Earth-Science Reviews 215 (2021) 103503 ELSEVIER Contents lists available at ScienceDirect Earth-Science Reviews journal homepage: www.elsevier.com/locate/earscirev I'ARTII-SCIi'NCK Invited review Phanerozoic paleotemperatures: The earth's changing climate during the last 540 million years Christopher R. Scotese a'*, Haijun Songb, Benjamin J.W. Mills0, Douwe G. van der Meerd a Department of Earth and Planetary Sciences, Northwestern University, Evanston, IL 60208, USA b State Key Laboratory ofBiogeology and Environmental Geology, School of Earth Sciences, China University of Geosciences, Wuhan 430074, USA c School of Earth and Environment, University of Leeds, Leeds LS2 9JT, UK d CNOOC International, 945 Bunker Hill Road, Houston, TX 77024, USA (D ARTICLE INFO ABSTRACT Keywords: paleoclimate paleotemperature Phanerozoic climate change climate history ice age icehouse hothouse Hirnantian Ice Age Permo-Carboniferous Ice Age End Triassic Extinction K/T Extinction K/T Impact Winter PETM Pleistocene Ice Age Future Global Warming a18o a13c This study provides a comprehensive and quantitative estimate of how global temperatures have changed during the last 540 million years. It combines paleotemperature measurements determined from oxygen isotopes with broader insights obtained from the changing distribution of lithologic indicators of climate, such as coals, evaporites, calcretes, reefs, and bauxite deposits. The waxing and waning of the Earth's great polar icecaps have been mapped using the past distribution of tillites, dropstones, and glendonites. The global temperature model presented here includes estimates of global average temperate (GAT), changing tropical temperatures (AT° tropical), deep ocean temperatures, and polar temperatures. Though similar, in many respects, to the temperature history deduced directly from the study of oxygen isotopes, our model does not predict the extreme high temperatures for the Early Paleozoic required by isotopic investigations. The history of global changes in temperature during the Phanerozoic has been summarized in a "paleotemperature timescale" that subdivides the many past climatic events into 8 major climate modes; each climate mode is made up of 3-4 pairs of warming and cooling episodes (chronotemps). A detailed narrative describes how these past temperature events have been affected by geological processes such as the eruption of Large Igneous Provinces (LIPS) (warming) and bolide impacts (cooling). The paleotemperature model presented here allows for a deeper understanding of the interconnected geologic, tectonic, paleoclimatic, paleoceanographic, and evolutionary events that have shaped our planet, and we make explicit predictions about the Earth's past temperature that can be tested and evaluated. By quantitatively describing the pattern of paleotemperature change through time, we may be able to gain important insights into the history of the Earth System and the fundamental causes of climate change on geological timescales. These insights can help us better understand the problems and challenges that we face as a result of Future Global Warming. 1. Introduction There have been several recent studies of the Earth's changing temperature during the Phanerozoic, e.g. Grossman (2012a&b); Veizer and Prokoph (2015); Scotese (2016); Mills et al. (2019); Henkes et al. (2018) ; Wing and Huber (2019); Verard and Veizer (2019); Song et al. (2019) ; Grossman and Joachimski, 2020). All of these studies are based on oxygen isotope measurements of 8180 in carbonate or apatite mi-crofossils. The first extensive compilation of oxygen isotope data assembled by Veizer and Hoefs (1976) and Veizer et al. (1986, 1999, 2000) has been used by other researchers to characterize temperature change at a variety of temporal and geographic scales (Royer et al., 2004; Prokoph et al., 2008; Grossman, 2012 a&b). Recently, two comprehensive, global Phanerozoic compilations of oxygen isotope paleotemperatures have been assembled (Song et al., 2019 and the PALects database, Grossman et al., 2018). Urey et al. (1951) and Emiliani (1955) were the first scientists to appreciate the utility of oxygen isotopes. For every 4.3° C increase in temperature, there is a 0.1% decrease in the amount of lsO (Epstein and Mayeda, 1953) used to produce the calcium carbonate of the forami-nifera's shell. By measuring the ratio of lsO to leO, also referred to as 8180, it is possible to estimate the temperature at which the calcium carbonate was produced. Savin et al. (1975), Savin (1977), Miller et al. (1987), Zachos et al. (2001, 2008) and Cramer et al. (2009) used the * Corresponding author. E-mail addresses: cscotese@gmail.com (C.R. Scotese), haijunsong@cug.edu.cn (H. Song). https://doi.Org/10.1016/j.earscirev.2021.103503 Received 14 September 2020; Received in revised form 10 December 2020; Accepted 4 January 2021 Available online 11 January 2021 0012-8252/Crown Copyright © 2021 Published by Elsevier B.V. All rights reserved. CR. Scotese et al. Earth-Science Reviews 215 (2021) 103503 changing ratios of lsO /leO in benthic foraminifera to describe the changing temperature of the world's oceans and the growth of the polar icecaps. For an excellent summary of the science of oxygen isotopes see Grossman (2012a). The oxygen isotope record is strongly biased towards tropical regions and measurements of temperature derived from localities above 40° N or S are sparse; therefore, paleotemperature curves that exclusively use 5lsO data (Song et al., 2019; Verard and Veizer, 2019; Veizer and Pro-koph, 2015; Grossman, 2012a&b; Royer et al., 2004 ) are measurements of tropical temperature, not global temperature. Only the Phanerozoic paleotemperature curves of Scotese (2016), Mills et al., (2019, 2020), Wing and Huber (2019), and Valdes et al. (2018) (Fig. 1), either adjust the isotopic temperatures to compensate for the difference between tropical and global temperatures or use non-isotopic temperature information (e.g., HadCM3 paleoclimate model simulations, Valdes et al., 2020). These curves are in good agreement and show the classic "double hump" Phanerozoic temperature history shown in Fig. 2 (Fischer, 1981, 1982, 1984; Frakes et al., 1992, fig. 11.1; Scotese et al., 1999; Sum-merhayes, 2015). Temperatures are high during the Early Paleozoic. Cooler temperatures prevail during the Late Paleozoic, followed by warmer Mesozoic and early Cenozoic temperatures, finally returning to cooler temperatures in the Late Cenozoic. This pattern is linked to the "Wegener Supercontinent" cycle (Nance et al., 2014; Van der Meer et al., 2014, 2017), namely: the breakup of the supercontinent, Pannotia, during the latest Precambrian (Powell et al., 1993; Powell, 1995; Scotese, 2009; Nance and Murphy, 2018), the formation of Pangea during the late Paleozoic (~300 Ma), its subsequent breakup (~200 Ma), and the assembly of the modern continental configuration. Specific features that the curves in Fig. 1 share in common are: 1) A broad sinusoidal pattern with peaks near 20 Ma, 50 Ma, 90 Ma, 245 Ma, 370 Ma, 420 Ma, and 500 Ma, with the highest peaks occurring at, 90 Ma, and 245 Ma. CLIMATE a V V V CAM ORD SI DE CAR. PER TR JUR CRET. CEN a Fig. 2. "Double Hump" pattern of Phanerozoic Climate (Fischer, 1981, 1982), I = icehouse, G = greenhouse. 2) Broad troughs link the 370 Ma and 245 Ma peaks and the 245 Ma and 90 Ma peaks. 3) The three youngest peaks (i.e. 90 Ma, 50 Ma, and 20 Ma) decrease, Phanerozoic Global Average Temperature 48" - Valdes etal. (2018) after Mills et al. (2019) " Wing and Huber (2019) vt c 540 520 500 480 460 440 420 400 380 360 340 320 300 280 260 240 220 200 180 160 140 120 100 80 60 40 20 0 millions of years Fig. 1. Estimates of Phanerozoic Global Average Temperature (GAT). Sources: Wing and Huber (2019), Valdes et al. (2018), Mills et al. (2019), and This study. 2 CR. Scotese et al. Earth-Science Reviews 215 (2021) 103503 stepwise, in height. 4) The coherence of the curves breaks down during the middle and early Paleozoic. Some curves rapidly trend toward higher isotopic temperatures. 5) The temperatures between 330 Ma and 290 Ma (Permo-Carbon-iferous Ice Age), are ~14°C on average, but fluctuate 6°C with a period of —10 million years. 6) A sharp "dip" occurs at in all curves at 445 Ma (Hirnantian Ice Age). The pre-Carboniferous portions of these temperature curves can be divided into two groups: the "hothouse" middle and early Paleozoic versions (Vaides et al., 2018; this study), and the "extreme hothouse" middle and early Paleozoic versions (Mills et al., 2019; Wing and Huber, 2019). It should be emphasized that the paleotemperature curve that we will present here is a "model" of Phanerozoic temperature change. No one line of evidence, such as, lithologic indicators of climate, oxygen isotopic measurements, or CO2 proxy information (Foster et al., 2017) is sufficient to estimate past temperatures. We combine multiple, independent lines of evidence to produce a synthetic, multi-disciplinary model that describes how global temperature may have varied during the last 540 million years. The temperature predictions made by this model can be used to form hypotheses that can be tested by the acquisition of new 8180 temperature data, by other independent measurements of paleotemperature (e.g.,Tex 86 or clumped isotope data), as well as the expected effects of temperature on the evolution of life and the environment. 2. Methods 2.1. Derivation of the Phanerozoic Global Average Temperature Curve Our Phanerozoic Global Average Temperature Curve has been produced by combining: 1) estimates of the changing pole-to-Equator temperature gradient obtained from lithologic indicators of climate (i. e. paleo-Koppen belts), and 2) estimates of tropical changes in temperature obtained from oxygen isotopes. The temperatures obtained from oxygen isotope measurements have been summarized and sometimes temperatures have been modified to better agree with the geological and paleontological record. We only schematically illustrate the extremely rapid (<1 million years), high amplitude temperature fluctuations that occurred during the coldest intervals. These dramatic temperature fluctuations are triggered by variations in insolation due to changes in the three Milanko-vitch parameters (obliquity, precession, and eccentricity). The effect of on short term climate are discussed in more detail in Kump et al. (1999c), Ruddiman (2001), and Hay (2016). 2.2. Estimates of Pole-to-Equator Temperature gradients derived by mapping Kbppen Climate Belts Long-term global temperature (>50 million years) is controlled by multiple tectonic and environmental processes that drive the Earth's climate from icehouse to hothouse conditions and vice-versa. The dominant factors are the level of greenhouse gases (principally CO2 ), which are regulated by volcanic outgassing and the draw down of CO2 due to weathering (van der Meer et al., 2014; Torsvik et al., 2020) the geographic configuration of the continents and ocean basins (paleoge-ography and paleoceanography), and the reflectivity of the Earth's surface (albedo). Many of these factors are interconnected by a complex network of positive and negative feedback loops that can accelerate or slowdown changes in long-term global temperature (Hay, 2016; Ruddiman, 2001). In this paper we use the geologic record, specifically the paleogeo-graphic distribution of lithologic indicators of climate to map how the Earth's Pole-to-Equator temperature gradient has changed through time. The expansion, contraction, and shifting position of the Earth's climatic belts provides important insights into the changes in the Earth's long-term climate history. Using modern temperature and rainfall records, we can map five very distinct climatic regions called "Koppen Climate Belts" (Fig. 3). The Koppen Climate Belts are defined by seasonal variations in temperature and precipitation (Koppen, 1918). These variations give rise to regional climates and create the mosaic of diverse environments that cover the Earth. These environments include: (A) tropical rainforests near the Equator, (B) desert belts at subtropical latitudes that transition into (C) warm temperate grasslands and forests. In the modern world, as we move poleward, warm-temperate regions are replaced by (D) seasonally warm/cold temperate regions and (E) finally frigid polar regions. These climatic zones extend over the oceans where they are primarily zonal. Each of these climatic zones is characterized by a distinctive, flora, fauna, land-cover, and geology. Modern Koppen Climatic Belts A - Equatorial Rainy, B - Subtroppical Arid, C- Warm Temperate, D- Cool Temperate, E - Polar Fig. 3. Modern Koppen belts and the average temperature of each belt. 3 C.R. Scotese et al. Earth-Science Reviews 215 (2021) 103503 Fig. 4. Early Permian (280 Ma) lithologic indicators of climate and continental paleo-Koppen Belts (Boucot et al., 2013). Fig. 5. Mid-Cretaceous (100 Ma) lithologic indicators of climate and continental paleo-Koppen Belts (Boucot et al., 2013). Using lithologic indicators of climate such as coals, evaporites, and glacial deposits, it is possible to map "paleo"- Kóppen Climate Belts for ancient time periods (Ziegler et al., 2003; Boucot et al., 2013; Fig.s 4 and 5). Over the past 20 years, a global database of over 8,500 lithologic indicators of climate was assembled (Boucot et al., 2013). For a thorough discussion of both lithologic and biological indicators of climate, see Parrish, 1998; Boucot et al., 2013; and Cao et al., 2019. Other important lithologic indicators of climate are: soil minerals such as bauxite, an aluminum ore which forms in warm, wet climates; calcrete, or caliche, which forms in semi-arid regions; and kaolinite which forms in regions with climates that are sometimes wet and sometimes dry (warm temperate climate belt). Dropstones, like tillites, are important indicators of frozen lakes or sea ice. A glendonite it is a pseudomorph of ikaite, a low temperature, hydrated polymorph of CaC03 that forms at temperatures <4° C. Recent experimental data suggests that it may also form at higher temperatures under special conditions (Tollefsen et al., 2020). The legend inset on Fig.s 4 and 5 summarizes the association of the various lithologic indicators of climate with warm/cool and wet/dry environmental conditions. Bauxites, as a general rule, reflect tropical-subtropical humid, monsoonal conditions. Their modern occurrence is almost entirely restricted to the Equatorial Wet Belt. The occurrence of bauxite deposits in northern Europe and Siberia during the late Jurassic, Cretaceous, Paleocene, and Eocene times (Boucot et al., 2013), is one of the strongest geological indications of warm and wet conditions at high latitudes. When we plot these lithologic indicators of climate on a set of paleogeographic maps we find that there have been times (e.g., Early Permian, Fig. 4) when the Earth's climate was much like the present-day "icehouse" world. An icehouse world is simply defined as a time when the Earth is covered by permanent ice at either pole. For permanent ice 4 C.R. Scotese et al. Earth-Science Reviews 215 (2021) 103503 18 °C Long-term Temperature from paleo-Koeppen maps 540 520 500 480 460 440 420 400 380 360 340 320 300 280 260 240 220 200 180 160 140 120 100 80 60 40 20 0 millions of years Fig. 6. Long-Term Phanerozoic temperature trend calculated by estimating the changing area of paleo-Koppen belts (see Supplementary Materials for data and details of calculations). to accumulate in the polar regions (>67° N & S) the temperatures must remain below freezing during the summer months. In other words, the global average temperature (GAT) must be less than around 18°C and the average annual temperature of the polar region must be below -5 °C. Though the tropics remain relatively warm (26° C) in an icehouse world, the polar regions are frigid (-5°C to - 50°C). There are also have been times when there was no ice above the polar circle - even during the winter (e.g., Late Cretaceous, Fig. 4). During these "hothouse" times, the average temperature of the Earth was generally above 20°C (68°F) and the polar regions were relatively warm (5°C to 15°C) and no ice could accumulate. It is a well-established fact that no polar ice existed during the Paleocene-Eocene Thermal Maximum (55.6 Ma, Mclnerney and Wing, 2011 ) or the Cenomanian-Turonian Thermal Maximum (93 Ma, Ziegler et al., 1985). The long-term temperature curve (Fig. 6) was calculated by mapping past extent of the five major Koppen belts using lithologic indicators of climate (tillites, dropstones, glendonites, high latitude mangroves, palms, and crocodiles; temperate coal, evaporites, calcretes, tropical coals, bauxites, and laterites (Boucot et al., 2013)). These paleoclimatic reconstructions, spaced at intervals of 5 million years, were used to characterize changes in global temperature along a spectrum of climatic states ranging from Severe Icehouse (GAT = 10° - 14° C) to Extreme Hothouse (GAT = 22° - 26° C; Fig. 5). The complete set of more than 100 paleo-Koppen maps illustrating the distribution of lithologic indicators of climate are provided in the Supplemental Materials (Part 3). The average temperatures of the modern Koppen belts was calculated using the global temperature model of Legates and Wilmott (1990). The Equatorial Rainy Belt has a Mean Annual Temperature (MAT) of 26°C; the Subtropical Arid Belt's MAT is 22°C; the Warm Temperate Belt's MAT is 16°; the Cool Temperate Belt's MAT is 5 °C; and the North & South Cold Polar Belts' MAT average -20° C. The Global Average Temperature (GAT) was then estimated by summing the areas of the 5 major Koppen climatic belts on each of the paleoclimatic reconstruction and then multiplying that area by the MAT of that Koppen belt. For example, for the modern world (Fig. 3): the Equatorial Rainy Belt (A) covers 23% of the Earth's surface, the Arid Belt (B) covers 28%, the Warm Temperate Belt (C) covers 20%, the Cool Temperate Belt (D) covers 20%, and the Polar Belt (E) covers 9%. The present-day global average temperature (GAT) = .23*(26°C) + .28*(22°C) +.20* (16°C)+.20*(5°C) + .09 *(-20°C), which simplifies to (5.98°C + 6.16 + 3.2 + 1.0-1.8) °C = 14.5°C. On the mid-Cretaceous paleoclimatic reconstruction (Fig. 5): the Equatorial Rainy Belt covers 25% of the Earth's surface, the Subtropical Arid belt covers 29%, the Warm Temperate Belt covers 44%, the Cool Temperate Belt covers 2%, and the Cold Polar Belt covers 0%. The GAT = .25*(26°C) + .29*(22°C) + .44*(16°C) + .02*(5°C) + 0.0, which gives 20° C. The Boreotropical belt, which is strictly a climatic feature of hothouse worlds is assigned the same average temperature as the Arid Belt. In a similar fashion, the global average temperature for the Early Permian (280 Ma, Fig. 4) was calculated as follows: GAT = .20* (26°C)+.29*(22oC)+.16*(16oC)+.30*(5°C)+.05*(-20oC), which equals 14.6°C . The global average temperature during the early Permian icehouse was similar to the present-day icehouse. Global temperatures were calculated in this manner for 100 Phanerozoic reconstructions of paleo-Koppen belts (one map ~ 5 million years). For an in-depth discussion of the data and methodology used see the Supplementary Materials. The resulting long-term global temperature curve is shown in Fig. 6. The estimates of global temperature obtained in this manner, do not provide a precise or detailed measurement of how the Earth's temperature has changed though time. The width of the Koppen belts are only approximate and are poorly known for older time periods, especially the Early Paleozoic. Though we assume a zonal pattern for the oceans, there are certainly major distortions caused by ocean currents and upwelling (Fig. 3). Most importantly, the temperatures assigned to each of the Koppen belts are based on modern icehouse values and may not reflect temperatures in past hothouse worlds. Despite these limitations, the Koppen approach does provide one important bit of information. This procedure tells how the Pole-to-Equator temperature gradient has changed through time. The relative widths of the equatorial wet and the subtropical arid belt do not change significantly through time because they are controlled by Hadley Cell Circulation. The changing in Pole-to-Equator temperature gradient is due nearly exclusively due to the changing width of the Warm Temperate, Cool Temperate and Polar Belts. In icehouse worlds, like the present-day, Pole-to-Equator temperature gradient is very steep. The temperature falls .75° -1° C per degree of latitude as we move towards the Pole (e.g., if we start at 30°C at the Equator, we end up with temperatures of -40C° to -60° C at the pole). During hothouse worlds (e.g., Cenomanian-Turonian Thermal Maximum, 93 Ma), the pole-to-Equator temperature gradient was much shallower, approximately .20° - .33° C per degree of latitude. That means that if we start out at 30° C at the Equator, the temperature at the Pole would still be well above freezing (0° to 12°C). Fig. 7 illustrates the possible range of ancient Pole-to-Equator gradients. The plus signs lie along the modern Pole-to-Equator gradient and describe how temperatures change as a function of latitude. In the present-day world, the temperature near the Equator is 26° C. The temperature remains nearly constant in the tropics (0° - 15° latitude), and then begins to decrease rapidly. Freezing temperatures are reached at 60° latitude; falling to -35° C at the Poles. The first column labelled "GAT" is the global average temperature obtained by summing the latitudinal temperatures along the corresponding Pole-to-Equator temperature gradient curve. For example, the modern world (GAT = 14.5°C) falls between the 14°C and 15°C curves. The other columns describe how the polar temperature, the temperature of the deep ocean and the latitudinal gradient (between 30° - 60° latitude), change with each "GAT' curve. These topics will be discussed later. The final column labelled, "Frequency % Phanerozoic", records 5 CR. Scotese et al. Earth-Science Reviews 215 (2021) 103503 Pole to Equator Temperature Gradients Fig. 7. The polar temperature and the Pole to Equator temperature gradient for different Global Average Temperatures (GAT). Polar Temperature = average temperature above 67 latitude (N&S), Deep Sea = the average temperature at the bottom of the oceans (after Valdes et al., 2020). Pole to Equator Gradient = the average change in temperature for every one degree of latitude measured between 30 and 60 latitude. The Pole to Equator temperature gradient is shallow near the Equator and steepens rapidly near the Pole. The plus signs are the combined average temperatures for the present-day northern and southern hemispheres. Frequency = the percent of the time during the Phanerozoic characterized by this Pole-to-Equator temperature gradient. All of these calculations are based on an average tropical temperature of 26C (15 N - 15 S). the frequency of these various GATs during the Phanerozoic. Past hothouse worlds with global temperatures ranging between 19° - 20°C have been the most frequent (~60%). Icehouse worlds with global temperatures similar to the modern world (>15°C) are relatively rare (—10%). It should be noted that some very shallow Pole-to-Equator gradients (GAT > 21°C), though hypothetically possible, have probably not been achieved. Conversely, Pole-to-Equator gradients associated with GATs below 10° C, have only been seen in the Snowball Earth worlds of the late Precambrian (Hoffman et al., 1998; Hoffman and Schrag, 2000; Hoffman and Schrag, 2002 ; Hoffman and Li, 2009). 55 50 Phosphatic brachiopod Conodont Fish Belemnite a Bivalve Brachiopod Planktonic foram Others Age (millions of years ago) Fig. 8. Raw and mean values of oxygen isotopes from phosphatic and carbonate fossils for reconstructing tropical sea surface temperatures over the past 500 million years (modified after Song et al., 2019). The scale of 518Ophos is used for phosphatic fossils, i.e., phosphatic brachiopod, conodont, and fish. The scale of 5180Carb is used for carbonate fossils, i.e., belemnite, bivalve, brachiopod, planktonic foraminifer, and others. Magenta curve represents the mean values of sea surface temperatures per million years. Shaded area represents 95% confidence intervals. 6 C.R. Scotese et al. Earth-Science Reviews 215 (2021) 103503 Fig. 9. Phanerozoic Isotopic Temperature of the Tropics (Song et al., 2019). A. = Each dot represents the average of all temperatures that fall within a given one million year interval. The best-fit curve was obtained using the Savitsky-Golay smoothing technique (window 11-15, degree 4). B. Change in Tropical Temperature (ATtrop). The black dots along the x-axis are the times when no data are available. Though these estimates of Global Average Temperature are based on a methodology rooted in the modern icehouse world, it is possible to make adjustments to these GATs so that they also reflect temperature changes in past hothouse worlds. As can be seen in Fig. 7, all of the Pole-to-Equator temperature gradient curves are set to 26° C at the Equator. This equatorial temperature is certainly lower than the equatorial temperatures during hothouse times. In the next section we describe how we used oxygen isotopic data to estimate the change in tropical temperatures through time. 2.3. Global Temperature Change (~10 - 20 million years) derived from Oxygen Isotopic Data In addition to the gradual change in global temperature recorded by the changing pole-to-Equator gradient, we know that the Earth's temperature has varied significantly over periods of 10 - 20 million years. The evidence for these temperature changes comes from the measurements of the ratio of lsO /leO (also referred to as 8180) in the shells and bones secreted by marine organisms (Fig. 8). The solid black line in Fig. 9A illustrates the isotopic temperature of seawater for the last 500 million years based on a compilation of >22,000 oxygen isotope measurements (Song et al., 2019). Isotopic values were converted to temperature using the equation of Lecuyer et al. (2013) for phosphate fossils and by using the equation of Hays and Grossman (1991) for carbonate fossils. Because the samples upon which this curve is based come entirely from tropical and subtropical latitudes (< 40° N&S), this curve is essentially an estimate of tropical temperatures through time. Each dot in Fig. 9A represents the average of all isotopic estimates of temperature that fall within a one million year interval. Error estimates of the average temperature calculation are given in Song et al. (2019). The best-fit curve in Fig. 9A was obtained using the Savitsky-Golay smoothing technique which fits successive sub-sets of adjacent data points with a low-degree polynomial by the method of linear least squares (Savitsky and Golay, 1964). We found the Savitsky-Golay method did a better job honoring the data points than either the high-order polynomial or LOESS techniques traditionally employed. Though the isotopic temperature curve is relatively flat for much of the Mesozoic and Cenozoic, the curve has a steep, linear negative slope during the early Paleozoic. The isotopic estimates of temperature for these times appear to be much higher (up to +24° C difference at 500Ma) than the temperatures obtained from lithologic indicators of climate. It should be noted that the isotopic temperatures illustrated in this curve do not take into account regional variations in the isotopic composition of seawater, or proposed Phanerozoic changes in the lsO leO ratio of seawater. Some researchers believe that there is a long-term trend in the average isotopic composition of seawater (Veizer et al., 2000; Prokoph et al., 2008; Verard and Veizer, 2019). Others vehemently disagree (Grossman, 2012a, 2012b). This debate has been going on for more than 30 years. It seems incredible that the average temperature of the tropical oceans could have been as high as 50° C (122° F) during the early Paleozoic and late Precambrian. This implies that diurnal or seasonal temperatures were much higher. Even in extreme cases, modern organisms cannot survive when seawater temperatures reach 40°C (104°F) (Fraenkel, 1960). Verard and Veizer (2019) propose that a systematic decrease in plate tectonic activity during the last 540 million years may have systematically increased the ratio of lsO to leO in seawater; however, they admit that the details of the correlation are poor and that tectonic activity prior to the formation of Pangea is not well-known. It may be possible that the 8180 composition of seawater has not remained constant, nor has it gradually changed in a quasi-linear fashion. Rather the 8180 could have varied non-monotonically through time. Upon reflection, what may be knowable is not the "absolute" 7 C.R. Scotese et al Earth-Science Reviews 215 (2021) 103503 temperature provided by the isotopic data, but rather the "relative" change in temperature at shorter timescales. For our purposes, the most useful information that can be gleaned from the isotopic temperature record is the change in the tropical temperatures through time (i.e., the A T°trop)) Fig- 9B. In order to calculate the A T°trop, we first estimated the long-term isotopic temperature signal by averaging the isotopic data using a 60 million year running-average. We chose 60 million years because it best approximated the long-term temperature signal from the Koppen analysis. The 60 million year running average was then subtracted from the best-fit isotopic tropical temperature estimates (Fig. 9A, black line) to obtain the A TVop value (Fig. 9B). 2.4. Modifying the Isotopic Estimates of Changing Tropical Temperatures Using Geological and Paleontological Constraints The isotopic temperature of seawater can be affected by multiple environmental conditions: hypersalinity or hyposalinity, the presence of continental ice caps, or the water depth at which sampled organisms originally resided (Grossman2012a,b). In this section we review the isotopic estimate of tropical temperature described in the previous section and make modifications based on geological and paleontological considerations. Fig. 10 illustrates the changing temperature of the Tropics (AT° trop) based on isotopic measurements (dotted line). Superimposed on the dotted line are colored lines that represent seven time intervals when the "raw" isotopic estimates of tropical temperature do not agree with geological or paleontological information and therefore, must be adjusted or modified. These seven intervals are: the Cambrian, 540-485 Ma; the latest Ordovician, 445 -443 Ma; parts of the Permo-Carboniferous, 340 - 275 Ma; the Triassic, 252 - 200 Ma; the Early Cretaceous, 125 Ma - 95 Ma; the latest Cretaceous, 80 - 65 Ma; and the Cenozoic (65 - 0 Ma). The procedure that we used to determine whether in isotopic measurements require modification was initially developed at a workshop convened by Scott Wing and Brian Huber at the Smithsonian National Museum of Natural History (April, 2017 & 2018). At those meetings a multidisciplinary group of more than 20 earth scientists convened at the Museum to refine a Phanerozoic paleotemperature curve that eventually became part of an exhibit on Global Warming in the Museum's Earth History Hall (Wing and Huber, 2019). During that meeting participants were invited, based on their geological and paleontological knowledge and expertise, to make additions and revisions to the proposed Phanerozoic temperature curve which was posted as a wall-sized display. The results of that collaborative effort are illustrated in Fig. 1 (Wing and Huber, 2019). The criteria used to modify or adjust the isotopic estimate of temperature included the addition of ephemeral events which have not been sampled in the isotopic record (e.g., KT Impact Winter, Permo-Triassic Thermal Maximum), the elimination of spurious thermal maxima during times of icehouse conditions (e.g., late Pennsylvanian thermal maximum), and the elimination of cooling events that would require large, permanent ice caps during times of hothouse conditions (e.g., latest Cretaceous cooling). This same basic procedure was used by the authors to modify and adjust the raw isotopic estimates of temperature so that they better agreed with geological and paleontological constraints. In the following section, we describe the changes that have been made to the raw isotopic estimates and the reasons for these changes. 2.4.1. Cambrian, 540 485 Ma. There is sparse isotopic data older than 500 million years (Bergmann et al., 2018b; Hearing et al., 2018; Henkes et al., 2018). The data that is available indicates that tropical temperatures were in excess of 40° C; as noted previously, this is problematic (Fraenkel, 1960). For a more detailed discussion of global temperatures during the Cambrian, see section 5.2. The dark red curve in Fig. 10 is a speculative estimate of temperature changes based on the carbon isotope record. They coincide with 10 proposed Cambrian d13C isotopic excursions (Zhu et al., 2006). The covariance of d13C and dlsO trends has long been noted (Wenzel and Joachimski, 1996; Jenkyns et al., 2002 ). Approximately 80% of the positive d13C excursions are correlated with warmer temperatures (hyperthermals). The correlation is generally attributed to the causal relationship between higher ocean temperatures, the formation of deep water anoxia, and the subsequent preservation of organic carbon. 2.4.2. Latest Ordovician (Hirnatian, 445-443 Ma) The isotopic estimates of tropical temperatures during the Hirnan-tian Ice Age (see section 5.3), show a notable dip of 3° C. We have exaggerated the dip in temperature to help explain the growth of the massive, Hirnantian south polar ice cap that extended well into the 8 C.R. Scotese et al. Earth-Science Reviews 215 (2021) 103503 Late Pennsylvanian (312-300 Ma) O South China • USA ,0° 330 325 millions of years I Ml o o- . lo^c0o ,8°.°*° o°. / Oo«oo£ I 0 o „ „0 Fig. 11. Comparison of isotopic tropical temperature data from the late Pennsylvanian of South China (open dots) and the USA (black dots) (Song et al., 2019). Ordovician tropics (~35° S). 2.4.3. Permo-Carboniferous Icehouse Evaporitic anomalies are known to affect the isotopic measurements made in the great subtropical epeiric seas of the Paleozoic (Mii et al., 1999). Because evaporation preferentially removed the lighter isotope of oxygen (leO), the observed isotopic temperature measurements obtained from organisms living in these evaporitic seas are erroneously too cool. Conversely, freshwater is rich in leO, so isotopic estimates of temperature made in areas that receive a large influx of freshwater are erroneously too warm. As Fig. 10 shows, a broad temperature peak (dotted line) sits in the middle of the Permo-Carboniferous Ice Age. A closer examination of the isotopic temperature record reveals that the isotopic data for the late Pennsylvanian come primarily from South China and the USA (Fig. 11). The isotopic temperatures for China (Chen et al., 2013, 2016) and the USA (Luz et al., 1984; Joachimski et al., 2006; Elrick and Scott, 2010; Rosenau et al., 2014) are very different. The average temperature for the USA data is ~ 25° C, whereas the average temperature for South China is ~13°C. This difference in temperature cannot be explained in terms of latitudinal position. 310 million years ago, both of these regions straddled the paleo-Equator (Scotese, 2014). The difference in isotopic temperatures is more likely due to environmental differences. In the USA, the influx of fresh water rich in leO from the rising Central Pangean mountain ranges may have altered the lsO / leO ratio, making the seawater "lighter", giving a false, warmer isotopic temperature. We invoke a similar explanation to explain the much smaller thermal anomaly in the early Permian (~285 Ma). As described above, evaporitic anomalies may cause isotopic temperatures to appear "too cool". We can speculate that this phenomenon may explain the cooler than expected temperatures during the early Carboniferous (~335 Ma). We have applied an ad hoc correction of 4-5° C to the mid-Pennsylvanian isotopic measurements to bring them into line with the geological observations. This adjustment keeps temperatures below the global temperature required to form permanent polar ice caps (see section 3.2). A more modest adjustment of 2-3°C was applied to the earliest Permian (295 - 285 Ma) because extensive south polar glacial deposits indicate the Permo-Carboniferous glacial maximum occurred during the latest Carboniferous - earliest Permian (see section 5.5). Conversely, the isotopic temperatures for the Visean appear to be much too cold suggesting that the Visean represented the depths of the Permo-Carboniferous Ice Age. This is not the case, the Visean was one of the warmest intervals of the Carboniferous (Mii et al., 1999). Consequently, isotopic temperatures have been increased a modest 2-3 °C. 2.4.4. Triassic and early Jurassic (252 - 200 Ma) Three adjustments to the average isotopic temperature were made for the Triassic. The Permo-Triassic Extinction was the greatest mass extinction of the Phanerozoic (see section 5.6). The kill mechanism was massive, global warming (see section 3.1.2). It is likely that average global temperatures reached 40° C (Sun et al., 2012). The sharp spike in temperature has has been added at the Permo-Triassic boundary . Global temperatures peaked during the earliest Triassic, fell sharply during the early Triassic (Sun et al., 2012), and extreme hothouse conditions did not ameliorate until the mid-Triassic (Ladinian; see section 5.6). The averaged isotopic temperatures for the late Triassic and early Jurassic are far too low and erroneously suggest possible icehouse conditions. This anomaly is probably due to the fact that some of the isotopic measurements come from belemnites and clams that once inhabited deeper, cooler waters (Dera et al., 2011; Wierzbowski and Joachimsky, 2007). To remedy this mismatch between isotopic and geologic evidence, global temperatures during the Triassic and early Jurassic were increased by 2-3 °C. Finally, the end of the Triassic was marked by a major extinction event that was most likely triggered by the eruption of the Central Atlantic Magmatic Province (see sections 3.2 and 5.6). There is no oxygen isotope record of this relatively brief event (1-2 million years, Ernst, 2014). To explain the extinction event and other paleontological evidence of global warming, we have added a narrow thermal spike at 200-201 Ma. 2.4.5. Early Cretaceous (130 - 105 Ma) and latest Cretaceous (80 - 65 Ma). The Late Jurassic - earliest Cretaceous is generally considered to be a relatively cool period with evidence of intermittent polar ice. Towards the end of this interval (—120 Ma), massive volcanic eruptions took place in the Central Pacific (Greater Ontong Java Plateau), along the 9 C.R. Scotese et al. Earth-Science Reviews 215 (2021) 103503 Fig. 12. Comparison of Cenozoic deep ocean isotopic temperatures and tropical temperatures A. gray line = low resolution tropical temperatures (this study), and the black line = high resolution tropical temperatures for the Cenozoic. B. isotopic temperatures from deep ocean, benthic foraminifera (Zachos et al., 2001, 2008; Westerhold et al., 2020). southern margin of Africa, and forming the Kerguelen Plateau (see section 3.2 for details). The oldest Cretaceous oceanic anoxic event, OAEla, the Selli/Goguel Thermal Maximum is coincident with these LIP events. Because this oceanic anoxic event is thought to have been triggered by significant global warming, we have added a thermal peak at —120 Ma. This early Aptian warming was followed by a "cold snap" in the late Aptian - early Albian (see section 5.8). The little "bump" that precedes the Cenomanian-Turonian Thermal Maximum represents the Paquier/Urbino Thermal Maximum (OAElb) (see Table 6). 2.4.6. Cenozoic (65 - 0 Ma) The classic record of deep ocean temperatures based on benthic foraminifera assembled by James Zachos (Fig. 12, Zachos et al., 2001, 2008; Westerhold et al., 2020) provides a framework for describing the temperature fluctuations during the Paleocene-Eocene Hothouse and the Late Cenozoic Icehouse (Koeberl and Montanari, 2009). The temporal resolution of the Phanerozoic isotopic temperature record from the tropics (Fig. 9A) is one control point per million years. While this resolution is adequate for older geological periods, it does not permit a detailed description of Cenozoic temperature changes. We can improve the Cenozoic portion of the temperature curve if isotopic information from deep ocean, benthic foraminifera are used (Zachos et al., 2001, 2008; Westerhold et al., 2020). The isotopic record from benthic foraminifera provides a nearly continuous record of deep sea temperature changes from beginning of the Cenozoic to the present-day and has a temporal resolution of — 100,000 years. As illustrated in Fig. 12, we have converted the Cenozoic deep ocean temperatures to a high resolution estimate of tropical temperatures by superimposing the high-resolution deep ocean temperature curve on our low-resolution estimate of tropical temperatures. This was done by adding —11°C to the pre-ice, Paleogene- early Oligocene portion of the deep ocean temperature curve (65 Ma - 28 Ma), as well as adding —14°C to the early and mid-Miocene portion of the deep ocean temperature curve (28 - 14 Ma). The late Miocene to Recent portions of the temperature curve (14 - 0 Ma) were transposed using linear approximation. The overall shape and amplitudes of the superimposed curves are in remarkable agreement (Fig. 12). The only significant mismatch is the late Eocene - early Oligocene (33 - 38 Ma) portion of the curve. The Oligocene portions of the transposed deep ocean temperature are 6° - 8° C cooler than isotopic tropical temperatures, which is certainly anomalous. We have chosen to use the updated high resolution Cenozoic tropical temperatures when building our model of Cenozoic global average temperatures. It should be noted that the methods that Hansen et al. (2008) have used to estimate global average temperatures from deep sea temperatures were not used. Though the results he obtained for the Neogene are identical to our estimates, his methodology overestimates global temperatures when applied to hothouse climates. 2.4.7. Summary of Modifications to the Isotopic Temperature Curve In summary, —30% of the Phanerozoic isotopic paleotemperature curve was modified using geological and paleontological criteria. In most cases, 2-3°C were added to or subtracted from the isotopic temperatures. Three notable exceptions are the Hirnantian Ice Age (-4°C), the Permo-Triassic Extinction Event (+7 °C), and the KT Impact Winter (-6°C). Two new thermal maxima were added (200Ma, CAMP; 124 Ma, OAEla), three problematic apparent thermal maxima were removed (315-295 Ma; 285-275Ma, and 40-30 Ma), and two anomalous apparent cool episodes were either reduced (80-65Ma) or eliminated (340-330 Ma). In addition, a high-resolution isotopic curve was used to represent Cenozoic temperatures (see section 5.9) and Cambrian carbon isotope excursions were used as a proxy for temperature change (see section 5.2). 10 C.R. Scotese et at 2.5. Combining the Estimates of the Changes in the Pole-to-Equator temperature gradient with the Revised ATj-ij curve derived from Oxygen Isotope Data In the final step of our methodology, we combine the estimates of Global Average Temperature obtained from the changes in the Pole-to-Equator temperature gradient (Fig. 6) with the revised changing temperature of the Tropics (AT0 trop) (Fig. 10). The combined geological and isotopic temperature curve (Fig. 13) is similar in many respects to the curve derived solely from isotopic data (Fig. 9A), but also differs in several important ways. 1. The Paleogene portion of the isotopic temperature curve is ~ 1.5°C cooler than the geologic temperature curve. 2. The isotopic temperature curve, overall, tends to indicate slightly warmer temperatures (e.g peak of the Cenomanian-Turonian thermal high). 3. The late Carboniferous through Triassic, and the late Jurassic -early Cretaceous portions of the isotopic temperature curve are 1.5° -2° C warmer than the geologic temperature curve. 4. The mid-Ordovician through Devonian portion of the isotopic temperature curve is significantly warmer than the geologic temperature curve (> 6° - 8°C). 5. The Cambrian and early Ordovician temperatures indicated by the isotopic temperature curve are nearly double the geologically-inferred (~ 50°C versus ~ 25°C, respectively). 3. Discussion A central thesis of this paper is that changes in the Earth's temperature during the Phanerozoic have been caused by factors that act on different time scales. This is not a new idea but rather goes back to the archetypical insights of Alfred Fischer (Fischer, 1981, 1982, 1984), who Earth-Science Reviews 215 (2021) 103503 recognized climatic oscillations and cycles in the biosphere. We recognize three major timescales of temperature change. Long-term changes in temperature (>50 million years) are due to global changes in the rates of volcanic CO2 degassing associated with seafloor spreading and sub-duction, as well as long-term changes in the weathering of continents. Long-term changes in temperature (>50 million years) are modeled by mapping the extent of ancient climatic belts (Kóppen belts) which vary in response to changes in the Pole-to-Equator temperature gradient. Medium-term changes in temperature (10 - 20 million years) can be deduced from changes in the isotopic temperature of the tropical seas. How well does our Phanerozoic temperature model (Fig. 13) match the geological record? In the next section we compare the composite Phanerozoic temperature curve with a variety of geological phenomena including: LIP eruptions, bolide impacts, the formation of permanent ice caps, and changes in tropical, deep ocen, and polar temperatures during the Phanerozoic. We also discuss top priorities for future research. 3.1. Short-Term Global Temperature Excursions (< 1 to several million years) due to LIP Eruptions and Bolide Impacts We know that very rapid excursions in temperature can be caused by a variety of tectonic and geological phenomena. For example, global temperatures can rise rapidly due to the release of vast amounts of CO2 into the atmosphere by eruptions from Large Igneous Provinces (LIP) (Clapham and Renne, 2019). In Fig. 14, the black rectangles represent the timing and relative magnitude of 21 major LIP eruptions (Ernst, 2014; Ernst and Youbi, 2017). Conversely, temperatures can fall rapidly, creating short-lived icehouse worlds, due to major bolide impact events (e.g., K/T boundary Impact Winter or possibly the Hirnantian Ice Age). Very large impact events (crater size > 150 km) may also trigger large igneous eruptions that create a period of global warming that follows impact-related cooling (e.g., Chicxulub impact event). In Fig. 14, the -1 34 Fig. 13. Phanerozoic Global Average Temperature (GAT), blackline = Global Average Temperature, dashed line = Long-term temperature change derived from changes in the pole-to-Equator temperature gradient calculated from the changing area of Koppen Climatic Belts (see Figure 6). When the Global Average Temperature is below 18C large polar icecaps can form. When the Global Average Temperature is above 18C large polar icecaps are unlikely to form. 11 CR. Scotese et al. Earth-Science Reviews 215 (2021) 103503 UP Magnitude (ÍO6 km2) per my Sil] a n Ch KDlB BKD2 SI CE © ® VH i ■ El Crater Diameter Siberian Traps Chicxulub CAMP G OJ SEAF Kerg Popigoi O cw ■ Ar a ■ SM ■ Q Kara PE ^ Afar ■|7bo/c Deccan hj KK " Mad ■ PK Mj car Nat CB 0 0 0 Mn NAIP 0 MS © CR 340 300 260 220 millions of years I— ISO —1— 140 1-r- 60 30 (km) 120 Fig. 14. Timing and magnitude of Large Igneous Provinces (rectangles) eruptions and bolide impacts (circles). The size of the rectangles indicates the relative eruptive intensity (106 km2/my) (left-hand scale). See Table 1 and 2 for abbreviations. Sources: Ernst (2014), Spray (2020). Table 1 Major Large Igneous Provinces Name Start End Area Relative Average Age Age 106km2 to Deccan Temperature Increase per million years Columbia River 16 15 0.24 13% 0.4 (CR) Afar 31 29 2 111% 1.7 North Atlantic 56 54 1.3 72% 1.1 Igneous Province (NAIP) Deccan 66 65 1.8 100% 3.0 Madagascar 93 91 1.6 89% 1.3 (Mad) Naturaliste (Nat) 100 99 1.2 67% 2.0 Kerguelen (Kerg) 120 118 2.3 128% 1.9 SE African 125 118 3.5 194% 0.8 (SEAf) Hikurangi (GOJ) 121 119 2.7 150% 2.3 Ontong Java 121 119 4.27 237% 3.6 (GOJ) Parana- 131 129 2 111% 1.7 Entendeka (PE) Karoo-Ferrar 183 181 3.6 200% 3.0 (KF) Central Atlantic 201 199 4 222% 3.3 Magmatic Province (CAMP) Siberian Traps 252 251 7 389% 11.7 Emeishan (Em) 260 259 0.25 14% 0.4 Tian Shan (TS) 330 329 1.7 94% 2.8 E. Europe (EE) 365 364 0.8 44% 1.3 Viluy Rift (VR) 375 374 0.8 44% 1.3 Kola-Dneiper 2 359 358 1.5 83% 2.5 (KD2) Kola-Dneiper 1 373 372 1.5 83% 2.5 (KD1) Kalkarindja 512 509 2.1 117 1.2 (Kalk) circles illustrate the timing and crater size of the largest, well-established impact events (Spray, 2020). 3.2. Large Igneous Provinces (LIPs) The slow and steady release of CO2 from the volcanism that accompanies subduction and seafloor spreading is one of the major factors that determine the Earth's thermal equilibrium. This thermal equilibrium changes slowly because the plates move slowly and the continents erode slowly. In contrast, LIPs erupt rapidly, release tremendous amounts of CO2 and consequently warm the Earth rapidly. The amount of global warming depends on the volume of CO2 released and the amount of CO2 in the atmosphere (i.e., climate sensitivity; Royer, 2016). A major LIP eruption during a period of icehouse climate will have a much greater effect than the same LIP eruption during a period of hothouse climate. The warming effects of LIPs will continue as long as voluminous eruptions continue, often for several million years. Beginning soon after eruption, the weathering of subaerial basalts may remove CO2 from the atmosphere reducing global temperatures to near pre-emption levels. The rate at which this occurs depends, in part, on the climatic setting of the LIP (Walker et al., 1981; Berner et al., 1983; Marshall et al., 1988; Raymo et al., 1988; Worsley and Kidder, 1991; Bluth and Kump, 1991; Otto-Bliesner, 1995; Gibbs et al., 1999; Berner, 2004; Nardin et al., 2011; Goddéris et al., 2012, 2014) Table 1 lists 21 of the largest Phanerozoic Large Igneous Provinces (Ernst, 2014; Ernst and Youbi, 2017). They are also plotted as black squares in Fig. 15. It is clear from a cursory look at Fig. 15 that there is a strong correlation between warm periods and the eruption of LIPs. 19 of the 21 major LIPS fall within periods of global warming. Most notable are the correlation between: 1) the West Siberian Traps and the Permo-Triassic Thermal Maximum (252 Ma); 2) the correlation between the Barremian-early Aptian Warm Period (125 - 116 Ma) and the synchronous eruptions that formed the Greater Ontong Java plateau (Taylor, 2006), the Southeast African oceanic plateaus, and the Kerguelen Plateau; 3) the Viluy rift and the mid-late Devonian Hangenberg Thermal Maximum (375 - 355 Ma); 4) the Karoo-Ferrar LIP and the Toarcian Thermal Maximum (180 Ma) and; 5) the late Paleocene early Eocene Warm Period (58 - 55 Ma) and the eruption of the North Atlantic Igneous Province (NAIP). The one major exception was the East Africa eruptions (Afar LIP) which took place during the period of intense of late Eocene -early Oligocene cooling. Reluctantly, we have excluded the 12 C.R. Scotese et al. Earth-Science Reviews 215 (2021) 103503 Large Igneous Province Eruptions | Major Bolide Impacts © Putative Bolide Impacts Change in Tropical Temperatures Siberian Traps "Khione" 0 Chicxulub Fig. 15. Comparison of the timing of LIPS (black squares, Table 1), large bolide impacts (circles with dots, Table 2), and putative large impact events (light gray circles with x's) with the changes in tropical temperature (AT^p). The size of the lettering is roughly proportional to the size of the LIP or bolide impact. See Table 1 and 2 for abbreviations. Sources: Ernst (2014), Spray (2020). Caribbean LIP (Sullivan et al., 2020) from consideration because the origin of the underlying lithosphere is certainly older (Jurassic?) than the purported age of the LIP (90 Ma), and the age of the surfacing vol-canics is not well-constrained. As has been noted by numerous authors, (Kidder and Worsley, 2010; Wignall, 2001, 2015; Rampino and Self, 2015; Bond and Grasby, 2017; Ernst and Youbi, 2017; Clapham and Renne, 2019; Rampino et al., 2019; McKenzie and Jiang, 2019; Schobben et al., 2019; Suarez et al., 2019), the global warming caused by the eruption of large LIPS has often resulted in mass extinctions. LIP eruptions correlate with extinction events during the middle Cambrian (Kalkarindji, 510 Ma); at the Frasnian-Famennian boundary (Kellwasser extinction, 373-372 Ma); during the latest Devonian (Hangenberg extinction, 359-358 Ma); at the Triassic-Jurassic boundary (CAMP, 201 - 200 Ma); the Toarcian extinction (Karroo - Ferrar LIP, 183 - 181 Ma); the Eocene-Oligocene boundary extinction (Afar- E. African Rift, 31-29 Ma); not to mention the two greatest extinction events of all time, the Permo-Triassic Mass Extinction (West Siberian Traps, 252 Ma - 251 Ma) and the K/T Extinction (Deccan Traps, 66 Ma - 65 Ma). Though the geochemistry of LIP lavas indicate that most LIPS are derived from deep mantle sources, there remains the possibility that a few of the largest LIPS (e.g., West Siberian Traps, CAMP, and the Greater Ontong Java Plateau ; Ernst, 2014; Ernst and Youbi, 2017) may have been produced by extremely large impact events (crater size > 200 km). Any craters produced by these gigantic impact events would probably have been obscured by the subsequent voluminous volcanic eruptions. 3.3. Bolide Impacts Table 2 lists 29 bolide impact events that generated craters greater than 25 km in diameter (Spray, 2020). Fig. 14 illustrates the relative size of these craters and the timing of the impact events. The amount of energy generated by an impact event is described by the equation (Hughes, 2003; Rampino, 2020): E = 9.1 x 1024 Da' where E = is the kinetic energy of the bolide impact in ergs, and D = the diameter of the resulting crater. Proposed values for the exponent, a, range from 2.59 (Hughes, 2003) to 3.89 (Melosh, 1989). In either case, it is clear that the energy associated with an impact event correlates with the cube of the impact crater diameter. The amount of aerosols, SO4, and particulate matter injected into the atmosphere is primarily a function of the energy associated with the impact event. Whether a bolide lands on continental crust or oceanic crust may also affect the character of the impact event. In either case, the cooling effect of impact events drops off rapidly as the size of the bolide decreases. It has been suggested that after the impact of the bolide that formed the Chicxulub crater (150 km), the global mean temperature would have dropped 7° - 12° C (Vellekoop et al., 2014). By comparison, the impact event that formed the Manicouagan crater in Quebec, which is approximately half the size of the Chicxulub crater, would have generated only 2.5 °C of cooling. Impact events that produce craters that are less than 25 km in diameter cause negligible amounts of global cooling (< 1°C). As illustrated in Fig. 15, there are remarkable number of Phanerozoic bolide impacts (circles) that coincidentally occurred during periods of major global cooling (18 of 22). The most notable are the Puchezh-Katunki impact (167 Ma, Middle Jurassic Cool Interval), the Morok-weng, Mjolnir and Tookoonooka impacts (145 - 128 Ma, Tithonian-early Barremian Cool Interval), and the Carswell impact (115 Ma, Aptian - Albian Cold Snap). Though the timing of these impact events and the periods of cooling appears to be entirely coincidental, we cannot rule out the possibility that these intense, short-lived cooling events may have triggered ice-albedo feedbacks (Park and Royer, 2011) that accelerated cooling processes already underway. A quick scan of Fig. 15 reveals that there are several relatively brief periods of rapid global cooling (-2°C to -12°C) that may have been triggered by a bolide impact, but no crater has been associated with these cooling events. This, of course, is not surprising due to the incompleteness of the geological record. The craters that formed during these impact events may have been eroded, tectonically destroyed, or buried beneath a thick pile of sediments. There are no positively identified impact craters in the oceanic basins (Gersonde et al., 1997). Impact events older than 180 million years would have taken place on oceanic crust that has been subducted. Given the proportion of the area of continental crust versus the area of oceanic crust, we would expect that, during the last 540, million years the deep oceans of the Earth were 13 CR. Scotese et al. Earth-Science Reviews 215 (2021) 103503 Table 2 Impact craters with diameters >25 km Name Location Age Diameter Impact Estimated (Ma) (km) Energy Relative to Chicxulub Temperature Drop (° C) Vredefort -2023 160 120% -12°C Chicxulub Yucatan 66 150 100% -10°C Sudbury -850 130 70% -8°C Popigai 35.7 +/- 2 90 27% -3°C Acraman 590 90 27% -3°C Manicouagan 214 +/- 85 23% -2.5°C (Man) 1 Morokweng 145 +/- 70 14% -1.5°C (Mor) .8 Kara 70.3 +/- 2.2 65 11% -1.2°C Beaverhead -600 60 9% - -1.0°C Tookoonooka 128 +/- 55 7% - -1.0°C (Took) 5 Charlevoix 342+/- 54 7% - -1.0°C (Ch) 15 Siljan 376.8 +/- 1.7 52 6% <-1.0°C Kara Kul (KK) < 5 52 6% <-1.0°C Montagnais 50.5 45 4% <-1.0°C (M) +/- .76 Woodleigh 364 +/- 40 3% <-1.0°C (W) 8 Saint Martin 220 +/- 40 3% <-1.0°C (SM) 32 Puchezh- 167 +/- 40 3% <-1.0°C Katunki 3 (PK) Mj0lnir (Mj) 142 +/-2.6 40 3% «-1.0°C Chesapeake 35.5 40 3% «-1.0°C Bay (CB) +/- .3 Araguainha 254.7 40 3% «-1.0°C (Ara) +/- 2.5 Carswell (Car) 115 +/-10 39 3% «-1.0°C Clearwater 290 +/- 36 2% <- .5°C West (CW) 2.0 Manson (Mn) 73.8 +/- .3 35 2% <- .5°C Keurusselka <1800 -30 2% <-.5°C Yarrabubba. -2000 30 2% <- .5°C Slate Islands -450 30 2% <-.5°C (SI) Shoemaker 1630 +/- 5 30 2% <- .5°C Mistastin (Ms) 36.4 +/- 4 28 1% «- .5°C Clearwater 460- 26 1% «- .5°C East (CE) 470 hit by > 20 bolides that could have generated impact craters >50 km in diameter. One or two of these oceanic impacts could have been as large, or larger, than the Chicxulub crater (>150 km). It is possible that as yet unknown bolide impacts may have caused some of the cooling events shown in Fig. 15. The most notable putative impact is the "Khione" impact event which could have triggered the latest Ordovician ice age (444 Ma) and mass extinction (Brenchley et al., 1994; Kump et al., 1999a; Sheehan, 2001; Finnegan et al., 2011). The "Khione" (key-OH-nay) impact event is named after the daughter of the Greek god, Boreas, god of winter. Khione is the goddess of snow. The Hirnantian icehouse is especially interesting because it was very shortlived, yet very intense. Ice sheets extended from the South Pole to a paleolatitude of nearly 35°S (Boucot et al., 2013). The Hirnantian Ice Age is anomalous because it occurred during a time of relatively high global temperatures (GAT = 20° - 22° C). Other putative oceanic impact events may have caused the rapid drops in global temperature that are observed during the early Devonian (Pragian, 407 Ma), latest Triassic (Rhaetian, 203 Ma), early Jurassic (Pliensbachian, 184 Ma), and mid-Oligocene (28 Ma). These speculative impacts are indicated by the gray circles with x's in Fig. 15. As noted earlier, Fig. 15 compares the chronology of LIP eruptions (squares) and bolide impacts (circles) with the changes in global temperature. In Tables 1 and 2 we have estimated the magnitude of the warming and cooling effects of these LIP eruptions and bolide impacts, respectively (Black and Gibson, 2019; Vellekoop et al., 2014; Kamber and Petrus, 2019; Suarez et al., 2019). Large LIP eruptions nearly all have significant warming effects (i.e. several degrees centigrade). The cooling effects of most impacts are negligible. Several impacts were large enough that they should have had a noticeable effect on global temperatures (Sudbury, Popigai, Acraman, and Manicouagan). The Chicxulub impact (K/T boundary) and the putative Khione-type impacts undoubtedly had profound effects on global climate, extinction, and the evolution of life. 3.4. Icehouse Worlds During the past 720 million years, the Earth's climate has transitioned from a frigid icehouse to a steaming hothouse seven times. Three of these icehouse periods took place during the late Precambrian (the Sturtian icehouse, 720 - 660 Ma; the Marinoan icehouse, 635 Ma; and the Gaskiers icehouse, 585 Ma (Ogg et al., 2016). During icehouse intervals, large permanent icecaps occupied either the northern or southern polar regions, or both. These times of icehouse conditions are recorded by glacial debris and tillites on land, dropstones in adjacent deep sea sediments that were released by melting icebergs, and glen-donites in shallow marine environments where the temperature of the bottom waters is within a few degrees of freezing (Boucot et al., 2013). Fig. 16 plots the geological record of tillites, dropstones, and glen-donites, as well as the area of permanent ice cover during the Phaner-ozoic (gray shading). Fig. 16 also plots the global average surface temperature (GAT) during these icehouse intervals (dotted line). As expected, there is a good correspondence between the times when there is geological evidence for icehouse conditions and the times when the GAT is below 18°C (dashed line). Permanent polar ice caps can only form if the global average temperature is less than 18°C. At that global temperature, the average annual temperature of the polar region (>67 ° N&S) is not warm enough during the summer months to melt the winter snow and ice. Without a complete summer thaw, snow and ice can accumulate and large ice caps can grow rapidly. Conversely, a large, permanent polar ice cap cannot form if the average global temperature is greater than 20°C (68°F). At that temperature, the average annual temperature of the polar regions is ~7°C, too warm for a permanent ice cap to form. A transition zone exists when global temperature ranges between 18°C and 20°C. Snow and ice will be present during the winter months and patches of permanent ice may develop close to the Pole or at high elevations (>500 m). The last three major icehouse intervals (Hirnantian Ice Age (450-444 Ma), Late Paleozoic Icehouse (360 - 285 Ma), and the Late Cenozoic Icehouse (45 - 0 Ma) represent 23% of Phanerozoic time. However, the paleotemperature model presented here predicts that there may have been additional time intervals when icecaps may have existed, albeit briefly. If we accept the correlation of temperature and area shown in Fig. 16 at face value, then the largest of these icecaps may have been the half the area of East Antarctica (434 - 385 Ma), the smallest icecaps may have been twice the area of Greenland (150 - 111 Ma). Several of these brief icehouse interludes are plausible because they occur during cool (but not frigid) intervals: 1) the Wenlock Cool Interval (434-326 Ma, 2) the Pragian Cool Event (411-406 Ma), (3) the Early-Middle Devonian Cool Interval (401-385 Ma), and 4) the Callovian Cool Event (166-164 Ma), Tithonian-Early Barremian Cool Interval (150-128 Ma). There are sparse, high latitude tillites, dropstones, and glendonites supporting the 14 CR. Scotese et al. Earth-Science Reviews 215 (2021) 103503 GAT 10S Hirnantian Ice Age Permo-Carboniferous Ice Age Siluro-Devonian Cool Period 60 Late Cenozoic Ice House millions km 30 Late Jurassic -early Cretaceous Cool Period Large Permanent Icecaps No Permanent Icecaps :r.: '■J « glacial deposits ?5 6 520 480 440 millions of years SO 54 U 320 280 1? 389v 0 Fig. 16. Phanerozoic Ice Ages, gray = global area of snow and ice cover (106 km2), dotted line = snow and ice predicted by Global Average Temperatures (GAT < 18G, Figure 12), numbers = number of glacial deposits (tillites, dropstones, and glendonites). Note inverted temperature scale (left side). Tropical Temperature Global Average Temperature Deep Ocean Temperature Average PolarTemperature Fig. 17. Tropical, Global Average, Deep Ocean, and Polar Temperatures, (a) red line = tropical temperature (15 N - 15 S), (b) black line = global average temperature (GAT), (c) blue line = deep ocean temperatures (after Valdes et al., 2020) (d) light blue line = polar temperature (> 67 N&S). Jurassic- Cretaceous, and Aptian - Albian cool periods (Boucot et al., 3.5. Tropical Temperatures 2013). Much of the Siluro-Devonian is considered to be a cool period, but there are no known glacial deposits from this time interval (Boucot Estimating the Phanerozoic history of tropical temperature is et al., 2013). straight-forward. Since we know the change in tropical temperature ( ATVop) Fig- 10), and the present-day average tropical temperature 15 CR. Scotese et al. Earth-Science Reviews 215 (2021) 103503 (26°C ; Legates and Wilmott, 1990), then the ancient tropical temperature is simply the sum of the present-day tropical temperature (26° C) and the change in tropical temperature (AtropT0). Tropical Temperature = 26° C + AT°trop For example, the predicted average temperature of the tropical seas during the Paleocene-Eocene Thermal Maximum (55.6 Ma) was 26°C + 5°C = 31°C. The predicted mean temperature of the tropical seas back to 540 Million years is 28° C. Tropical temperatures, for the most part, are modeled to have fluctuated between 25°C and 32°C (Fig. 17), with the exception of two periods of extreme warmth (Permo-Triassic (~39°C) and Cenomanian-Turonian (~34°C)) and three intervals of extreme cold (Hirnantian(~23°C), Permo-Carboniferous Ice Age (~23°C), K/T impact winter (~22°C)) It seems rather remarkable that the mean temperature of the polar regions back to 540 Million years is~ 0°C. This implies that the polar regions have hovered between ice-free and ice-covered conditions. Average Polar temperatures, for the most part, have fluctuated between 12°C and -12°C, with the exception of three periods of extreme warmth (Cambrian hothouse (~14°C), Permo-Triassic extinction (~17°C), and the Cenomanian-Turonian Thermal Maximum (~13°C)) and four intervals of extreme cold (Hirnantian Ice Age (~-32 °C), Permo-Carboniferous Icehouse (—24 °C), the KT Impact Winter (~-30°C) and Late Cenozoic Icehouse (~-20 °C). Throughout most of the Phan-erozoic the average temperature of the southern polar region (-4°C) has been considerably colder than the average temperature of the northern hemisphere (4 °C). Only for a brief interval in the Late Permian was the northern hemisphere colder than the southern hemisphere. 3.6. Deep Ocean Temperatures During hothouse times, the temperature gradient in the oceans is reduced and the temperature of deep ocean waters can be as high as 20°C (Valdes et al., 2020). During icehouse times, cold bottom-water is produced adjacent to the polar ice caps and this cold, salty water sinks to the depths of the ocean basins and refrigerates the oceans. In the modern world, the cool North Atlantic and North Pacific bottom waters, and the colder circum-Antarctic bottom-waters form at latitudes of 60° N&S (Emiliani, 1954). We assume that this is the approximate latitude where cold deep ocean waters would have originated in the past. We realize that factors, such as the formation of warm, hyper-saline bottom-water in the wide epeiric seas of the middle and early Paleozoic (Chamberlin, 1906; Brass et al., 1982) may also played an important role. Though this estimate of deep ocean temperatures is simplistic and imprecise, it generally agrees with the deep ocean temperatures obtained through climate modeling (Valdes et al., 2020). 3.7. Polar Temperatures Polar Temperatures were calculated by averaging the temperatures in the polar region above 67° latitude. The results are shown in Fig. 17. For example if the Global Average Temperature (GAT) is 14.5°C (modern value), then the average temperature of the polar region (above 67° latitude) is a very cold -20 °C. On-the-other-hand, during a hothouse interval like the PETM, the average polar temperature was above 8°C (Fig. 17). 3.8. Top Priorities for Future Research There are several important issues and areas of research that need further investigation, these include: variations in the isotopic record due to insufficient sampling in time and space and the fundamental changes in the Earth System that have driven temperature change. 3.8.1. Insufficient Sampling in Time and Space As is the case for any geological investigation, due to the incompleteness of the geological record, there are bound to be gaps in the record of isotopic temperature. The temporal record of isotopic measurements compiled by Song is ~80% complete. That is to say, average isotopic temperatures have only been calculated for 400 of the 500 one million years intervals of the Phanerozoic. The data density can be directly viewed in Fig. 9. The black dots along the x-axis in Fig. 9B are the times for which there are no isotopic temperature measurements. The missing data is not randomly distributed. There are large data gaps during the: late Eocene (36 - 40 Ma), early Albian (110 - 115 Ma), Kimmeridgian -Tithonian (157-143 Ma), Norian (213 - 229 Ma), late Cisuralian (278 - 289 Ma), early Visean (336 - 346 Ma), Emsian (396 -406 Ma), earliest Devonian (417- 420 Ma), and early and middle Ordovician (455 - 495 Ma). We recommend that these gaps in the data record be priority targets for future research. The sampling of isotopic temperature measurements is also geographically sparse. Most of the sample sites are highly clustered and there are few sample sites from paleolatitudes above 40 N or S. Though the data set can be accurately characterized as "tropical", the possibility Global Average Temperature Strontium Flux Ratio Coincident Trends 540 520 500 480 460 440 420 400 380 360 340 320 300 280 260 240 220 200 180 160 140 120 100 80 60 40 20 0 millions of years Fig. 18. Comparison of Phanerozoic global average temperatures (black line) with strontium flux ratio relative to the present-day flux (dashed line). Red line highlights when the trends are coincident. 16 CR. Scotese et al. Earth-Science Reviews 215 (2021) 103503 remains that for some time periods the isotopic samples may be biased either towards the Equator, or away from the Equator. As we saw in the case of the late Pennsylvanian data, it is also important to identify environmental variables such as salinity or ocean circulation that may affect the isotopic temperature. One of our future research goals is to carefully analyze the isotopic temperature database for these kinds of geographic and environmental biases (Judd et al., 2020). 3.8.2. Modeling the Fundamental Causes of Global Temperature Change Correcting the isotopic artifacts and improving the temporal and geographic sampling of isotopic data are important, but they do not address fundamental questions regarding the ultimate causes of changes in the Earth's temperature. Long-term changes in temperature (>50 million years) and medium-term changes in temperature (10-20 million years) are primarily driven by the slow and inexorable changes brought on by plate tectonics and continental weathering (van der Meer et al., 2014; Berner, 1994; Brune et al., 2017). To the first order, the amount of CO2 released into the atmosphere is proportional to the rates of seafloor spreading and subduction (van der Meer et al., 2014). The isotopic composition of the oceans, in particular the ratio of 87Sr/86Sr , can be used as a proxy for the rate of sea floor spreading (van der Meer et al., 2017). New 86Sr is introduced into the oceans by the weathering of new basaltic oceanic lithosphere which is rich in non-radiogenic, depleted upper mantle-derived 86Sr. On-the-other-hand, continental sources are richer in 87Sr, which is a decay product of 87Rb. Faster periods of seafloor spreading introduce more 86Sr into the oceans. This decreases the ratio of 87Sr to86Sr of seawater. In Fig. 18, the non-radiogenic strontium flux ratio (van der Meer et al., 2017) ratio relative to the present (dashed line) is compared to our estimate of Phanerozoic global temperatures (solid line). There appears to be a remarkably good correlation between the global temperature and non-radiogenic strontium flux. For 65% of the Phanerozoic, when the ratio of 87Sr/86Sr in the oceans rose or fell (red line), global temperatures also rose or fell. The correlation is especially good for the last 180 million years (breakup of Pangea) and during the Permo-Carboniferous Ice Age (assembly of Pangea, 320 Ma - 255 Ma). For other time intervals, the isotopic and temperature records are in partial agreement (e.g., Cambrian, 540 - 500 Ma; Devonian - early Carboniferous, 400 - 345 Ma). However, there are also significant time intervals when the isotopic and temperature record are completely out-of-synch (marked by X's). Most notable are the Hirnantian Ice Age and the Triassic Hothouse. This suggests that there are other factors besides the volcanic degassing of CO2 that drive global temperature change. E.g. during the Hirnantian Ice Age, a massive bolide impact and subsequent positive ice-albedo feedbacks might have cooled the Earth despite the fact that plate tectonic activity was high at the time. During the Triassic, on-the-other-hand, the increased degree of aridification due to extreme hothouse temperatures (>40°C) and mega-monsoonal desertification across the interior of Pangea may have slowed down chemical weathering and allowed CO2 to build up in the atmosphere which increased global temperatures. We can speculate that other geological processes that occur with rates of change that match the temperature peaks and troughs might also have produced the observed pattern of Phanerozoic temperature change. For example, mountain building and unroofing, periods of Global Average Temperature 32° 30° 28° 26° 24° 22" 20° 18" 16° 14" 12° chronotemps Cambrian-Ordovician Hothouse Late Devonian Hothouse KjM Late Paleozoic Icehouse Permo-Triassic The 32 30° 28° 26° 24° 22° 20° 18° 16° 14" 12° CAMBRIAN ....... ORDOVICIAN DEVONIAN 1 1 1 ° i i i CARBONIFEROUS Fig. 19. A Paleozoic Paleotemperature Timescale. white = warm time intervals, black = cool time intervals. Solid black line = Global Average Temperatures (GAT), < 18°C = large permanent, icecaps, > 18°C = no large, permanent icecaps. Light gray jagged lines = a schematic representation of >50 glacial/interglacial cycles during the Permo-Carboniferous. Timescale = International Chronostratigraphic Chart V2020/01. Refer to Table 3 for more information about each chronotemp and abbreviations. 17 CR. Scotese et al. Earth-Science Reviews 215 (2021) 103503 Global Average Temperature chronotemps c13 [O IASS 11... I .. 111. JURASSIC ■ l.... I...... 1111 i i 111 111 11 11 i l i CRETACEOUS PALEOCENE Fig. 20. A Mesozoic Paleotemperature Timescale. white = warm time intervals, black = cool time intervals. Solid black line = Global Average Temperatures (GAT), < 18°C = large, permanent icecaps, > 18°C = no large, permanent icecaps. Timescale = International Chronostratigraphic Chart V2020/01.Refer to Table 3 for more information about each chronotemp and abbreviations. ophiolite obduction and subsequent chemical weathering, the opening of oceanic gateways, and some evolutionary events (e.g., the emergence of land plants) all occur on timescales comparable to the rates of change seen in the A TVop curve. A careful and systematic analysis of the evolutionary, paleogeographic, tectonic, paleoceanographic, and environmental changes that may have driven global temperatures will continue to be a fruitful area of research. Part II. Phanerozoic Paleotemperature Timescale 4. A Phanerozoic Paleotemperature Timescale The main goal of our research has been to construct a continuous record of temperature change during the last 540 million years (Fig.s 19-21). In the second half of this paper, we document this record of temperature change with specific paleotemperature events that are analogous to the "golden spikes" (GSSP) of the geological timescale (Gradstein et al., 2012). In some ways, this paleotemperature timescale is similar to the linear magnetic reversal time scale that was assembled at the beginning of the plate tectonic revolution more than 50 years ago. However, instead of normal and reverse magnetic polarity events, the paleotemperature timescale is defined by "warming" and "cooling" temperature intervals called "chronotemps". Our preliminary attempt to identify these warm and cool intervals is illustrated by the cool (black) and warm (white) subdivisions of the global average temperature curve (Fig. 19-21). The Phanerozoic paleotemperature timescale can be subdivided into 24 pairs of warm and cool intervals. It should be noted that the resolution of this version of the timescale is 1 million years; rapid temperature excursions shorter than a million years (e.g. PETM, K/T Impact Winter) are not well-resolved. The cool sections of the temperature timescale are labeled cl - c24. Some of these cooling events are well-known like the Pleistocene Ice Age (cl) or the Hirnantian Ice Age (c22). Other cool intervals are more speculative (c6, Aptian-Albian Cold Snap) and require further investigation. Similarly, the warm intervals are labeled wl - w24. Some of these warming events are well-known like the Permo-Triassic Thermal Maximum (wl3) and the Toarcian Warm Interval (w9). Other warm intervals are more speculative. Also, there is no strict equivalence between two cool intervals or two warm intervals. Two cool intervals may have very different minimum, maximum, and average temperatures (Tables 3 - 7). The same is true for the warm intervals. Each warm and cool time interval is illustrated in Fig.s 19-21 and supporting references are given in Tables 3-7. These tables provide precise information about the name of each interval, the start time and end time, the Tropical Temperature, the Polar Temperature, and the Global Average Temperature (GAT) of each interval, as well as the timing of major LIP eruptions and bolide impacts. Key citations for each chronotemp are also provided. Each interval may contain several discrete short-term temperature events or excursions. For example during the Paleocene-Eocene hothouse (W5, PEH), there were several thermal maxima, the PETM (W5.9, Paleocene-Eocene Thermal Maximum), EETM (W5.6, Early Eocene Thermal Maximum), and METM (W5.1, Middle Eocene Thermal Maximum). Conversely, a few temperature intervals may be grouped into a larger, geologically meaningful temperature interval; for example, the Late Paleozoic Icehouse (C13-C17) consists of the cool intervals: C13 (Late Permian Cooling), C14 Mid-Permian Cool Interval, C15 (Permo-Carboniferous Icehouse), C16 (Early Mississippian Cooling), and C17 (Famennian-Earliest Tournaisian 18 CR. Scotese et cd. Earth-Science Reviews 215 (2021) 103503 Global Average Temperature Mid-Cretsceous Hothouse chronotemps c6 c5 c4 c3 c2 cl ■ II ™ 1 ■ ■ I I I I I I I I I I I I I I I I I I ] I I I I I R ETACEO US 11111111 j 11111 PALEOGENE 1 I I I I I l. L.J-JL-L-Lx PALEOCENE EOCENE OLIGOCENE NEOGENE MIOCENE CRScot«e,2020 Fig. 21. A Cenozoic Paleotemperature Timescale. white = warm time intervals, black = cool time intervals; Light gray jagged lines = a schematic representation of >50 glacial/interglacial cycles during the Plio-Pleistocene. Solid black line = Global Average Temperatures (GAT), ), < 18°C = large, permanent icecaps, > 18°C = no large, permanent icecaps. Timescale = International Chronostratigraphic Chart V2020/01. Refer to Table 3 for more information about each chronotemp and abbreviations. Ice Age), as well as the intervening warm intervals. We have tried to standardize the names and abbreviations used to describe the various chronotemps depending on the length of a temperature event (icehouse, ice age, cool interval, cool event), its magnitude (e.g. thermal maximum, glacial maximum), or if the chronotemp represents a transition between warm or cool time intervals (e.g. warming or cooling). Often discrete temperature intervals are connected by prolonged periods of warming or cooling (e.g Eocene-Oligocene Transition). We have used a gray scale gradient in Fig.s 19-21 to schematically represent the rate of cooling or warming. We do not use the term "climate" as part of the naming convention to avoid confusion with broader climatic considerations (i.e., precipitation, temperature gradients, equability, etc.), and have avoided the use of other somewhat pejorative terms such as "optimum". Also, whenever possible we have used a specific stage name (e.g., Cenomanian-Turonian Thermal Maximum) rather than a more generic geological description (e.g., Cretaceous Thermal Maximum). We also prefer the term "interval" rather than "period" because of the geological connotation of the latter term. "Events" tend to be temperature phenomena that last only a few million years or less. "Excursions" are events that depart from the norm, i.e, a warm event during a cool interval. Unfortunately, in our attempt to rationalize and standardize the naming conventions, older more established names that have precedence have been mostly discarded (e.g., the Mid-Miocene Climatic Optimum (MMCO) has become the Middle Miocene Thermal Maximum (MMTM). What are the advantages to dividing the timescale into discrete temperature events and giving each of them a number and a name? Some may argue that our state of knowledge is still too primitive and that attempting to build a paleotemperature timescale is premature. We would argue that by constructing a paleotemperature timescale, even if this first version is imperfect, we now have a structure that can be built upon, refined, corrected, and expanded. Also having a paleotemperature timescale is essential if we are to understand the tempo and mode of climate change during Earth history. By characterizing, in a quantitative way, the pattern of paleotemperature change through time, we may be able to gain important insights into the history of the Earth System and the fundamental causes of climate change. 5. Climate Modes of the Phanerozoic 5.1. Introduction In this section we will review the evidence used to construct the Phanerozoic Paleotemperature Timescale (Tables 3-7, Fig.s 19-21). The 24 pairs of warm and cool temperature intervals listed in Tables 3-7 can be grouped into eight distinct "climate modes" (after Frakes, 1989; Frakes et al., 1992; Vaughan, 2007). Four of these climatic groups represent four warm periods in Earth history (Cambro-Ordovician, Siluro-Devonian, Triassic, and mid Cretaceous-Paleogene), three climatic groups represent cold periods in Earth History (Late Ordovician-Silurian, Late Paleozoic, and Late Cenozoic), and one "cool" period made up of mixture of cool and warm intervals (Jurassic - early Cretaceous). These warm and cool modes are similar to the eight warm and 19 Table 3 Paleotemperature Timescale: Late Ordovician - Silurian Icehouse and Cambro-Ordovician Hothouse, Chronotemps C21 - C24 (ICS Time Scale vl026/04) Climate Mode Chrono AT Tropical Polar GAT Start End Start End * Also Known As Bolide Impacts Large Igneous Provinces Key References Temp C° C° C° C° Name Abbr. Age Age Stage Stage C21-C22 C21-C22 -0.8 25.2 -11.9 15.4 Late Ordovician-Silurian LOSI 458 426 early Sandbian earliest Ludlow Icehouse Late Ordovician - C21 -0.1 25.9 -6.4 17.2 Wenlock Cool Interval WCI 434 426 earliest earliest Ludlow 39,40,42 Wenlock Silurian Icehouse C21.1 -1.0 25.0 -7.5 16.3 Mulde Cool Event MCE 429 428 late Wenlock late Wenlock 39,42 426 - 458 Ma C21.2 0.1 26.1 -5.3 17.6 Ireviken Cool Event ICE 433 432 early Wenlock early Wenlock 39,41,42 AT = 32 W22 1.1 27.1 -5.9 18.3 Llandovery Warming LW 439 434 middle late Llandovery 38,39,40,41,42,46,49 Llandovery C22 -1.6 24.4 -15.6 13.9 Late Ordovician Icehouse LOI 458 440 early Sanbdian early Llandovery Slate Islands 38,4143,44,45,46,47,49 450 Ma C22.1 -3.3 22.7 -29.6 9.3 Hirnantian Ice Age HIA 445 441 early early Llandovery HICE, 43,44,45,46,47,49 Hirnantian Hirnantian Extinction Cambro - W23-C24 2.7 28.7 5.3 22.1 Cambro-Ordovician COH 541 458 basal earliest Sandbian Hothouse Cambrian Ordovician W23 0.3 26.3 -4.0 18.1 Darwinian Warm Interval DWI 470 458 earliest earliest Sandbian Clearwater 43,44,45,47,48,49 Dapingian East -465 Ma Hothouse C23 1.2 27.2 1.6 20.5 Early Ordovician Cooling EOC 485 470 early earliest Dapingian 43,44,45,47 Tremadocian W23-W24 W24 3.1 29.1 6.3 23.8 Middle-Late Cambrian MLCH 521 485 latest late Furongian 111,112, Hothouse Terraneuvian 458 - 541 Ma W24.1 2.6 28.6 3.7 23.0 Top of Cambrian Warm TCWE 486 485.5 latest latest Furongian TOCE 111,112, Event Furongian AT = 83 W24.2 1.8 27.8 2.1 22.3 Steptoean Cool Event SCE 497 494 early Paibian late Paibian SPICE 111,112, W24.3 3.7 29.7 6.2 24.4 Drumian Warm Event DWE 505 504.5 earliest earliest Drumian DICE 111,112, Drumian W24.4 3.2 29.2 7.7 24.2 Redlichiid-Olenid ROEWE 509 508.5 earliest "Stage earliest "Stage 5" ROECE 111,112, Extinction Warm Event 5" W24.5 4.5 30.5 9.5 25.6 Archaecyathid Extinction AEWE 511 510.5 middle "Stage middle "Stage 4" AECE Kalkarindja 111,112, Warm Event 4" 512-509 Ma W24.6 2.3 28.3 9.1 23.5 Mingxinsi Cool Event MCE 514 512 earliest "Stage 4" early "Stage 4" MICE 111,112, W24.7 3.0 29.0 11.8 24.1 Cambrian Arthropod CARCE 521 517 earliest "Stage middle "Stage 3" CARE 111,112, Radiation Cool Event 3" C24 4.6 30.6 12.2 22.6 Early Cambrian Cool ECCI 541 521 early late Terreneuvian 111,112, Interval Terreneuvian C24.1 4.2 30.2 12.8 24.3 Shiyantou Warm Event SWE 525 524.5 late late Terreneuvian SHICE 111,112, Terreneuvian C24.2 7.1 33.1 13.4 22.1 Zhujiaqing Cool Event ZCE 529 526 middle late Terreneuvian ZHUCE 111,112, Terreneuvian C24.3 7.1 33.1 13.4 22.1 Basal Cambrian Warm BCWE 541 536 earliest early Fortunian BACE 111,112, Event Fortunian Abbreviations: * MMCO = Mid-Miocene Climatic Optimum, OCE = Oligocene Cooling Event, EOT = Eocene Oligocene Transition, MECO = Middle Eocene Climatic Optimum, ETM = Eocene Thermal Maximum, OAE = Oceanic Anoxic Event, MDE = Middle Danian Event, LDE = Late Danian Event, DAN = Danaina, KT = Cretaceous/Tertiary, K-Pg = Cretaceous - Paleogene, D-a through D-m = Dera et al. (2011), ETE = End Triassic Event, R-TRI through Trll = Retallack (2013), NCIE = Norian Carbon Isotope Excursion, FP1 through FP4 = Fielder et al. (2008) Permian glacial interval, FC1 through FC4 = Fielder et al. (2008) Carboniferous glacial interval, D-I through D-XV = Davydov et (2010), M-Cl and M-C2 through Mii et al. (2013) Carboniferous, M-D = Mii et al. (2013) Devonian, HICE= Hirnantian Carbon Isotope Excursion, TOCE = Top Cambrian Carbon Excursion, SPICE = Steptoean Positive Carbon Isotope Excursion, DICE = Drumian Carbon Isotope Excursion, ROECE = Redlich.-Olenel. Extinction Carbon Excursion, AECE = Archaeo. Extinction Carbon Excursion, MICE = Mingxinsi Carbon Excursion, (conünued on next page) Table 3 (continued) o Climate Mode Chrono AT Tropical Polar GAT Start End Start End * Also Known As Bolide Large Key References Impacts Igneous Provinces CARE = Cambrian Arthropod Radiation Carbon Excursion, SHICE = ShiYantou Carbon Excursion, ZHUCE = Zhujiaqing Carbon Excursion, BACE = Basal Cambrian Carbon Excursion, GIA = Gaskers Ice Age, MIA = Marinoanlce Age, SBE = Snowball Earth Key References: I Davydov and Cozar, 2017; 2 Montanez et al 2006; 3 Davydov , 2014 ;4 Montanez and Poulsen, 2013; 5 Davydov et al., 2010; 6 Schmitz and Davydov, 2012; 7 Fielding et al., 2010; 8 Grossman et al., 2008; 9 Fielding et al., 2008; 10 Buggisch et al., 2008; II Mii et al., 1999; 12 Bruckschen et al., 1999; 13 Retallack, 2013; 14 Waterhouse and Shi, 2013;; 15 Mii et al., 2012; 17 Chen et al., 2013; 18 Frank et al., 2015; 19 Haig et al., 2014; 20 Henderson et al., 2012 21 Korte et al., 2008; 22 Korte et al., 2005; 23 Mii et al., 2013; 24 Montanez et al., 2007; 25 Retallack et al., 2011; 26 Zeng et al., 2012; 27 Sun et al., 2012; 28 Podlaha et al., 1998; 29 Jenkyns et al., 2002; 30 Malchus and Steuber, 2002 31 Weirzbowski and Joachimski, 2007; 32 Dera et al., 2011; 33 Vickers et al., 2019; 34 Joachimski et al., 2004; 35 Van Gelderin et al., 2006; 36 Joachimski et al., 2009; 37 Racki, 2020; 38 Azmy et al., 1998; 39 Wenzel et al., 2000, 40 Brand et al., 2006; 41 Munnecke et al., 2010; 42 Trotter et al., 2016; , 43. Trotter and Barnes, 2008; 44 Young et al., 2009; 45. Munnecke et al., 2010; 46 Finnegan et al., 2011; 47 Lenton et al., 2012; 48 Quiron and MacLeod 2014; 49 Bond and Grasby, 2017; 50 Brett et al., 2009 51 Trotter et al., 2015; 52 Preto et al., 2010; 53 Sun et al., 2016; 54 Rigo and Joachimski, 2010; 55 Miller et al., 2017; 56 DalCorso et al., 2018ab; 57 Bernardi et al., 2018; 58 O'Brien et al., 2017; 60 Alsenz et al., 2013; 61 Brassel 2009; 62 Erba et al., 2015; 63 De Lurio and Frakes, 1999; 64 Frakes and Francis, 1988; 65 Grasby et al., 2017; 66 Vickers et al., 2019; 67 Erba et al., 2004; 68 Herrle et al., 2015; 69 Barral et al., 2017; 71 Farnsworth et al., 2019; 72 Mutterlose 2009; 73 Jenkyns et al., 2012; 74 Navarro-Ramirez et al., 2017; 75 Ladant and Donnadieu, 2016; 76 McArthur et al., 2007; 77 Pirrie et al., 1995; 78 Pirrie et al., 2004; 79 Price, 1999; 80 Price and Mutterlose, 2004 81 Price and Nunn, 2010; 82 Price and Passey, 2013; 83 Puceat et al., 2003; 84 Ogg et al., 2016; 85 Vandenberghe et al., 2012; 86 Miller et al., 2020; 87 Gasson et al., 2020; 88 Shackelton et al., 1984; 89 King, 2016; 90 Scotese, 2020 91 Upchurch et al., 1999; 92 Upchurch et al., 2015; 93 Kiehl and Shields, 2013; 94 Kump and Pollard, 2008; 95 Hay et al.,2016) 96 Vellekoop et al., 2014, 2016; 97 Li and Keller, 1998; 98 Wilf et al., 2003; 99 Punekar et al., 2014; 100 Petersen et al., 2016 101 Zachos et al., 2001, 2008; 102 Rea et al., 1990; 103 Kennett and Stott, 1991; 104 Mclnerney and Wing, 2011; 105 Zeebe et al., 2009; 106 Lawver and Gahagan, 2003; 107 Lisiecki and Raymo, 2005; 108 Archer, 2009; 109 Hilgen et al., 2012; 110 Pillans and GIbbard, 2012; 111 Peng et al., 2012; 112 Zhu et al., 2006; 113 Westerhold et al., 2020;114 (Alley and Frakes, 2003); 115 Wilf et al., 2003 , 116 Gasson and Kiessling, 2020 Table 4 Paleotemperature Timescale: Late Paleozoic Icehouse and Siluro-Devonian Hothouse, Chronotemps C13 - W21, (ICS Time Scale V1026/04) Climate Mode Chrono AT Tropical Polar GAT Start End Start End * Also Known Bolide As Impacts Large Igneous Provinces Key References Late Paleozoic Icehouse C13-C17 253 - 365 Ma AT = 112 Siluro - Devonian Hothouse W18-W21 365 - 426 Ma AT = 61 Temp C13-C17 C13 W14 C14 W15 C15 C15.1 C15.2 C15.3 C15.4 C15.5 C15.6 C15.7 C15.8 W16 C16 W17 C17 C17.1 W18-W21 W18 W18.1 W18.2 W18.3 C° 0.0 1.9 1.7 0.1 -1.3 -1.6 C° 26.0 27.9 27.7 26.1 24.7 24.4 -1.4 24.6 -0.6 -1.2 -2.7 -3.6 -1.7 -0.8 0.5 1.4 4.0 3.6 2.9 1.9 4.4 5.1 7.0 4.9 25.4 24.8 23.3 22.4 24.3 25.2 26.5 27.4 30.0 29.6 28.9 27.9 30.4 31.1 33.0 30.9 C° -9.0 -0.2 -3.3 -7.3 -13.8 -16.2 C° 16.7 20.2 19.3 17.2 14.7 13.9 -20.2 13.1 -22.9 -22.9 -18.1 -18.9 -7.6 -4.0 -1.3 0.6 1.5 -2.1 -2.2 -0.7 4.2 6.2 8.1 6.1 13.0 12.6 12.6 11.7 15.8 17.2 18.8 20.0 22.3 21.1 20.6 20.2 23.2 24.2 26.1 24.0 Name Late Paleozoic Icehouse Latest Permian Cooling Emeishan Thermal Excursion Mid-Permian Cool Interval Artinskian- Kungurian Warming Permo- Carboniferous Icehouse Permo- Carboniferous Glacial Maximum Permo- Carboniferous Ice Age Latest Gzhelian Thermal Excursion Mid-Gzhelian Cool Event Bashkirian-Moscovian Cool Event Mid-Bashkirian Cool Event Serpukhovian-Bashkirian Cool Event Early Serpukhovian Cooling Visean Warm Interval Early Mississippian Cooling Early Tournaisian Warm Interval Famennian-Tournaisian Ice Age Hangenberg Cool Event Late Silurian -Devonian Hothouse Late Devonian Hothouse Late Kellwasser Event Kellwasser Thermal Maximum Early Kellwasser Event Abbr. Age Age LPI 365 253 LPC 259 253 ETX 262 259 MPCI 272 262 AKW 287 272 PCI 330 287 MBCE SBCE ESC VWI EMC ETWI FTIA HC E SLDH LDH LKE KTM EKE Stage late Famennian Wuchipingian middle Capitanian early Roadian late Artinskian early Serpukhovian Stage latest Changsingian latest Changhsingian F-P4 late Capitanian Kimura Event, R-P7 middle Capitanian F-P3, D-XV, R-P6 latest Kungurian R-P4,P5 middle Artinskian M-C3 Araguainha 255 Ma PCGM 299 293 earliest Asselian early Sakmarian PCIA 313 291 early Moscovian late Sakmarian LGTX 300 299 MGCE 302 301 BMCE 315 312 318 317 324 322 330 326 342 330 353 342 358 353 365 358 361 358 426 365 378 365 372 372 372.5 372.5 373 373 latest Gzhelian middle Gzhelian latest Bashkirian middle Bashkirian latest Serpukhovian early Serpukhovian early Visean late Tournaisian early Tournaisian late Famennian late Famennian early Ludlow middle Frasnian latest Frasnian latest Frasnian late Frasnian latest Gzhelian middle Gzhelian early Moscovian F-P1.P2; D-XII, XIII,XIV;R-P2, P3 —50 glacial/interglacial cyclothems R-Pl D-X F-C4, D-VIII middle Bashkirian F-C3, D-VII earliest Bashkirian F-C2, D-VI middle Serpukhovian F-Cl.D-V M-C2 latest Visean early Visean middle Tournaisian M-D earliest Tournaisian D-I,II latest Famennian late Famennian late Famennian earliest Famennian latest Frasnian latest Frasnian M-Cl, D-III,IV Charlevoix 342 Ma Woodleigh 364 Ma Late Devon. OAE, Enkeberg, Condroz Emeishan 260-259 Ma Tian Shan 330-329 Ma Kola Dneiper 358 Ma E. European 365-364 Ma Kola Dneiper 1 373-372 Ma Viluy Rift 375-374 Ma 3,7,14,17,18 3.12,13,17,25 3,8,7.13,14,15,18,21,22,23,24 3,11,13,14,15,19,20,21,22,23,24,26 3,4,5,14,15,16,21,22,23 2,4,5,15,18,20,21,22,23,24,26 2,4,5,6,7,9,13 3,5,8,12,13 2,3 5,9 6.9 1,3,5,9 1,4,5.8,9,10,11 4,5,8,10,11,12 4,5,8,11,12 4,5,8,10,11,12 5.10 34.35.36 34.35.37 34,35,36,37 37, (continued on next page) Earth-Science Reviews 215 (2021) 103503 higher latitude waters of Baltica and Gondwana. In the "cool" temperature model, global temperatures increased during the middle and late Cambrian as the continental cratons were flooded and the resulting decrease in the planetary albedo warmed the Earth (Landing and Westrop, 2004; Landing, 2012). Temperatures continue to fall during the middle Ordovician (Trotter et al., 2008). Trotter and Barnes (2008) proposed that sustained, cooler tropical temperatures (26°-30°C) provided a more hospitable environment for evolutionary innovation and may have been the impetus for the Great Ordovician Biodiversification Event (GOBE; Paris et al., 2004). As noted earlier (section 3.1.2.1), the ~2° temperature fluctuations during the Cambrian shown in Fig. 19 are speculative. They coincide with 10 proposed Cambrian d13C isotopic excursions (Zhu et al., 2006). It should be noted that the global synchroneity of some of these excursions (in particular TOCE and REOCE) has been disputed (Landing, personal communication). The covariance of d13C and dlsO trends has long been noted (Wenzel and Joachimski, 1996; Jenkyns et al., 2002 ). Approximately 80% of the positive d13C excursions are correlated with warmer temperatures (hyperthermals). The correlation is generally attributed to the causal relationship between higher ocean temperatures, the formation of deep water anoxia, and the subsequent preservation of organic carbon. 5.3. Late Ordovician - Silurian Icehouse (C21 - C22, 426-458 Ma), Table 3, Fig. 19 The most spectacular, short-term cooling event of the Phanerozoic is the Hirnantian Ice Age (445 - 441 Ma, Brenchley et al., 1994; Sheehan and Coorough, 1990; Sheehan, 2001; Finnegan et al., 2011). Whether the Hirnantian Ice Age (C22) was preceded by a prolonged, stepwise cool-down is still debated. Biostratigraphic (Brenchley et al., 1994) and geochemical data (Finnegan et al., 2011) indicate that the maximum glacial advance was very short-lived (less than one million years; Ling et al., 2019). The repercussions of the Hirnantian Ice age were felt worldwide and left indelible signatures in the geochemical, paleonto-logical (Sheehan and Coorough, 1990; Finnegan et al., 2012), and sedimentary rock record . The widespread occurrence of latest Ordovician tillite deposits (Beuf et al., 1971; Boucot et al., 2013) indicates that ice sheets covered ~50% of Gondwana and extended to latitudes of ~35 S. The enormous extent of snow and ice cover increased the Earth's albedo and triggered ice-albedo feedbacks that rapidly cooled the Earth. The asymmetric latitudinal disposition of the Earth's land masses during the late Ordovician, namely the fact that no large continents occupied the northern hemisphere, probably prevented the Earth from slipping into another "Snowball Earth". Only thin sea-ice accumulated on the freely circulating, open northern oceans. This meant that during the late Ordovician, the northern hemisphere remained relatively warm and prevented a runaway Snowball Earth-like global freeze. Multiple explanations have been proposed for the cause of the Hirnantian Ice Age. Crowley and Baum (1991, 1995) suggested that the growth of the Gondwana ice cap was facilitated by the combination of increased precipitation and cold temperatures along the northern margin of Gondwana. Other authors have invoked increased chemical weathering of young mountains (Taconic ranges) or recently obducted ophiolites that lead to a drawdown in atmospheric CO2, which promoted globally cooler temperatures (Kump et al., 1995, 1999b,c; Swanson-Hysell and Macdonald, 2017; Landing, 2018). This effect may have been enhanced by the evolution of simple, non-vascular land plants (Lenton et al., 2012). A more spectacular explanation, as discussed in section 3.3, is that the Hirnantian Ice Age was triggered by a bolide impact as large or larger than the Chicxulub impact. There is no geologic record of this massive impact on the continents. Though the impact site may be buried under younger continental strata, it is more likely that the Khione impact targeted the late Ordovician ocean basin and that the impact Table 5 Paleotemperature Timescale: Jurassic - Early Cretaceous Cool Interval and Triassic Hothouse, Chronotemps C7 - W13 (ICS Time Scale vl026/04) Climate Mode Chrono AT Tropical Polar GAT Start End Start End * Also Known As Bolide Impacts Large Igneous Provinces Key References Temp C° C° C° C° Name Abbr. Age Age Stage Stage Jurassic - Early C7-C9 1.7 27.7 3.5 21.0 Late Jurassic - Early Cretaceous Cool Interval JKCI 199 128 latest Hettangian early Barremian Cretaceous C7 0.5 26.5 1.3 19.5 Tithonian-early Barremian Cool Interval TEBCI 150 128 early Tithonian early Barremian Tookoonooka 128 Ma 28,29,32,58,64,66,67,114 Cool Interval Mj0lniri 142 Ma C7-C9 Morokweng 145 Ma 128 - 199 Ma C7.1 3.0 29.0 3.3 21.9 Faraoni Thermal Excursion FTX 131 131 middle Hauterivian middle Hauterivian Parana-Entendeka 131-129 Ma 84, AT = 71 C7.2 -0.3 25.7 0.2 18.6 Weissert Thermal Excursion WTX 136 136 middle Valanginian middle Valanginian 84, C7.3 0.6 26.6 2.1 19.8 Early Tithonian Cooling ETC 150 149 early Tithonian middle Tithonian W8 1.4 27.4 3.3 20.6 Kimmeridgian Warm Interval KWI 164 150 earliest Oxfordian early Tithonian Kimmeridgian Anoxic Event 28,29,30,32,33,58,61, 66,67,73,76 W8.1 1.2 27.2 3.0 20.5 Early Tithonian Cool Event ETC E 152 152 early Tithonian early Tithonian D-m W8.2 2.9 28.9 5.0 22.2 Latest Oxfordian Thermal Maximum LOTM 158 157 late Oxfordian latest Oxfordian D-l 32, W8.3 0.0 26.0 2.1 19.3 Middle Oxfordian Cooling MOC 161 158 middle Oxfordian late Oxfordian D-k W8.4 1.4 27.5 3.1 20.7 Late Callovian - early Oxfordian Warming cow 164 161 late Callovian early Oxfordian D-j C8 2.9 28.9 4.2 22.0 Middle Jurassic Cool Interval MJCI 174 164 earliest Aalenian earliest Oxfordian Puchezh-Katunki 167 Ma 28-32,60,61,64-67,73, 76,79-82 C8.1 0.1 26.1 2.1 19.4 Callovian Cool Event CCE 166 164 early Callovian late Callovian D-i 28,29,32 C8.2 3.5 29.5 5.0 22.7 Late Bathonian Warming LBW 167 166 late Bathonian latest Bathonian D-h 32, C8.3 3.7 29.7 5.1 22.9 Late Toarcian - Bajocian Cooling LTBC 178 168 late Toarcian late Bajocian D-e,f,g 28,29,32 W9 4.0 30.0 5.2 23.1 Toarcian Warm Interval TWI 183 174 early Toarcian latest Toarcian D-c,d Karoo-Ferrar 183-181 Ma 29,32,64 W9.1 Early Toarcian Thermal Maximum ETTM 183 180 early Toarcian late Toarcian Toarcian_OAE,D-c 28,32 C9 1.5 27.5 4.9 21.2 Early Jurassic Cool Interval EJCI 199 183 earliest Sinemurian late Pliensbachian D-a,b 28,29,32 Triassic Hothouse W10-W13 5.1 31.1 9.0 24.9 Triassic Hothouse TH 253 199 latest Changhsingian latest Hettangian W10-W13 W10 3.7 29.7 7.2 23.4 End Triassic Thermal Event ETTE 201 199 latest Rhaetian earliest Hettangian Triassic/ Jurassic_OAE, ETE CAMP 201-199 Ma 28,29,32 199 - 253 Ma W10.1 4.4 30.4 8.0 24.1 Central Atlantic Magmatic Province CAMP 200 200 latest Rhaetian latest Rhaetian 28,29,32 AT = 54 CIO 2.4 28.4 6.1 22.1 Rhaetian Cool Interval RCI 209 201 earliest Rhaetian late Rhaetian 51, Wll 5.6 31.6 8.7 25.2 Late Norian Warm Interval LNWI 217 209 middle Norian late Norian R-Trll Manicouagan 214 Ma 13,51 Wll.l 6.0 32.0 8.5 25.5 Manicouagan Cool Excursion MCX 214 214 middle Norian middle Norian Cll 3.7 29.7 8.2 23.6 Early Norian Cool Interval ENCI 227 217 earliest Norian middle Norian Saint Martin 220 Ma 51,57 W12 4.3 30.3 8.8 24.2 Mid-Carnian Warm Interval MCWI 234 227 middle Carnian latest Carnian R-Trl0 Wragellia LIP 230 Ma 13,51,52,53,54,55,56,57 W12.1 4.8 30.8 9.4 24.7 Carnian Pluvial Episode CPE 233 230 middle Carnian middle Carnian NCIE 1-4 13,51,52,53,54,55,56,57 C12 5.0 31.0 9.3 24.8 Ladinian-Carnian LCC 242 234 earliest Ladinian middle Carnian R-Tr8-9; Ladinian 51,53,54,55,57 Cooling Humid Interval (continued on next Earth-Science Reviews 215 (2021) 103503 crater was subsequently subducted. This hypothesis was first proposed in 1986 by Kent Colbath who noted the rapid extinction of tropical acritarch genera in the southern Appalachian basin and their subsequent replacement with more cosmopolitan taxa in the earliest Silurian (Colbath, 1986). The extinction of warm, shallow water tropical faunas during the latest Ordovician and their replacement by cool-water benthic taxa or cosmopolitan planktonic taxa has been confirmed by numerous other studies (Sheehan, 2001). The Hirnantian extinction event ranks second in terms of taxonomic severity (~46% marine genera extinct). Only the Permo-Triassic Extinction (58% marine genera extinct; Sepkoski Jr., 1996; Bambach et al., 2004) was more cataclysmic. Though some snow and ice in the polar regions of South America (central Brazil and Bolivia) lingered into the earliest Silurian (Grahn and Caputo, 1992; Grahn and Gutierrez, 2001; Diaz-Martinez and Grahn, 2007), rapid warming during the Llandovery (W22) was followed by a period of slow ecological recovery. The remainder of the early-middle Silurian was characterized by a moderate pole-to-equator temperature gradient which was cool by early Paleozoic standards but warmer than today's world (Moore et al., 1994). Most early and middle Silurian tropical taxa were widely distributed (i.e., cosmopolitan) with the exception of the high latitude, cool-water brachiopods, Clarkeia (southern hemisphere; Boucot, 1990; Benedetto and Sanchez, 1996) and Tuviella (northern hemisphere; Cocks, 1972). Three carbon isotopic excursions occurred during the middle and late Silurian. Two of these excursions (Ireviken and Mulde) bracket the Wenlock and may have been caused by the growth of ephemeral, south polar ice caps (Brand et al., 2006). 5.4. Silwo-Devonian Hothouse (W18 - W21, 365-426 Ma) Table 4, Fig. 19 The late Silurian through late Devonian hothouse includes four warm intervals (W21 - W18) separated by three relatively cool episodes (C20-C18). There is no geological evidence for permanent polar icecaps during any of the cool intervals (Boucot et al., 2013). The following summary of the Siluro-Devonian warm mode is based on the isotopic evidence summarized by Joachimski et al. (2009), the biogeographic information synthesized by Stock (2005), the lithologic indicators of climate assembled by Boucot et al. (2013), and a comprehensive synthesis of the Frasnian-Famennian crisis by Racki (2020). The Siluro-Devonian Warm Interval spans the early Ludlow to early Lochkovian time interval (Joachimski et al., 2009; Munnecke et al., 2010; Trotter et al., 2016). It is followed by a pair of brief cooling and warming events during the Pragian (C20) and early Emsian (W20), respectively. The Pragian Cool Event coincides with a widespread fall in sea level (~100m) during the earliest Devonian and the resulting unconformity marks the end of the Tippecanoe Supersequence. This fall in sea level is seen on most continents, but is especially prominent on North America (Schuchert, 1910, 1955; Sloss, 1963). After a brief warming event in the early Emsian, temperatures continued to fall during the remainder of the Emsian and into the middle Devonian (Eifelian - middle Givetian; Joachimski et al., 2004, 2009; van Gelderin et al., 2006). The Early-Middle Devonian Cool Interval (C19, GAT = 18.7°C) was followed by a brief, rapid rise in temperature during the late Givetian (W19, Givetian Thermal Maximum, GAT = 21.2°C). Late Givetian warming coincided with the Taghanic onlap (Brett et al., 2009), the start of a global transgressive cycle that would culminate with the highest sea level of the Paleozoic (late Frasnian; Haq et al., 1987; Haq and Schutter, 2009). Global temperatures cooled briefly during the early Frasnian but rose steadily through the remainder of the Frasnian, culminating in the Kellwasser Thermal Maximum (W18.2, 372.5 Ma). A recent synthesis of Frasnian and Famennian isotopic, paleontologic, and tectonic information (Racki, 2020) provides numerous detailed insights into the paleoclimatic events that took place across the Frasnian-Famennian boundary. According to Racki (2020), the Kellwasser Event, also known as the C.R. Scotese et al. Earth-Science Reviews 215 (2021) 103503 Frasnian-Famennian Extinction Event, can be subdivided into two pairs of rapid warming/cooling events. The first event, the Early Kellwasser Event (373 Ma), may have been triggered by the eruption of the Vilui and Kola LIPS (see Table 1). Shortly thereafter, continued volcanic eruptions (Pripyat-Dneiper-Donets LIP) triggered a second episode of rapid warming and cooling (Late Kellwasser Event; 372 Ma). An alternate hypothesis invoking a bolide impact has also been proposed (McLaren, 1970, 1983). The combined Kellwasser Extinction Event was the fourth most severe extinction event when measured in terms of ecological severity (McGhee et al., 2013). It is associated with highest sea level of the Paleozoic (Johnson, 1988; Johnson et al., 1985; Sand-berg et al., 2000) as well as with a dramatic increase in 87Sr/86Sr composition of seawater (Zhang et al., 2020). For a more detailed account of the Frasnian-Famennian Extinction Event see McGhee Jr. (1996, 2005), Racki (2005), and Sandberg et al. (2000, 2002). The isotopically inferred temperature events outlined above are in good agreement with the paleoclimatic conditions inferred from pale-obiogeography and lithologic indicators of climate. According to Boucot et al. (2013), the global climatic gradient during the Early Devonian through to the Eifelian was "moderate", resulting in cool but not freezing temperatures in temperate latitudes. Like much of the earlier Paleozoic, the high southerly latitudes of Gondwana were characterized by an exclusively clastic facies with unweathered mica flakes and a distinctive cool-water fauna (Malvinokaffric province). Conodonts (Girard et al., 2005) and reef-like stromatoporoids (Stock, 2005), which preferred warm waters, were absent from the cooler, temperate shallow seas of central and southern Gondwana. Global temperatures warmed in the Givetian and Frasnian; and the Malvinokaffric province was eliminated, replaced by warmer water faunas. Faunas at low latitudes became less endemic and more cosmopolitan (Stock, 2005). Temperatures continued to warm throughout the Frasnian and Famennian, until the onset of the Late Famennian Ice Age (discussed in the next section). 5.5. Late Paleozoic Icehouse (C13 - C17, 253-365 Ma) Table 4, Fig. 19 Spanning more than 100 million years, the Late Paleozoic Icehouse was the longest interval of cold climates during the Phanerozoic. Polar ice existed, more or less continuously, at either the South or North Poles with the exception of brief warming events during the Mississippian (Tournaisian and Visean Warm intervals, W17 and W16, respectively), latest Pennsylvanian (Latest Gzhelian Thermal Excursion, CI 5.3,) and late Permian (Emeishan Thermal Excursion, W14). Global average temperatures ranged from 13° C during the depths of the Permo-Carboniferous Ice Age to ~22°C during the Early Tournaisian Warm Interval, the warmest time period during the Carboniferous or Permian. The Late Paleozoic Icehouse began in the latest Devonian with the formation of a short-lived, medium-sized south polar icecap (Caputo et al., 2008). This ice cap, centered in Brazil, stretched from the proto-Andean mountains, eastward to Rio de Janeiro and Gabon, and northeastward into equatorial west Africa and Niger. The Famennian Ice Age (C17) lasted less than 5 million years and is correlated with the widespread extinction of marine faunas. This late Devonian extinction event ranks fourth in terms of taxonomic severity (50% genera lost; McGhee et al., 2013). During the early Mississippian (Tournaisian), global climate first warmed (W17), then cooled dramatically (C16; Grossman et al., 2008). Throughout the remainder of the Mississippian, Gondwana moved steadily northward across the South Pole. Ice sheets contracted, then expanded, moving southward into south-central Argentina. The south polar ice cap nearly vanished during the Visean Warm Interval (W16), retreating to the highlands of the proto-Andes of Bolivia and western Argentina, as well as into the remnants of the Pan-African mountain ranges in southeastern Brazil and southwest Africa (Fielding et al., 2008a, 2008b). These latest Devonian and Mississippian machinations were just a prelude to the impending Permo-Carboniferous Ice Age (313 - 291 Ma.) Beginning in the early Serpukhovian (latest Mississippian, 330 - 326 Ma), the world began to cool once again. However, this time it would remain cool, with Global Average Temperatures (GAT) less than 18°C for nearly 80 million years. Starting in the earliest Pennsylvanian (Bashkirian, 323.2 Ma), a large ice cap began to develop at the South Pole and grew equatorwards reclaiming the southern reaches of the Amazon basin, crossing south-central Africa, and extending across Antarctica and most of Australia (Mory et al., 2008). Though often portrayed as a single, large Ice sheet (Scotese et al., 1999), the South Polar Ice Cap was actually composed of several growing, glacial nuclei (Fielding et al., 2008a, 2008b; Fielding et al., 2010; Montahez and Poulsen, 2013) that coalesced and expanded equatorwards reaching latitudes of 35° south by the mid-Pennsylvanian (315.2 Ma, Moscovian). Beginning in the early Moscovian (~313 Ma), the growing South Polar Ice Cap began to rhythmically wax and wane in synchrony with the changes in the shape of Earth's orbit (Milankovitch cycles; Milan-kovitch, 1920; Kump et al., 1999b; Hay, 2016) that regulated the amount of solar energy received by the Earth. The expansion and contraction of the South Polar Ice Cap caused cyclical changes in sea level. Paired transgressions and regressions, separated by —400,000 years, pulsed across the broad, flat continental cratons generating more than 55 repeating sedimentary packages called "cyclothems" (Wanless and Weller, 1932). Cyclothems have been mapped across the mid-continent of North America (Illinois and Mid-Continent Basin; Heckel, 2013) and, to a lesser extent, the Appalachian basin. Cyclothems have also been identified in similar aged rocks in Donets Basin of the Ukraine (Montahez et al., 2007, 2016). The cyclothems are direct evidence of Permo-Carboniferous glacial-interglacial cycles (Wanless and Shepherd, 1936; Heckel, 1994, 2008; Montahez and Poulsen, 2013) and are identical in nature to glacial-interglacial sedimentary sequences deposited during the Pleistocene Ice Age (CI.3). The oldest cyclothems in the mid-continent are middle Pennsylvanian in age (Desmoinesian /early Moscovian, ~313 Ma; Heckel, 2013). During one particularly warm interglacial episode, the Latest Gzhelian Thermal Excursion (C15.3), global temperatures may have warmed sufficiently to have temporarily reduced or eliminated the South Polar Ice Cap (Davydov et al., 2010). However, whatever the cause, the warmth did not last and glacial-interglacial cycles resumed in earnest. The youngest cyclothems are early Permian in age (late Sak-marian, 291 Ma). Cyclothems bracket the coldest phase of the Permo-Carboniferous Ice Age (C15.2, GAT = 13°C) and disappear shortly after the Permo-Carboniferous Glacial Maximum (C15.1, 299 - 293, early Asselian - early Sakmarian; Montahez and Poulsen, 2013). The deep oceans were "refrigerated" during much of the Carboniferous and early Permian (Valdes et al., 2020). Cold bottom-waters generated by the seasonal melting of the South Polar Ice Cap filled the oceans with near-freezing waters from the bottom up. Along some continental margins, this cold bottom-water was carried to the surface by upwelling. Glendonites formed in shallow marine sediments where the temperatures were < 4°C (De Lurio and Frakes, 1999). Glendonites occur in association with the shallow water carbonates of the Rundle Group of western Canada which was located near the equator during the Visean (Brandley and Krause, 1994). At south polar latitudes, numerous glendonites are found in fine-grained clastic sediments deposited along the margins of Gondwana (Boucot et al., 2013). In addition, the thermal continuity of these cool shelf and deep shelf environments allowed early Permian marine faunas to migrate freely between cool temperate southern latitudes, across the tropics, and into cool temperate northern latitudes (Waterhouse and Shi, 2013; Shi, 2001). By the middle Artinskian (285 Ma), global temperatures had warmed (W15) signalling the end of the Permo-Carboniferous Ice Age (Ziegler et al., 1997). Though the large South Polar Ice Cap was gone for good, intermontane glaciers inhabited the highlands of eastern Australia, the Trans-Antarctic Ranges of Eastern Antarctica (Frank et al., 2015), and the far northern, mountainous reaches of Siberia until the end of the Permian. 26 Table 6 Paleotemperature Timescale: Mid Cretaceous - Paleogene Hothouse, Chronotemps W5 - W7, (ICS Time Scale V1026/04) Climate Mode Chrono AT Tropical Polar GAT Start End Start End * Also Known As Bolide Impacts Large Igneous Key References Provinces Temp C° Mid Cretaceous - W5-W7 3.0 Paleogene W5 3.4 Hothouse W5.1 3.4 W5-W7 W5.2 3.4 39.4 - 128 Ma W5.3 1.8 AT = 88.6 W5.4 4.8 W5.5 5.1 W5.6 5.1 W5.7 4.9 W5.8 4.1 W5.9 5.0 W5.10 2.0 C5 1.3 C5.1 1.6 C5.2 3.8 C5.3 -4.6 C° C° C° Name 28.7 7.1 22.6 Mid-Cretaceous- Paleogene Hothouse 29.4 8.1 23.4 Paleocene - Eocene Hothouse 29.4 6.9 23.1 Middle Eocene Thermal Maximum 29.4 6.9 23.1 Late Lutetian Thermal Maximum 27.8 4.7 21.4 Middle-Late Eocene Cooling 30.8 9.6 24.8 Azolla Cool Event 31.1 10.1 25.2 Early-Middle Eocene Warm Interval 31.1 10.1 25.1 Early Eocene Thermal Maximum 30.9 10.3 25.0 Early Eocene Thermal Event 3 30.1 9.6 24.3 Early Eocene Thermal Event 2 31.0 10.7 25.2 Paleocene-Eocene Thermal Maximum 28.0 7.9 22.3 Thanetian Thermal Event 27.3 7.1 21.6 Late Cretaceous- Early Paleocene Cool Interval 27.6 7.6 21.9 Late Danian Cooling 29.8 9.9 24.1 KT Impact Thermal Excursion 21.4 1.4 15.7 KT Impact Winter 30.3 8.8 24.2 Mid-Cretaceous Hothouse Abbr. Age MKPH 128 PEH 59 METM 41 LLTM 41.5 MLEC 48.5 ACE 48. EMEWI 56 EETM 54 EETE3 53 EETE2 54 PETM 55.6 TTE 59 LKEPCI 82 Age 39.4 39.4 40.5 41.5 34 48.5 46 49 53 54 55.6 59 59 Stage early Barremian earliest Thanetian earliest Bartonian latest Lutetian late Ypresian late Ypresian earliest Ypresian middle Ypresian middle Ypresian early Ypresian earliest Ypresian earliest Thanetian early Campanian Stage middle Bartonian middle Bartonian early Bartonian latest Lutetian latest Priabonian late Ypresian early Lutetian late Ypresian middle Ypresian early Ypresian earliest Ypresian earliest Thanetian earliest Thanetian Montagnais 51 Ma EECO ETM3, X-event ETM2, ELMO Paleocene/Eocene_OAE, MDE ELPE, PCIM peak, MPBE Manson 74 Ma LDC 63 62 late Danian KTITX 66 65 latest Maastrichtian KTIW 66 66 latest Maastrichtian MKH 111 82 early Albian latest Danian earliest Danian latest Maastrichtian early Campanian LDE, top C27N DAN-C2 58,84,85,89,101 NAIP 62-54 Ma 84,85,86,87,89,101,113 84,85,86,87,89,101,113 113, 84,85,86,89,101 89, 84,85,86,89,101,113 84,85,86,87,89,101,113 84,85,86,89,101,113 84,85,86,89,101,113 84,85,86,87,89,101, 102-105,113 84,85,89,101,113 58,59,69,71,75,83-85, 91,92,101 84,85,113 83,84,85,89,113,115 KT Extinction. K-Pg Extinction Chicxulub 66 Ma Deccan 66-65Ma 83,96,97,98,99,100,113 58,68,69,71,93,94,95 W6.1 6.6 32.6 11.9 26.7 Coniacian Thermal Maximum W6.2 8.2 34.2 13.0 28.2 Cenomanian- Turonian Thermal Maximum W6.3 3.6 29.6 7.5 23.3 Breistroffer Thermal Maximum W6.4 2.1 28.1 6.0 21.9 Event 6 Thermal Event W6.5 1.5 27.5 5.4 21.3 Petite Verol Thermal Event W6.6 1.2 27.2 5.0 20.9 Amadeus Thermal Maximum W6.7 1.2 27.2 4.9 20.9 l'Arboudeyesse Thermal Event W6.9 1.4 27.4 5.1 21.1 Leenhardt Thermal Event W6.10 1.3 27.3 4.8 20.9 Paquier/Urbino Thermal Maximum OAE3 87 87 latest Coniacian OAE2 94 93 latest Cenomanian OAEld 101 100 latest Albian EV6 102.5 102.5 late Albian PVTE 105 105 late Albian OAElc 106 106 late Albian ATE 109 109 middle Albian LTE 110 110 middle Albian OAElb 111 111 early middle Albian latest Coniacian earlyiest Turonian OAE3 OAE2 (Bonarelli) latest Albian OAEld late Albian late Albian late Albian middle Albian middle Albian early middle Albian OAElb OAElc, Jassines Madagascar 93- 58,68,69,71 91 Ma Naturaliste 100- 84, 99 Ma (continued on next page) C.R. Scotese et al. Earth-Science Reviews 215 (2021) 103503 1 s f: 6 =3 6 =3 H h s < 2 O n P :> s ™ ■§ § s e " ! > h p p a g 5 —: o; s o; 6 o; -b iH -b Z w h X h LO CN CO LO d CN Ctf CN CN CO CN CO CN CN CN CN CX) o CO ,_ o LO CN CO CO CN CO CO CN d CO a* LO CO CN CO CX) CN LO CN in CN d CO d CO d CO CO CN CN CO LO CX) CO CO o o CO CN d C6 C6.1 C6.2 W7 W7.1 W7.2 W7.3 W7.4 W7.5 Though the coldest portion of the Late Paleozoic Icehouse was over, the Permian remained a "cool" world, with a moderate pole-to-equator temperature gradient (.75°C per 1 degree of latitude). An exception was the Emeishan Thermal Excursion (W14), which occurred 260 million years ago during the late Capitanian. As the name implies, this brief thermal excursion was the result of the eruption of the Emeishan flood basalts in southwestern China (Ernst, 2014; Rampino and Shen, 2019). The Emeishan eruption was precursor to the much more massive West Siberian eruptions that would end the Late Paleozoic Icehouse, terminate the Paleozoic Era, and cause the greatest extinction event of all time (Erwin, 1993, 1995, 2006). 5.6. Triassic Hothouse (W10 - W13, 253 - 199 Ma), Table 5, Fig. 20 At the beginning of the Mesozoic Era, the Earth was an extreme hothouse world with average tropical temperatures approaching 40° C. It now widely accepted that the extreme global warming the ended the Late Paleozoic Icehouse was caused by the voluminous eruption of the West Siberian Large Igneous Province (LIP)(Ernst, 2014). During a brief interval (— 1 million years) at the Permo-Triassic boundary (252.1 Ma; Kamo et al., 2003), more than 15 million km3 of basaltic lava flowed from rifts beneath the West Siberian basin (Reichow et al., 2005, 2009; Saunders et al., 2005) and the Putorana and Tunguska plateaus. This outpouring buried more than 50% of Siberia (7 million km2) under a mantle of basaltic lava 1-4 km deep (maximum 6.5 km). The primoridial, mantle-derived CO2 released by these eruptions was supplemented by additional CO2 derived from the combustion of thick, buried coal deposits (late Carboniferous - early Permian) that lay along the subterranean path of the erupting volcanics. The Permo-Triassic Mass Extinction is marked by a <313C spike (Gruszczynski et al., 1989; Grossman, 1994; Scholle, 1995). The combined primordial and coal-derived CO2 (—5000 gigatons) more than doubled atmospheric levels of CO2. As a result, tropical temperatures surged from ~25°C prior to the eruption to nearly 40°C. This "Super Hothouse" global warming parched life on land and led to the formation of an anoxic "Strangelove" ocean (Hsu et al., 1985; Kump, 1991; Hotinski et al., 2001; Zhang et al., 2001; Grice et al., 2005; Heydari et al., 2008). As the atmosphere warmed rapidly, the greatest effect was felt near the poles. The cold polar air masses warmed and the cold surface waters at high latitudes, which were the oceanographic engine that had previously filled the Late Permian oceans with cold oxygenated bottom water, began to warm and the engine stalled. As the oceans warmed, they became stratified and began to stagnate. Organic carbon accumulated in the deep ocean, using up any available oxygen. Anoxia set in (Song et al., 2014). The first ocean to become anoxic from top to bottom was the landlocked PaleoTethys Ocean (Wignall and Hallam, 1992; §engor and Atayman, 2009). The NeoTethys soon became anoxic; the deeper more isolated portions of Panthalassa (between the Panthalassic mid-ocean ridge and western North America, -were the last to become toxic. The poisonous waters from PaleoTethys, NeoTethys, and Panthalassa spilled onto the shelves and the Late Permian benthic, neritic, and planktonic fauna began to die out. The oceans belched great volumes of CH4 into the atmosphere and the level of atmospheric oxygen fell. As temperatures rose even higher, the food chain collapsed and many complex marine ecosystems were wiped out forever. On land the conditions were just as extreme. The equatorial regions baked. Average summer temperatures in the subtropical regions of Pangea exceeded 40 degrees C. Near the poles, temperatures were warm and seasonal changes were nearly eliminated. The movement of warm air masses and ocean currents into the polar regions made any attempt at winter cooling impossible. The interior of Pangea, 1000's of kilometers distant from any source of water, was an intolerably hot, abiotic desert. A few isolated habitats may have been refugia from the killing heat. Mountains that poked up into the westerly trade winds still received abundant rainfall during the winter months. The young peaks of the 28 Table 7 Paleotemperature Timescale: Late Cenozoic Icehouse, Chronotemps Wl ■ C4, (ICS Time Scale V1026/04) Climate Mode Chrono AT Tropical Polar GAT Start End End ^ Also Known As Bolide Impacts Large Igneous Provinces Key References Temp C° C° C° C° Name Abbr. Age Age Stage Stage Post-Anthro Wl 5.0 31.0 -3.6 19.5 Post-Anthropogenic PAW 2300 CE 10 Ka PAW Anthropocene Anthropocene Global Warming Crisis 108,90 pocene Warming Warming C1-C4 0.5 26.5 -8.6 17.2 Late Cenozoic Icehouse LCI 39.4 0 middle Bartonian Modern 84,85,8 6,89,109,110,113 Late Cenozoic CI 0.0 26.0 -20.2 14.2 Pleistocene Ice Age PIA 2.7 Ma 12 Ka latest Pliocene late Pleistocene 84,86,8 9,107,110,113 Icehouse Cl.l -1.5 24.5 -21.7 13 Younger Dryas YD 12.5 Ka 12 Ka latest Pleistocene earliest Holocene 84,89,107,110,113 C1-C4 C1.2 -3.5 22.5 -23.7 11.0 Last Glacial Maximum LGM 21 Ka 20 Ka late Pleistocene late Pleistocene 84,86,8 9,107,110,113 0 - 39.4 Ma C1.3 -2.5 23.5 -22.7 12.0 Pleistocene Glacial/ Interglacial Interval PGI 781 Ka 11.7 Ka middle Pleistocene earliest Holocene 84,86,8 9,101,107,110,113 AT = 39.4 C1.4 -3.5 22.5 -23.7 11.0 N. Hemisphere Ice Rafted Debris Influx NHIRD 2.7 Ma 0 middle Pliocene Modern 86,87,8 8,89,110 W2 0.6 26.6 -18.4 15.0 Pliocene Warm PWI 5.3 2.7 earliest Zanclean latest Piacenzian Kara Kul <5 Ma 84,85,8 6,87,89,101, Interval 107,110,113 W2.1 0.3 26.3 -19.2 14.7 Pliocene Thermal Maximum PTM 3.3 3 early Piacenzian middle Piacenzian M2 84,85,8 6,87,89,101,110,113 C2 0.9 26.9 -15.1 16.1 Late Miocene Cool Interval LMCI 11 5.3 early Tortonian latest Messinian MI-5, 6 84,86,8 9,101,109,113 C2.1 0.2 26.2 -18.0 14.9 Messinian Glacations MG 6.3 5.5 middle Messinian late Messinian 109, W3 1.3 27.3 -10.4 17.4 Middle Miocene MMWI 18 11 middle Burdigalian early Tortonian Mi-lb, 2, 3 Columbia River 84,86,87,89,101,109,113 Warm Interval 16-15 Ma W3.1 0.9 26.9 -13.7 16.4 Tortonian Thermal Maximum TTM 11 11 early Tortonian early Tortonian 113, W3.2 2.3 28.3 -9.3 18.5 Middle Miocene MMTM 16 14 latest Burdigalian latest Langhian MMCO, Monterey Event 84,86,87,89,101,109,113 Thermal Maximum C3 0.8 26.8 -7.4 17.7 Early Miocene Cool EMCI 23 18 early Aquitainian middle Burdigalian Mi-la 84,89,101,109,113 Interval C3.1 -0.1 25.9 -7.3 17.1 Early Miocene Cool Mi-1 23 23 earliest Miocene earliest Miocene 84,89,101,109,113 Event W4 1.1 27.1 -4.7 18.6 Late Oligocene Warm LOWI 27 23 early Chattian latest Chattian 84,85,8 9,101,113 Interval W4.1 1.6 27.6 -4.7 19.0 Late Oligocene Warm LOWE 25 25 early Chattian middle Chattian Astergerinoides 84,85,8 9,89,101,113 Event guerichi influx C4 -0.3 25.7 -0.8 18.4 late Eocene -Oligocene Cooling LEOC 39.4 27 middle Bartonian early Chattian Popigai 36 Ma Mistastin 37 Ma 84,85,8 6,87,89,101,113 C4.1 -1.6 24.4 -5.1 16.5 Late Oligocene Cool Events LOCE 28 25 early Chattian middle Chattian Svalbardella influx #3, Oi-2b,2c 84,85,8 6,89,101,113 C4.2 -1.7 24.3 -2.9 16.9 Early Oligocene Cool EOCE 34 28 early Rupelian late Rupelian Svalbardella influx Afar 31-29 Ma 84,85,8 6,89,101,113 Events #2, Oi-la,2,2*,2a C4.3 -1.6 24.4 -1.8 17.3 Eocene/Oligocene Rapid Cooling EORC 34 33 latest Priabonian earliest Rupelian EOT, Oi-1 84,85,8 6,87,89,101,106,113 C4.4 1.3 27.3 4.0 20.8 Late Eocene Cool Event LECE 39.4 39.4 middle Bartonian middle Bartonian Svalbardella influx #1 84,85,8 9,101,113 C.R. Scotese et al. Cape Mountains in South Africa, where our burrowing mammal-like reptilian ancestors clung to life, was such a refugia. In the oceans, species that could tolerate the warmer sea surface temperatures or that could retreat to the cooler, deeper environments on shelf-edge and slope, also survived. Most importantly, some species that were able to reproduce rapidly and (e.g. microgastropods) were able to take advantage of the favourable conditions that sporadically appeared. The Permo-Triassic Mass Extinction dwarfs all other mass extinction events. During the Permo-Triassic Mass Extinction, 57% of all marine families went extinct (Sepkoski Jr., 1989), and an estimated 96% of all marine species were extinguished (Raup, 1979), though time percentage has been revised downwards to 81% by Stanley (2016). At the individual level, this means that 90 - 99% of all living things were wiped out. The Permo-Triassic Mass Extinction killed more species in low latitudes and led to a reduced latitudinal diversity gradient in the Early Triassic (Song et al., 2020). Though the extinction event appears to be geologically "instantaneous" for most taxonomic groups (especially brachiopods, bryozoans, crinoids, tabulate and rugose corals), arguments have been made that other taxa were already in decline (e.g. trilobites, graptolites, conodonts). For these "perched" fauna, the massive warming event was the coup de grace. Details of the rapid rise and fall of temperatures during the Early Triassic (252-247 Ma) have been described by Sun et al. (2012). The Early Triassic Extreme Hothouse (W13) was not a single spike, but rather a series of ups and downs. The Permo-Triassic Thermal Maximum (W13.4, 253-251 Ma) was followed by a 4°C fall in temperature (W13.3, Dienerian Cooling, 251-249 Ma), a 6°C rise in temperature (W13.2, Latest Smithian Thermal Maximum, 249-248 Ma), and finally a 6°C fall in temperature by the end of the Early Triassic (W13.1, Latest Olenekian Cooling, 248-247 Ma). The severity of the extinction event is revealed by subsequent "gaps" in the fossil record. Plants and animals were not able to quickly reestablish the complex ecosystems that prevailed before the extinction event. There are no significant coal deposits (i.e. no complex rainforest ecosystems) until 20 million years after the mass extinction event ("coal gap"; Veevers et al., 1994; Looy et al., 1999). The Permo-Triassic Mass Extinction Event wiped out diverse tropical and temperate rainforest flora. It is interesting to note that the forests and scrublands were replenished by a new stock of plants (pteridosperms and conifers) that evolved from xerophytic ancestors (Zechstein flora) that were "pre-adapted" to the hot, dry conditions that prevailed in the Early Triassic (Looy et al., 1999; Hochuli et al., 2010). In the early Triassic oceans, there were similar reef and chert "gaps". In the shallow marine seas and on far-flung oceanic atolls, there were no coral reefs (Stanley, 2003). Reef ecosystems did not become reestablished until the Middle Triassic (Ladinian), approximately 14 million years after the Permo-Triassic mass extinction wiped out tabulate and rugose corals (Fliigel, 2002). These new reefs were built by a new type of coral animal (scleractinians) whose Permian ancestor was a soft-bodied "anemone-like" anthozoan that did not build limestone reefs. The radiolarian plankton were similarly wiped out, resulting in the absence of bedded-chert deposits in the deep oceans. Global temperatures fell during the Middle Triassic, reaching moderate hothouse temperatures (24° C) by the end of the Anisian (Trotter et al., 2015). These less extreme temperatures facilitated the re-establishment of tropical and temperate rainforests, coral reefs, and radiolarian plankton. The cooling trend continued through the remainder of the Middle Triassic (Ladinian) and into the early Late Triassic (CI 2, Ladinian-Carnian Cooling, 242-233 Ma; Trotter et al., 2015). Dense forests of the tree-like lycopod, Pleuromania, grew near the poles (Ziegler et al., 1994). The pattern of growth rings in Pleuromania indicates that these plants grew at rates 10-100 times faster than modern trees (Taylor et al., 2000). This suggests that though the climate was highly seasonal, light rather than temperature was the limiting growth factor. During the middle Carnian, temperatures warmed (Rigo et al., 2007; Earth-Science Reviews 215 (2021) 103503 Trotter et al., 2015; Sun et al., 2016), rainfall increased and the once extensive carbonate platforms along the margins of southwestern Tethys were flooded by clastic deposits carried onto the continental shelf by hyper-active river systems (Dal Corso et al., 2018a, 2018b). This dramatic change in climate is known as the "Carnian Pluvial Event" (Wl 2.1, 233-230 Ma; Ruffell et al., 2015; Ogg, 2015). It has been proposed that these climatic events were triggered by the formation of the Wrangellian oceanic plateau and associated volcanic islands (Greene et al., 2008; Greene et al., 2009a, 2009b, 2010). The C02 released by the eruption of these basalts warmed the atmosphere which in turn intensified the Pangean monsoonal weather system (Parrish, 1993) and brought more moisture to the interior of the continent. This rainfall fed new river systems that transported vast amounts of sand, silt, and mud to oceans. The arrival of huge amounts of sediments shut-down carbonate factories, promoted deep water anoxia, and increased environment stress. During the Carnian Pluvial Event (CPE), important ammonite groups became extinct (Balini et al., 2010), conodonts went through a major crisis (Rigo et al., 2007; Martinez-Perez et al., 2014), and other groups such as bryozoans and crinoids show a sharp decline (Simms and Ruffell, 1989). On the continents, new floras evolved that were adapted to wetter conditions (Roghi et al., 2010; Preto et al., 2010; Mueller et al., 2016a, 2016b ). After this event, modern conifers and bennetitaleans evolved (Willis and McElwain, 2014; Kustatscher et al., 2018) and dinosaurs emerged as the dominant and most diverse terrestrial fauna (Benton, 1993; Bernardi et al., 2018). The Mid-Carnian Warm Interval (W12, 234-227 Ma), which included the Carnian Pluvial Interval, was followed by the Early Norian Cool Interval (Cll, 227-217 Ma; Trotter et al., 2015). Global temperatures spiked again during the late Norian, (Wll, 214-209 Ma; Trotter et al., 2015), before cooling off slightly at the end of the Triassic (C10, Rhae-tian Cool Interval, 209-201; Trotter et al., 2015). The dip in temperatures during the late Norian can be attributed to a brief cooling excursion triggered by the Manicouagan impact (Wll.l 214 Ma; Spray, 2020). At the end of the Triassic and into the earliest Jurassic (201.3 Ma), flood basalts erupted across a large portion of North America, South America, Africa, and southern Europe (estimated area ~ 10 million km2; Marzoli et al., 1999; McHone and Puffer, 2003; Hames et al., 2003). During a brief episode of —600,000 years), the Central Atlantic Magmatic Province (CAMP) released approximately 80,000 gigatons of C02 into the atmosphere (Whiteside et al., 2010; Torsvik et al., 2020). This exhalation of greenhouse gases increased global temperatures 3°-6° C (McElwain et al., 1999; Beerling and Berner, 2002; Korte et al., 2009; Dera et al., 2011) giving rise to the brief End Triassic Thermal Event (W10, 201-199Ma). This rapid episode of global warming triggered the now familiar cascade of catastrophic environmental changes: increased terrestrial weathering and erosion Ahlberg et al. (2003), changes in ocean chemistry (i.e., mercury poisoning; Sanei et al., 2012) oceanic acidification (Hautmann, 2004; Hautmann et al., 2008; Clark-son et al., 2015), photic zone euxinia, falling marine productivity, and possible deep ocean anoxia (Isozaki, 1997). The End Triassic Extinction (ETE) is ranked third in terms of taxonomic severity, with an estimated 73% extinction of marine genera (McGhee et al., 2013). The once widespread seed fern Dicrodium disappeared from the fossil record (van de Schootbrugge et al., 2009). There was a 90% species turnover in the terrestrial megaflora (McElwain and Punyasena, 2007). A "fern-spike" suggests widespread changes in vegetation across Europe and North Atlantic (van de Schootbrugge et al., 2007; Bonis et al., 2010). A number of non-marine clades of vertebrates went extinct, though marine reptiles and fish actually flourished (McCune and Schaeffer, 1986; Benton et al., 2013; Friedman and Sallan, 2012). The End Triassic Extinction was the most severe extinction crisis ever experienced by scleractinian corals (Fliigel, 1994; Fliigel and Kiessling, 2002; Kiessling et al., 2002) and was catastrophic for reef communities. Bivalves and ammonoids, though they were hit hard by the End Triassic Extinction, had been in decline throughout the latest 30 C.R. Scotese et al. Triassic (Rhaetian; Hallam, 2002). For excellent summaries of the events surrounding the End Triassic Extinction, see Preto et al. (2010), Whiteside and Grice (2016), Bond and Grasby (2017), and Torsvik et al. (2020). 5.7. Jurassic - Early Cretaceous Cool Interval (C7 - C9, 199-128 Ma), Table 5, Fig. 20 Not all climate modes can be neatly pigeon-holed as either a steaming hothouse world or a frigid icehouse world. The 71 million year interval from the beginning of the Jurassic (Hettangian, 199 Ma) through the Early Cretaceous (early Barremian, 128 Ma) is made up of more than 16 distinct warm and cool events (Table 5). Cooler events make up ~67% of this time interval and justify the designation "Jurassic - Early Cretaceous Cool Interval", though the appellation, "Jurassic -Early Cretaceous Mixed Interval" would work as well. The Jurassic - Early Cretaceous Cool Interval is made up of three cool, but not cold, intervals separated by two warm intervals. Both of these warmer intervals, the Toarcian Warm Interval (W9) and the Kimmeridgian Warm Interval were hothouse worlds with average global temperatures of 22°C. The average global temperature during the Jurassic - Early Cretaceous Cool Interval was a relatively mild 17.5°C, which is close to the average temperature for the entire Phanerozoic (18°C). The Early Jurassic Cool Interval (C9, 199-183 Ma) and the Middle Jurassic Cool Interval (C8, 174 - 164 MA) are both characterized by moderate pole-to-equator temperature gradients with annual average polar temperatures several degrees above freezing. Dropstones and glendonites found in northeastern Siberia (Boucot et al., 2013) are evidence of cold winters in the northern hemisphere during much of the early and middle Jurassic (late Pliensbachian - early Bathonian). The winter snow and ice, however, did not persist. The warm summers melted any seasonal snow and ice, preventing the growth of permanent ice caps in either hemisphere (Sellwood and Valdes, 2006). Distinctive Jurassic plant groups occupied different latitudinal climatic zones. The greatest diversity of plants occurred at mid-latitudes (40° N&S) where forest were composed a mixture of ferns, cycads, sphenopsids (like horsetails), pteridosperms (seed ferns), and conifers (Rees et al., 2000). The equatorial regions, which were sparsely populated, were hot and dry and dominated by xeromorphic (arid-adapted) forms, small-leafed conifers, and cycads. Large-leafed conifers and deciduous gingkos grew in the polar regions (Rees et al., 2000). These plant groups suggest that the tempemperatures were seasonally cold, but not frigid. Estimates of tropical seawater temperatures from dlsO measurements are 4° to 6 °C cooler than modern tropical temperatures (Dera et al., 2011) indicating a moderate pole-to-equator temperature gradient. The Early Jurassic Cool Interval (C9) was followed by the Toarcian Warm Interval (W9), during which temperatures spiked 4°- 8°C (Dera et al., 2011). This warm pulse has been attributed to increased atmospheric CO2 arising from the eruption of the Karroo-Ferrar LIP (Pan-khurst et al., 1998, 2000; Elliot et al., 1999; Courtillot and Renne, 2003; Jourdan et al., 2005; Ernst, 2014; Ernst and Youbi, 2017). The global nature of this warming event is recorded by widespread oceanic anoxia (Toarcian OAE; Jenkyns et al., 2002; Jenkyns, 2010; van de Schoot-brugge et al., 2013). Much has been made of the extremely arid conditions and unbearable heat of equatorial and subtropical Pangea (30N to 30 S). By modern standards, the climate at low latitudes during the Toarcian Warm Interval was unbearably hot. Average equatorial sea surface temperatures which were the warmest of the Jurassic exceeded 30° C and temperatures in the interior of Pangea often exceeded 40° C (Crowley, 1994). The intense summer heating of the large land areas north and south of the Equator strongly deflected the Intertropical Convergence Zone (ITCZ) during the summer months. This modification of the basic Hadley Cell circulation pattern has been called "mega-monsoonal" atmospheric Earth-Science Reviews 215 (2021) 103503 circulation (Parrish, 1993). Megamonsoons resulted in drier conditions along the Equator and the formation of a broader Subtropical Arid Belt. Though the equatorial and subtropical regions of Pangea were arid during the Early Jurassic, most of the world was covered by lush, habitable forest vegetation (Rees et al., 2000). Abundant Early Jurassic coal deposits are found in both the warm and cool temperate belts (up to 60 N& S) (Boucot et al., 2013). Bauxite deposits, indicative of warm and wet conditions, are found in Europe and west central Asia (Kazakhstan) (30N - 45N). The early Jurassic Toarcian hothouse (W9) was followed by a relatively cool period during the middle Jurassic (Jenkyns et al., 2002; Dera et al., 2011). In a rare turnabout, during the early Jurassic, the South Polar region appears to have been warmer than the North Polar region. Dinosaurs and small flying reptiles, which would have had difficulty surviving freezing winters, inhabited the central Antarctica Trans-antarctic Ranges (Hammer and Hickerson, 1996); whereas dropstones and glendonites, indicative of freezing winter conditions, have been reported from Arctic Siberia (Boucot et al., 2013; 70 N - 80 N). Global temperatures warmed again during the late Oxfordian and Kimmeridgian (W8, Kimmeridgian Warm Interval, Podlaha et al., 1998; Jenkyns et al., 2002; Dera et al., 2011). The "mega-monsoonal" atmospheric circulation that characterized the earlier Mesozoic (Parrish, 1993) was still in place, but showed signs of weakening. As the newly formed intra-Pangean ocean basins (Central Atlantic Ocean and Western Indian Ocean) widened, they brought new sources of moisture to the interior of Pangea. The influx of moisture from these young ocean basins dampened the severe, seasonal swings in temperature and precipitation that had plagued the dessicated core of Pangea since the Late Permian. For other climate model results for the Late Jurassic see Moore et al. (1992), Valdes and Sellwood (1992), Valdes (1994), Sellwood and Valdes (2006). The Kimmeridgian Warm Interval (164-150 Ma) was followed by a long interval characterized by cool, but not frigid, temperatures (Tithonian-early Barremian Cool Interval, C7, 150 - 128 Ma). The Tithonian-early Barremian Cool Interval was punctuated by isolated warm events (Weissert (136 Ma) and Faraoni (131 Ma) thermal excursions). Global Climate during the earliest Cretaceous (Berriasian to Barremian) can be characterized as something "in-between" a hothouse and an icehouse (Frakes et al., 1992). The average global temperature was about 17° C. This was five degrees warmer than the Late Cenozoic Icehouse (CI, 12° C), but seven to nine degrees cooler than the Mid-Cretaceous - Paleogene Hothouse (W5-W6). Evidence for more temperate climatic conditions is based on the occurrence of dropstones, glendonites, and a few tillites (pebbly mudstones, Boucot et al., 2013) in polar latitudes that co-occur with evidence of temperate forests (coal, plant fossils) and dinosaurs. Dropstones of Early Cretaceous age (Ber-riasian/Barremian) are widespread in South Australia, Queensland, New South Wales, and the Northern Territory of Australia (Boucot et al., 2013). Glendonites occur in South Australia and New South Wales . In the northern hemisphere there are dropstones in Siberia and Svalbard, and glendonites in northern Siberia, Svalbard and the Arctic Islands (Grasby et al., 2017; Brassell, 2009; Frakes and Francis, 1988; Frakes and Francis, 1990,Francis and Frakes, 1993; Frakes et al., 1995; De Lurio and Frakes, 1999; Vickers et al., 2019). The best interpretation for this mixture of cool and warm climatic indicators is that it was cold enough in the winters for lakes and rivers to freeze over. Snow covered the ground and there were glaciers at higher elevations. In the summer months it was warm enough to support the growth of lush vegetation and an influx of dinosaurs migrating in from warmer regions. It is interesting to note that the Kimmeridgian Anoxic Event is the only potential oceanic anoxic event for the time interval spanning the middle Jurassic (Aalenian, 174 Ma) to the early Barremian (128 Ma). It does not appear to be as widespread as the Cretaceous OAE events. The lack of anoxic basins during the earliest Cretaceous seems quite unusual 31 C.R. Scotese et al. in light of the fact that there were many "restricted" marine basins that would have been ideal habitats for anoxia to develop. The lack of OAEs may have been due to the fact that the bottom waters during the middle and late Jurassic through to the earliest Cretaceous were relatively well-oxygenated. The occurrence of glendonites at high latitudes during much of the Early Cretaceous indicates that cool, oxygen-rich bottom waters were being generated at polar latitudes preventing the bottom waters in lower latitudes from becoming anoxic. 5.8. Mid Cretaceous - Paleogene Hothouse (W5 - W7, 39.4-128 Ma), Table 6, Fig. 20 If one imagines where the current phase of anthropogenic global warming is heading, one immediately thinks of the hothouse worlds of the Late Cretaceous and Eocene (Huber, 1998; Huber et al., 2000). During the Mid Cretaceous - Paleogene Hothouse global temperatures were indeed much warmer than the present-day (GAT = 28° C during the Late Cretaceous versus GAT = 15°C for the Modern). It remains to be seen whether we will succeed in warming the Earth to that degree, but at least we now know what a warmer world would look like. The Mid Cretaceous - Paleogene Hothouse is one of the best documented paleotemperature intervals. Table 6 lists some of the key references for this time interval and summarizes their primary conclusions regarding regional and global temperatures. The best single source for information about the Cretaceous portion of this hothouse interval is the O'Brien et al. (2017) summary of sea surface temperatures (SSTs) based on oxygen isotope and TEXs6 temperature estimates. The TEXs6 technique uses the lipid chemistry of the cell membrane of a common group of pelagic protokaryotes (Thaumarchaeota) to estimate temperatures (Schouten et al., 2002). The paleotemperatures are derived by measuring the ratio of key lipids (crenarchaeols). It has been noted that TEX86 temperature estimates tend to be —50% higher than dlsO temperature estimates (O'Brien et al., 2017; Fig. 8). Approximately 90% of the available TEXs6 paleotemperature estimates for the Cretaceous have been obtained from samples that are Aptian or younger in age. Moreover, there are very few dlsO temperature estimates for times older than the mid-Albian (O'Brien et al., 2017). Fortunately, as noted earlier, geological evidence (glendonites, dropstones, and rare tillites) from the Early Cretaceous helps to fill in these data gaps. Both oxygen isotope and TEXe6 measurements identify an early Barremian - middle Aptian warm interval (W7, 128 - 118 Ma, GAT = 22° C), which was followed by a cooler period during the late Aptian-early Albian (C6, 118 - 111 Ma, GAT = 19°C), which preceded the rapid ramping up to a thermal maximum during the latest Cenomanian-earliest Turonian (W6.2, 94 - 93 Ma, GAT = 28°C). According to O'Brien et al. (2017), temperatures cooled gradually during the remainder of the Late Cretaceous reaching a minimum of ~21°C in the late Maastrichtian, just prior to the KT impact event. The Mid Cretaceous - Paleogene Hothouse (W6) began in the latest Barremian - earliest Aptian (—128 Ma) with two thermal events, the Haupblatterton Thermal Event (W7.5; Mutterlose et al., 2009) and the oldest Cretaceous oceanic anoxic event, OAEla, the Selli/Goguel Thermal Maximum (W7.4; Erba et al., 2015; Herrle et al., 2015; O'Brien et al., 2017). Average global temperatures during the Mid Cretaceous -Paleogene Hothouse was a ~23° C. Surface waters in the Cool Temperate regions (SST = 21-23°C) were only slightly cooler than the superheated tropical seas (29° C; O'Brien et al., 2017). Oceanic bottom waters were also much warmer than the present-day (9° - 17° C; Valdes et al., 2020). The Selli/Goguel Thermal Maximum (OAEla) is the first of nearly a dozen potential thermal spikes that characterize this time interval (see Table 3). The nature and origin of these OAEs has been much debated (Schlanger and Jenkyns, 1976; Arthur and Sageman, 1994; Meyer and Kump, 2008). Previous notions that OAEs were simply the result of rapid rises in sea level (Arthur and Sageman, 1994) or due to the stagnation of the ocean basins caused by thermohaline density stratification, have Earth-Science Reviews 215 (2021) 103503 fallen out of favour. There are two current schools of thought. The first argues that the OAE's were "unusual", synchronous, global events. According to this argument, "catastrophic" tectonic events triggered "unusual" atmospheric, biologic, geochemical, and oceanographic conditions that promoted extensive deepwater anoxia that resulted in the formation of carbon-rich black shales (Total Organic Carbon often > 30%). Proponents of this school of thought argue that the following scenario may explain the widespread occurrence of the carbon-rich black shales associated with the early Aptian Selli/Goguel Thermal Maximum (OAEla): • The eruption of the mid Cretaceous superplume (Larson, 1991; Larson and Erba, 1999) radically changed atmospheric and oceanic chemistry. • Greenhouse gases from the erupting lavas, (i.e. CO2), warmed the Earth. • Increased warmth accelerated chemical weathering on land; consequently, a greater flux of nutrients was carried to the oceans. • Land-derived nutrients, together with a higher concentrations of biolimiting metals made available by increased hydrothermal activity associated with the extensive submarine eruptions (Duncan and Huard, 1997; Jones and Jenkyns, 2001) promoted greater marine productivity resulting in more carbon deposition. • The increased productivity depleted the available supply of oxygen in the water column, which led to basin-wide anoxic or dysoxic conditions. • Water-column anoxia, in turn, favoured the preservation of the carbon by inhibiting bacterial decay and carbon recycling. • The results were widespread carbon-rich black shales (Demaison and Moore, 1980) A second school of thought argues that the OAEs do not represent unusual or catastrophic, global events, but rather represent "business as usual". In other words, a certain constellation of biologic, geochemical, tectonic, atmospheric, and oceanographic conditions favour the development of local, basin-wide anoxia. The Cretaceous were unusual only in the sense that this constellation of favorable conditions was more likely to occur than one might have been expected. In essence the paleogeography of the Aptian/Albian (and Cenomanian/Turonian) was especially favourable for the formation of highly productive, anoxic basins that promoted "nutrient trapping" (Meyer and Kump, 2008). These two schools of thought epitomize the classical conflict between "catastrophic" versus "uniformitarian" explanations of Earth processes. As in the case of most false dichotomies, each hypothesis may hold part of the answer. Each hypothesis may explain different aspects of the Earth System processes that produce oceanic anoxic events. The "catastrophic" hypothesis may be the best explanation for the rare, but truly global, "mega" OAE events (i.e. Selli/Goguel Thermal Maximum (OAEla) and Cenomanian-Turonian Thermal Maximum (OAE2), whereas the "uniformitarian" hypothesis may be a better explanation for the more frequent, regional, and less intense OAE events ( OAElb, OAElc, OAEld, OAE3). The Aptian-Albian Cold Snap (C6, 118-111 Ma) separates the Selli/ Goguel Thermal Maximum (OAEla) from the remaining Late Cretaceous OAE's. For a brief interval in the late Aptian and Early Albian, the global climate cooled off sufficiently for winter snow and ice to return to the northern and southern polar regions (Puceat et al., 2003; Jenkyns et al., 2012; Erba et al., 2015; Herrle et al., 2015; O'Brien et al., 2017). Glendonites are reported from Ellesmere Island, Axel Heiberg Island, Svalbard, northern Greenland, and east-central Australia (Eromanga Sea) indicating that cool bottom waters once again had chilled the deep ocean basins (Frakes and Francis, 1988; Grasby et al., 2017; VIckers et al., 2019). The warmest Cretaceous temperatures occurred during the Cenomanian-Turonian Thermal Maximum (W6.2, 94 - 93 Ma). Second only to the Permo-Triassic Thermal Maximum (W13), the global average 32 C.R. Scotese et al. Earth-Science Reviews 215 (2021) 103503 temperature reached 28°C and the pole-to equator temperature gradient was flattened with a temperature differential of only ~20°C degrees between the polar region (13°C) and the tropics (34°C). Not even a hint of ice existed at the poles during the Cenomanian-Turonian Thermal Maximum (Ziegler et al., 1985). The presence of tropical plants and dinosaurs on Antarctica (Dettmann, 1989; Cantrill and Poole, 2012) and above the Arctic Circle indicates that temperatures rarely fell below freezing even during the winter months (Wolfe and Upchurch, 1987; Parrish and Spicer, 1988; Spicer et al., 2008). Recent descriptions of angiosperm leaf floras from Antarctica indicate that similar warm and wet conditions existed near the South Pole during the Late Cretaceous (Hayes et al., 2006). In general, during times of hothouse conditions, the equatorial and subtropical belts expand slightly poleward; the Polar and Cool Temperate belt are replaced by an expanded Warm Temperate belt that brings tropical conditions to latitudes above 50° north and south (Paratropical Belt of Boucot et al., 2013; the megathermal rainforests of Morley, 2011). Despite the overwhelming geological and paleontological evidence for warm polar regions during the Mid-Cretaceous Hothouse, early climate simulations tended to "run cold" and had a difficult time modeling these warmer polar temperatures (Barron and Washington, 1982). Various attempts have been made to modify the input parameters to the climate models to produce simulations more consistent with the geological data. Initial attempts to fix this problem used extremely elevated levels of greenhouse gases to warm the poles (15x modern CO2; Bice and Norris, 2002). However, there is no geological support for CO2 concentrations in the Cenomanian/Turonian much above 5x the modern value (van der Meer et al., 2014). The extreme high levels of CO2 needed to keep the polar regions ice-free would necessarily make terrestrial and shallow marine habitats at low latitudes uninhabitable (Jacobs et al., 2005). Another way to make the polar regions warmer is to modestly increase the concentration of greenhouse gases and also modify the land cover in polar regions to a darker, denser vegetation (Upchurch et al., 1999). The darker vegetation has a lower albedo and consequently more solar energy is absorbed at the surface. In this model, positive feedbacks between high-latitude forests, the atmosphere, and the ocean all contribute to significantly warmer temperatures at high latitudes during the Late Cretaceous (Upchurch et al., 1999). A third explanation invokes a Late Cretaceous "Super-Gulf Stream" that vigorously carried warmth from the Equator to the Poles (Brady et al., 1998). Though intuitively appealing, an analysis of the dynamics indicates that it is not possible to carry enough heat poleward using ocean currents alone. The atmosphere must also play an important role. In addition, much like today, the paleogeography of the Late Cretaceous presents a nearly landlocked polar region that would have been isolated from Gulf Stream-like ocean currents. One of the more promising approaches has been to change the high-altitude cloud parameterization that is used in climate models like the Community Climate System Model version 3 (CCSM3, Kiehl and Shields, 2013). The high albedo of low-altitude cumulus clouds reflects incoming sunlight back to space, which cools the Earth. Wispy, high-altitude clouds, on the other hand, reflect thermal energy back to the surface of the Earth resulting in net global warming (Kump and Pollard, 2008). Fewer "warm clouds" form in the modern world because anthropogenic atmospheric pollution reduces the amount of warm cloud condensation nuclei. When cloud parameters characteristic of pristine regions are introduced into the climate model, significant additional warming occurs, especially in polar regions (Upchurch et al., 2015). Combined with a modest elevation in the concentration of atmospheric CO2 (2x - 6x modern levels), the modelled temperature of polar regions remains above freezing throughout most of the year. The most radical hypothesis that has been proposed to explain the warm polar climates of the Late Cretaceous involves a fundamental rethinking of the way the atmosphere circulates. One of the basic features of the modern atmosphere is Hadley Cell circulation. In the Hadley Table 8 Model of the Budget of Atmospheric C02 that Predicts Future Global Warming (1) TACE(t) GtC = (P(t) x E(t) toe x FF%(t) x cl) +NNF(t) where, TACE(t) - Total Annual Carbon Emissions, (GtC) GtC - gigatons of carbon P(t) - Global population at time (t) in billions of people E(t) - Average Energy Use of each person at time (t), so-called "Energy Footprint", = 1.7 tons of oil in 2000 or 11.54 Barrels of oil, where 1 ton of oil = 6.67 barrels of oil toe - billion tons of oil equivalent FF%(t) - percentage of total global energy that comes from fossil fuels at time (t). cl - conversion factor (1)1 btoe = .85 GtC, where btoe = billion tons of oil equivalent NFF(t) - combined Non-Fossil Fuel Sources of C02 (i.e, cement, deforestation, agriculture) (2) N(t) = TACE(t) * %Rmv(t) where, N(t) - Amount of carbon in gigatons removed annually by the oceans and the land, %Rmv(t) - % of annual carbon emissions removed by Nature (historical value ~ 45%). (3) NACE(t) GtC = TACE(t) -N(t) -SQ(t) GtC where, NACE(t) - Net Annual Carbon Emissions, (GtC) SQ(t) - amount of C02 (GtC) removed from the atmosphere by carbon sequestration technology, assuming one sequestration plant can sequester 1 million tons of C02 per year. (4) C02(t) ppm = C02 (t-50)ppm + (NACE(t) x 50 x c2 ppm) where, C02(t )ppm - atmospheric concentration of C02 (ppm) at time, (t) CO2(t-50 )ppm - atmospheric concentration of C02 (ppm) 50 years earlier 50 - 50 years (time interval between control points) c2 - conversion factor (2), 1 GtC = .47 ppm C02 (5) Temperature(t) = 14.5°C + (4.5°C x (ln(C02(t) ppm/369 ppm) / ln(2))) where, Temperature(t) - global average temperature in the year, (t) 14.5°C - Global Average Temperature in the year, 2000. 4.5°C - the value for climate sensitivity (CS) used in this exercise ppm - parts per million 369 ppm - atmospheric concentration of C02 in the year, 2000 CS - Climate Sensitivity, ranges in value from 1.5 - 4.5 1 GtC02 = .27 GtC 1 GtC=3.67 GtC02 1 ton oil = 6.67 barrels of oil Cell, warm air rises at the Equator, moves poleward, cools and descends over the subtropical desert belt (-35° N&S). In Hay's model (Hay, 2008; Hay, 2016), this simple, well-organized convective flow is replaced by a chaotic system of super-cyclonic eddies, which are like mega-hurricanes. Hundreds of these mega-hurricanes would have annually transferred vast amounts of heat from the Equator to the Poles during the Late Cretaceous. Though an intriguing and out-of-the-box proposition, no climate model can currently simulate this complex alternative to Hadley Cell circulation. After reaching peak Cretaceous temperatures during the Cenomanian-Turonian Thermal Maximum, temperatures gradually fell during the remainder of the Cretaceous. Maximum sea surface temperatures did not drop below 30°C until late in the Santonian (84 Ma) or early in the Campanian (O'Brien et al., 2017). This gradual cooling may have been punctuated by several, ephemeral cooling events at ~85Ma, -76 Ma, and -71 Ma (Miller et al., 1999, 2004, 2005a, 2005b) as evidenced by dlsO temperature estimates from planktonic foraminifera. Also, an enigmatic dropstone deposit of Campanian - Maastrichtian age (75 - 70 Ma) has been reported from the region of the Anadyr River in Chukotka (Ahlberg et al., 2002). The modest, but steady, fall in temperatures during the Late Cretaceous was catastrophically interrupted by the arrival of the bolide that produced the 150 km diameter impact crater near the town of Chicxulub 33 C.R. Scotese et al. Earth-Science Reviews 215 (2021) 103503 Atmospheric C02 ppm 1300 Fig. 22. IPCC Estimates of Global Warming compared to the Results Predicted by the Carbon Budget Model (this study). The blue curve represents the projected increase in atmospheric C02 based on the continued burning of fossil fuels at the "modern rate" (1970 - 2020). The black curves are the C02 levels used in the IPCC models RCP 2.5, RCP 4.5, RCP 6.0, and RCP 8.5. The red curve is the projected C02 levels obtained in our model (between 750 ppm and 800 ppm at 2200 CE). Estimates from RCP 2.5 and RCP 4.5 are too low. The estimate of C02 used in RCP 8.5 is much too high. Estimates of C02 used in RCP 6.0 are about right. (Devil's Tail) in northern Yucatan (Alvarez et al., 1980; Schulte et al., 2010; Hildebrand et al., 1991). The Chicxulub impact is the largest known bolide impact of the Phanerozoic (Spray, 2020). The most likely scenario is that the impact event vaporized 3000 megatons of crustal material and injected this fine particulate matter high into the atmosphere. This material fell back to Earth forming a global "clay layer". The K/T boundary clay layer contains several unusual stratigraphic markers: 1) an iridium anomaly (Alvarez et al., 1980; Smit, 1999; Miller et al., 2010), 2) microtektites (Yancey and Guillem-ette, 2008), 3) shocked quartz (Bohor et al., 1987; Smit, 1999), and 4) soot (from forest fires) that connect it directly to the Chicxulub impact event. (It should noted that the authors prefer to use the term "K/T" 34 I-I 8 540 520 500 480 460 440 420 400 380 360 340 320 300 280 260 240 220 200 180 160 140 120 100 80 60 40 20 0 millions of years Fig. 23. Projection of Future Global Warming onto Phanerozoic Temperature Time Scale. The likely amount of Post-Anthropogenic Warming (PAW) (red line). The boxes indicate times in the geological past when global temperatures were within the range of predicted PAW. When the Global Average Temperature is below 18° C large polar icecaps can form. When the Global Average Temperature is above 18°C large polar icecaps are unlikely to form. 34 C.R. Scotese et al. Earth-Science Reviews 215 (2021) 103503 rather than the more precise "K-Pg" to describe events occurring at the Cretaceous -Paleogene boundary. The term "K/T" has precedence, is still widely used, is more familiar to most readers, and is therefore a clearer descriptor than the technically correct, but more obscure and less euphonius, "K-Pg". ) While suspended in the atmosphere, this delicate shroud of material blocked the sun and turned day into night - a night that lasted several months to a year. Without sunlight, plants on land and plankton in the oceans died. Small and large herbivores gradually starved. Without herbivores to prey on, predators then starved - all the while, snow continued to fall (probably for several decades). As a consequence of the collapse of the food chain, ~75% of all species were wiped out (Sepkoski Jr., 1996). The effect of this extinction event on global ecosystems was second only to the great Permo-Triassic Extinction (McGhee et al., 2013). The ensuing "Impact Winter" scenario plunged the Earth into a frigid deep-freeze comparable to the coldest glacial stages of any Phanerozoic ice age. The drastic cooling, however, was short-lived (Vellekoop et al., 2014, 2016) and was followed by an equally short-lived period of global warming triggered by the final, massive eruption of the Deccan LIP (Ernst, 2014; Keller et al., 2017). The first eruptions of the Deccan LIP predate the Chicxulub impact by 1-2 million years (Chenet et al., 2008; Keller, 2011, 2014). It has been proposed that an earlier impact event (Shiva impact) triggered the Deccan eruptions (Chatterjee et al., 2006), however this hypothesis has not received much support. It seems likely, however, that the Chicxulub impact did influence or enhance Deccan volcanism. It has been noted by several authors (E. Shoemaker, pers. comm.) that the impact site in Yucatan is nearly antipodal to the eruption site of the Deccan LIP in India. Though the antipodal paleolatitudes are identical (26° N vs 26° S), the antipodal paleolongitudes are offset by several thousand kilometers. Nevertheless, it seems plausible that Shockwaves from the impact passed through the earth and were recon-centrated beneath the Deccan hotspot stimulating more voluminous eruptions (Richards et al., 2015; Renne et al., 2015). In any event, the excess atmospheric CO2 from the Deccan eruptions caused a 4-8° spike in global temperatures (Petersen et al., 2016; Bond and Grasby, 2017) that ushered in the Paleogene Warm interval. Following the KT impact Winter, global temperatures warmed during the Paleocene, reaching a maximum during the early-middle Eocene (Paleocene - Eocene Hothouse, W5, 62 - 39.4 Ma). This period of global warmth was probably triggered by CO2 injected into the atmosphere by massive volcanic eruptions in the North Atlantic Igneous Province (NAIP; Ernst, 2014). It was in this global, tropical "Garden of Eden" that our mammalian ancestors diversified, crossed newly-erupted volcanic land bridges between northern Europe and Greenland, and expanded across the globe (Wallace, 2004). The classic record of deep ocean temperatures based on benthic foraminifera assembled by James Zachos ( Zachos et al., 2001, 2008; also Westerhold et al., 2020) provides a framework for describing the temperature fluctuations during the Paleocene-Eocene Hothouse and C.R. Scotese et al. the Late Cenozoic Icehouse (Koeberl and Montanari, 2009). These deep ocean temperatures have been converted to global average temperatures using the technique described in section 3.3. The three most prominent features of this detailed temperature record are: 1) the Paleocene- Eocene Thermal Maximum (W5.8, 55.6 Ma; Rea et al., 1990; Kennett and Stott, 1991; Wing et al., 2003; Mclnerney and Wing, 2011), 2) the Early Eocene Thermal Maximum (W5.5, -50 Ma), and the Middle Eocene Thermal Maximum (W5.1, -41 Ma). These are certainly "event" driven changes in climate. The one explanation for these spikes in temperature is the rapid release of massive amounts of methane hydrates (clathrates) from the deep sea (Zeebe et al., 2009). Methane is a powerful greenhouse gas and the release of gigatons of methane into the atmosphere would have produced the observed rapid rise in global temperature. Other hypotheses, summarized by Wing and Mclnerney (2011, p 494), suggest that the excess amount of greenhouse gases may have come from wildfires, volcanic intrusions into organic-rich sediments, drying of epicontinental seas, or the thawing of Antarctic permafrost. Most recently, Gutjahr et al. (2017) and Jones et al. (2019) have proposed that massive eruptive episodes associated with the North Atlantic Igneous Province (NAIP) provided the excess CO2 responsible for these thermal maxima. A prominent feature of the early Eocene Hothouse is the broad, dome-shaped rise and fall in temperature that defines the Middle Eocene Warm Interval (W5.4, 56-46 Ma). The shape of this curve indicates that systematic changes were taking place over millions of years (Huber and Caballero, 2011). These changes in global temperature were probably driven by gradually changing paleogeographic, plate tectonic, or pale-oceanographic conditions. The final notable feature of the Zachos Curve is the rapid fall in temperature at the end of the Eocene (C4.3, Eocene Oligocene Rapid Cooling, 33-34 Ma). It has been proposed by many authors that this cooling event was driven by the development of a through-going Circum-Antarctic Current and the subsequent isolation and refrigeration of Antarctica. 5.9. Late Cenozoic Icehouse (CI - C4, 1880 CE - 39.4 Ma), Table 7, Fig. 21 "And now there came both mist and snow And it grew wondrous cold And ice, mast-high, came floating by As green as emerald." Samuel Taylor Coleridge, Rime of the Ancient Mariner (1798) As illustrated in Fig. 21, the cooling trend that began after the Early Eocene Thermal Maximum (W5.5, 52-50 Ma) continued during the middle and late Eocene, though it was interrupted briefly by the Middle Eocene Thermal Maximum (W5.1, 41 Ma). What Earth System event initiated the Late Cenozoic Icehouse? The generally accepted explanation is that the collision of India, which took place at the height of the Early Eocene Thermal Maximum (—50 Ma; Molnar and Tapponnier, 1975; Rowley, 1996, 1998), triggered a cascade of events that lead to global cooling (Raymo and Ruddiman, 1992). The collision of India with south-central Asia resulted in the rapid uplift of the Himalaya Mountains. By —35 Ma (late Eocene - early Oligocene), the Himalayas and Tibetan plateau had achieved 90% of their modern height (Rowley and Currie, 2006), though the Tibetan plateau was less than half its present area. These young mountains were in the path of the Asian monsoon, which brought warm temperatures and abundant moisture. This lead to rapid mechanical and chemical weathering. The flux of calcium to the world's oceans increased and these calcium cations combined with carbonate ions to form limestone which drew-down the amount of CO2 in the atmosphere. The gradual decrease in atmospheric CO2 cooled the Earth. As global temperatures began to cool, the high altitude glaciers on the Gamburtsev mountains and Trans-Antarctic Ranges (> 1000 m) coalesced into ice sheets. At the Eocene-Oligocene boundary (C4.3, 34-33 Ma), temperatures plunged precipitously signalling the first major Earth-Science Reviews 215 (2021) 103503 accumulation of ice on Antarctica. Permanent sea ice also formed around the periphery of Antarctica generating increasing amounts of cold bottom water. Icebergs coursed through the southern oceans and the first record of ice-rafted debris appeared in the deep-sea record. The global cooling also forced plants to adapt to the harsher winter conditions (Wolfe, 1971, 1978, 1992, 1994). Paleoceanographic events in the Southern Hemisphere, in particular the formation of the Circum-Antarctic Current, played a key role in the chronology of these events (Kennett, 1995). The Drake Passage is the seaway that flows between the southern tip of South America (Patagonia) and the northern tip of the West Antarctic Peninsula (Palmer Peninsula). In the early Mesozoic, these two regions were part of a continuous Andean mountain range. The ligation between Patagonia and the Palmer Peninsula was tested when Gondwana began to rift apart in the late Jurassic. Despite being stretched and extended as the Weddell Sea opened during the Cretaceous, Patagonia and the Palmer Peninsula were not completely separate until the late Eocene (—40 Ma). The best estimate for the age of the opening of the Drake Passage is 45 Ma (shallow water connection, <1000m) to 35 Ma (deep water connection, >1000 m) (Livermore et al., 2005). An age of 41 Ma for the opening of the Drake Passage is based on the change in neo-dymium isotope ratios from sediments on the Agulhas Ridge that suggests an influx of shallow Pacific water (Scher and Martin, 2006). When the Southeast Indian Ocean between southern Australia and Antarctica (Wilkes Land) began to open in the Late Cretaceous, the eastern end of the rifted margin made a right-angle bend, hooking southward around Tasmania. This strike-slip boundary, the Tasman Fracture Zone, effectively closed off the eastern end of the Southeast Indian Ocean during the earliest phases of opening. When Australia began to move rapidly northward during the Late Eocene (40 Ma - 35 Ma), the overlapping bits of the Australian and Antarctic plates (South Tasman Rise and North Victoria Land, respectively) separated, allowing deep waters from the Indian Ocean and South Pacific Ocean to mix. The first deep water connection between the Indian Ocean and the South Pacific was through the Tasman Straits. Both the Drake Passage and the Tasman Gateway (Kennett et al., 1974; Exon et al., 2004; Kennett and Exon, 2004) were fully opened by the early Oligocene (34-30 Ma; Lawver and Gahagan, 2003). As a consequence, the Circum-Antarctic Current was able to isolate Antarctica from the world's oceans resulting in the "refrigeration" of Antarctica (C4.2, 34-28 Ma). The rapid growth of the Antarctic icecap during the Early Oligocene produced a major regression at the Rupelian/ Chattian boundary (28.1 Ma). During the Oligocene, the massive Antarctic ice cap grew and shrank in fits and starts. These chaotic transitions are recorded in a dozen cooling and warming events (C4.1 and C4.2). Global temperatures remained cool until the end of the Oligocene when there was a slight warming (W4, 27-23 Ma). Though parts of West Antarctica were still forested during the Oligocene, mountain glaciers grew in the highlands of the Palmer Peninsula, reaching the ocean by the latest Oligocene. In the Northern Hemisphere there were local glaciers in Svalbard and central Greenland, but no permanent icecap. It is also worth noting that by the Early Oligocene, the collision between the Arabian peninsula and Iran was nearly complete. Though all the ocean floor had been subducted, a shallow seaway, the proto-Persian Gulf, filled the foredeep of the Zagros mountains of Iran and Iraq. During brief highstands of sea level during the late Oligocene and early Miocene, this shallow seaway connected with the deeper waters of the eastern Mediterranean. The closure of this westernmost extension of Tethys in the earliest Oligocene eliminated the westward flowing "Subtropical Eocene NeoTethys" (STENT) current. Some authors have speculated that this may have contributed to global cooling during the early Oligocene (Toggweiler and Bjornsson (2000), Hotinski and Togg-weiler, 2003; Jovane et al., 2009). At the start of the Miocene, only Antarctica, straddling the South Pole, was covered by a permanent icecap. Cool conditions prevailed in the northern hemisphere as well, but there was no permanent ice. Global 36 C.R. Scotese et al. temperatures warmed in the middle Miocene (W3, 18-11 Ma), the Antarctic ice cap shrank and snow and ice disappeared from the northern hemisphere. By the end of the Miocene (~ 5 Ma), conditions had once again cooled (C2, 11 - 5.3Ma) and a permanent icecap had begun to form in the Arctic. The initiation of the Greenland ice sheet in the middle-late Miocene corresponds with the minimum sea level for the Miocene (-50 m; Miller et al., 2020). The growth of the Arctic icecap may have been triggered by two paleogeographic-paleoceanographic changes: 1) In the late Neogene (4-5 Ma), the Panama volcanic archipelago rose above sea level, creating the Panama land bridge. This land bridge connected North and South America, permitting the interchange of fauna and flora (Marshall et al., 1982). More importantly, this land bridge served as a blockade, isolating the equatorial Atlantic and Pacific Oceans. Warm, equatorial Atlantic waters were diverted northward (the Gulf Stream). The warm waters of the Gulf Stream warmed the Arctic regions, but more importantly, provided a new source of moisture which fed winter snows. Increased snowfall lead to the growth of glaciers and the Arctic ice sheets grew large. 2) In the eastern hemisphere, the northward movement of Australia during the early Miocene and the collision of Australia with Southeast Asia during the middle-late Miocene (—10-15 Ma) blocked equatorial circulation between the western Pacific Ocean and the Indian Ocean. As a consequence, the distance that ocean waters circulated along the Equator was shortened resulting in a net cooling of tropical surfaces waters. This reduction in the Earth's thermal budget, in turn, may have led to increased cooling at the poles. Global temperatures continued to plummet during the Pliocene. During the late Pliocene (—3.3 Ma), global temperatures warmed sufficiently to melt enough of the North and South polar icecaps so that sea level rose (—60 m), flooding the continental margins (W2.1, Pliocene Thermal Maximum). The Pleistocene Ice Age (CI) began in the latest Pliocene (2.7 Ma) and ended 11,700 years ago (Younger Cl.l Dryas). Driven by the tilt of the Earth (obliquity) and the shape of the Earth's orbit (eccentricity), the polar ice sheets waxed and waned causing global sea level to rise and fall (Miller et al., 2020). When ice sheets were at their maximum, sea level was —100 meters lower than today. During the interglacial part of the cycle, when much of the polar ice caps had melted away, sea level was —70 meters higher than today. Sea level has risen and fallen more than 50 times during the last 2 million years. This cycle of sea level change has been recorded in the changing ratio of the 180/160 preserved in benthic/planktonic forami-nifera (Lisiecki and Raymo, 2005; Raymo et al., 2018). This fluctuating record begins nearly 2.7 million years ago; for the first —2 million years, the cycle of icecap growth and retreat was modulated by the changing tilt of the Earth's axis (40,000 year obliquity cycle). Starting nearly one million years ago, the frequency of ice cap formation slowed from every 40,000 years to every 100,000 years, as the changing, eccentric shape of the Earth's orbit became the major forcing function. It should be noted that most of these marine isotopic sequences (MIS) match a similar cycle of changing CO2 concentration recorded in mile-long ice cores from Greenland and Antarctica (Barnola et al., 1987; Petit et al., 1999; Alley, 2000; EPICA Community Members, 2004; Jouzel et al., 2007). We are currently about halfway through a typical glacial/interglacial cycle. If humans did not inhabit the Earth, about 20,000 years from now, global temperatures would have once again begun to fall and ice sheets would have expanded into the oceans surrounding Antarctica and would have descended from the Arctic to begin a slow and steady march across the northern continents. However, this will not happen. The Earth has entered a "super-interglacial". The injection of CO2 into the atmosphere as a consequence of the burning of fossil fuels has warmed the Earth more than 1°C and will continue to warm the Earth for another 300 years (—2300 CE). In the next section, we discuss how long anthropogenic global warming will continue and how warm the Earth will become. Earth-Science Reviews 215 (2021) 103503 5.10. Post-Anthropogenic Warming (Wl) (2300 CE - 10,000 years in the future), Table 7, Fig. 21 It is well-established fact that burning fossil fuels releases CO2, a greenhouse gas, into the atmosphere, causing the Earth to warm faster than it would naturally (Archer, 2005; Archer, 2009,Archer et al., 2009; Kidder and Worsley, 2012; Steffen et al., 2019). How warm will it get? How quickly will it warm to those new levels? In order to answer these questions, we need to know future trends in 1) global population (United Nations, 2019), 2) the average "energy footprint" per global resident (British Petroleum, 2019), 3) the future mix of energy sources (oil, gas, coal, hydroelectric, nuclear, and renewables; British Petroleum, 2019; Rogner, 2012; Shell, 2018), 4) the rate at which the natural environment will absorb excess CO2 emissions (Tans, 2009), and 5) the warming effect of greenhouse gases (climate sensitivity; Royer, 2016; Farnsworth et al., 2019). Though these parameters are not known with any certainty, we can nevertheless make an informed estimate of future trends and produce a reasoned prediction of the potential amount of global warming during the next 300 years (Scotese, 2020). We have estimated the amount of future global warming using a straightforward carbon budget model that predicts the changing amount of atmospheric CO2. This CO2 budget model is described in Table 8. Using these carbon budget equations, a dynamic model was constructed that predicts the amount of global warming during the next 300 years. Assuming that the Global Average Temperature (GAT) in 2000 was roughly 14.5°C (58° F), and the concentration of atmospheric CO2 was 369 ppm. This model predicts that in 2200 the concentration of atmospheric CO2 will be —777 ppm, which is a little more than double the concentration of CO2 in the year 2000 and the global temperature will rise about 5° C from 14.5°C (58°F) to 19.5°C (67° F). This change in temperature is indicated by the red star in Fig. 21. As has been widely reported (Collins et al., 2013), most future warming takes place during the mid-to-late twenty-first century. Rapid global warming is nearly unavoidable because two key parameters -global population and energy use per capita - will rapidly rise during the next several decades. According to the model, in the year 2100, global climate will have warmed beyond the 2°C limit recommended by the Paris Accords (Fig. 22). In 2200, after reaching a maximum temperature of 19.5°C, the global climate will begin to cool. This is due in part to natural processes, but most of the modelled decrease in global temperature is the result of proposed human intervention (i.e. carbon sequestration; Shell, 2018). If 10,000 carbon sequestration plants are built in the next 300 years, then by 2300 the combined action of these carbon sequestration plants will have removed 1.8 trillion gigatons of CO2 from the atmosphere and the temperature will stabilize at 19.5°C (Shell, 2018). If no carbon sequestration plants are built, then the projected future temperature will be closer to 20.5 C°. Both of these predictions are in line with IPCC estimates. What will the world be like after the Warming? This is the question that is probably the most on people's minds. One can answer this question two ways. The first approach, taken by the Intergovernmental Council on Climate Change (IPCC) and Burke et al. (2018), has been to run hundreds of climate simulations that predict global climatic conditions using various global warming scenarios (Fig. 22). These results are very detailed and make explicit predictions (IPCC, 2007, 2018, 2019; Collins et al., 2013). Another way to understand our warmer future world is to compare the various possible warming outcomes to the climates of the past. The area shaded in red on Fig. 23 represents the temperatures that lie within the range of the global warming predicted by the model presented here (16.5°C - 19.5°C). You can see that this shaded region is much cooler than some of the hothouse climates of the past. It is very good news that future global warming will likely not reach these extreme hot house temperatures. Along the time-axis in Fig. 23, several boxes highlight time intervals in the deep past that had similar global climates. Of these, 37 C.R. Scotese et al. the Late Eocene - Miocene Icehouse is probably the best analog to the climate we might experience after this period of Post-Anthropogenic Warming (PAW). The geography of the early Oligocene world (30 million years ago) and the early Miocene world (15 Ma) resembled the modern world (Fig. 24). The continents were essentially in the same places and the oceans were about as wide as they are today. A large ice cap covered Antarctica, which was isolated from the other continents by the Circum-Antarctic Ocean. The northern hemisphere lacked any permanent ice cover, though snow covered the northernmost reaches of the continents during the winter. Unlike today's climate, the climate during the Oligocene and early Miocene was a little warmer near the Equator, and a diverse fauna thrived in tropical rain forests and in the oceans. Modern-sized desert belts separated the tropics from an expanded warm temperate belt that stretched across the northern hemisphere. The northern sub-polar regions, as well as Australia, were both wetter and warmer. These were worlds where land mammals thrived, diversified, and spread across the continents. Whales and enormous sharks ruled the seas. Plants also diversified. Grasslands covered the steppes and savannas and tropical and temperate forests provided a diverse set of habitats. All in all, the world was a reasonably nice place to live. Given enough time (5,000 -10,000 years; Archer, 2005; 2009a,b), we might expect the same equable conditions to prevail after the Warming. The other five climatic "matches", though they had similar ranges of global temperature, had very different geographies. The climatic zones on a global Pangea would have been very different than the modern climatic belts. In conclusion, we are leaving our Ice Age heritage behind. A new, warmer world awaits us. The problem we face is not so much where we are headed, but rather how we will get there. The time for decisive action is quickly slipping away. Our stern challenge is to adapt to the rapid period of global warming that will take place during the next 100 years. To do this we must immediately take action to reduce our CO2 emissions and take the steps needed to mitigate the undeniable damage that will be done to our environment, civilization, and society in the near future. 6. Summary and Conclusions It has been long recognized that the Earth's climate, in particular the average global temperature, has alternated between "icehouse" and "hothouse" states. More than 70 years ago, Umbgrove (1947), and later Fischer (1981, 1982, 1984) recognized that these climatic "modes" varied on short-term, medium-term, and long-term timescales. During the past 20 years, beginning with the comprehensive pioneering work of Veizer et al. (1999) and followed by the insightful syntheses of Royer et al. (2004), Zachos et al. (2001, 2008), Grossman (2012a, 2012b), Boucot et al. (2013), Veizer and Prokoph (2015), O'Brien et al. (2017), Henkes et al. (2018), Valdes et al. (2018), and most recently Song et al. (2019), we now stand at the threshold to a deeper, more complete understanding of both the tempo and mode of global temperature change during the Phanerozoic. The goal of this essay has been to synthesize all of the available evidence for temperature change during the last 540 million years into a single, coherent, well-constrained global temperature curve (Fig. 19-21, Table 3-7). Though the details may vary, there is good agreement between the temperature curve we present and earlier work (see Fig. 1). We have produced this model of Phanerozoic temperatures by deconstructing temperature change into three components: 1) long-term (>50 million years) temperature changes, 2) medium-term (10 - 20 million years), and 3) short-term temperature changes (a few million years or less). The Earth's long-term temperature change is controlled by multiple tectonic and environmental processes that drive the Earth's climate from icehouse to hothouse conditions, and vice versa. These changes can be described by mapping the changing Pole-to-Equator gradient revealed Earth-Science Reviews 215 (2021) 103503 by paleogeographic distribution of lithologic indicators of climate such as: tillites, dropstones, glendonites, high latitude mangroves, palms, and crocodiles; temperate coal, evaporites, calcretes, tropical coals, bauxites, and laterites, (see Fig.s 4 and 5; Boucot et al., 2013). Long-term temperature changes occur due to changes in the Earth's "climatic equilibrium" which is controlled by the level of greenhouse gases in the atmosphere (principally CO2), the geographic configuration of the continents and ocean basins (paleogeography and paleoceanography), the effectiveness of erosion and chemical weathering, and the reflectivity of the Earth's surface (albedo). Many of these factors are interconnected by a complex network of positive and negative feedback loops that can accelerate or decelerate changes in long-term global temperature (Hay, 2016; Ruddiman, 2001). Fig. 13 illustrates our best estimate of the long-term temperature of the Earth during the past 540 million years. Medium-term (10-20 million years) changes in the Earth's temperature are revealed by the record of isotopic temperatures measured in carbonate and phosphatic fossils. Recent global compilations of isotopic measurements of temperature (Grossman, 2012a&b, Veizer and Prokoph, 2015; Song et al., 2019; Grossman and Joachimski, 2020) have allowed us to generate a first-order model that describes the variation in tropical and polar temperatures for the last 540 million years (Fig. 17). There are still numerous uncertainties regarding the interpretation of isotopic temperatures. For example, the isotopic composition of seawater can vary due to excess evaporation or the local influx of freshwater, as well as the removal of leO from the oceans due to the growth of vast continental ice sheets. Also, the role that geographic and oceanographic variations play in 8180 variability has yet to be rigorously examined. Despite these caveats, we believe that the temperature record obtained from oxygen isotope data is robust and that a testable, first-order signal is generally recognizable (Fig. 9). The causes of medium-term fluctuations in temperature are many and complex. We have not yet unravelled these forcing functions, but they are likely to include tectonic and geologic events such as: continental collisions and subsequent mountain building and unroofing, periods of ophiolite obduction and subsequent chemical weathering, the opening of oceanic gateways, and some mega-evolutionary events such as the evolution of land plants. One of the predictions made by our analysis of medium-term temperature changes is that there may have been several "mini-ice ages", interspersed between the major glacial episodes (e.g. Hirnantian Ice Age, Late Paleozoic Icehouse, and Late Cenozoic Icehouse; Fig. 16). Short-term changes in the Earth's temperature are probably the most demonstrable because they are related to well-documented geologic events, namely the eruption of enormous volcanic provinces (LIPs) and the impact of large bolides. In this essay, we have documented that the eruption of ~20 large LIPs are strongly correlated with times of warmer global temperatures (Fig. 15Tables 3-7). Conversely, it appears that large bolide impacts are well-correlated with cooler periods in Earth history (Fig. 15, Tables 3-7), though a cause and effect relationship is less certain. It remains possible that a few of the most dramatic cooling events in Earth history may have been caused by as yet unrecognized large impact events (e.g. "Khione" event, Hirnantian Ice Age). In this essay, we combined the long-term, medium-term, and short-term changes in global temperature to produce the Phanerozoic Temperature Timescale (Fig. 19-21, Tables 3-7). The Phanerozoic Temperature Timescale is divided into 24 "chronotemps", or distinct warm and cool intervals. The youngest chronotemp (Wl) is the current period of anthropogenic warming. The oldest postdates the Cryogenian "Snowball Earth". Tables 3-7 list the names, timing, and temperatures associated with each chronotemp. It can certainly be argued that constructing a Phanerozoic Temperature Timescale is premature and an effort filled with error and un-proven assumptions. Though this is probably true, we believe that the effort is worthwhile, as Voltaire said, "II meglio e Vinimico del bene". We now have a structure that can be built upon, refined, corrected, 38 C.R. Scotese et al. Earth-Science Reviews 215 (2021) 103503 expanded, and compared to other temperature models that use data sets that we did not use (i.e., clumped isotopes, TEXe6)- Moreover, having a paleotemperature timescale is essential if we are to understand the tempo and mode of climate change during the past 720+ million years. By characterizing, in a quantitative way, the pattern of paleotemperature change through time, we may be able to gain important insights into the history of the Earth System and the fundamental causes of climate change. These insights may be helpful guide as we traverse an uncertain path into the future. Declaration of Competing Interest None Acknowledgements CRS would like to thank the past industrial sponsors of the PALEO-MAP Project. H.S. is supported by the National Natural Science Foundation of China (41821001). B.J.W.M. is supported by the UKNatural Environment Research Council (NE/S009663/1). This work is part of the PhanTASTIC project led by Scott Wing and Brian Huber at the Smithsonian Institution National Museum of Natural History. The workshop was supported by a gift from Roland and Debra Sauermann. CRS especially benefitted from discussions of past climate with Dana Royer, Linda Ivany, Paul Valdes, Dan Lunt, Colin Summerhayes, Jan Veizer, and Ethan Grossman. CRS would like to thank Joseph Westrich, Colin Summerhayes, David Archer, Daniel Horton, and Peter Lange for their valuable scientific and critical input on the future global warming essay. Seth Stein and his class at Northwestern University (EPS 362) provided valuable feedback on my essay describing future global warming and convinced me to rewrite my initial essay as a do-it-yourself global warming calculator. Class members Peter Puleo, Tim Coston, and Chutian Chen provided especially helpful reviews. 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