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Climatic and Environmental Challenges: Learning from the Horn of Africa

David Ambrosetti
Jean-Renaud Boisserie
Deresse Ayenachew
et al.

Links between Solid Earth, Climate Changes, and Biodiversity through Time: Insights from the Cenozoic

Pierre Sepulchre

Texte intégral

1The Cenozoic is the most clearly defined geological era in terms of climate and life history. Since the late 60’s, oxygen isotopic values have been measured on benthic foraminifera shells coming from deep-sea records. These values give insights about the evolution of deep-sea temperatures and continental ice volume during the last 65 million years. More than ten years ago, Zachos et al. (Zachos et al., 2001) compiled over 40 published isotopic records to depict carbon and oxygen isotopic trend through the Cenozoic. This seminal study has been recently updated and is classically used as the paleoclimatic reference in most climate and paleoecological studies. From this record one can divide the Cenozoic into warm periods (Paleocene and Eocene), also called “Greenhouse climates”, and cooler periods (Oligocene, Miocene and Pliocene) during which permanent icesheets developed at the poles.

The Paleocene-Eocene Thermal Maximum

  • 1 ppmv stands for parts per million volume. It describes very small quantities of gaz concentration. (...)

2Temperatures were warm during the Paleocene; polar regions were ice-free and characterized by subtropical flora and fauna. As an example, the late Paleocene flora that grew on the former land bridge connecting Antarctica to South America (Seymour Island fossil records), at latitude 63°S, is described as a “paratropical forest growing in a warm and rainy climate” (Reguero et al., 2002). Atmospheric CO2 reconstructions are uncertain for this period but suggest pCO2 higher than 1000 ppmv1.

3At ~55.5 Ma, a large isotopic excursion is observed in the δ18O and δ13C records from benthic microfossils. This excursion is associated with a 4-6°C deep-sea warming, a 5°C increase in tropical Sea Surface Temperatures (hereafter SST) and an up to 8-10°C increase in high-latitude SST (Zachos et al., 2003). This event has been called the Paleocene-Eocene Thermal Maximum (PETM) and is now used to define the Paleocene-Eocene boundary, as it is contemporaneous with the Paleocene-Eocene transition observed in the North American Land Mammal Series. The event lasted no more than 20,000 years, and a return to normal values occurred after 170,000 years, making the PETM the most abrupt climate change recorded in Earth history. Hypotheses to explain the mechanisms and the source of the massive input of depleted carbon in the global system, ultimately leading to this carbon isotope excursion and global warming, are still debated. The usual mechanisms driving isotopic excursions (volcanism, weathering) are too slow (with too low values for carbon originating from volcanic activity) to be responsible for the PETM. The consensual hypothesis invokes a massive release of isotopically light carbon (relatively little 13C) in the ocean system, through methane injection. Where does the methane come from? Some authors invoke volcanic activity in the northeastern Atlantic triggering a massive methane release (Svensen et al., 2004); others suggest that huge amounts of methane hydrates (clathrates) stored at the bottom of the oceans were destabilized 55.5 Ma ago (for a still unknown reason) (Dickens et al., 1997).

4The PETM is contemporaneous with a major extinction event in the deep ocean. Between 30 and 50% of the benthic foraminifera community were wiped out within less than 10,000 years. The massive input of methane in the water at all depths could have triggered low oxygenation and increased water corrosity, both through methane oxidation. Marine productivity may have been altered as well. However, none of these changes was global; otherwise extinct cosmopolitan species may have survived in refugia and repopulated afterwards. The remaining factor is the global warming that occurred in the ocean; but the mechanism linking it to foraminifera extinction is not yet understood. Recent modeling studies suggest that the extinction was due to multiple environmental changes, all related to perturbation of the carbon cycle (Winguth et al., 2012).

  • 2 65 to 75 % of extant fauna follows Bergman’s rule.

5On the continents, temperatures increased by ~5°C during the PETM. This warming impacted continental flora which responded with inter- and intra-continental dispersals (Wing et al., 2005) and major compositional turnovers (20 % of US Gulf Coast palynoflora became extinct after the PETM (Harrington & Jaramillo, 2007)). The Paleocene-Eocene transition is also crucial for the fauna, as three major modern orders of mammals, namely Primates, Artiodactyla (e.g. deer, cows, camels) and Perissodactyla (e.g. horses), appear coincidently in the fossil record at that time in Europe, Asia and North America, while some Paleocene taxa (Champsosaurus, Plesiadapis) became extinct. The warming led to a reduction in size of several taxa (herbivorous ungulates, Primates, Artiodactyla), conforming to the famous ecological Bergman’s rule, that specifies that species have smaller body size at greater ambient temperatures2 (Secord et al., 2012). Apparent dwarfing of some mammal taxa was even evidenced at the PETM, and might be linked to the increased pCO2 which would have affected plant growth, digestibility and ultimately herbivore growth (Gingerich, 2006).

From the Eocene to the Oligocene: Descent into the Icehouse3

  • 3 This title is directly inspired from E. Thomas Research Focus published in Geology (2008).

6The δ18O record indicates that shortly after the PETM, a climatic optimum was reached, making the early Eocene the warmest period of the Cenozoic. Recent records suggest that climate variability existed during this period, causing climate shifts from warm to very warm (hyperthermal) states. At the end of the early Eocene a global cooling began that culminated at the Eocene-Oligocene transition, at 32 Ma. At that time continental drift had made Earth look closer to its present-day appearance: the Drake Passage, separating Antarctica and South America, had opened following an earlier opening of the Tasmanian gateway (separating Antarctica from Australia). India had drifted northward to collide with Asia, thereby starting the uplift of the Tibetan Plateau.

7The isotope shift observed at 34 Ma in the δ18O record has been explained both by the inception of the Antarctic ice sheet (at least on the western edge of the continent) and by global cooling (Hansen et al., 2008). Initiation of the Antarctic ice sheet could have been linked to two factors which likely interacted. First, the opening of the Drake Passage would have profoundly changed global ocean circulation and allowed the Circum-Antarctic Current to be activated, ultimately isolating the Antarctic continent. The second factor is the lowering of pCO2 throughout the Eocene, from more than 1,000 ppmv to less than 750 ppmv (Pearson et al., 2009), a threshold that climate model studies have shown to be crucial to initiate a glaciation over Antarctica (DeConto & Pollard, 2003). No conclusive explanation for the CO2 reduction during the Eocene exists but many consider that the uplift of the Himalayas played a major role. With the Himalayan orogeny, a lot of organic carbon from the continent was buried through sedimentation in the deltaic regions such as the Bengal fan, thereby activating the marine biological pump and enhancing atmospheric CO2 consumption. This mechanism was active during the whole Himalayan orogeny and consumed 2 to 3 times more CO2 than silicate weathering (Galy et al., 2007).

8In the oceans, Eocene cooling has been associated with a major extinction in tropical nannoplankton, foraminifera, bivalves and gastropods (Prothero, 2004). On the continents, Antarctic vegetation shifted from an evergreen forest dominated by Nothofagus sp. to a tundra vegetation (Thorn & DeConto, 2006). Neotropical fossil palynoflora show a continuous decrease in diversity, likely linked to global warming or to available area for species (Jaramillo et al., 2006). In North America, broad-leaved deciduous forests replaced paratropical forests within 500 thousand years, while several indicators show widespread increase in aridity. The consequences of climate change for North American faunal biodiversity are still a matter of debate, but recent studies suggest that the rise of the Oligocene fauna was likely linked to the immigration of Asian taxa following the sea-level drop induced by growth of the Antarctic ice sheet (Figueirido et al., 2012). In Eurasia, climate cooling at the Eocene-Oligocene boundary is associated with a major faunal turnover called “Grande Coupure”, during which 60% of endemic European fauna (82% of placental mammals) went extinct; subtropical taxa declined and new taxa (e.g., rhinoceroses) migrated from Asia. The vertebrate records from Mongolia also show a strong turnover with Perissodactyla replaced by lagomorphs and rodents (Yuanqing et al., 2007), consistent with a trend to aridification indicated by the lithology (Dupont-Nivet et al., 2007; Kraatz & Geisler, 2010) and by changes in rodent dental patterns (Gomes Rodrigues et al., 2012).

9After the Eocene-Oligocene transition, the global climate was marked by strong fluctuations, as testified by variations in the benthic foraminera δ18O record. A continuous record from the equatorial Pacific shows that these variations involved events in the carbon cycle and glaciations. Detailed analyses of the signal revealed that these events followed fluctuations in the earth’s orbit, i.e. linked to periodical variations of insolation (cyclical changes in eccentricity and obliquity of the earth’s orbit) (Pälike et al., 2006). Over these variations, there was a global warming trend between ~27 Ma and 24 Ma, with increases in deep-ocean temperatures and a reduction of the ice volume accumulated just after the Eocene-Oligocene transition. Afterwards, several records show an increase in deep-sea δ18O values at the Oligocene-Miocene transition. This increase corresponds to a transient cooling event (called Mi-1) involving a major increase in continental ice volume and a 2°C decrease in deep-ocean temperature, ~23 million-years ago (Lear et al., 2004).

Miocene climate changes and vegetation revolution

10One of the main characteristics of the Miocene period is the reorganization of the regional climates in tandem with important changes in paleogeography, including mountain uplifts and seaway opening and closing. China provides a detailed record of environmental changes for the Oligocene and the Miocene. Several indicators including coal, pollen assemblages, saline lake deposits, and pedogenic carbonates, as well as the analyzed of soil-forming processes, show that Chinese atmospheric circulation was dramatically modified ~22 million-years ago. Paleoenvironments organized in a typical zonal pattern during the Oligocene, with southern humid belt and subtropical desertic areas, were replaced by monsoon-like regimes, including two-season rainfall patterns and inland arid regions (Guo et al., 2008). At that time the Asian continent was particularly affected by tectonic-driven changes, with the ongoing uplift of the Himalayas and Tibetan Plateau, as well as with the shrinkage of the Paratethys Sea (including the creation of the African-Eurasian land connection by the late Miocene). Numerous climate modeling studies have been carried out to understand how these tectonic parameters drove climate changes through the Neogene. The idea that an uplifted Tibetan Plateau was necessary to maintain strong seasonal temperature contrasts between the ocean and the continent has been put forward since the 80’s (Ruddiman & Kutzbach, 1989), though recent studies suggest that the Himalayas even without the Plateau were a sufficient driver to set up a monsoonal circulation (Boos & Kuang, 2010). Other studies have suggested that the withdrawal of the Paratethys seaway helped the monsoon onset (Ramstein et al., 1997; Fluteau et al., 1999), which has been further confirmed by different models (Guo et al., 2008).

11Another example of the tectonics-climate-ecology link is South America during the Miocene. Although debate continues to rage concerning the exact timing and speed of the Andes uplift, it is clear that this unique mountain range, covering over 7,000 km south to north, with peak elevations in excess of 6,000 m, got uplifted from 25 Ma to 5 Ma. This uplift had several consequences on South American environments. Nowadays, the Andes act as a massive orographic barrier that deflects zonal winds and associated moisture fluxes coming from the Pacific and the Atlantic oceans. Climate zonation of South America shows widespread rain-shadows, as in Patagonia, where there is a strong rainfall contrast between the very humid western slope of the Andes that collect all the westerly-winds rainfall and the eastern slope which is very arid. Interestingly, the palynological records of eastern Patagonia show a transition, during the Miocene, from forested environments (Nothofagaceae, Podocarpaceae) to arid-adapted floras (Asteraceae, Chenopodacieae), that can be plausibly explained by the uplift of the southern Andes and the onset of the rain-shadow effect (Barreda & Palazzesi, 2007). The Andean uplift also affected the Amazon basin in many ways: the uplift of the northern cordilleras closed water channels between the continent and the Pacific and led to changes in river routing. From northward-flowing wetland systems (Pebas Lake), the Amazon drainage pattern was reorganized into a river system close to the present-day one, i.e. a long river originating in the Andes foothills and flowing into the tropical Atlantic. The causal relationship between these environmental changes and the outstanding present-day biodiversity of the Amazonian basin is a huge and fascinating research topic. Recent studies, combining fossil records, phylogenetic data, geology, sedimentary and paleoclimatological information (Hoorn et al., 2010), suggest that the Andean uplift drove climate changes (Sepulchre et al., 2010). Phylogenetics on plants have shown that the Andes favoured the appearance of island-like habitats, thereby fostering very rapid diversification (Hughes & Eastwood, 2006) and helped the dispersal of neotropical taxa to the South (Antonelli et al., 2009).

12The Miocene was a period of important paleoecological changes in Europe as well. During the Mid-Miocene Climate Optimum (MMCO), European climate became warm enough to allow the expansion of subtropical forests and associated hominoids in Europe. Numerical experiments with climate and vegetation models have recently shown that this climate change in Europe depended on particular values of pCO2 (between 560 and 700 ppmv), but was also linked to the size of Antarctic ice sheet that impacted the Atlantic ocean circulation, surface temperatures and moisture transported towards Europe (Hamon et al., 2012).

13As with South America, tectonics changes were also active in eastern Africa during the Miocene. Several geological indicators show that the Eastern African Rift System (EARS) got uplifted from ~20 to 3 Ma, inducing the building of a 6000-km-long elevated orographic barrier, mostly oriented north-south and bordered by peaks reaching between 1500 and 5100 m. The study of tooth enamel δ13C from the East African records shows a marked shift towards more positive values of δ13C between 10 and 8 Ma. This shift, together with paleontological data, suggests a change in ungulate feeding habits, switching from tree-leaves to grasses. Numerical simulations with the high-resolution General Circulation Model LMDz suggest that the gradual uplift of the EARS, by deflecting northward the moisture flux coming from the Indian Ocean, drove the aridification of eastern Africa and likely the subsequent changes in vegetation cover, the tropical forest being replaced by savannah and open grassland (Sepulchre et al., 2006).

14These changes took place with the context of a larger paleoecological revolution, namely the worldwide forest-to-grassland transition and the subsequent expansion and rise of the C4 grasslands displacing C3 vegetation. The terms C4 and C3 have to do with two physiologically distinct photosynthesis pathways used by plants; which one is dominant depends on growing season temperature and atmospheric CO2 concentration. Basically, at high temperatures and low CO2 concentration, the Rubisco enzyme of C3 plants barely distinguishes O2 from CO2, which leads to photorespiration, ultimately lowering the photosynthesis yield. Conversely, a CO2-storing strategy helps C4 to plants prevent photorespiration, but this process has an energy cost that is higher than the C3 photorespiratory cost at low temperatures and high CO2, while C4 photosynthesis is also less effective in low-light conditions. The critical thresholds of temperatures and CO2, defining the point where one pathway becomes dominant over the other, can be calculated. Under present-day pCO2, this temperature threshold lies between 20 and 25°C. In tropical environments the pCO2 threshold is ~500 ppmv (Osborne, 2008). The present-day geographical distribution of C3/C4 corresponds quite well to these constraints: C4 grasslands are found mainly in subtropical and tropical areas, while forested areas are dominated by the C3 pathway (Edwards et al., 2010).

15Recent phylogenies (reviewed in Edwards et al., 2010) suggest that C4 photosynthesis likely evolved several times, but that the earliest probable origin was during the early Oligocene, ca. 32-30 Ma. Phytoliths record the presence of C4 grasses in North American grasslands by 19 Ma, but their dominance over C3-plants came only several million years later (Edwards et al., 2010). Likewise a worldwide study (Pakistan, Great Plains of North America, East Africa) of δ13C measured on paleosols and fossil teeth of herbivores shows that C4 plants started to dominate low-latitude ecosystems on a global scale only around 8 million years ago (Uno et al., 2011).

16As C4 photosynthesis is favoured under warm temperatures and low CO2, authors have long suggested that the evolution towards this system and the subsequent rise of C4-plants were linked to the Cenozoic decrease in atmospheric pCO2 (Cerling et al., 1997). As recent pCO2 reconstructions show a marked drop in pCO2 at the Eocene-Oligocene (Pearson et al., 2009; Pagani et al., 2011), a link between pCO2 and evolution towards the C4 pathway at that time is highly probable. However, there is no clear declining trend in pCO2 during the Miocene (Pagani et al., 2009), while C4 abundance in vegetation composition increases (Edwards et al., 2010). Thus, other processes have been explored to explain the rise of C4-grasslands during the Miocene. Recent hypotheses involve two climate forcing factors, namely rainfall and the occurrence of fire (Osborne, 2008). One possibility is that the abrupt intensification of monsoon systems, which corresponds to the development of a seasonal climate (i.e. with a wet season and a dry season), could have increased wood-plants mortality and favoured C4-grassland expansion during the Miocene. An alternative viewpoint, based on δ18O records from Asian freshwater bivalve shells, suggests that the overall annual amount of rainfall could have decreased (without particular seasonal changes) during the Miocene, and that enhanced aridity favoured grasslands. Both mechanisms would have increased the frequency and intensity of droughts, thereby favouring grassland expansion. Development of seasonal climates is also thought to have increased the frequency of fires, by sustaining high biomass production during the wet season and reducing fuel moisture and generating frequent lightning strikes that can ignite fires during the dry season. Fires kill wood-plants and open up the habitats, favouring high-light conditions in which C4 grasses overtake C3 grasses, increasing primary productivity and production of dead foliage. As dead foliage of grasses is an efficient fuel for fires, there is a positive feedback ultimately leading to C4-grassland expansion (Keeley and Rundel, 2005; Scheiter et al., 2012). In order to confirm that these paleoecological processes drove the diversification of C4 as well, future studies will have to include both paleoenvironmental and phylogenetical information.

17For years, the Mio-Pliocene transition to grasslands in eastern Africa was thought to have been the main driver of hominid evolution. Indeed, until 1994 and the discovery of the australopithecine “Abel” in the Chad basin, all hominids were exclusively recorded in eastern or southern Africa, and palaeontologists were thinking in terms of the ‘East Side Story’ scenario, in which early hominids evolved in the savannah east of the Rift Valley, while early African apes remained in the tropical forest, west of the Rift (Brunet, 2010). However, older hominids were subsequently discovered west of the rift, with Sahelanthropus tchadensis being the earliest at ~7 Ma (Brunet et al., 2002). These discoveries led to reconsideration of the eastern African evolution hypothesis and to further analyses of paleoenvironments and paleoclimates of western Africa. The fossil fauna and sedimentological evidence associated with S. tchadensis showed that this hominid evolved in a perilacustrine environment, i.e. a mosaic of environments close to the present-day Okavango inland delta (mosaic of swamps, forests, savannahs and desert areas). Records also show repeated oscillations between dry and humid environments culminating in megalakes (Vignaud et al., 2002). Therefore the new challenges concerning paleoecology and paleoclimate for this period are first to understand what were the drivers of such climatic oscillations and second to quantify the links between changing environments and hominid evolution.

Pliocene warm period and the role of seaways

18During the last five million years, Earth’s climate completed its transition to a full “icehouse” state, i.e. a state in which ice sheets are present at both poles. Considering the benthic δ18O record, the Pliocene (5.3 Ma-1.8 Ma) can be divided into three sub-periods: the early Pliocene warm period (5.3 Ma-3 Ma), the mid-Pliocene warm interval (ca. 3 Ma), and the late Pliocene climate deterioration until the Pleistocene (Haywood et al., 2009). Sea surface temperatures estimated from Mg/Ca measurements and δ18O records show that the early Pliocene was 3 to 5°C warmer than present-day climate at mid-latitudes. Moreover the early Pliocene was also characterized by a weak (1.5 ± 0.9°C) SST gradient between the western and eastern Pacific, typical of sustained, permanent, El Niño-like conditions (Wara et al., 2005). The forcing mechanisms of the Pliocene warmth were several and are still very much debated; they include higher level of pCO2, enhanced ocean heat transport and strong ice-albedo feedback (Haywood et al., 2009). The Pliocene warmth was followed by a climate deterioration that ultimately led to the onset of the northern hemisphere ice sheet and glacial-interglacial cycles. It has long been thought that the closure of the Central American Seaway (CAS, also called Panama isthmus) could have driven this onset by enhancing the thermohaline circulation, increasing atmospheric moisture transport to high northern latitudes and allowing heavy snowfall on Greenland (Haug et al., 2001; Bartoli et al., 2005). However, this view is changing as (1) climate modeling experiments showed that pCO2 played a larger role in ice sheet initiation than did tectonics (Lunt et al., 2008) and (2) both geological and biological evidence demonstrates that the final constriction of the seaway occurred during the Miocene (Montes et al., 2012; Montes et al. 2015; Bacon et al. 2015). Ongoing research aims at providing constraints on the timing of the isthmus closure as the connection between North and South America was a crucial biogeographical event, with the so-called Great American Biotic Interchange (GABI), i.e. a major exchange of flora and fauna between the two continents. Among land mammals, the GABI was, surprisingly, a very asymmetric process. Indeed, the faunal evolutionary pathways clearly diverge between the two continents, as taxa of northern origin successfully diversified in South America, whereas southern taxa showed little or no diversification once in North America. This pattern has two consequences on extant fauna; one is that more than half (53%) of extant South American genera descend from taxa of northern origin. The other is that tropical North America was conquered by taxa from Amazonia, whereas the temperate North American fauna shows only 4 surviving genera from South America (Webb, 2006). The final closure of the CAS has also been invoked to explain the major extinction of Caribbean reef corals and mollusks at the end of the Pliocene. Mechanisms involved include the collapse of planktonic productivity with a closed seaway. However a detailed study of extinction rates of mollusks and corals has shown that a 2-million-year time lag between the environmental changes and the actual extinctions, suggesting that environmental change was not the sole driver of the extinctions (O’Dea et al., 2007).

19On the opposite side of the globe, another ocean gateway is also thought to have played a major role in Pliocene climate and environment changes, namely the Indonesian seaway, involving the Indonesian throughflow, or ITF. Tectonics shows that New Guinea and Australia lay 2-3° south of their present-day position during the Pliocene, while Halmahera, an island that reduces the Pacific flow in the Indian Ocean, was much smaller at that time. Ocean circulation simulations including this paleogeographic information suggest that these changes of the ITF during the Pliocene brought colder waters from the South Pacific to the Indian Ocean, thereby reducing evaporation and associated rainfall over eastern Africa. Authors also propose that the ITF likely stimulated the growth of ice sheets, by reducing atmospheric heat transport to high latitudes (Cane & Molnar, 2001). Aridification of eastern Africa is very well documented throughout the Miocene (see previous section) and the Pliocene, thanks to pollen records (Bonnefille et al., 2004). It is thought that this trend influenced early mammals and human evolution in eastern Africa, especially Australopithecus afarensis.

20This short synthesis of climate-biosphere interactions during the Cenozoic shows the complexity of interactions between life and abiotic factors. Whether climate and tectonics have been the main drivers of paleoecological changes through time or not is an ongoing question in Earth and Life sciences; still the Cenozoic seems to demonstrate it.

21Although the impact of modern humans on climate and ecosystems has no equivalent, in terms of range and speed, in the geological record, it is tempting to compare modern-day loss of species richness and human-driven climate change to what happened in the past. Climatologists attempt to compare anthropic climate change to the PETM in order to assess the environmental consequences of massive carbon input in the atmosphere (Hansen et al., 2008). Biologists now wonder whether Earth’s sixth mass extinction may be on its way or not. As more and more evidence suggests an affirmative answer (Barnosky et al., 2011), studying what happened in the past appears to be more and more relevant for biodiversity conservation purposes.


Antonelli, A., Nylander, J. A. A. et al., 2009, Tracing the impact of the Andean uplift on Neotropical plant evolution, Proceedings of the National Academy of Sciences, 106 (24), 9749.

Bacon, Christine D., Silvestro, Daniele, Jaramillo, Carlos, Tilston Smith, Brian, Chakrabarty, Prosanta, Antonelli, Alexandre, 2015, Biological evidence supports an early and complex emergence of the Isthmus of Panama, Proceedings of the National Academy of Sciences, 112 (19), 6110-6115.

Barnosky, A. D., Matzke, N., et al., 2011, Has the Earth’s sixth mass extinction already arrived?, Nature, 471 (7336), 51-57.

Barreda, V., Palazzesi, L., 2007, Patagonian vegetation turnovers during the Paleogene-early Neogene: Origin of arid-adapted floras, The Botanical Review, 73 (1), 31-50.

Bartoli, G., Sarnthein, M. et al., 2005, Final closure of Panama and the onset of northern hemisphere glaciation, Earth and Planetary Science Letters, 237 (1), 33-44.

Bonnefille, R., Potts, R. et al., 2004, High-resolution vegetation and climate change associated with Pliocene Australopithecus afarensis, Proceedings of the National Academy of Sciences of the United States of America, 101 (33), 12125-12129.

Boos, W. R., Kuang, Z., 2010, Dominant control of the South Asian monsoon by orographic insulation versus plateau heating, Nature, 463 (7278), 218-222.

Brunet, M., 2010, Two new Mio-Pliocene Chadian hominids enlighten Charles Darwin’s 1871 prediction, Philosophical Transactions of the Royal Society B: Biological Sciences, 365 (1556), 3315-3321.

Brunet, M., Guy, F. et al., 2002, A new hominid from the Upper Miocene of Chad, Central Africa, Nature, 418 (6894), 145-151.

Cane, M. A., Molnar, P., 2001, Closing of the Indonesian seaway as a precursor to east African aridification around 3-4 million years ago, Nature, 411 (6834), 157-162.

Cerling, T. E., Harris, J. M. et al., 1997, Global vegetation change through the Miocene/Pliocene boundary, Nature, 389 (6647), 153-158.

DeConto, R. M., Pollard, D., 2003, Rapid Cenozoic glaciation of Antarctica induced by declining atmospheric CO 2, Nature, 421 (6920), 245-249.

Dickens, G. R., Castillo, M. M. et al., 1997, A blast of gas in the latest Paleocene: Simulating first-order effects of massive dissociation of oceanic methane hydrate, Geology, 25 (3), 259-262.

Dupont-Nivet, G., Krijgsman, W. et al., 2007, Tibetan plateau aridification linked to global cooling at the Eocene-Oligocene transition, Nature, 445 (7128), 635-638.

Edwards, E. J., Osborne, C. P. et al., 2010, The origins of C4 grasslands: Integrating evolutionary and ecosystem science, Science, 328 (5978), 587-591.

Figueirido, B., Janis, C. M. et al., 2012, Cenozoic climate change influences mammalian evolutionary dynamics, Proceedings of the National Academy of Sciences, 109 (3), 722-727.

Fluteau, F., Ramstein, G. et al., 1999, Simulating the evolution of the Asian and African monsoons during the past 30 Myr using an atmospheric general circulation model, Journal of Geophysical Research, 104 (D10), 11995-11912, 11018.

Galy, V., France-Lanord, C. et al., 2007, Efficient organic carbon burial in the Bengal fan sustained by the Himalayan erosional system, Nature, 450 (7168), 407-410.

Gingerich, P. D., 2006, Environment and evolution through the Paleocene-Eocene thermal maximum, Trends in ecology & evolution, 21 (5), 246-253.

Gomes Rodrigues, H., Marivaux, L. et al., 2012, Expansion of open landscapes in Northern China during the Oligocene induced by dramatic climate changes: Paleoecological evidence, Palaeogeography, Palaeoclimatology, Palaeoecology, 358–360, 62–71.

Guo, Z., Sun, B. et al., 2008, A major reorganization of Asian climate by the early Miocene, Climate of the Past, 4 (3), 153-174.

Hamon, N., Sepulchre, P. et al., 2012, Growth of subtropical forests in Miocene Europe: The roles of carbon dioxide and Antarctic ice volume, Geology, 40 (6), 567-570.

Hansen, J., Sato, M. et al., 2008, Target atmospheric CO2: Where should humanity aim?, Open Atmos. Sci. J., 2, 217-231

Harrington, G. J., Jaramillo, C. A., 2007, Paratropical floral extinction in the Late Palaeocene-Early Eocene, Journal of the Geological Society, 164(2): 323-332.

Haug, G. H., Tiedemann, R. et al., 2001, Role of Panama uplift on oceanic freshwater balance, Geology, 29 (3), 207-210.

Haywood, A. M., Dowsett, H. J. et al., 2009, Introduction. Pliocene climate, processes and problems, Philosophical Transactions of the Royal Society A: Mathematical, Physical and Engineering Sciences, 367 (1886), 3-17.

Hoorn, C., Wesselingh, F. et al., 2010, Amazonia through time: Andean uplift, climate change, landscape evolution, and biodiversity, Science, 330 (6006), 927-931.

Hughes, C., Eastwood, R., 2006, Island radiation on a continental scale: exceptional rates of plant diversification after uplift of the Andes, Proceedings of the National Academy of Sciences, 103 (27), 10334-10339.

Jaramillo, C., Rueda, M. J. et al., 2006, Cenozoic plant diversity in the Neotropics, Science, 311 (5769), 1893-1896.

Keeley, J. E., Rundel, P. W., 2005, Fire and the Miocene expansion of C4 grasslands, Ecology Letters, 8 (7), 683-690.

Kraatz, B. P., Geisler, J. H., 2010, Eocene-Oligocene transition in Central Asia and its effects on mammalian evolution, Geology, 38 (2), 111-114.

Lear, C. H., Rosenthal, al., 2004, Late Eocene to early Miocene ice sheet dynamics and the global carbon cycle, Paleoceanography, 19 (4), 1-11.

Lunt, D. J., Foster, G. L. et al., 2008, Late Pliocene Greenland glaciation controlled by a decline in atmospheric CO2 levels, Nature, 454 (7208), 1102-1105.

Montes, C., Cardona, A. et al., 2012, Evidence for middle Eocene and younger land emergence in central Panama: Implications for Isthmus closure, Geological Society of America Bulletin, 124 (5-6), 780-799.

Montes, C., Cardona, A., Jaramillo, C., Pardo, A., Silva, J. C., Valencia, V., Ayala, C. et al. 2015, Middle Miocene closure of the Central American Seaway, Science, 348 (6231), 226-29.

O’Dea, A., Jackson, J. B. C. et al., 2007, Environmental change preceded Caribbean extinction by 2 million years, Proceedings of the National Academy of Sciences, 104 (13), 5501-5506.

Osborne, C. P., 2008, Atmosphere, ecology and evolution: what drove the Miocene expansion of C4 grasslands?, Journal of Ecology, 96 (1), 35-45.

Pagani, M., Caldeira, K. et al., 2009, The role of terrestrial plants in limiting atmospheric CO2 decline over the past 24 million years, Nature, 460 (7251), 85-88.

Pagani, M., Huber, M. et al., 2011, The role of carbon dioxide during the onset of Antarctic glaciation, Science, 334 (6060), 1261-1264.

Pälike, H., Norris, R. D. et al., 2006, The heartbeat of the Oligocene climate system, Science, 314 (5807), 1894-1898.

Pearson, P. N., Foster, G. L. et al., 2009, Atmospheric carbon dioxide through the Eocene-Oligocene climate transition, Nature, 461 (7267), 1110-1113.

Prothero, D. R., 2004, Did impacts, volcanic eruptions, or climate change affect mammalian evolution?, Palaeogeography, Palaeoclimatology, Palaeoecology, 214 (3), 283-294.

Ramstein, G., Fluteau, F. et al., 1997, Effect of orogeny, plate motion and land-sea distribution on Eurasian climate change over the past 30 million years, Nature, 386 (6627), 788-795.

Reguero, M. A., Marenssi, S. A. et al., 2002, Antarctic Peninsula and South America (Patagonia) Paleogene terrestrial faunas and environments: biogeographic relationships, Palaeogeography, Palaeoclimatology, Palaeoecology, 179, 189-210.

Ruddiman, W., Kutzbach, J., 1989, Forcing of late Cenozoic Northern Hemisphere climate by plateau uplift in southern Asia and the American west, Journal of Geophysical Research, 94 (D15), 18409-18418,18427.

Scheiter, S., Higgins, S. I. et al., 2012, Fire and fire-adapted vegetation promoted C4 expansion in the late Miocene, New Phytologist, 195 (3), 653-666.

Secord, R., Bloch, J. I. et al., 2012, Evolution of the earliest horses driven by climate change in the Paleocene-Eocene thermal maximum, Science, 335 (6071), 959-962.

Sepulchre, P., Ramstein, G. et al., 2006, Tectonic uplift and Eastern Africa aridification, Science, 313 (5792), 1419-1423.

Sepulchre, P., Sloan, L. C. et al., 2010, Modelling the response of Amazonian climate to the uplift of the Andean mountain range, Amazonia: Landscape and Species Evolution, 211-222.

Svensen, H., Planke, S. et al., 2004, Release of methane from a volcanic basin as a mechanism for initial Eocene global warming, Nature, 429 (6991), 542-545.

Thorn, V. C., DeConto, R., 2006, Antarctic climate at the Eocene/Oligocene boundary‚ climate model sensitivity to high latitude vegetation type and comparisons with the palaeobotanical record, Palaeogeography, Palaeoclimatology, Palaeoecology, 231 (1), 134-157.

Uno, K. T., Cerling, T. E., et al., 2011, Late Miocene to Pliocene carbon isotope record of differential diet change among East African herbivores, Proceedings of the National Academy of Sciences, 108 (16), 6509-6514.

Vignaud, P., Duringer, P. et al., 2002, Geology and palaeontology of the Upper Miocene Toros-Menalla hominid locality, Chad, Nature, 418 (6894), 152-155.

Wara, M. W., Ravelo, A. C. et al., 2005, Permanent El Niño-like conditions during the Pliocene warm period, Science, 309 (5735), 758-761.

Webb, S. D., 2006, The great American biotic interchange: Patterns and processes, Annals of the Missouri Botanical Garden, 93 (2), 245-257.

Wing, S. L., Harrington, G. J. et al., 2005, Transient floral change and rapid global warming at the Paleocene-Eocene boundary, Science, 310 (5750), 993-996.

Winguth, A. M. E., Thomas, E. et al., 2012, Global decline in ocean ventilation, oxygenation, and productivity during the Paleocene-Eocene Thermal Maximum: Implications for the benthic extinction, Geology, 40 (3), 263-266.

Yuanqing, W., Jin, M. et al., 2007, Major events of Paleogene mammal radiation in China, Geological Journal, 42, 415-430.

Zachos, J. C., Pagani, M. et al., 2001, Trends, rhythms, and aberrations in global climate 65 Ma to present, Science, 292, 686-693.

Zachos, J. C., Wara, M. W. et al., 2003, A transient rise in tropical sea surface temperature during the Paleocene-Eocene Thermal Maximum, Science, 302 (5650), 1551-1554.


1 ppmv stands for parts per million volume. It describes very small quantities of gaz concentration. 1 ppmv means 1 part in 1,000,000 parts that is 0.0001%.

2 65 to 75 % of extant fauna follows Bergman’s rule.

3 This title is directly inspired from E. Thomas Research Focus published in Geology (2008).


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