodynamic Constraints on Terrestrial Biomes

A little-explored factor in the biogeography of terrestrial biomes is a key thermodynamic constraint on climate. This constraint is that saturation water vapor pressure is a function of air temperature (Figure 8a). Specifically, the amount of water that air can hold is exponentially related to air temperature in a relationship known as the Clausius–Clapeyron Relation (Figure 8a). Consequently, cold air holds little water (1 g H₂O·m⁻³ air at −20 °C). In cold environments, air forced to saturation on cooling yields minimal condensation and precipitation. For this reason, ice-free landscapes nearest the poles are considered polar deserts (“Cold Deserts” in Figure 7). On the other hand, warm air has the potential to hold large amounts of water (up to 30 g H₂O m⁻³ at 30 °C). As a result, air in warm environments can range from being very dry to very moist (Figure 8b).

Globally, the distribution of mean annual precipitation with annual mean surface air temperature reflects this thermodynamic constraint (gray dots in Figure 8b). Because a greater amount of moisture is potentially available at higher temperatures, warmer climates have a wider range of precipitation values than colder climates (Figure 8b). As a result, tropical and temperate regions have a wider variety of biomes than boreal and Arctic regions (Figure 8b,c). The range of terrestrial annual net primary production (NPP) also follows this constraint, with a broader range of NPP values occurring at higher temperatures driven by a wider range of water balance coefficients (Figure 8d).

4.2.1. Precipitation Regime Expansion Under Increasing Temperatures

With global increases in surface temperature under climate change, the historical domain of annual mean temperatures in Figure 8b expands to the right. Within this new temperature domain comes an exponential increase in the water holding capacity of the atmosphere and, therefore, the potential for much higher annual precipitation. This expands the climate envelope along both temperature and precipitation axes to include climate regimes not currently present globally (dark red arrow in Figure 8b). Such an expansion of the climate envelope suggests an emergence of new, no-analog tropical bioclimates and biomes [35,79].

Dynamic global vegetation model simulations project that under future climate change, novel bioclimates will develop in the wet tropics by the end of this century [188]. With the development of new environmental niches, an important question is how they might become occupied [189,190]. For tropical forests, the key is whether tropical species have either narrow physiological limits readily exceeded with warming, with a loss of forest diversity [164,191,192], or can acclimate to novel high temperatures and, therefore, readily fill new tropical niches [193,194,195,196].

4.2.2. Uneven Shape of the Climate Space Envelope

The envelopes of precipitation versus temperature and water balance versus temperature have a dip between 10 and 20 °C (green double arrows in Figure 8b,d). This reflects that there is a limited number of areas on land with high precipitation within this subtropical–warm temperate temperature range. These land areas generally (1) lack high-elevation mountains that intercept moisture and generate orographic precipitation, (2) lie on the dry side of mountain ranges, such as in the lee of the Andes, and (3) lack strong flow from a moisture source (exceptions are southeastern regions of the United States and China). These conditions arise from global atmospheric circulation in relation to continental geography. These flow patterns are shifting with climate change ([21], Section TS.2.3), potentially expanding the temperature–precipitation climate envelope in this temperature range and, in doing so, creating bioclimates not currently present on continents.

Figure 8. Water phase-change thermodynamics as a constraint on the distribution of precipitation versus temperature, terrestrial biomes, and net primary production (NPP). (a) The Clausius–Clapeyron Relation: saturation water vapor pressure with respect to liquid water as a function of air temperature. When forced to saturation on cooling, water vapor condenses forming droplets or ice crystals and precipitates (following the blue double arrow to the left). In reverse, condensate evaporates and air dries on warming (blue double arrow to the right). (b) The global distribution of mean annual precipitation versus annual mean surface air temperature over land (gray dots) with an overlay of major terrestrial biomes, fashioned after Whittaker [150]. (c) The original Whittaker [150] biome diagram with more finely differentiated biomes than in (b). “Sclerophyllous woodland” corresponds to Mediterranean-type biomes. (d) Simulated terrestrial annual NPP as a function of annual water balance (annual precipitation—potential evapotranspiration) and annual mean surface air temperature for the globe. In (a–d), red-dashed vertical lines indicate where air temperature = 0 °C. The dark red arrow with “?” (upper right in b) suggests a future expansion of terrestrial climate space with warming and a resulting greater possible range in precipitation. The green double arrows in (b) and (d) indicate the reduced number of terrestrial locations with temperatures between 10° and 20 °C, another area of potential future expansion of climate space that is currently unrealized on continents. In (b), temperature and precipitation data are 1961–1990 means for terrestrial locations (excluding Antarctica) gridded at 10’ latitude/longitude (~18.5-km) resolution. [(a) After Barry and Chorley [125]. (b) Adapted from Principles of Terrestrial Ecosystem Ecology, Chapin et al., 2nd ed., 2011 [175]; used with permission of Springer Nature BV; permission conveyed through Copyright Clearance Center, Inc. (c) Adapted from Biogeography, Lomolino et al., 4th ed. [197], © 2010 Sinauer Associates, Inc., used with the permission of Oxford University Press-Books (US & UK), conveyed through Copyright Clearance Center, Inc; their figure after Whittaker [150]. (d) Adapted from Churkina and Running [185], reproduced with permission from SNCSC].

4.3. Planetary Atmospheric Circulation, Climate Regimes, and Terrestrial Biomes

4.3.1. Climate Regimes

As introduced earlier (Section 4.1.1), climate is spatially organized into regimes, each with characteristic conditions (Box 3). Within these regimes, bioclimates have typical ranges of temperature, precipitation, and other surface climate parameters, both in annual means and seasonal patterns [149,198]. These parameters change little in some reaches and then transition sharply from one climate regime to another (Figure 9). In addition, these regimes have distinctive patterns of interannual and decadal variability, such as those driven by teleconnections linked to, for example, ENSO and the Pacific Decadal Oscillation (PDO) [199,200,201,202]. Such within-region similarities arise from a coherent set of climate processes that consistently operate across a region (Box 3).

Box 3. The nature of terrestrial climate regimes.

Terrestrial climate regimes are characterized by the following:

A range of physical conditions considered at sub-daily to multidecadal timescales. Key elements of terrestrial climate regimes include temperature, precipitation, solar radiation, cloud cover, and wind.
General uniformity of climatic properties and processes over a region. This spatial coherence extends across a range of timescales (e.g., hourly to decadal) such that regimes are characterized as much by daily event structure and teleconnections as they are by their long-term annual and seasonal means.
Boundaries between regimes are linked to strong between-region gradients in climatic variables such as temperature and moisture or to a specific key bioclimatic factor. An example of the latter is the seasonal occurrence of freezing temperatures, which limits the distribution of tropical vegetation.
Linkage to broad-scale atmospheric and adjacent ocean circulation dynamics. In this way, global circulation dynamics and continental and ocean basin physiography are the defining context for a regime.
Relatively well-defined spatial shifts in surface climate at regime boundaries. These transitions are generally on the order of 100 km wide and are linked to changes in climate dynamics (e.g., [128,132,203]).
Close correspondence to biotic regions. This is reflected in that many climate classification systems are delineated by (or at least guided by) biomes or smaller biogeographic divisions (e.g., [152,153,204]). Conversely, many ecological regionalization schemes are defined by a consistency in climate regime [149,198,205].
Significant within-regime spatial heterogeneity. While regional-scale processes pervade across a climate regime and contrast with those in neighboring regimes, there is significant within-regime spatial heterogeneity. Within the context of overall regime features, this heterogeneity arises from (1) broad latitudinal/longitudinal gradients due to varying intensity of prevailing atmosphere circulation dynamics and (2) finer-scale changes in temperature, precipitation, and other surface variables that reflect heterogeneity in, for example, physiographic features.
While climate regimes are generally ascribed to a given geography, boundaries can shift significantly at multidecadal, centennial to millennial, and longer timescales [133,206,207].

The concept of climate regimes is basic to the design of international and national biogeosphere observation programs, such as GLORIA [208] and, in the United States, NEON [209].

The distribution of climate regimes and corresponding terrestrial biomes across continents does not follow strictly zonal patterns (Figure 7). Rather, their boundaries vary with longitude. This variation is driven by (1) prevailing global atmospheric circulation patterns that vary strongly with longitude [210] and (2) continental and ocean basin physiography, which further modify these circulation patterns. The former includes broad prevailing cyclonic and anticyclonic circulation systems that vary in strength with longitude (e.g., L’s and H’s, respectively, in Figure 3a). Land features include barriers to moisture (e.g., high mountains) and distance from moisture sources. The results of these physical processes are reflected in latitude- and longitude-dependent gradients of precipitation, surface temperature, and surface wind regimes, as shown in Figure 9 and Figure 10.

Exploring these figures reveals the complex structure of the climate system. In Figure 9a, precipitation changes most sharply from the tropics (the blue band representing the ITCZ) to the subtropics (orange domains of the Subtropical Highs). On the other hand, temperature gradients are strongest in mid- and subpolar latitudes (generally 30−60/70° N and S), where temperatures transition from roughly +20 °C to −5 °C (orange to light blue in Figure 9b). In addition, Figure 10c reveals the latitudinal and longitudinal complexities of surface winds transporting heat and moisture over oceans and continents and transferring momentum to the surface (e.g., in driving ocean currents). Surface winds exhibit organized structures that dominate one region and transition to distinctly different patterns in adjacent domains. This contrast is most evident over the oceans, such as from convergent lines near the equator in the ITCZ to spiral divergence in the Subtropical Highs (Figure 10).

4.3.2. Biomes and Circulation Features

The complex fields of global precipitation, surface temperatures, and surface winds (Figure 9 and Figure 10) have a strong correspondence with the distribution of terrestrial biomes (Figure 7). This arises from fundamental links between terrestrial biomes and key features of global atmospheric circulation (Figure 3 and Figure 4a).

Tropical and subtropical biomes are tied to the ITCZ and Subtropical Highs. These links are generally as follows:

In equatorial rainforests (e.g., Amazonian and Congolian forests; Figure 7), the ITCZ’s rain band is generally overhead year-round, resulting in high annual rainfall with some seasonality (Figure 9a).
The transition from tropical wet to dry forests and savannas to subtropical xeric shrublands (Figure 7) is set by rainy season length. That, in turn, is set by the seasonal swing of the ITCZ’s rain band (Figure 3). The ITCZ seasonal position generally “follows the sun” with a lag. In dry tropical forests, rains are limited to summer months, with yet fewer months in savannas and dry shrublands as ITCZ influence tapers off [212]. This creates a biogeographical pattern mirrored about the equator of dry–wet–dry bands of vegetation, which is most notable in western and central Africa as one travels southward from the Sahel through the Congolian rainforest to the Kalahari Desert (Figure 7).
Other wet tropical forests (Figure 7) are found beyond the realm of the ITCZ. These forests are tied to the Trade Winds and summer monsoons (Figure 3). When uplifted by highlands, these strong moist flows produce orographic precipitation. As part of the Hadley Circulation, the position and intensity of the Trade Winds are linked to ITCZ dynamics. The strength of summer monsoonal flow is governed by the intensity of summertime mid-continental heating in contrast to cooler nearby subtropical or tropical oceans. Some Trade Wind forests are separated by a dry zone from their equatorial counterparts (Figure 7). Consequently, these forests have developed highly endemic floras and faunas [213]. Examples are the Atlantic Forest of southeastern Brazil and Western Ghats of southwestern Indian, which are disjunct from the Amazon and Burmese–Southeast Asian forests, respectively.

Low-latitude linkages with strong longitudinal dependence are as follows:

Subtropical biomes shift with longitude from hot deserts on the west to subtropical moist forests to the east (Figure 7). This change depends on location relative to Subtropical Highs. Subtropical deserts are dominated by subsiding air from the center of Subtropical Highs, suppressing precipitation (Figure 3, Figure 4a and Figure 9a). Examples are the Sonoran and Atacama Deserts. On the eastside of continents, subtropical forests receive summertime moisture from poleward circulation around the western side of the Highs (Figure 3 and Figure 9a). This flow can extend inland, supplying moisture to warm temperate forests, such as in the southeastern United States and East Asia (Figure 7). Summer monsoons can also drive this moist flow inland.
Mediterranean-type biomes (“sclerophyllous vegetation” in Figure 7) have summer-dry and winter-wet climates. These are generally on the west side of continents and poleward side of Subtropical Highs. In summer, they are dominated by the dry subsiding air of the Subtropical High (Figure 3). In winter, the High retreats equatorward, replaced by Mid-latitude Westerlies. The Westerlies bring mid-latitude cyclonic storms and wintertime precipitation. This is illustrated by the summer-to-winter equatorward shift of Subtropical Highs and Mid-latitude Westerlies (e.g., Figure 3b to 3a for Californian and Figure 3a to 3b Chilean Mediterranean biomes).

Tropical forest, subtropical forest, subtropical desert, and Mediterranean climates are dynamically linked as their distributions are controlled by the ITCZ and Subtropical Highs—the ascending and descending branches of the Hadley Circulation, respectively (Figure 4a).

The distribution of mid- and high latitude terrestrial biomes is linked to positions of the Polar and Arctic Fronts. These connections are as follows:

Transitions from subtropical to warm and cold temperate biomes, whether forest or drylands (Figure 7), are determined by the seasonal position of the Polar Front (Figure 4b). In the Northern Hemisphere, the more poleward transitions to boreal forest, forest–tundra woodland, and Arctic tundra (Figure 7) are controlled by the seasonal swing of the Arctic Front (Figure 4b). Polar and Arctic Front seasonality is reflected in seasonal shifts in the relative dominance of subtropical, mid-latitude, subpolar, and polar air over these regions and generally follows changes in radiative balance (Figure 2b). The seasonal occurrence of subfreezing weather tied to colder air masses exerts a strong control over these biome [149]. Across these biomes, plant and animal communities change in their composition depending on species tolerances to freezing.
In the Polar Highs, subsiding cold, dry air (cA in Figure 4b) spreads equatorward from polar regions, forming Arctic and Antarctic Fronts. In the Northern Hemisphere summer, the prevalence of this air corresponds to southward transitions from polar deserts through Arctic tundras to the boreal treeline (Figure 7). This treeline is the northern limit of the boreal woodland (coniferous forest–tundra, Figure 7) and roughly corresponds to the mean summer position of the Arctic Front [130,131,214]. The boreal forest’s southern limit generally follows the mean wintertime position of the Arctic Front. In the Southern Hemisphere, a corresponding Polar High dominates the Antarctic continent. Where ice-free, Antarctic terrain is a polar desert, such as in the Dry Valleys of Antarctica. The continental extremely cold, dry air extends equatorward to just off the continent’s edge throughout the year. There, it meets maritime subpolar air, forming the Antarctic Front.

Extratropical linkages with strong longitudinal dependence are as follows:

In the Northern Hemisphere, extratropical biome transitions vary strongly with longitude. Temperate, boreal, and Arctic biomes lie father north on the west sides of the northern continents and to the south on eastern sides (Figure 7). This pattern is set by Northern Hemisphere ocean–continent contrasts. Climates on the western side of these continents are strongly influenced by the moderating effects of adjacent oceans. This maritime influence is brought inland by westerly atmospheric flow and is illustrated by the northward extent of brown, green, and light-blue temperature color bands in Europe in Figure 9b. On the other hand, winter climates on the eastern side of these continents are strongly influenced by atmospheric circulation from cold continental interiors (e.g., darker blues in northeastern Asia in Figure 9b). This west-to-east contrast in air mass source is the key determinant of longitudinal variation in the otherwise latitude-controlled distribution of temperate, boreal, and Arctic biomes [128,129,130,131,132,133,134,135] (Figure 7).
Across temperate regions, west-to-east biome transitions follow a wet–dry–wet sequence. These transitions run from forests to grasslands or shrub deserts and then back to forests. This pattern is observed from western to eastern temperate North America and from western Europe through central Eurasia to East Asia (Figure 7). In boreal North America, the transition is from xeric to mesic needleleaf forests (e.g., western versus eastern Canada, not distinguished in Figure 7) [215]. These moisture gradients arise primarily from mid-latitude cyclonic storms losing their moisture as they track from west to east. Moisture loss is from orographic precipitation in western mountain systems (e.g., in North and South America) and with distance from upstream moisture sources (Eurasia). Transitions back to mesic biomes occur when storms are reinvigorated as they travel further eastward. There, they draw on new sources of moisture from adjacent oceans (such as the Gulf of Mexico and the China Sea).
In Siberia, deciduous needleleaf (Larix spp.) boreal forests (B(d) in Figure 7) replace evergreen needleleaf boreal forests that are found at the same latitudes in western Eurasia and Canada. The distribution of Larix forest is tied to the wintertime presence of extremely cold Arctic air extending as far south as Mongolia and northern China. In summer, on the other hand, the deciduous forest climate is like that of the evergreen forest. Both forests are to the south of the mean summertime location of the Arctic Front [214], such that there is sufficient growing season warmth to support tree growth [216]. In winter, the Siberian deciduous boreal forest climate is more like the climates of the polar desert and Arctic tundras than of the evergreen forest. Deciduousness and other adaptations make Larix species well adapted to the extreme cold [217,218].

All biome climates are dynamically linked. Temperate biomes, boreal forests, and Arctic tundra climates are connected through the location and strength of meridional temperature gradients and the Mid-latitude Jet Streams and embedded cyclonic storms. Extratropical biome climates are tied to both subtropical and polar climates as the meridional temperature gradients result from flows of warm moist air out of lower latitudes and cold air from higher latitudes.

5. Earth System Dynamics and Ocean Biomes

For marine biomes, we focus primarily on the open-ocean euphotic zone and coastal systems. The euphotic zone, also known as the epipelagic, is generally the upper 200 m of the ocean within which sunlight is sufficient for photosynthesis [219]. The light limit for net primary production is generally 1% of surface solar radiation. Epipelagic production is the foundation for most marine food webs [219]. Moonlight additionally drives epipelagic photobiologic cycles and ecology [220].

Below the epipelagic, the mesopelagic has some minimal sunlight and is rich in marine life, especially consumers. This layer is known as the “Twilight Zone” and generally extends from 200 to 1000 m in depth.

In some biomes, strong vertical differences in temperature and salinity occur in the mesopelagic, resulting in a sharp vertical density gradient (known as a pycnocline). This gradient limits vertical mixing, leading to seasonal or year-round stratification. By limiting vertical mixing, stratification leads to the depletion of nutrients in the photic zone through uptake by primary producers. Nutrients released through decomposition in lower layers of the ocean are made available in the photic zone on deep mixing or upwelling.

Marine biomes are strongly determined by ocean circulation. Primary drivers of surface currents are winds and thermohaline density gradients. In the next two sections (Section 5.1 and Section 5.2), we describe the dynamics of these two circulation types and how they control marine ecology. Ocean carbon sequestration processes are discussed in Section 5.3. We cover coastal biomes in Section 5.4.

5.1. Wind-Driven Ocean Circulation Systems and Biomes

5.1.1. Physical Regimes

Ocean biomes have been classified in various ways according to surface environmental regimes, surface biodiversity patterns, and horizontal and vertical circulation dynamics [221,222,223,224,225,226,227,228,229] (Box 4). Physical regimes can be delineated in terms of water masses defined by characteristic physical properties (temperature and salinity) and source [230]. Here, we define epipelagic ocean biomes in terms of the following:

Global-scale wind-driven surface currents (Figure 6);
Vertical mixing regime;
Sea surface temperatures;
Resulting patterns of nutrient availability and net primary productivity (Figure 11).

Vertical mixing is driven by winds acting on the surface (i.e., by wind stress) or by processes driving upwelling and downwelling. In general, upwelling brings to the surface cold, deeper water that is nutrient rich and deoxygenated (both arising from decomposition at depth) (Figure 12). Primary upwelling processes include the following:

Intermediate or deep water forced to the surface when they encounter basin barriers (e.g., sea mounts and continental margins);
Offshore winds driving surface waters away from the coast, replaced by upwelling water;
Sustained winds parallel to continental margins or a combination of parallel wind systems, resulting in surface water divergence and upwelling of deeper water.

This last process is called Ekman transport and is described in Box 5, along with how it results in upwelling (or downwelling).

Box 4. The nature of epipelagic ocean biomes.

Key environmental factors that determine the ecology of epipelagic ocean biomes are as follows [224,230,231,232,233,234,235]:

Sea surface and vertical profiles of temperature, salinity, pH, and dissolved oxygen;
Primary production, indicated by surface chlorophyll a concentration (Figure 11);
Seasonality of net solar radiation at the surface, influencing primary production and surface-layer heating;
Euphotic zone depth (as a function of transparency) and nutrient concentrations (reactive N, Fe, PO₄³⁻, and Si);
Ocean basin geography, e.g., western and eastern basin boundaries, continental shelf geometry, and other bathymetric features that intercept currents (e.g., islands and sea mounts);
Surface currents and their eddies, which are wind driven and constrained by basin geography [232,236];
Sea-ice dynamics;
Downwelling and upwelling;
Mixing regime as a function of solar heating, air and sea temperatures, wind stress, surface currents, sea-ice cover, and basin geography and as modified by upwelling and downwelling.

Mixing regime determines whether stratification is permanent (subtropics), seasonal (mid-latitudes), or unstratified (subpolar). Polar waters are seasonally or permanently stratified depending on seasonal persistence of sea ice.

An epipelagic biome classification scheme is illustrated in Box 4, Figure 1 [237]. This classification is based on where (1) primary productivity is light limited in winter (mid- and high latitudes), (2) productivity is nutrient limited year-round (subtropical gyres), or (3) there is strong annual net downward surface heat flux (equatorial latitudes). These divisions generally correspond to water masses defined by temperature, salinity, and source region [230].

Box 4. Figure 1. Epipelagic biomes as delineated by Henson et al. [237] based on light, nutrient, and heating attributes. Mid- and high-latitude biomes, which are light limited in winter, productive in spring to early summer, and nutrient limited in summer, are the High-latitude North Pacific (labeled 1), Southern Ocean Pacific (5), High-latitude North Atlantic (6), Southern Ocean Atlantic (10), and Southern Ocean Indian (14). Subtropical gyre biomes, which are nutrient limited year round, are the Oligotrophic North Pacific (2), Oligotrophic South Pacific (4), Oligotrophic North Atlantic (7), Oligotrophic South Atlantic (9), and Oligotrophic Indian (13), where oligotrophic means low productivity. Low-latitude, high-heat-flux biomes are the Equatorial Pacific (3), Equatorial Atlantic (8), Arabian Sea (11), and Bay of Bengal (12). [From Beaulieu et al. [238], based on Henson et al. [237]. Used under a Creative Commons CC BY 4.0 license (modifications: white labels, southern biome coloration extended based on Henson et al.)].

Ocean biome-defining attributes also include characteristic interannual and interdecadal variability modes that are linked to planetary ocean and atmospheric dynamics. These modes arise from air–sea interactions and intra- and inter-basin connectivity. These dynamics include (1) multiannual oscillations, such as El Niño–Southern Oscillation (ENSO), the North Atlantic Oscillation (NAO), and the Indian Ocean Dipole (IOD), (2) multidecadal regime shifts, such as the Pacific Decadal Oscillation (PDO) and Atlantic Multidecadal Oscillation (AMO), and (3) interactions among these modes [145,147].

Figure 11. Global distribution of major epipelagic and near-coastal marine biomes as reflected in mean annual surface ocean primary productivity (as measured by satellite chlorophyll a concentration, ~1998–2003 average). Indicated western boundary currents are Gulf Stream (GS) and Brazil (BRC), Agulhas (AC), Kuroshio (KC), and Eastern Australia (EAC) currents. Eastern boundary currents associated with major subtropical coastal upwelling zones (in gray-blue ovals) are California (CAC), Humboldt (HC), Canary (CNC), and Benguela (BNC) currents. Key downwelling zones for oceanic deep-water formation (in red ovals) are in the North Atlantic (NAD; in the Labrador and Nordic Seas), Ross Sea (RSD), and the Weddell Sea (WSD) (based on ref. [239]). In the Southern Ocean, the Subtropical Convergence and Antarctic Convergence and Divergence zones are circumpolar. [Adapted from NASA/EOS [240]. Public domain].

Downwelling processes include the following:

Ekman transport as above, but with orientations resulting in surface convergence and downwelling (Box 5);
Surface currents intercepting barriers;
Density-driven sinking of surface waters, forming intermediate and deep water.

Density-driven sinking results from (1) inputs of cold or high-salinity surface water or (2) strong surface cooling from heat loss to the air (discussed further in Section 5.2). Important areas of deep-water formation are indicated by red ovals in Figure 11, including, for example, off the Labrador coast.

Box 5. Ekman transport and ocean upwelling.

Wind-forced movement of surface waters depends on the duration of that wind stress. Over short periods (e.g., less than 6 h), the surface water column generally flows in the same direction as the wind. If, on the other hand, the duration is long enough for the Earth’s rotation to come into play, then the effect of wind stress on the water surface is redirected by the Coriolis effect. In the Northern Hemisphere, surface currents are pushed 45° to the right of the wind direction (to the left in the Southern Hemisphere). Each layer pushes the next layer down further to the right (left), with a drop in speed due to friction. This is the Ekman Spiral (Box 5, Figure 1a). The net, depth-integrated Ekman transport of water is 90° to the right in the Northern Hemisphere (Box 5, Figure 1a,b), to the left in the Southern Hemisphere.

Box 5. Figure 1. Ekman transport of surface waters in the Northern Hemisphere: (a) Ekman Spiral vertical profile and (b) aerial view. In (a,b), surface wind direction is shown by dark blue arrows and surface flow by white arrows. In (a), subsurface water movement is indicated by the succession of smaller arrows turning with depth in a spiral. The surface current is 45° to the right of the wind direction, whereas the net depth-integrated transport of water is 90° to the right (large light-blue arrow). Directions in (a,b) are reversed in the Southern Hemisphere. In (a), note that as depths approach 100 m, water movement is turned 180°, flowing in a direction opposite to that of the wind; below that, the influence of the wind is dissipated. [(a) Adapted from: https://oceanservice.noaa.gov/education/tutorial_currents/media/supp_cur04e.html; accessed 6 October 2024. Public domain].

While ocean basin bathymetric forcing, offshore winds, and thermohaline currents also play a role [241,242,243,244,245], Ekman transport is a key mechanism for ocean upwelling (Figure 11). Along western continental coastlines, for example, equatorward winds establish westward Ekman transport that pulls surface waters away from the coast, which are replaced by deeper, nutrient-rich water (Box 5, Figure 2a,b; Figure 12a). These are critical upwelling zones in the subtropics and mid-latitudes of both hemispheres, such as in the California, Humboldt, Canary, and Benguela currents (Figure 11).

Upwelling also occurs in the open ocean, where winds establish Ekman transport in opposite directions. Along the equator, easterly Trade Winds to the north of the equator push surface waters northward, and those to the south push waters southward, that is, in both cases, away from the equator, resulting in divergence and upwelling (Box 5, Figure 2c). Upwelling in the Southern Ocean at the Antarctic Divergence (Figure 11) arises from adjacent subpolar westerly and easterly winds that set up divergent Ekman transport flows (Box 5, Figure 2d). When the wind arrangement is reversed, the result is convergence and downwelling.

Box 5. Figure 2. (a,b) Ekman transport-forced westward coastal upwelling along western coastlines in (a) the Northern and (b) Southern Hemispheres driven by persistent along-shore equatorward winds. Similarly, poleward winds along eastern seaboards result in eastward upwelling (not shown). (c) Upwelling along the equator in the Pacific Ocean, as indicated by high chlorophyll a concentrations, caused by divergent Ekman transport from easterly Trade Winds on both sides of the equator. (d) Upwelling in the Southern Ocean, as indicated by nitrate concentrations, caused by divergent Ekman transport from opposing winds driving West Wind Drift (WWD) and East Wind Drift (EWD) currents. In (a,b), the compass red arrows point to the north. [(a,b) Adapted from Lichtspiel [246]. Public domain. (c,d) From Webb [247]; used under a Creative Commons CC BY 4.0 license].

5.1.2. Biomes, Primary Production, and Wind-Driven Circulation Features

Phytoplankton productivity varies by biome largely according to nutrient and light availability (Figure 11 and Figure 12a). Zooplankton biomass primarily follows from primary productivity patterns [248], as do higher-order consumer productivity and biodiversity. Marine productivity and physical regimes are strongly linked to ocean circulation dynamics and air–sea interactions.

For low productivity biomes, these links are as follows:

Subtropical gyres (Figure 6)—Productivity in the center of subtropics gyres is nutrient limited. Weak winds associated with subsiding air of overhead Subtropical Highs result in strong, year-round stratification, limiting deep mixing and nutrient availability (Figure 12a). The gyres’ low productivity is indicated by dark blue in Figure 11. The gyres are anticyclonic, driven by easterly Trade Winds near the equator and Mid-latitude Westerlies at their poleward extent (Figure 10). To the east and west, these circulation systems are constrained by ocean basin geography, resulting in eastern and western boundary currents, respectively.
Downwelling zones—Where strong downwelling of surface water circulates phytoplankton out of the photic zone, productivity becomes light limited. Downwelling also oxygenates otherwise hypoxic deep water, supporting decomposition.

Nutrient-enriched biomes include the following, roughly in order of increasing productivity:

Western boundary currents—Upwelling replenishes nutrient levels of western boundary currents (Figure 6 and Figure 11). Mentioned earlier, these include the Gulf Stream and Brazil, Agulhas, Kuroshio, and Eastern Australia currents. Upwelling mechanisms include (1) westward equatorial currents intercepting continental margins and (2) Ekman transport driven by poleward winds on the west side of Subtropical Highs [241] (Box 5).
Equatorial upwelling zones—Equatorial surface currents have relatively high productivity due to the upwelling of nutrient rich water (Figure 6, Figure 11 and Figure 12a). This upwelling is driven by easterly Trade Winds on opposite sides of the equator, resulting in divergence from opposing Ekman processes (Box 5, Figure 2c).
Subpolar gyres and West Wind Drift currents—North Pacific and North Atlantic Subpolar Gyres are cyclonic, driven by mid-latitude westerly winds to the south and subpolar easterly winds to the north and additionally constrained by ocean basin geography (Figure 6 and Figure 11). These temperate and subpolar waters are highly productive due to nutrients supplied via cold-season wind-driven vertical mixing (Figure 12a and Figure 13a). In spring, algal blooms result from seasonally increasing light levels, in addition to continued nutrient input from mixing. Mixing continues until summer stratification is established. In the Southern Ocean, mid-latitude westerly winds drive the Antarctic Circumpolar Current (West Wind Drift, Figure 6), largely unconstrained by basin geometry. These winds force deep vertical mixing in the winter and spring to depths on the order of 500–1000 m (Figure 13b), supplying nutrients (Figure 12a) supporting high levels of productivity (Figure 11).
Southern Ocean upwelling zones—Other highly productive regions in the Southern Ocean are upwelling zones near the Antarctic continent (Figure 11). Upwelling occurs (1) in the Antarctic Divergence, (2) where poleward flowing deep water intercepts the continental margin, and (3) along the coast, owing to strong offshore winds originating from the Antarctic interior [245,249]. In the Antarctic Divergence, upwelling is driven by adjacent westerly and easterly winds that result in opposing Ekman transport flows (Box 5, Figure 2d). Upwelling of cold, hypoxic deep-water lowers surface temperatures and decreases dissolved oxygen along the continent (Figure 9b and Figure 12b) and, along with mixing in the Antarctic Circumpolar Current, strongly increases nutrient levels in the Southern Ocean (Figure 12a).
Eastern boundary current upwelling zones—Coastal upwelling zones along western continental margins in the subtropics and mid-latitudes are highly productive. These are primarily along the western coasts of California, Peru, Chile, West Africa, and Namibia (Figure 11) and are associated with corresponding eastern boundary currents—the California, Humboldt, Canary, and Benguela currents. Upwelling is driven by Ekman transport that is forced by equatorward Trade Winds on the east side of the Subtropical Highs (Box 5, Figure 2a,b). These upwelling zones often have a strong signature in maps of nutrient-rich and deoxygenated epipelagic water (Figure 12a,b).
Continental runoff—High-nutrient continental runoff enhances the productivity of waters offshore river deltas that drain high-precipitation basins. For example, enrichment of nutrient-poor ocean waters by Amazon, Congo, and La Plata (Argentina and Uruguay) rivers is prominent in Figure 11.

Figure 12. Epipelagic (a) nitrate concentration (μmol/kg) and (b) oxygen levels (as % of saturation) at a depth of 65 m. Interpolated from in situ observations. These patterns are reflected in those of primary and secondary production (Figure 11 and [248], respectively). [From World Ocean Atlas 2023, https://www.ncei.noaa.gov/access/world-ocean-atlas-2023f/. Accessed 26 July 2025. Public domain.].

Figure 13. Ocean maximum (95th percentile) mixed-layer depth for (a) mid-March and (b) mid-September. Climatology based on two decades of observations (roughly 2000–2020). Areas that are dark blue are deeply mixed; yellow to light green areas are shallowly mixed (stratified). Note that the color bar scale is logarithmic (doubling every 2 contours; white outside the contours = no data). [From Johnson and Lyman [250], © 2022 American Geophysical Union; used with permission from John Wiley and Sons.].

As with terrestrial biomes, the physical regimes of epipelagic biomes are tied to each other and those of the atmosphere. This is through the direct connection between Subtropical High and Trade Wind dynamics and low productivity in subtropical gyre centers versus high productivity in boundary current and equatorial upwelling regions. Extratropical linkages are through mid-latitude westerly and subpolar easterly winds that drive deep seasonal mixing and high productivity in the Antarctic Divergence and Northern Hemisphere subpolar gyres. These connections result in characteristic interannual to multidecadal variability patterns driven by air–sea interactions (Box 4) [145,147]. These variability modes are the source of planetary teleconnections that link the dynamics of marine and terrestrial biomes globally [199,200,201,202,251].

5.2. The Thermohaline Circulation

The Thermohaline Circulation is a global overturning circulation driven by contrasts in surface water density (Figure 14) [239,252]. It is also referred to as the “ocean conveyor belt.” Components include the following:

Wind-driven surface currents;
Deep-water formation by downwelling of dense surface seawater;
Deep- and bottom-water currents;
Upwelling of these deeper waters.

Both cooling (hence “thermo”) and increased salinity (“haline”) contribute to increases in surface density. The circulation is multi-basin, connecting marine biomes of the Pacific, Atlantic, Indian, and Southern oceans (Figure 14). Figure 15 gives a view of the oceans as one body of water as a way to visualize this connectivity.

The Atlantic Ocean sector of the Thermohaline Circulation is the Atlantic Meridional Overturning Circulation (AMOC). The AMOC ties Atlantic epipelagic and coastal biomes together and to global ocean dynamics via the Thermohaline Circulation. The AMOC consists of northward warm surface currents in the North and South Atlantic (red lines, Figure 14) and deep-water return flow (blue lines, Figure 14) [239]. Mechanisms involved are as follows:

In the high-latitude North Atlantic, near-surface warm, high-salinity waters of the Gulf Stream become denser with (1) cooling due to heat lost to the atmosphere and (2) increased salinity owing to brine rejection as sea ice forms. In the Nordic and Labrador Seas, this water sinks, forming Atlantic Deep Water (yellow ovals, Figure 14).
This deep water flows at depth through the North and South Atlantic and Southern Ocean. It returns to the surface in the Antarctic Circumpolar Current (Figure 14) and in subtropical upwelling areas (blue-gray ovals, Figure 11) [243,244].
At the surface, these waters flow through the subtropics and tropics where they warm and, owing to high evaporation rates, increase in salinity (green shaded areas in Figure 14). This warm surface return flow is concentrated in the western North Atlantic, forming the Gulf Stream (Figure 6). The circulation then returns to the high-latitude North Atlantic.

Transporting heat poleward, the Gulf Stream and its northeastern extensions—the North Atlantic and Norwegian Currents—amplify Northern Hemisphere ocean–continent contrasts, thereby intensifying the longitudinal variation in terrestrial biome distribution discussed earlier (Section 4.3.2). These currents strongly modify the atmosphere by heating the air and increasing its moisture content. That air brings warmth to mid- and high-latitude Europe. Consequently, temperate biomes of western Europe extend farther poleward than those in East Asia and eastern North America, as previously noted (Section 4.3.2). For example, the temperate–boreal transition mixed forest (Figure 7) extends to roughly 60° N in Scandinavia but only to around 50° N in Manchuria (China) and Nova Scotia (Canada).

The Thermohaline Circulation, wind-driven currents, and global atmospheric circulation form an interconnected dynamic system. Its dynamics respond in complex ways to altered forcing, as addressed later in Section 6.

5.3. Ocean Carbon Pumps

The Thermohaline Circulation is instrumental in the ocean’s long-term sequestration of CO₂. Atmospheric CO₂ is incorporated into surface waters as dissolved inorganic carbon, (DIC; dissolved CO₂ and carbonate equilibrium species) and particulate organic carbon (POC; including living microbes, detritus, and fecal pellets). In regions of deep-water formation, epipelagic DIC and POC are subducted (“pumped”) into deep and bottom waters of the North Atlantic and Southern Ocean (Figure 14) [256,257]. Carbon sequestered at depths of ~1000 m remains isolated from the surface at 100-year timescales and on the order of 1000 years for depths >2000 m [257,258,259].

The transport of DIC occurs via a “solubility” pump (referring to the solubility of CO₂ in water). It entails the vertical mixing of high-DIC water to deeper, low-DIC layers. The transport of POC is a “biological pump”, as it involves the vertical mixing and sinking of live and dead particulate organic carbon. This includes diel and seasonal vertical migrations of zooplankton and other fauna between the epipelagic and mesopelagic [219].

Where deep water returns to the surface through upwelling, dissolved and gaseous CO₂ generated from the decomposition of POC is outgassed to the atmosphere. In the net, deep, long-term sequestration is positive [256]. In a biogeochemical loop, the upwelling of deep water also brings into the euphotic zone water rich in nutrients supporting algal CO₂ fixation. Deep sequestration of dissolved and particulate carbon provides, respectively, key physiochemical and biogeochemical negative feedbacks that stabilize GHG-forced climate change [256] and serve as a net-zero mechanism (Section 2.3). However, processes controlling carbon deep-water sequestration and return to the surface are highly nonlinear, and their possible dynamics under climate change are not well understood [111,260].

5.4. Coastal Biomes

Biogeographical realms for coastal and continental shelf environments are broadly organized into tropical, temperate, and polar divisions [223]. Within tropical and temperate divisions, we define four major coastal biomes identified by key habitat-forming (or “foundation”) species groups. These are reef-forming coral, mangrove, seagrass, and kelp species. As coastal foundation species, they provide three-dimensional habitats for highly diverse invertebrate and vertebrate (e.g., fish) communities [261,262,263,264,265]. These environments are generally high in light and high in nutrients (from upwelling or continental runoff) and, therefore, are highly productive.

Figure 16. Global distribution of (a) coral reefs (in red), (b) mangrove forests (green), (c) seagrass meadows (black), and (d) kelp forests, as represented by the distribution of the order Laminariales (black). Ocean surface temperature regimes are also shown in (c). [(a) From Teh et al. [266]; used under a Creative Commons CC BY 4.0 license. (b) From Giri et al. [267], © 2010 Blackwell Publishing Ltd.; used with permission from John Wiley and Sons. (c) From Orth et al. [268], reproduced with the permission of the American Institute of Biological Sciences/Oxford University Press. (d) Reprinted from Jayathilake and Costello [269], © 2020, with permission from Elsevier].

The corresponding biomes (coral reefs, mangrove forests, seagrass meadows, and kelp forests) are pan-oceanic. Their distribution is broadly determined by coastal sea temperatures or, for mangroves, air temperatures. Coral reefs and mangrove forests lie primarily in tropical waters (Figure 16a,b). Coral reefs occur near or offshore in fringing and barrier reefs, while mangrove forests occur in estuaries and along shores. Seagrass beds are found more broadly, from the tropics to temperate and subpolar regions (Figure 16c), typically in coastal lagoons [270]. Kelp forests occur in (1) colder waters of temperate and subpolar regions and (2) cold-water upwelling zones in the subtropics, such as in the Humboldt and Benguela currents (Figure 16d). These temperature regimes are tied to surface currents and upwelling that control the position and strength of ocean fronts, roughly represented by the ocean temperature regimes shown in Figure 16c.

Two other important coastal foundation species groups are reef-forming bivalves and salt marsh species (graminoids and succulents) [271]. Bivalve reefs are found from the tropics to temperate regions, generally on rocky shores [271,272]. Salt marshes are found worldwide but are limited in their latitudinal range more so by competition with mangroves in the tropics and disturbance by ice action in the Arctic than directly by climate ([273], Chap. 3).

6. Changing Climate Dynamics

In the next two sections, we explore how atmospheric and ocean circulation dynamics are changing in response to anthropogenic forcing. These changes have direct consequences for regional bioclimates and biomes.

6.1. Changing Atmospheric Circulation

From Section 4, we have insights into how each feature of global atmospheric circulation exhibits complex nonlinear behavior and how this complexity structures the terrestrial biosphere. We also understand that, as part of the Earth System’s heat engine, dynamics of these features are tied together and with those of oceanic circulation (Section 3). Accordingly, atmospheric circulation features change in concert in response to the Earth’s growing positive energy imbalance. The increased energy content of the Earth System [1] has amplified global atmospheric circulation as a means to transport additional surplus energy poleward [122,274].

Since the mid- or late 20th century, the Hadley Circulation has intensified. This is reflected in the following:

A narrowed ITCZ and increased precipitation in the tropical rain belt [275,276,277].
Poleward expansion of the Hadley Cells, on the order of 1° latitude/decade in both hemispheres [278,279,280].
Increase in strength and poleward extent of the Subtropical Highs [42,279,280,281,282].
On the eastern flanks of the Subtropical Highs, their strengthening has further suppressed precipitation in subtropical deserts. Expansion of the Highs has also broadened the deserts poleward [42,283] and decreased precipitation in adjacent, transitional subtropical–mid-latitude dry climates (e.g., in some Mediterranean-climate regions) [42,284]. Coupled with increased tropical precipitation, these changes have sharpened the wet–dry contrast across lower latitudes [276,285].
On the western flanks of the Subtropical Highs, east–west shifts in their extent have altered the poleward flow of moisture into adjacent continents, such as in East Asia and southeastern United States [44,286,287,288]. This has altered precipitation regimes of corresponding wet subtropical and warm temperate forests, with regional responses dependent on the nature of the shift [286,287].
Poleward shifts in the distribution of tropical cyclones (e.g., hurricanes and typhoons) tied to Hadley Circulation expansion [289,290,291,292]. Higher tropical sea surface temperatures and increased water content of the tropical atmosphere over oceans [293,294] have additionally reinforced changes in tropical cyclone storm tracks [291,294].
Vertical extension of the Hadley Cells, raising the height of the tropical tropopause [280,295]. The vertical and horizontal expansion of the Hadley Circulation reflects a taller and broader overturning circulation in the tropics.

Hadley Circulation intensification has additionally resulted in changes in mid- and high-latitude circulation. These include the following:

Poleward shifts in the positions of mid-latitude temperature gradients (i.e., Polar Fronts), Jet Streams, and mid- and high-latitude cyclonic storm tracks and associated precipitation patterns [276,279,296,297,298]. These shifts are forced by Subtropical High expansion.
An increase in the intensity of mid-latitude and Arctic cyclonic storms. This is in terms of increases in the frequency of stronger storms, storm duration, and precipitation rates and totals [298,299]. Shifts in summertime Arctic Fronts will be moderated by dynamics tying their position to that of the Arctic Ocean and feedbacks with boreal woodlands [130].

The strength of these patterns varies by season, longitude, and Northern versus Southern Hemispheres. Many of the observed patterns have been evaluated over a roughly 40-year record, whereas others are over as many as ~65 years. To date, some individual trends are statistically significant. Taken together, they show the emergence of an anthropogenically forced climate change signal above long-term natural variability.

Long-term shifts in Hadley Circulation strength and position also drive an energy imbalance between Northern and Southern Hemispheres. This imbalance is linked to differences in heating of and heat energy transport between the two hemispheres [300]. Greater warming in the Arctic (“Arctic amplification”) and northern land masses, relative to lower and southern latitudes, has shifted the ITCZ further northward since the late 20th century [285,301,302,303]. The northern warming bias may be countered in the future by a more slowly acting cessation of the Atlantic Meridional Overturning Circulation, resulting in a cooling of the North Atlantic and Northern Hemisphere in general (discussed more in Section 6.2.3) and, with it, a southward shift in the ITCZ [304].

These complex response dynamics and resulting changes in surface climate are manifested in changes in biome structure and function. For example, shifts in the ITCZ alter the hydrological cycle of tropical regions with latitudinal, longitudinal, and seasonal changes in the tropical rain belt’s position and daily precipitation event structure (i.e., rainfall intensity and frequency) [300,305]. Such shifts impact the persistence, structure, and function of wet and dry tropical forests and savannas [306]. Similarly, on the western sides of continents, poleward shifts in Subtropical Highs and mid-latitude storm tracks increase the aridity of subtropical and temperate drylands, altering their ecology [307].

6.2. Changing Ocean Dynamics

From Section 5, we understand how ocean circulation exhibits complex nonlinear behavior and how these dynamics structure the marine biosphere. The oceans play four major roles in ongoing climate change. First, the oceans have absorbed the greatest portion of heat energy retained by the Earth System owing to anthropogenic forcing. This amounts to roughly 90% of heat energy gained since the 1970s [1]. Second, carbon uptake by the ocean, including via carbon pumps, acts to stabilize the climate system. Third, ocean circulation transports heat poleward as part of the Earth’s heat engine. Fourth, the Thermohaline Circulation drives inter-ocean basin deep circulation, which has a role in determining marine and terrestrial climate regimes globally and regulating the energy balance between Northern and Southern Hemispheres. These interacting dynamics are changing the physical and biotic nature of marine and coastal biomes.

6.2.1. Physical Regime

Ocean surface circulation and mixed-layer dynamics are sensitive to atmospheric and cryospheric changes that include the following:

Tropospheric circulation patterns and resulting changes in driving winds;
Net radiative flux at the surface and sensible and latent heat exchange with the atmosphere, controlling sea surface temperature, evaporation, and seawater freezing (causing brine rejection) and these, in turn, affecting density [308];
Sea-ice cover limiting these surface energy fluxes [308];
Freshwater inputs from precipitation, sea-ice melt, continental glacier and ice sheet melt, and continental runoff, affecting seawater salinity [309].

The resulting observed ocean changes include the following:

Greater sea surface temperatures and ocean heat content, with the rate of warming regionally variable and accelerating in recent decades [48,310,311];
Increased (decreased) near-surface salinity in high (low)-salinity regions related to regional changes in evaporation and freshwater inputs [21,48];
Longer and more extensive marine heat waves [312];
Rising sea level from thermal expansion and continental glacier and ice-sheet melt [313];
Altered broad-scale ocean circulation patterns, which include intensification and poleward shifts in the position of western boundary currents and subtropical and subpolar gyres [45,123,314,315,316];
Regional changes in stratification, nutrient availability, and oxygen levels depending on strength of surface heating, upwelling and downwelling processes, and freshwater inputs [48,111,309].

As with the observed changes in atmospheric circulation, some of these patterns to date show the emergence of an anthropogenic signal above long-term natural variability [123]. Collectively, they present a picture of ongoing oceanic change ([21] (Section TS.2.4)), [48].

Ocean and atmospheric changes are closely linked not only through changes in wind speed and direction, which drive ocean circulation change [316], but also through air–sea exchanges of energy, water vapor, and momentum. Greater sea surface temperatures in the subtropics shift the mid-latitude ocean temperature gradient poleward [291]. Warmer mid-latitude sea temperatures, in turn, warm the air so that the troposphere’s mid-latitude temperature gradient also shifts poleward [140]. This shift in air–sea heat exchange contributes to the poleward extension of the Hadley Circulation and to poleward shifts in Mid-latitude Jet Streams and storm tracks, as discussed earlier (Section 6.1) [291,317].

6.2.2. Marine Biome Change Linked to Changing Ocean Climate Dynamics

These physical regime changes affect (1) euphotic zone nutrient availability, (2) the solubility carbon pump, and (3) biotic processes [231]. Biotic changes include those in primary production, other photobiologic cycles, the biological carbon pump, fisheries health, sea bird populations, and coral bleaching [48,231,318,319,320,321,322,323]. Changes in circulation regimes have led to species range changes, which have altered food webs, restructured marine communities, and triggered the emergence of novel associations.

For example, along the eastern flank of the North Atlantic Subpolar Gyre, decreasing productivity and zooplankton community shifts have occurred since the mid-19th century. These changes reflect a northwestward poleward shift in gyre circulation and a resulting transition from subpolar to subtropical regimes in the east [45]. Similarly, in the Arctic Ocean, an increasing influence of North Atlantic and North Pacific waters over recent decades has shifted pelagic, continental shelf, and sea-ice communities from Arctic to more subpolar (boreal) associations (“borealization” of the Arctic Ocean) [324,325]. These include changes in vertebrate, invertebrate, and algal communities. In addition, warming, circulation changes, and sea-ice loss are removing barriers to dispersal across the Arctic Ocean. This has permitted, or may soon permit, faunal and floral exchange between the North Pacific and North Atlantic [326,327].

In the mid-latitudes, the pelagic spring algal bloom is sensitive to changes in ocean dynamics and atmospheric forcing. Spring bloom productivity and algal community structure are controlled by competing factors associated with the depth of mixing [328]. These are as follows:

Light limitation, with deeper mixing circulating phytoplankton to below the euphotic zone, versus –
Release from nutrient limitation, with deeper mixing entraining nutrient-rich deeper water into the euphotic zone.

Which factor dominates varies by water mass properties (i.e., subpolar versus subtropical waters) and interannual–interdecadal variability modes (e.g., NAO and PDO in the Northern Hemisphere) [328,329]. Both factors are sensitive to changes in ocean climate. Changes in limiting factors and spring bloom timing affect the phenology and structure of the phytoplankton community. These, in turn, alter those of the zooplankton community and higher trophic levels, such as fish and seabird populations [328,329]. In addition, shifts in phytoplankton community size structure affect the biological pump’s effectiveness in POC transport into deep waters [330,331].

Light penetration in the euphotic zone has decreased in recent decades across many ocean biomes. This is likely due to increased productivity in response to global circulation changes and higher sea surface temperatures [321]. This has, in the main, reduced photic zone depth (“global ocean darkening”) [321].

Marine biosphere responses are in the context of other stressors. These include increased nutrient loading (altering nutrient ratios), acidification, deoxygenation, overharvesting (altering population and food web structure), alien species introductions, and synthetic pollutants (including plastics) [48,66,231,332,333,334]. Cumulative or synergistic interactions with climate change result in complex ecological responses [335].

6.2.3. Atlantic Meridional Overturning Circulation (AMOC) Climate Sensitivity

The AMOC is a climate system tipping element [77]. If forced beyond a threshold (tipping point), system-amplifying processes rapidly shift its overturning dynamics to a shutdown state [336]. This is first manifested as a slowdown of the AMOC. Locations where AMOC stability is particularly sensitive to perturbation are where North Atlantic Deep Water forms in the Labrador and Nordic Seas (Figure 14). Here, perturbation comes from increased inflows of freshwater, which counter high-density seawater formation and sinking.

Increases in freshwater inflows originate from (1) increased river discharge into the Arctic, (2) increased import of Arctic Ocean sea ice into and subsequent melting in the subpolar North Atlantic, and (3) Greenland Ice Sheet melt runoff and iceberg discharge [337]. High-latitude warming is accelerating these three processes. Other positive Earth System feedbacks involved are (1) decreased Arctic Ocean sea-ice cover (increasing solar heating of surface water), (2) changes in atmospheric circulation (modifying sea-level pressure patterns and surface winds), and (3) increased precipitation (a direct freshwater input) [338,339,340,341,342].

The decrease in subpolar North Atlantic deep-water formation results in a slowdown and reorganization of the AMOC. A slowdown weakens the transport of cold northern deep water into the South Atlantic, causing a warm anomaly in Southern Ocean surface waters. The AMOC slowdown also reduces the Gulf Stream’s transport of warm, high-salinity subtropical waters into the northern North Atlantic, giving rise to a cold anomaly. Together, cooling in the North Atlantic and warming in the South Atlantic creates a “bipolar see-saw” dynamic in the AMOC [337,343].

Observed trends indicate that the AMOC has significantly weakened since the middle of the last century (in one estimate by 15% [344]) [345,346,347]. No analogous slowdown is seen in the paleorecord covering the past 1000 years [46]. Increased influx of freshwater from riverine, glacial, and sea-ice sources has contributed to the slowdown [60,337,346]. Since the beginning of the 20th century, the North Atlantic has cooled by up to −0.3 °C/century (−0.003 °C/year) (Figure 17a) [337,347]. More recently, Labrador Sea surface waters have cooled rapidly by as much as −2.5 °C over 25 years (−0.1 °C/year) (Figure 17b) [46,337]. Region-dependent changes in ocean productivity and shifts in plankton community ranges have accompanied the slowdown [45,348].

The AMOC shut down repeatedly during the Late Pleistocene, on the order of every 10,000 years starting around 120,000 years ago [239,350]. This occurred with strong cooling of the North Atlantic and warming of the Southern Ocean [239]. AMOC restoration was relatively rapid, generally spanning several thousand years, during which the see-saw pattern reversed to a warmer North Atlantic and colder Southern Ocean as observed today. The bipolar see-saw reflects the redistribution of heat energy between the two hemispheres accomplished by alternating modes of deep circulation [239,343,351]. The AMOC has been relatively stable for the last 8000 years [21,352].

Climate model simulations under future GHG emission scenarios suggest that a shutdown (“collapse”) of the AMOC could occur in the next 15–300 years, depending on emission scenario and across GCMs, with a “best” estimate of 50 years [353,354]. An AMOC shutdown would result in the weakening and redirection of the Gulf Stream, with cooling over Europe. A potential cooling in Europe of 3–8 °C would more than counter broader hemispheric warming trends through this century [355]. Such a cooling response is projected to be stronger in winter, with severe cold outbreaks [356,357]. Projected changes also include altered Jet Stream dynamics, affecting storm tracks and precipitation over Europe [356,357].

The sensitivity of the AMOC to high-latitude freshwater inputs is an example of strong nonlinear dynamics in the Earth System resulting in rapid changes in regional climate dynamics and hydrology and significant socioecological consequences [21,358]. The AMOC is part of a tipping cascade starting from (1) increased Arctic air temperatures accelerating the melting of the Greenland Ice Sheet, (2) slowdown of the AMOC, and (3) a resulting southward shift in the ITCZ [25,60]. That shift would lead to a drying of the Amazon Forest and reduced African and Asian monsoons, each with their own tipping points [25,60]. In this way, the AMOC is a core global tipping point setting off other tipping elements, which would, through positive feedback loops, act together to increase global surface temperatures and alter global precipitation patterns [353].

6.2.4. Coastal Biome Climate Sensitivity

Coastal foundation species groups and, therefore, their dependent communities are sensitive to climate change directly and indirectly through complex interactions among climate impacts and other stressors (Figure 18) [271]. Climate-related controls over coastal biomes are linked to conditions in both the adjacent ocean (including water temperature, wave action, and sea level) and inland watersheds (e.g., inputs of freshwater, nutrients, and sediments). Recent climate anomalies have revealed rapid system altering and extensive system-dependent responses of coral, seagrass, mangrove, and kelp environments to climate change [265,359,360,361]. These have included widespread coral bleaching events, loss of low-latitude seagrass meadows and kelp forests, and expansion of mangroves, all tied to warming conditions ([70], Section TS.5).

7. Summary and Discussion: Main Lessons for Conservation Planning and Resource Management

The next three sections, Section 7.1, Section 7.2 and Section 7.3, integrate material from this paper to address the key questions posed in Section 1.1. We conclude by discussing frameworks for decision making under climate uncertainty (Section 7.4).

7.1. Overview of Earth System Dynamics Driving Biome Distributions

The Earth System exerts dynamic and thermodynamic control over the structure and function of the biosphere. These controls arise out of the Earth System functioning as a heat engine transporting excess heat energy from low to high latitudes. The resulting organization and dynamics of global atmospheric and oceanic circulation define continental climates and ocean regimes. These circulation structures and mechanisms set the stage for the world’s biomes. We summarize these linkages for terrestrial and ocean biomes in the next sections.

7.1.1. Terrestrial Biomes and Atmospheric Dynamics

The distribution of terrestrial biomes is related to broad structures of the troposphere. These are primarily the ITCZ, Subtropical Highs, the easterly Trade Winds, westerly Mid-latitude Jet Streams, Polar and Arctic/Antarctic Fronts, and Polar Highs.

The location and intensity of the ITCZ, Subtropical Highs, and Trade Winds are directly tied to the strength of the Hadley Circulation. This is powered largely by the conversion of latent heat energy to kinetic energy in the ITCZ. Likewise, the position and intensity of the Mid-latitude Jet Streams and associated cyclonic storms are tied to the strength of mid-latitude temperature gradients. These gradients are formed from converging air masses flowing poleward from the Subtropical Highs and mid-latitude sources and equatorward from mid- and high-latitude sources. The strength of the poleward flows out of the Subtropical Highs is linked to the intensity of the Hadley Circulation. Consequently, low-, mid-, and high-latitude circulation systems are strongly interconnected.

This structure is mirrored in the distribution of terrestrial biomes. In the tropics, from the equator poleward, the progression of wet and dry tropical forests to savannas is tied to the seasonal swing of the ITCZ rain belt. Subtropical deserts occur in the dry, subsiding air of the Subtropical Highs. Warm and cold temperate biomes are delineated by the position of the Polar Front by season. In the Northern Hemisphere, the distribution of boreal forests, boreal woodlands, and Arctic tundra is similarly marked by the seasonal swing of the Arctic Front. Within temperate regions, the distribution of humid versus dry biomes is strongly influenced by circulation patterns over continents, the distance from moisture sources, and the interaction of these flows with mountain ranges (with temperate forests on the windward side and temperate deserts in their lee). These circulation patterns include mid-latitude cyclonic storms, summer monsoons, and poleward flow of warm, moist air around the western side of Subtropical Highs.

7.1.2. Marine Biomes and Ocean Dynamics

Epipelagic biome structure and function are tied to upper ocean circulation regimes. These regimes include subtropical gyres, their western and eastern boundary currents, subpolar gyres, and upwelling and downwelling zones. Major coastal biomes are coral reefs, mangrove forests, seagrass meadows, and kelp forests. These are broadly determined by water (or shore air) temperatures and linked to surface current and upwelling regimes.

Ocean circulation is forced by prevailing winds at the surface and thermohaline density currents and constrained by ocean basin geometry. Dependence on wind forcing ties ocean responses to atmosphere circulation changes. Dependence on density currents ties the Thermohaline Circulation change to changes in continental freshwater runoff (including from ice sheets), subpolar sea-ice melt, precipitation inputs, and air–sea interactions (including energy fluxes and evaporation).

Epipelagic community structure (e.g., food webs) is tied to euphotic zone physical and chemical properties (sea temperatures, light penetration, salinity, nutrient levels, etc.) and primary production. These properties are strongly determined by winds and surface heating, which drive the seasonality and depth of vertical mixing. In the low-wind regimes of Subtropical Highs, strong, year-long stratification limits productivity in the subtropical gyre centers. Forced by the Mid-latitude Westerlies, deep seasonal mixing supports high productivity in the subpolar gyres and mid-latitude West Wind Drift currents.

Upwelling and downwelling mechanisms include (1) wind-driven Ekman transport, (2) mechanical forcing (such as upwelling forced by ocean basin bathymetry), and (3) density-driven circulation. By bringing up nutrient-rich deep water, upwelling supports high productivity, especially in subtropical gyre eastern boundary currents, equatorial waters, and the Southern Ocean. Density currents drive the global Thermohaline Circulation, including the Atlantic Meridional Overturning Circulation (AMOC). AMOC involves (1) deep-water formation by the downwelling of high-density water in the high-latitude North Atlantic, (2) upwelling in the Southern Ocean, and (3) return surface flow, including the Gulf Stream, which brings warm water to the northeastern North Atlantic. The Thermohaline Circulation influences the global distribution of marine biomes and, by moderating adjacent continental climates, also affects the distribution of terrestrial biomes.

7.2. The Dynamics of Change

Additional heat energy retained by the Earth System from increases in GHGs (along with changes from other human forcings) goes to (1) increasing heat stored in the atmosphere, oceans, and land and cryosphere ice loss and (2) increased energy transferred among and within these system components. These involve transfers of kinetic and potential energy (i.e., circulation dynamics) and sensible and latent heat energy (i.e., thermodynamics).

These system-level changes drive three major, interconnected changes in global dynamics and thermodynamics. These are as follows:

Higher temperatures of the troposphere, ocean surface and deep waters, and land and a net loss of sea and land ice, including permafrost;
Greater potential atmospheric water vapor content resulting from warmer air temperatures;
Altered global circulation of the atmosphere and oceans.

The nature and magnitude of these changes have the potential to disrupt regional terrestrial and marine climates. This is through four mechanisms. First, greater land-surface air and sea surface temperatures drive regional shifts in mean climate and increase the frequency of extremes, including terrestrial and marine heat waves. The loss of land and sea ice reduces surface albedo, initiating a positive feedback. This snow/ice albedo feedback increases absorption of solar radiation and accelerates regional warming.

Second, altered global circulation of the atmosphere shifts the location and intensity of precipitation generation tied to the ITCZ, Trade Winds, Subtropical Highs, monsoons, mid-latitude cyclonic storms, and orographic forcing. Atmosphere circulation changes also redirect and modify air masses, altering temperature and moisture regimes.

Third, changes in ocean circulation alter the position, dynamics, and physical attributes of mid-ocean gyres, boundary currents, upwelling and downwelling zones, and coastal bioclimates. In addition to circulation changes, higher sea surface temperatures and greater atmospheric water vapor content alter air–sea energy transfers. These affect locations for (1) the development of tropical and mid-latitude cyclonic storms and (2) centers of action for major interannual–interdecadal climate variability modes (e.g., ENSO and NAO) and resulting teleconnections.

Fourth, with higher temperatures in hot climates, the atmosphere can hold more water vapor, potentially leading to greater precipitation than currently seen in the tropics. As a result, the climate envelope for terrestrial biomes will expand along both temperature and precipitation axes. This expansion will lead to the emergence of novel low-latitude climate regimes.

Together, these dynamics highlight changes in the very nature of regional climates brought on by the Earth’s increasing energy imbalance. This is mediated by the way energy redistribution is carried out by the climate system through global atmospheric and ocean circulation and energy, mass, and momentum exchanges among Earth System components. Regional climates will become unrecognizable in terms of, for example, altered frequency distributions of weather events and modes of climate variability, reorganization of relationships among bioclimatic variables, and the redistribution and recharacterization of air and water masses.

7.3. Biome and Socioecologic Futures

The linkages between biomes and global circulation dynamics and thermodynamics mean that changes in the location and intensity of circulation features will have strong effects on biome distribution and ecology. Observed planetary shifts are altering terrestrial climate regimes and oceanic physical regimes to be outside of corresponding envelopes of natural variability. This puts at risk biosphere elements that have evolved to function within a safe operating space tied to the natural range of variability. At risk are the structure and function of species populations, ecosystems, seascapes, and landscapes [17,24,80,362,363,364,365]. That is, in a biospheric context, climate has become nonstationary. We are seeing the emergence of novel physical environments, biotic communities, and species interactions [15,94,365,366,367]. Broken, tight climate-dependent relationships among species are early warning signs of system disruption, as observed for co-evolved species in phenological synchrony [161,364,368].

This environmental instability will have socioeconomic consequences, including for hard and soft infrastructure and ecosystem services [51,55,56,57,58,59,94,369]. Current water resource and emergency management infrastructure, for example, is already poorly equipped for handling megadroughts and extreme flooding [370]. In human history, times when regional temperatures ranged outside corresponding envelopes of normal variability were periods of ecological disruption, stressed agricultural production, and social strife from resource competition, leading to inflation, famine, warfare, and population decline [63,371].

Climate and the biosphere—and their capacity to provide ecosystem services—are undergoing rapid change, with new equilibria expected to be long in coming. Given the Earth System’s ongoing energy imbalance [1,372], signs of unprecedented change [39,42,44,45,46], and the threat of tipping cascades being triggered by anticipated climate change [60,353], the climate system and biosphere may already be committed to a path of transformation [21,67,373]. The low likelihood of climate stabilization under even the most ambitious GHG-reduction pathways points to broad-scale biospheric reorganization by the end of this century. Such bioclimatic instability and biospheric restructuring are not readily predictable due to the complex and highly connected nature of the Earth System [374]. Consequently, biome structural and functional changes will be unanticipated, if not counterintuitive [16].

7.4. Action Under Uncertainty

Global biospheric change—characterized by rapid change, instability, emerging novel conditions, and hard-to-predict future dynamics—poses a challenge for biodiversity conservation, resource management, and socioeconomic adaptation [55,63]. Such efforts need to employ frameworks and tools for boosting ecological and socioecological adaptive capacity and handling uncertainty [375]. Two general frameworks for exploring climate change impacts and supporting decision making at local and regional scales are as follows:

Top-down—An impact assessment approach driven “from the top” by global climate model simulations under future socioeconomic scenarios (such as those setting GHG emission pathways). These climate projections may be downscaled by regional climate models (dynamical downscaling) or by statistical methods [376,377,378]. Resulting climate scenarios are subsequently used to drive resource models, including those projecting ecosystem, hydrologic, and socioecological impacts (e.g., [379,380]). These projections can inform planning and decision making at regional and management-unit scales.
Bottom-up—A vulnerability approach assessing climatic resilience and resistance of “at-the-bottom” resources of concern, such as species, landscapes, seascapes, and ecosystem services [78,381,382]. This is an integrative approach that puts these vulnerabilities in the context of other stressors and focuses on local dynamics [381,383]. Local strategies also need to account for cross-scale interactions that tie local processes to regional threshold dynamics, such as wildfires and desertification [374].

A cascade of uncertainties characterizes top-down impact assessments, with each step in the assessment adding uncertainty [384,385]. To quantify this uncertainty, top-down protocols commonly use a suite of climate projections from multiple Earth System Models driven by a range of future emission scenarios (“climate model ensembles”) [81]. An additional approach is to pool results from multiple impact models, including those which vary by subdiscipline (e.g., species and ecosystem ecology), spatial and temporal scales, and method (such as empirical and mechanistic models) [386,387]. In this way, top-down approaches estimate uncertainty that is quantifiable but likely underestimated [384].

Bottom-up approaches acknowledge the limitations of climate and ecological models in capturing competing positive and negative feedbacks in the Earth System, many of which are not well understood or not included in models [11,105,106,110,260]. Instead, bottom-up approaches endeavor to assess uncertainty that might be unquantifiable but is knowable [384]. Beyond these sources of uncertainty is what is unknowable and outside our capacity to foresee [384,385]. Assessments can blend the two approaches [387,388,389].

Key tools for bottom-up vulnerability assessments include those that follow below. These further explore uncertainty or do not explicitly account for uncertainty:

Expert elicitation—Expert synthesis of established knowledge, including indigenous and traditional ecological knowledge, to gain insights into the vulnerability of species and systems to climate change and to devise strategies to reduce or cope with this risk [387,390,391].
Scenario planning—Instead of relying on climate and ecological model projections for systems whose complexities are inherently difficult to simulate, scenario planning uses a framework to envision ecological consequences across a spectrum of probable, as well as less probable but still plausible, futures (including conceivable “surprises”). This explores a range of “what if” climate outcomes from those that are moderately to severely disruptive to species and ecosystems [78,388,392]. Scenario planning can also incorporate paleo- and recent ecological histories to put the future in the context of prior system trajectories [79,389,393].
Least-regrets, win–win, and safe operating space strategies—Actions that in addition to addressing climate vulnerabilities have non-climate related benefits, such as in reducing other environmental threats and promoting socioeconomic well-being [394,395,396]. The reduction of local non-climate stressors (including habitat loss, pollution, and invasive species) works toward keeping systems within their safe operating space in face of difficult-to-control global stressors, such as climate change [30].
Physical integrity of landscapes—For terrestrial systems, a focus on the maintenance and restoration of physical landscape attributes (e.g., connectivity and hydrologic function) to provide for (1) the movement of species when habitats change and (2) the development of intact ecological and evolutionary processes following extreme climate disruption, regardless of what ecosystems may arise there [397,398,399].
Adaptive management and adaptive pathway planning—Adaptive management emphasizes reactive flexibility, relying on monitoring and periodic reassessment to adjust strategies in response to observed changes and new knowledge [389,400,401,402]. Complementary to this, adaptive pathway planning is proactively anticipatory, preparing for multiple ecological trajectories and establishing contingency plans to shift strategies as the future unfolds [389,403].

Resource managers and conservation practitioners are well positioned to implement bottom-up vulnerability-based strategies due to their focus on site to regional scales, their detailed knowledge of species and ecosystems of concern, and the need to act without complete information. In addition, inclusion of local communities in vulnerability-focused management programs can bring mutual benefits [404]. These benefits are strongest when management strategies draw on traditional knowledge and align with a community’s divergent needs and social values [396,405]. Such programs can be designed in concert with local climate adaptation plans and socioeconomic development initiatives [403,406,407].

Author Contributions

Conceptualization and writing—original draft preparation, T.G.F.K.; writing—review and editing, T.G.F.K. and K.F. All authors have read and agreed to the published version of the manuscript.

Funding

This research received no external funding.

Data Availability Statement

The original contributions presented in this study are included in the article.

Acknowledgments

T.K. gives special thanks to James V. A. Conkey, William L. Franklin, Peter J. Richerson, Thomas M. Powell, Roger A. Pielke, Sr., Terri T. Schulz, and Betsy Neely for their insights into global ecology, Earth System dynamics, and biodiversity conservation in the context of climate change. Many thanks to Terri Schulz, Joseph E. Holmes, Mark C. Otto, and Christopher F. Knud-Hansen, and three anonymous referees for their valuable comments on the manuscript. Many thanks to Chad Stouffel for IT support.

Conflicts of Interest

The authors declare no conflicts of interest.

Appendix A. Glossary

The first use of a term in the text starting with Section 2 is given in italics.

Air–sea interactions—The flux (transfer) of energy, mass, and momentum between the ocean surface and overlying atmosphere. Energy fluxes include net radiation and the exchange of sensible and latent heat energy (defined below). Mass fluxes include those of water vapor, other gases (e.g., CO₂ and O₂), aerosols (e.g., sea spray), and particulate deposition. Momentum flux is transferred by wind to the water surface (see “Wind stress”).

Anticyclonic—Rotation that is clockwise in the Northern Hemisphere and anticlockwise in the Southern Hemisphere. In the atmosphere, flow around a high-pressure center is anticyclonic.

Bioclimate—Attributes of surface climate that directly control biotic processes.

Biome—The coarsest level of ecological organization within the biosphere (see Box 2 and Box 4).

Boundary currents—Ocean surface currents located on the eastern and western sides of ocean basins. In the subtropical gyres, they flow equatorward on the eastern sides of oceans and poleward on the west. Directions are reversed in the subpolar gyres.

Brine rejection—The process of salt being expelled as ice crystals form during sea-ice formation. This increases the salinity of surrounding sea water.

Carbonate equilibrium—The chemical equilibrium of CO₂ dissolved in water: [CO₂ + H₂O] ⇌ H₂CO₃ ⇌ HCO₃⁻ ⇌ CO₃²⁻. Collectively, these chemical species are known as dissolved inorganic carbon (DIC).

Chaotic behavior—Chaotic systems are governed by deterministic processes (that is, laws of physics) but exhibit seemingly random behavior. Outcomes are sensitive to initial conditions. Examples are weather systems, global circulation dynamics, and modes of climate variability.

Circulation dynamics—The effect of geophysical forces or apparent forces on the motion of the atmosphere or ocean. These include those from pressure gradients, density gradients (e.g., of temperature and, in the case of the oceans, also of salinity), and the Coriolis effect (defined below). Thermodynamics (defined below) and mechanical forcing (e.g., from flow intercepting mountain ranges or sea mounts) also play a role.

Coriolis effect—The effect of the planet’s rotation on air and water in motion. This apparent force turns winds and currents to the right in the Northern Hemisphere and to the left in the Southern Hemisphere. It is called an apparent force because no physical force acts on a moving air or water mass; instead, the observed deflection results from the Earth’s rotation beneath it. It is effective at length and timescales that are long enough for the Earth’s rotation to play a role (timescales on the order of ≥18 h) and is a function of latitude, increasing with distance from the equator.

Cyclonic—Rotation that is anticlockwise in the Northern Hemisphere and clockwise in the Southern Hemisphere. In the atmosphere, flow around a low-pressure center is cyclonic.

Dynamic global vegetation model (DGVM)—See Box 2.

Earth System—Earth System components are the atmosphere, oceans, lithosphere, cryosphere, and biosphere. Dynamics of and interactions among components determine climatic, biogeophysical, and biogeochemical processes [408]. Lithosphere elements that interact with other Earth System components include the land surface to bedrock and ocean sediments and basement.

Earth System Model (ESM)—An advancement from global climate models (GCMs), ESMs include both physical climatic, biogeophysical, and biogeochemical (e.g., carbon cycle) processes.

Ekman transport—See Box 5.

Envelope of natural (or normal) variability—A defined range of environmental variability that governs ecosystem processes, where “normal variability” is stationary in time and “natural variability” is that prior to anthropogenic forcing. In this paper, natural variability and normal variability are equally taken to be both stationary and applied to the pre-industrial era. These may be defined for specific domains, such as globally or for regions. They are also referred to in the literature as “historical range of variability” or similar [29,409]. Ecosystem processes include biological and physical (e.g., hydrologic) dynamics.

Extratropical—Outside of the tropics, generally referring to 30–90° N/S latitudes. It is alternatively taken to encompass temperate, subpolar (boreal), and polar regions.

Flux—Transfer of an entity across a unit surface area per unit time. See “Air–sea interactions.”

Front—A boundary in the atmosphere or oceans separating contrasting air or water masses, respectively. The contrast may involve temperature, atmospheric moisture, sea water salinity, or other physical attributes.

Functional type—Species that perform similar functions in an ecosystem, independent of their taxonomy. It is generally defined in terms of a specific ecological function, such as a role in trophic webs or biogeochemical dynamics.

Geophysical fluid dynamics—See “Circulation dynamics.”

Gyre (oceanic gyre)—A large-scale, generally circular pattern of wind-driven currents in ocean basins.

Heat engine—A system in which kinetic motion is driven by the release of thermal energy.

Hysteresis—A system property where the state of a system is path dependent, meaning it depends on its history and is not uniquely determined by its current forcing.

Latent heat (of transformation)—Energy related to water phase changes, such as heat absorbed or released during evaporation and condensation, respectively. Also, heat absorbed or released during ice melting and freezing.

Legacy effect—Persistent influence of past conditions.

Meridional—Following lines of longitude and, therefore, crossing lines of latitude. Winds moving north or south are said to be “meridional.” The opposite of “zonal.”

Mid-latitude—Generally pertaining to 30–60° N/S latitudes, although broader definitions are not uncommon. In some parts of the globe, mid-latitude atmosphere and ocean dynamics extend to 70° latitudes or higher.

Niche (ecological niche)—A multivariate environmental domain that is suitable for a species to persist. The full domain (fundamental niche) may or may not be available to and occupied by the species at any point in time (realized niche). An available environment may not be occupied owing to, for example, dispersal limitations and competitive and trophic interactions with other species.

Nonlinearity—A system property where changes in output are not proportional to changes in inputs. Some nonlinearities are simple and readily modeled, for example with logarithmic and exponential functions. On the other hand, in highly interconnected systems, nonlinear behavior arises from complex feedbacks among components. These feedbacks can be positive (amplifying) or negative (stabilizing). Such systems can be self-organizing, exhibiting emergent properties—characteristics not directly derived from forcings. These include complex structures and dynamics. Under altered forcing, their behavior can be counterintuitive and difficult to predict. In the Earth System, nonlinear dynamics range from stability under strong forcing (resilience and resistance) and oscillations to teleconnections, hysteresis, chaotic behavior, and tipping points (abrupt transitions between equilibrium states) [22,25]. (See also “Earth System” and glossary entries for the last four terms).

Nonstationarity (in time)—A random process whose statistical properties vary with time. The opposite of stationarity (defined below).

Overturning circulation—In the atmosphere, a thermally forced vertical circulation with rising motion where heating dominates and sinking where cooling dominates. In the oceans, an overturning circulation is driven by density, sinking where temperatures are coldest and salinity is greatest.

Pycnocline—A vertical density gradient in ocean waters due to differences in temperature and salinity. See “Stratification”.

Safe operating space—The range of environmental stressors within which a biotic system (e.g., species population, community, or ecosystem) is sustainable. System stressors may act independently or synergistically with accumulative or compounding effects, respectively. Reducing one stressor can increase a system’s capacity to adapt to another [30].

Sensible heat—Thermal energy transferred from one body to another through conduction.

Socioecological system—Linked ecological and socioeconomic systems that control material and energy flows in the Earth System. Includes human dimension factors such as culture, technology, behavioral psychology, and societal inertia [103,410,411].

Stationarity (in time)—A random process whose statistical properties (e.g., its mean, variance, and autocorrelation structure) remain constant over time. Stationarity can also include stability in cyclic or quasiperiodic processes. In this paper, we consider climatic stationarity to encompass natural variability (without anthropogenic forcing) at daily, seasonal, multidecadal, and centennial to millennial timescales (see “Envelope of natural variability”).

Stratification—Ocean stratification is the layering of seawater caused by vertical changes in density, where density is a function of temperature and salinity. A less-dense surface layer overlies denser deep water and is typically on the order of 200 m deep. A pycnocline (vertical density gradient) separates the two layers. The seasonal persistence of stratification depends on the density gradient and strength of winds acting on the water surface. Stratification limits the depth of mixing, nutrient availability for primary production, dissolved oxygen content at depth, and carbon sequestration.

Synoptic scale—In meteorology, a length scale ranging from several 100 to several 1000 km, with a corresponding timescale of one to several days. It often refers to the scale of mid-latitude cyclonic storms.

Teleconnections—Connections between atmospheric and ocean dynamics in one region and climate variability in another remote region. These generally operate at intra- and interdecadal scales and are commonly linked to climate system modes of variability, such as El Niño–Southern Oscillation (see Box 4).

Thermally direct circulation—See “Overturning circulation.”

Thermodynamics—The physics of converting one form of energy to another, such as between thermal and kinetic energy. This also includes various forms of thermal energy exchange, such as sensible and latent heat (defined above).

Thermohaline—Where both temperature and salinity contribute to the density of seawater.

Tipping element (tipping point, tipping cascade, and core global tipping point)—A tipping element is a part of the Earth System capable of multiple stable states. With sufficient forcing, a tipping element crosses a threshold (tipping point), beyond which its current equilibrium behavior is lost. At that point, amplifying processes rapidly shift regional or hemispheric climates to a new stable state. The shift may be essentially irreversible. “Essentially” means that such a change in state is likely irreversible on human timescales. Tipping processes are hysteretic, such that, if reversible, the return to the initial state follows a different path. The response domain may extend far beyond the element itself due to dynamically linked tipping elements (a tipping cascade). In the context of increasing surface temperatures, a core global tipping point is one that triggers other tipping elements through a cascade of tipping elements, which then act together in positive feedback loops to increase global surface temperatures [25,77,353].

Tropopause—The upper boundary of the troposphere, where the atmosphere transitions to the stratosphere (see next).

Troposphere—The lower atmospheric layer, approximately 7–18 km thick, extending from the surface to the tropopause. Most weather affecting surface climate occurs within this layer.

Wind stress—Wind acting on a water surface imparting momentum from the atmosphere to the ocean (or other water body), driving surface currents and vertical mixing.

Zonal—Following lines of latitude. The opposite of “meridional.” Winds moving west or east are said to be “zonal”.

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Figure 1. (a) Summer (June, July, and August) land temperature anomalies for extratropical Northern Hemisphere (30–90° N) for 1–2023 CE. Anomalies are relative to the 1850–1900 mean. Yellow = mean ensemble reconstructed record (1–2023), red = instrumental record 1850–2023, and gray = 95% range of uncertainty of reconstructed record based on the variance among ensemble members. Blue and red bars above the x-axis indicate generalized time spans of cold and warm Northern Hemisphere climate anomaly periods, respectively, based on refs. [36,37,38] (LALIA = Late Antique Little Ice Age). These periods are shown for comparison with the recent period (1850–2023). (b) Comparison of the frequency distribution of temperature anomalies for the recent record (red, 1850–2023) and the past ~2000 years (yellow). Exceptionally cold and warm summers are marked with black vertical dashed lines labeled by year. The blue-to-red arrow and corresponding vertical dashed lines indicate a shift in the period means from 1–1849 (blue) to 1850–2023 (red) (based on text in [39]). [Adapted from Esper et al. [39]; reproduced with permission from SNCSC (Springer Nature Customer Service Center)].

Figure 3. Mean global circulation of the atmosphere in January (a) and July (b). H and L represent semi-permanent high- and low-pressure centers, respectively. Gray contours show sea-level pressure (hPa), red arrows are prevailing near-surface winds, and the purple line is the mean position of the ITCZ (or, in (b), the summer Monsoonal Front over land in Africa and Asia). The latitudinal extent of the ITCZ is typically 5 to 15° broader than its mean position. Dark blue lines are the mean position of wintertime Mid-latitude or Polar Jet Streams (derived from [125]). [Adapted from The Atmosphere: An Introduction to Meteorology, 14th ed., Lutgens et al. [126], © 2019 Pearson Education, Inc.; used with permission].

Figure 4. (a) Elements of equator-to-pole heat energy transfer by global atmospheric circulation shown in meridional (latitudinal) cross-section for the Northern Hemisphere winter. The Hadley Cell, Intertropical Convergence Zone (ITCZ), Subtropical High (STH), Mid-latitude Jet Stream, Polar High, and near-surface zonal winds (easterlies and westerlies) are indicated, as are Arctic and Polar Fronts. Elements are idealized and highly exaggerated in the vertical direction. (b) Mean distribution of tropospheric air masses over North America and neighboring oceans and corresponding fronts (blue lines with triangles). Air masses are continental Arctic (cA), continental polar (cP), maritime polar (mP), continental tropic (cT), and maritime tropic (mT). Referring to the latter two air masses as “tropic” comes from that they lie along the Tropic of Cancer; however, these are subtropical in nature (their sources being largely between 15 and 30° N latitude). While “polar” air masses (mP and cP) are named so by convention, these are mid-latitude or subpolar in nature, with their sources generally from 30 to 60° N latitude. For the rest of the paper, we use terms for air masses that are consistent with their source region (i.e., subtropical, mid-latitude, subpolar, and polar for Arctic/Antarctic air). The Polar Front separates subpolar air (cP) and mid-latitude air (largely mP) from warm subtropical air masses (mT, cT), whereas the Arctic Front separates Arctic air (cA) from subpolar/mid-latitude air. [(a) Adapted from Barry and Chorley [125]; used with the permission of Taylor & Francis Informa UK Ltd.—Books, with permission conveyed through Copyright Clearance Center, Inc. Their figure is adapted from Palmén [127]; used here with permission from the John Wiley and Sons/Copyright Clearance Center. (b) Adapted from: https://commons.wikimedia.org/wiki/File:Airmassesorigin.png. Accessed 24 October 2024. Original source: U.S. National Oceanic and Atmospheric Administration. Public domain].

Figure 5. Full-disk snapshot of atmospheric water vapor, illustrating the complex nature of air masses in terms of moisture content. The image is over the Atlantic Ocean, western Europe, and Africa (obscured) for 2 September 2010. Bright areas are tops of deep cumulus convection in the ITCZ with high water vapor, whereas dark areas are of dry air. [Image ©2010 EUMETSAT: https://earthobservatory.nasa.gov/features/Water. Accessed on 7 July 2024. Used under Creative Commons license CC BY-SA 3.0 IGO.].

Figure 6. Global wind-driven ocean surface currents. Warm currents are in red, with cold in blue. Gyres with abbreviated labels include the North Pacific (NP-SPG) and North Atlantic (NA-SPG) Subpolar Gyres. Abbreviated labels for currents include the North (N Eq C) and South (S Eq C) Equatorial Currents, Equatorial Countercurrent (Eq CC), North Pacific Current (N Pac C), and Norwegian Current (NC). The NA-SPG includes the North Atlantic Current (N Atl C) on the south and east, the Irminger Current (IC) to the north, and the East Greenland (EGC) and Labrador (Lab C) Currents on the northwest and west, respectively. In the Southern Ocean, the West Wind Drift is also called the Antarctic Circumpolar Current (ACC), and the East Wind Drift along the continent’s coast is also known as the Antarctic Subpolar Current. [Adapted from: https://www.noaa.gov/jetstream/ocean/circulations/jetstream-max-major-ocean-currents. Public Domain. Accessed 12 December 2024.].

Figure 7. Global distribution of terrestrial biomes. In the Northern Hemisphere, the Cold Desert (gray) and the Lichen and Moss Tundra (light purple) are also referred to as the High Arctic, whereas the Dwarf Shrub and Meadow Tundra (medium purple) is the Low Arctic. Coniferous Forest–tundra is also known as boreal woodland. Taiga is boreal forest. The Siberian boreal region in cyan and labeled “B(d)” in red is Deciduous Taiga (deciduous needleleaf boreal forest, dominated by Larix species), in contrast to Evergreen Taiga (evergreen needleleaf boreal forest) to the west in Eurasia (in dark cyan). Mixed Forest is mixed cold-deciduous broadleaf and evergreen needleleaf forest (i.e., temperate–boreal transitional forest). Subtropical Moist Forest (bright blue) includes warm-temperate evergreen forests (e.g., in southeastern United States and China). For the most part, Sclerophyllous Vegetation (pink) corresponds to Mediterranean-type biomes. [Adapted from: Ökologix, https://commons.wikimedia.org/wiki/File:Vegetationszonen.png; used under Creative Commons CC BY-SA 3.0 license (modifications: “B(d)” labels added). Accessed on 3 June 2022].

Figure 9. Global climatologies of (a) annual mean daily precipitation and (b) annual surface temperature (sea surface temperatures over oceans and near-surface air temperatures over land and sea ice). In (b)’s color bar, note that the narrow yellow and cyan bands accentuate gradients between roughly 15–17 °C and 2–4 °C, respectively, relative to smoother gradients in other parts of the color bar. [(a) Courtesy of H. Yang, Alfred Wegener Institute. (b) Adapted from http://berkeleyearth.org/archive/land-and-ocean-data/. Accessed on 16 August 2022. Used with permission from Berkeley Earth, licensed under Creative Commons BY-NC 4.0 International].

Figure 10. Global near-surface (10 m) wind regimes over oceans and land (1980–2009 mean). (a) Point clouds illustrating four major wind regimes over land (i–iv) based on wind speed (distance from a plot’s center point) and consistency in wind direction (indicated by the distribution of points around the compass graphic). Prevailing wind direction is indicated by the radial line on the compass graphic. (b) Mapped distribution of points belonging to the four wind regime classes in (a) by color on a plot of directional anisotropy (consistency in wind direction) versus wind speed. (c) Prevailing near-surface winds represented by streamlines, with white arrow heads indicating direction. Where streamlines converge, there is upward motion, and where they diverge, sinking motion. Streamline color over land indicates wind regime in (a,b). [Image from Kling and Ackerly [211]; used with permission of Springer Nature BV, conveyed through Copyright Clearance Center, Inc.].

Figure 14. The Thermohaline Circulation. Near-surface currents (red lines) flow towards four main deep- and bottom-water formation regions (yellow ovals) in the northern North Atlantic (Labrador and Nordic Seas) and the Southern Ocean (Ross and Weddell Seas). Deep-water (blue lines) and bottom-water (purple) currents circulate at depth, eventually becoming near-surface water due to upwelling and a reduction in density (from warming and freshwater inputs [253]). Regions with high sea surface salinity (above 36‰) are indicated with green shading and low salinity (below 34‰) with blue shading. [Adapted from Rahmstorf [239]; reproduced with permission from SNCSC].

Figure 15. Map of The World Ocean depicting the oceans as a single body, through which the Thermohaline Circulation circulates (Figure 14). The projection is the Spilhaus One World Ocean coordinate system centered over Antarctica [254]. Blue shading indicates ocean floor bathymetry. [Adapted from [255]. Used with permission].

Figure 17. (a) Global sea surface temperature change over the record 1900–2017 (units = °C/117 years) based on interpolated observations. (b) North Atlantic sea surface temperature trend 1993–2018 (°C/year) based on satellite observations. In (b), in the core of the cold anomaly, the recent 25 year trend is on the order of −0.1 °C/year (dark blue). [(a) From https://www.climate.gov/media/13692. Accessed 3 November 2024. Public domain. (b) From DG MARE [349]; used under a Creative Commons CC BY 4.0 license (modification: legend inserted).].

Figure 18. Interactions among direct and indirect impacts of climate change on coastal biome foundation species. Also indicated are accumulative and synergistic interactions with non-climatic environmental stressors. Yellow circles with an “!” indicate extreme-event effects. [From Wernberg et al. [271]; used under a Creative Commons CC BY 4.0 license].

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Kittel, T.G.F.; Ferron, K. Nonlinear Earth System Dynamics Determine Biospheric Structure and Function: I—A Primer on How the Climate System Functions as a Heat Engine and Structures the Biosphere. Climate 2026, 14, 38. https://doi.org/10.3390/cli14020038

AMA Style

Kittel TGF, Ferron K. Nonlinear Earth System Dynamics Determine Biospheric Structure and Function: I—A Primer on How the Climate System Functions as a Heat Engine and Structures the Biosphere. Climate. 2026; 14(2):38. https://doi.org/10.3390/cli14020038

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Kittel, Timothy G. F., and Kelly Ferron. 2026. "Nonlinear Earth System Dynamics Determine Biospheric Structure and Function: I—A Primer on How the Climate System Functions as a Heat Engine and Structures the Biosphere" Climate 14, no. 2: 38. https://doi.org/10.3390/cli14020038

APA Style

Kittel, T. G. F., & Ferron, K. (2026). Nonlinear Earth System Dynamics Determine Biospheric Structure and Function: I—A Primer on How the Climate System Functions as a Heat Engine and Structures the Biosphere. Climate, 14(2), 38. https://doi.org/10.3390/cli14020038

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Nonlinear Earth System Dynamics Determine Biospheric Structure and Function: I—A Primer on How the Climate System Functions as a Heat Engine and Structures the Biosphere

Abstract

1. Introduction

1.1. Key Questions

1.2. Review Objectives

1.3. Terms Defined

2. Biosphere Stability and Earth System Dynamics

2.1. Stationary Envelope of Normal Variability

2.2. System Reorganization

2.3. Climate Stabilization

3. The Planetary Heat Engine and Global Atmospheric and Oceanic Circulation

3.1. The Climate System as a Heat Engine

3.2. Global Atmospheric and Ocean Circulation Driven by the Planetary Heat Engine

3.2.1. Planetary Circulation Systems

3.2.2. Global Heat Energy Transfers Accomplished by Planetary Circulation Systems

4. Earth System Dynamics and Terrestrial Biomes

4.1. Terrestrial Biomes and Mean Climate

4.1.1. Nonstationarity of Biome–Mean Climate Relationships

4.1.2. Climate Nonstationarity and Empirical Ecological Models

4.2. Phase-Change Thermodynamic Constraints on Terrestrial Biomes

4.2.1. Precipitation Regime Expansion Under Increasing Temperatures

4.2.2. Uneven Shape of the Climate Space Envelope

4.3. Planetary Atmospheric Circulation, Climate Regimes, and Terrestrial Biomes

4.3.1. Climate Regimes

4.3.2. Biomes and Circulation Features

5. Earth System Dynamics and Ocean Biomes

5.1. Wind-Driven Ocean Circulation Systems and Biomes

5.1.1. Physical Regimes

5.1.2. Biomes, Primary Production, and Wind-Driven Circulation Features

5.2. The Thermohaline Circulation

5.3. Ocean Carbon Pumps

5.4. Coastal Biomes

6. Changing Climate Dynamics

6.1. Changing Atmospheric Circulation

6.2. Changing Ocean Dynamics

6.2.1. Physical Regime

6.2.2. Marine Biome Change Linked to Changing Ocean Climate Dynamics

6.2.3. Atlantic Meridional Overturning Circulation (AMOC) Climate Sensitivity

6.2.4. Coastal Biome Climate Sensitivity

7. Summary and Discussion: Main Lessons for Conservation Planning and Resource Management

7.1. Overview of Earth System Dynamics Driving Biome Distributions

7.1.1. Terrestrial Biomes and Atmospheric Dynamics

7.1.2. Marine Biomes and Ocean Dynamics

7.2. The Dynamics of Change

7.3. Biome and Socioecologic Futures

7.4. Action Under Uncertainty

Author Contributions

Funding

Data Availability Statement

Acknowledgments

Conflicts of Interest

Appendix A. Glossary

References

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