Use of Foraminifera in Climate Science
Summary and Keywords
The understanding of past changes in climate and ocean circulation is to a large extent based on information from marine sediments. Marine deposits contain a variety of microfossils, which archive (paleo)-environmental information, both in their floral and faunal assemblages and in their stable isotope and trace element compositions. Sampling campaigns in the late 19th and early 20th centuries were dedicated to the inventory of sediment types and microfossil taxa. With the initiation of various national and international drilling programs in the second half of the 20th century, sediment cores were systematically recovered from all ocean basins and since then have shaped our knowledge of the oceans and climate history. The stable oxygen isotope composition of foraminiferal tests from the sediment cores delivered a continuous record of late Cretaceous–Cenozoic glaciation history. This record impressively proved the effects of periodic changes in the orbital configuration of the Earth on climate on timescales of tens to hundreds of thousands of years, described as Milankovitch cycles. Based on the origination and extinction patterns of marine microfossil groups, biostratigraphic schemes have been established, which are readily used for the dating of sediment successions. The species composition of assemblages of planktic microfossils, such as planktic foraminifera, radiolarians, dinoflagellates, coccolithophorids, and diatoms, is mainly related to sea-surface temperature and salinity but also to the distribution of nutrients and sea ice. Benthic microfossil groups, in particular benthic foraminifera but also ostracods, respond to changes in water depth, oxygen, and food availability at the sea floor, and provide information on sea-level changes and benthic-pelagic coupling in the ocean. The establishment and application of transfer functions delivers quantitative environmental data, which can be used in the validation of results from ocean and climate modeling experiments. Progress in analytical facilities and procedures allows for the development of new proxies based on the stable isotope and trace element composition of calcareous, siliceous, and organic microfossils. The combination of faunal and geochemical data delivers information on both environmental and biotic changes from the same sample set. Knowledge of the response of marine microorganisms to past climate changes at various amplitudes and pacing serves as a basis for the assessment of future resilience of marine ecosystems to the anticipated impacts of global warming.
Microfossils are common constituents of marine deposits and may dominate the lithology of the sediment. Examples are the widely distributed calcareous nannofossil-dominated chalks of the late Cretaceous, the Cenozoic foraminiferal oozes at low to intermediate latitudes, and the Cenozoic diatom oozes of the Southern Ocean (Kennett, 1982). In the modern oceans, vast areas are covered by pelagic sediments, which primarily consist of microfossil remains from single-celled planktic organisms (Dutkiewicz, Müller, O’Callaghan, & Jónasson, 2015). A few examples of typical microfossil associations in the sand fraction of marine sediments are shown in Figure 1. In tropical to temperate regions, the deep-sea floor above the calcite compensation depth (CCD) is commonly covered by calcareous nannofossil-foraminiferal ooze, while siliceous microfossils (diatoms and radiolarians) are dominant below the CCD, and in high-productivity areas (Berger & Herguera, 1992; Dutkiewicz et al., 2015). Benthic organisms are more abundant in continental slope and shelf ecosystems, and in marginal marine settings, where their remains can represent the dominant microfossil group (Fig. 1). Under suitable preservation conditions, such as hypoxia, neritic marine sediments commonly contain a significant amount of organic microfossils, particularly the cysts of dinoflagellates (dinocysts), which provide complementary environmental information on near-coastal oceanography and oxygenation (Sluijs, Pross, & Brinkhuis, 2005).
Even a small deep-sea sediment sample commonly contains a large number of microfossils. The diversity, morphology, species composition, and geochemistry of their skeletal remains from ocean sediment cores has delivered a wealth of information on changes of oceanic climate and circulation during the Mesozoic and Cenozoic time periods. Older marine sediments from land sections commonly also contain a variety of microfossils. Important Paleozoic microfossil groups include acritarchs and radiolarians (Riegel, 2008; Servais et al., 2016), larger foraminifera (e.g., fusulinids: BouDagher-Fadel, 2018), and ostracods (Crasquin & Forel, 2013). The evolution of the different groups is closely linked to climate and sea-level changes, and is punctuated by the impact of mass extinctions and subsequent radiation phases (Falkowski et al., 2004). Stable isotope records of foraminifera document the Cretaceous and Cenozoic climate evolution in detail (Zachos, Pagani, Sloan, Thomas, & Billups, 2001; Friedrich, Norris, & Erbacher, 2012; Holbourn et al., 2018). The close reflection of orbital changes in the data series provides a powerful tool for cyclostratigraphic age control of marine sediments (Imbrie et al., 1984; Lisiecki & Raymo, 2005; Hinnov & Hilgen, 2012).
Marine micropaleontological research in the 21st century is commonly conducted in interdisciplinary teams and delivers quantitative information on the physical and biogeochemical properties of past oceans (e.g., Kucera, Schneider, & Weinelt, 2006). These studies are accompanied by and integrated with research on the biodiversity, molecular phylogeny, biomineralization and biology of the different microfossil groups (e.g., Armbrust, 2009; De Nooijer, Spero, Erez, Bijma, & Reichart, 2014; Morard et al., 2016; Bernhard, Geslin, & Jordan 2018). In the early 21st century, a new tradition of communication skills is developing for the realization of joint transdisciplinary projects between proxy- and model-based research and research placed within a societal context.
This article aims at providing a concise overview on the use of marine micropaleontology in climate science. Emphasis is laid on applications of foraminifera and their geochemical test composition with reference to other microfossil groups from the late Mesozoic and Cenozoic eras. Benthic and planktic foraminifera are common constituents in marine sediments representing a wide range of shallow to deep marine ecosystems. Due to their comparably large size, foraminifera are easy to study and probably represent the prime group of microfossil applications in climate science. The various topics addressed reflect a personal choice of the author and do not provide an adequate representation of other microfossil groups and applications for the Paleozoic and early Mesozoic eras.
Development of Marine Micropaleontology in Climate Science
Historical Milestones in Marine Micropaleontology
The historical development of marine micropaleontology is closely connected to technical innovations and the systematic exploration of the oceans. The invention of the first compound microscope by the Dutch optician Zachariasse Janssen in the year 1590 and subsequent advancement of optical microscopes allowed for the study of sand-sized and smaller objects in great detail.
Among the first images of marine microfossils is a foraminifer depicted in a letter of Antoni van Leeuwenhoek from the year 1700, which can be clearly assigned to Elphidium, a widespread genus of intertidal and inner-neritic ecosystems. More systematic studies on microfossils followed, including the work of Bianchi (1739), Soldani (1780, 1791) and Fichtel and Moll (1803), in which foraminifera were considered as “micro-mollusks,” specifically, microscopic ammonites (Romano, 2015). The class foraminifera was finally introduced by d’Orbigny (1826).
A great advancement in the study of marine microfossils came with the scientific exploration of the oceans during numerous ship-based expeditions. Sampling campaigns in the late 19th and early 20th centuries were dedicated to the inventory of life in the oceans, sediment types, and microfossil taxa. The documented results of campaigns such as the expedition of the H.M.S. Challenger (Fig. 2) in the years 1872 to 1876 still provide the systematic basis for many modern micropaleontological studies. The scientific results of this expedition were published in no fewer than 50 volumes. One volume was dedicated to the description and illustration of radiolarians by Ernst Haeckel (1887), another volume to the documentation of foraminifera by Henry B. Brady (1884). The microfossil collections of Ernst Haeckel (radiolarians) and Christian G. Ehrenberg (a micropaleontologist of the 19th century who studied diatoms) are still matter of reexamination (Tanimura & Aita, 2009). The foraminiferal taxonomy of Brady (1884) was revised twice, by Barker (1960) and Jones (1994), who updated the taxonomy to modern standard and added valuable historical information on the Challenger expedition, collections, and the activities of contemporary scientists. Various other expeditions followed and over the years accumulated information on the taxonomy and distribution patterns of marine microfossils. These expeditions included the Gauss expedition (1901–1903), the British Terra Nova expedition (1910), the German Meteor expedition (1925–1927), and the Swedish Albatross expedition (1947–1948), just to mention a few examples.
With the initiation of the Deep-Sea Drilling Project (1968–1983; international phase from 1975) and subsequent international ocean drilling programs (ODP 1985–2003, IODP since 2003), sediment cores were systematically drilled in all ocean basins and the recovered materials were used to boost the understanding on Cretaceous and Cenozoic climatic, environmental, evolutionary, and ocean circulation changes. Among many others, micropaleontological highlights include the documentation of extinction and recovery dynamics of the ecosystem across the Cretaceous–Paleogene boundary interval (Culver, 2003; Coxall, D’Hondt, & Zachos, 2006; Alegret, Thomas, & Lohmann, 2012; Lowery et al., 2018), and evidence of the impact of hyper-thermal conditions, ocean acidification, and deoxygenation during the Paleocene–Eocene thermal maximum (and other hyperthermal events) on planktic and deep-sea ecosystems (Thomas, 1989; Gibbs, Bown, Sessa, Bralower, & Wilson, 2006; Thomas, 2007; Jennions, Thomas, Schmidt, Lunt, & Ridgwell, 2015; Schmidt, Thomas, Authier, Saunders, & Ridgwell, 2018) (see “Resilience and Recovery Potential of Marine Ecosystems with Respect to Perturbations”).
Modern marine micropaleontology relies to a large extent on field studies but increasingly involves laboratory experiments, such as cultivation of plankton and benthic microfossil groups under controlled conditions (e.g., Hemleben & Kitazato, 1995; Spero, Bijma, Lea, & Bemis, 1997; Kitazato & Bernhard, 2014; Schlüter et al., 2014), applications of high-resolution computer tomography (e.g., Caromel, Schmidt, & Rayfield, 2017), and the modeling of population dynamics and biogeographic patterns (e.g., Weinmann, Rödder, Lötters, & Langer, 2013).
Plankton Evolution, Biostratigraphy, and Ecology
The biogeochemical cycling of organic and inorganic carbon in the oceans and the functioning of marine ecosystems on different trophic levels are fundamentally dependent on photosynthesizing prokaryotes and eukaryotic phyto- and zooplankton (Longhurst, 1991; Falkowski & Knoll, 2007). Important phytoplankton groups such as calcareous nannofossils, autotrophic dinoflagellates, and diatoms, first originated in the Mesozoic, and—in spite of several extinction events—experienced marked radiation phases during the late Cretaceous, Paleogene, and Neogene (Knoll & Follows, 2016; Wiggan, Riding, Fensome, & Mattioli, 2018) (Fig. 3). Evolutionary turnover and productivity pulses of radiolarians exhibit a complex pattern and induced a stepwise decrease of dissolved silica levels during the Phanerozoic (Racki & Cordey, 2000). As a first approximation, the evolution of planktic foraminifera parallels that of phytoplankton and reveals particularly high diversities during the late Cretaceous, Eocene, and middle Miocene (McGowran, 2012). The appearance and dispersal of planktic calcifiers profoundly changed the CaCO3 saturation state of the ocean, leading to the establishment of a calcite compensation depth in the deep ocean (Zeebe & Westbroek, 2003). Changes in plankton diversity retrace the long-term sea-level development, suggesting a close relationship between planktic ecosystems, sea-surface temperature (SST) and the area of flooded continental shelves (e.g., Bown, Lees, & Young, 2004; Falkowski et al., 2004) (Fig. 3). The complex interplay between evolutionary and ecological processes also influences the body size and morphology of marine plankton, although the relations are not fully understood (Gibbs et al., 2006; Schmidt, Lazarus, Young, & Kucera, 2006).
The rapid floral and faunal turnover of phyto- and zooplankton and its wide distribution in marine sediments are the basis for manifold biostratigraphic and paleo-biogeographic applications. The origination and extinction patterns of marker taxa are used to define biostratigraphical zones. These biozones are specified by the International Commission on Stratigraphy, based on the temporal range of a single taxon or the combination of first and last appearance data (FAD, LAD) of several taxa (Wade, Pearson, Berggren, & Pälike, 2011; Gradstein, 2012) (Fig. 4). Regional biostratigraphic schemes were originally established in cooperation with oil and gas exploration in the early 20th century, and in the following decades reached a high level of sophistication with the generation of global biostratigraphic and chronological schemes in the frame of the international drilling campaigns (summary in Wade et al., 2011). Since the late 20th century, the application of astrochronological approaches has allowed for the refinement of biochrons and has greatly improved the applicability of marine proxy records for accurately dated paleoclimate reconstructions (e.g., Raffi et al., 2006; Hinnov & Hilgen, 2012). Useful ecobiostratigraphic information can also be retrieved from temporal changes in the abundance patterns of certain microfossil taxa during the Quaternary, such as the planktic foraminifer Globorotalia menardii in the Atlantic Ocean (Ericson, Ewing, Wollin, & Heezen, 1961), or the radiolarian Cycladophora davisiana in the Southern Ocean (Brathauer, Abelmann, Gersonde, Niebler, & Fütterer, 2001).
Field studies using plankton tows, filtering of water samples from various depth levels, and sediment traps yielded manifold insights into the distribution and ecology of different systematic groups and provide the basis for the paleo-ecological interpretation of planktic microfossils (e.g., Bork et al., 2015). The majority of phytoplankton species inhabit the mixed layer of the oceans, as long as sufficient light and nutrients are available, and account for almost half of the net primary production on Earth (Field, Behrenfeld, Randerson, & Falkowski, 1998; Falkowski & Knoll, 2007). Some zooplankton groups, such as planktic foraminifera and radiolarians, live in parts in the mixed layer, but also inhabit deeper water levels. These groups avoid turbid coastal conditions, because they pass through vertical habitat changes during their life cycle and are commonly confined to narrow salinity ranges (Lazarus, 2005; Schiebel & Hemleben, 2017). Accordingly, the proportion of planktic foraminifera to the total number of foraminifera in the sediment, often also referred to as plankton/benthos ratio, reflects the depositional water depth or distance to the coast (or both). This proxy can be applied as a simple but powerful tool to the approximation of paleo-water depth at the time of deposition (Gibson, 1989; Van der Zwaan, Jorissen, & de Stigter, 1990) (Fig. 5). As a rule of thumb, planktic foraminiferal tests account for approximately 50% of the total foraminiferal tests in sediments from the shelf break, and their proportion decreases above and increases below that level (Fig. 5). This relation is also affected by changes in shelf and slope geometry, and changes in productivity and oxygen content (Berger & Diester-Haass, 1988; Van Hinsbergen, Kouwenhoven, & van der Zwaan, 2005) (see “Quantitative Reconstruction of Relative Sea-Level Change”).
Most modern plankton taxa and assemblages are closely associated with SST resulting in zonal distribution patterns of coccolithophorids (McIntyre & Bé, 1967; Ziveri, Baumann, Böckel, Bollmann, & Young, 2004), diatoms (Cermeño & Falkowski, 2009), radiolarians (Moore, 1978; Lazarus, 2005), and planktic foraminifera (Bé, 1977; Schiebel & Hemleben, 2017). Similar biogeographic patterns have been reconstructed for past oceans (e.g., Mutterlose, Bornemann, & Herrle, 2005; Woods et al., 2014). Species–SST relationships are particularly well expressed in planktic foraminifera (Fig. 6), for which reason this group provided the most accurate and widely applicable transfer functions for late Quaternary SST reconstructions (Kucera et al., 2005) (see “Quantitative Reconstruction of Surface-Water Temperature and Salinity”).
The distribution of phytoplankton also responds to nutrient availability, as impressively illustrated by the distribution of chlorophyll in the surface ocean, based on satellite remote sensing (NASA Earth Observatory). In high-productivity regimes, plankton communities are typically dominated by diatoms and other siliceous microfossils (e.g., Gersonde, Crosta, Abelmann, & Armand, 2005). In low-latitude upwelling regions, such as the Arabian Sea, diatoms seem to compete with coccolithophorids for nutrients, resulting in seasonal and spatial plankton successions (Schiebel et al., 2004).
Benthic-Pelagic Coupling and Ecology of Benthic Foraminifera
Deep-sea benthic ecosystems are linked to the surface ocean via organic matter fluxes serving as the basic food resource for the organisms at the sea floor and within the sediments. This dependence is described as benthic-pelagic coupling (Graf, 1989), in which climate-related productivity changes in the surface ocean are transferred to the deep-sea realm (Cronin & Raymo, 1997) (Fig. 7). In the open ocean, approximately 10–40% of the organic carbon produced by photo- and zooplankton is exported from the photic zone, and only 0.01–1% arrives at the sea floor (Betzer et al., 1984; Berger & Wefer, 1990; Henson, Sanders, & Madsen, 2012). This proportion depends on water depth, settling velocity, and microbial decomposition rate in the water column, thus on the efficiency of the biological pump (Passow & Carlson, 2012). Deep-sea ecosystems are further influenced by oxygen availability, which is controlled by the ventilation of subsurface water masses, water temperature, and the microbial oxygen consumption. In the contemporary oceans, the majority of benthic ecosystems are well ventilated. Strong oxygen minimum zones (OMZ) develop at intermediate depth below high-productivity areas, for example in the Arabian Sea and the western boundary currents off Africa and America (Helly & Levin, 2004).
Oceanographic data document a worldwide OMZ expansion during the late 20th and early 21st centuries responding to global climate warming (Keeling, Körtzinger, & Gruber, 2010; Schmidtko, Stramma, & Visbeck, 2017). The expected changes will likely have significant biological impacts, such as vertical compression of benthic habitats (Stramma, Levin, Schmidtko, & Johnson, 2010).
Changes in benthic-pelagic coupling through time can be monitored through the investigation of microfossils, because various groups represent specific planktic and benthic ecosystems and different trophic levels. In this context, benthic foraminiferal faunas and their stable isotope and trace element signals can serve as proxies for the documentation of past natural oxygen variability and food fluxes (e.g., Zahn, Winn, & Sarnthein, 1986; Van der Zwaan et al., 1999; Murray, 2006; Jorissen, Fontanier, & Thomas, 2007; Hoogakker, Elderfield, Schmiedl, McCave, & Rickaby, 2015, Hoogakker et al., 2018) (see “Quantitative Reconstruction of Surface Primary and Export Productivity, Organic Matter Fluxes, and Oxygen”).
The distribution of deep-sea benthic foraminifera was originally related to specific water depth intervals (Bandy & Chierici, 1966), which appeared to be associated with physical and chemical characteristics of distinct water masses (e.g., Schnitker, 1974). Based on this relationship, shifts of water mass boundaries during glacial and interglacial changes were tentatively reconstructed, although the underlying ecological mechanisms remained elusive (e.g., Streeter & Shackleton, 1979). Subsequent ecological studies revealed that the species composition and microhabitat structure of deep-sea benthic foraminifera respond to changes in food availability and oxygen concentration at the sea floor (summary in Jorissen et al., 2007), and also to near-bottom current strength (Schönfeld, 2002).
Different benthic foraminifera inhabit specific microhabitats on and below the sediment surface and are able to change their microhabitat in response to changing biogeochemical conditions (e.g. Corliss, 1985; Mackensen & Douglas, 1989; Linke & Lutze, 1993). The simplified, general ecology of deep-sea foraminifera is best described by the so-called Trophic-Oxygen model (or TROX model), which considers the counteracting influences of food and oxygen, and resulting biogeochemical niches (Jorissen, de Stighter, & Widmark, 1995; Fontanier et al., 2002) (Fig. 8). In oligotrophic and well-oxygenated environments, the fauna is of low to intermediate diversity and mainly comprises epifaunal (epibenthic) taxa. In mesotrophic environments, faunal diversity is at a maximum and a variety of epifaunal and infaunal (endobenthic) niches are developed. Eutrophic, oxygen-limited ecosystems are inhabited by a low-diversity fauna with high standing stock and dominance of deep-infaunal taxa, which are adapted to dysoxic conditions (Jorissen et al., 1995). Laboratory experiments revealed a diverse metabolic capacity of benthic foraminifera and demonstrated that various species are able to respire nitrate through denitrification in order to sustain their respiration even under anoxic conditions (Risgaard-Petersen et al., 2006; Piña-Ochoa et al., 2010).
Further ecological studies demonstrated that benthic foraminifera have specific requirements concerning the quality of organic matter as food source (Caralp, 1989; Koho et al., 2008), and some taxa respond to seasonal pulses of phytodetritus to the deep sea (Gooday, 1988; Ohga & Kitazato, 1997; Gooday & Rathburn, 1999; Heinz, Kitazato, Schmiedl, & Hemleben, 2001). Most deep-sea benthic foraminiferal taxa have excellent dispersal capacities through the dissemination of propagules (Alve & Goldstein, 2010), accounting for broad distributional ranges. Their biogeography is controlled by the combination of ocean history, such as formation of gateways, the specific evolutionary histories, and various environmental factors (Gooday & Jorissen, 2012). Specifically, zonal patterns in the diversity and species composition of Cenozoic deep-sea benthic foraminifera seem to be mainly linked to surface productivity and related food fluxes (Thomas & Gooday, 1996).
Neritic and littoral environments are more environmentally variable than deep-sea environments, because they are influenced by strong geographic and bathyal gradients in light, temperature, salinity, pH, substrate, and current strengths (Culver, Woo, Oertel, & Buzas, 1996; Sen Gupta 1999; Murray, 2006). These parameters are commonly associated with water depth or elevation and are reflected in the vertical zonation of benthic foraminifera (e.g., Scott & Medioli, 1978; Milker et al., 2009) and ostracods (Cronin, 2015) in shallow-marine and coastal environments. Accordingly, these microfossil groups prove useful in sea-level reconstructions at various timescales, although alteration of fossil assemblages by taphonomic processes has to be considered (Murray & Alve, 1999; Berkeley, Perry, Smithers, Horton, & Taylor, 2007). Coastal ecosystems are susceptible to anthropogenic impacts, such as pollution and eutrophication-induced hypoxia, and benthic foraminifera, ostracods, and dinoflagellates prove useful as biomonitoring tools (Thomas, Gapotchenko, Varekamp, Mecray, & Buchholtz ten Brink, 2000; Ruiz et al., 2005; Gooday et al., 2009; Zonneveld et al., 2012; Alve et al., 2016).
Stable Isotope Records and Changes in Climate and Ocean Circulation
Stable isotope analyses have been carried out on a variety of calcareous, siliceous, and organic microfossils. In this context, the stable oxygen and carbon isotope signatures of foraminifera, expressed in δ18O and δ13C, are widely used in paleoceanography and paleoclimatology (Ravelo & Hillaire-Marcel, 2007). The geochemical composition of foraminiferal test calcite basically reflects various environmental factors during calcification, although the reasons for species-specific controls on isotope fractionation during biomineralization, so-called “vital effects,” are still not fully understood (e.g., De Nooijer et al., 2014).
The foraminiferal δ18O signal primarily reflects the combined influences of ice volume, temperature, and salinity (summaries in Rohling & Cooke, 1999; Pearson, 2012). Accordingly, δ18O compilations document the hyper-thermals of the late Cretaceous and early Paleogene and subsequent Antarctic and Arctic glaciation histories in great detail (Zachos et al., 2001; Zachos, Dickens, & Zeebe, 2008; Friedrich et al., 2012; Holbourn et al., 2018) (Fig. 9).
Cultivation experiments revealed a strong fractionation of oxygen isotopes during calcification under the influence of variable temperature at constant isotopic composition of water (Epstein, Buchsbaum, Lowenstam, & Urey, 1953). This observation was used to establish equations for temperature reconstructions (summary in Bemis, Spero, Bijma, & Lea, 1998). Foraminiferal δ18O records contain significant variability in the orbital bands of eccentricity (100, 400 kyr), obliquity (41 kyr), and precession (19, 24 kyr), demonstrating the impact of orbital forcing on global climate and amplification of these signals within the Earth system (Imbrie et al., 1984) (Fig. 10). The characteristic pattern of stacked δ18O records is widely used to evaluate the past dynamics of ice volume, sea level, and temperature (e.g., Shackleton, 1987; Siddall et al., 2003), but also to develop age models for Pliocene and Pleistocene (Imbrie et al., 1984; Lisiecki & Raymo, 2005) and older sediment successions (Grossman, 2012). During the Quaternary, temperature and salinity varied comparatively little in deep and bottom waters, thus deep-sea benthic foraminiferal δ18O records for this time interval can be interpreted as a first approximation of continental ice volume (Waelbroeck et al., 2002).
On geological time-scales, δ13C records from marine carbonates have been widely used in stratigraphy, because the δ13C signature of dissolved inorganic carbon in the ocean reflects the portioning between organic carbon and carbonate and is therefore directly linked to the global carbon cycle and the terrestrial and marine biosphere (Saltzman & Thomas, 2012).
The δ13C signal of planktic foraminifera primarily reveals information on the air–sea exchange of CO2, the photosynthesis–remineralisation cycle, and stratification of the surface water (Rohling & Cooke, 1999). Among other factors, the δ13C offsets between different taxa yield insights into species-specific vital effects, symbiont activity (Spero, Lerche, & Williams, 1991), and calcification depths, which characterizes certain depth habitats in the mixed surface layer and thermocline (Mulitza et al., 1999). In marginal basins, δ13C-based habitat reconstructions of planktic foraminifera may be biased by input of river run-off leading to decreased δ13C values in epipelagic taxa and potential habitat changes (Rohling et al., 2004).
The δ13C record of deep-sea epibenthic foraminifera is widely used for the reconstruction of changes in intermediate and deep-water circulation (e.g., Duplessy et al., 1988; Pahnke & Zahn, 2005; Mackensen, 2008). This application is based on the microbial decay of organic matter in the water column, which releases 12C and results in decreasing δ13C values of dissolved inorganic carbon in the water mass while spreading in the ocean (Charles & Fairbanks, 1992). Epifaunal and infaunal foraminifera reveal specific δ13C offsets, which can be related to metabolic and porewater effects (Grossman, 1987; McCorkle, Keigwin, Corliss, & Emerson, 1990).
The combined analysis of the δ13C signals from taxa with different microhabitat preferences (specifically, epifaunal and infaunal) retraces the porewater δ13C gradient in the sediment, which depends on the organic matter flux rate and bottom water oxygen content (McCorkle & Emerson, 1988). Accordingly, the evaluation of multispecies δ13C records opens applications for a variety of paleoceanographic reconstructions, such as changes in deep-water oxygenation (e.g., Schmiedl & Mackensen, 2006; Hoogakker et al., 2015, 2018) and surface water productivity (e.g., Zahn et al., 1986; Schilman, Almogi-Labin, Bar-Matthews, & Luz, 2003).
Applications of Foraminifera in Climate Science
Quantitative Reconstruction of Marine Environmental Parameters
One of the prime challenges of marine micropaleontology is the delivery of quantitative information on marine environmental parameters and processes in the past, such as changes in sea-level, sea-surface temperature and salinity, oxygen content and organic matter fluxes, pH, and current velocities. Such data do not only enhance the accuracy of paleoclimate reconstructions but can be also used for the validation of results from earth system model experiments. The discussion will now address a selection of current aspects of marine micropaleontology with a particular focus on foraminifera-based research.
Quantitative Reconstruction of Relative Sea-Level Change
Eustatic sea-level changes are ultimately linked to climate variability and have shaped the morphology, sediment facies, and ecosystems of continental shelves and coastal areas on various timescales. During the early 21st century, quantitative reconstructions of past global sea-level changes have been refined utilizing the δ18O records of deep-sea benthic foraminifera from the open ocean (e.g., Waelbroeck et al., 2002), but also the δ18O records of planktic foraminifera from marginal basins such as the Red Sea and Mediterranean Sea (Siddall et al., 2003; Rohling et al., 2014; Grant et al., 2014). The marginal basins respond to sea-level changes by reduced exchange with the open ocean through narrow gateways, leading to amplified changes in surface water salinity, which in turn affect the foraminiferal δ18O values. The generated time series yielded insights into the magnitude of glacial low-stands, contrasts between different interglacial high-stands, and even subtle sea-level changes during millennial-scale climate variability (Grant et al., 2014).
Quantitative sea-level estimates can also be obtained from the proportion of planktic foraminifera in the total foraminiferal fauna (see also “Plankton Evolution, Biostratigraphy, and Ecology”). This simple approach was modified by removal of the proportion of infaunal benthic foraminifera, which can vary independently from water depth and instead depend on local organic matter fluxes and oxygen content (Van der Zwaan et al., 1990; Van Hinsbergen et al., 2005). The obtained exponential function fits data sets from various oceans (Van der Zwaan et al., 1990) and proves widely applicable, but uncertainties remain relatively high (Fig. 11).
Much of the existing knowledge on late Holocene sea-level change and its coastal impacts is gained from the application of microfossil-based transfer functions on benthic foraminifera, ostracods, benthic diatoms, and testate amoebae from salt marshes (e.g., Scott & Medioli, 1978; Scott, Medioli, & Schafer, 2001; Kemp & Telford, 2015). The transfer function relates the usually unimodal distribution patterns of different taxa in a training data set to the desired environmental parameter (e.g., elevation, water depth, etc.) using regression methods such as Partial Least Squares (PLS), Weighted Averaging (WA), or the combination of both (WA-PLS) (Fig. 12). Application to well-preserved fossil assemblages from sediment cores delivered a wealth of accurate sea-level estimates, which extended the historical tide gauge records significantly back in time (e.g., Nydick, Bidwell, Thomas, & Varekamp, 1995; Horton & Edwards, 2006; Kemp et al., 2011; Zong & Sawai, 2015; Kemp et al., 2017). As one of the main results, the relative sea-level records confirm accelerated sea-level rise since the late 19th century.
Similar statistical approaches have been extended to shelf environments and used for the reconstruction of Holocene sea-level changes in the Mediterranean Sea (Rossi & Horton, 2009; Milker, Schmiedl, & Betzler, 2011). The prediction errors of the established transfer functions were in the order of 5% to 10% of the water depth range considered in the training data set. The reconstructed Holocene sea-level histories for shelf environments in the western Mediterranean Sea match independent reconstructions confirming the good performance of shelf foraminifera-based transfer functions (Fig. 13). Also, transfer functions for water-depth estimates were applied to a range of other geological problems, including the estimation of vertical movements in the course of prehistoric megathrust earthquakes on the Pacific east coast (Milker et al., 2016) and quantification of neotectonic processes in the eastern Mediterranean (Milker et al., 2017).
Quantitative Reconstruction of Surface-Water Temperature and Salinity
Sea-surface temperature (SST) and sea-surface salinity (SSS) represent essential parameters for the understanding of past climate and ocean circulation changes. In the 1970s and 1980s, the first quantitative reconstruction of global SST distribution for the Last Glacial Maximum (LGM) was realized based on regression analyses of planktic microfossils in the framework of the joint international program Climate: Long Range Investigation, Mapping, and Prediction (e.g., CLIMAP project members, 1976). Surprisingly, the CLIMAP results suggested substantial regional contrasts in the glacial cooling pattern (Mix, Bard, & Schneider, 2001). Improved training data sets and transfer functions led to amended reconstructions for the Atlantic Ocean in the project Glacial Atlantic Ocean Mapping (GLAMAP; Sarnthein et al., 2003). The accuracy of SST reconstructions was further improved by an interdisciplinary study which combined data from various plankton groups (diatoms, radiolarians, dinoflagellates, and planktic foraminifera) and geochemical proxies in the project Multiproxy Approach for the Reconstruction of the Glacial Ocean Surface (Kucera et al., 2006; MARGO project members, 2009). Based on these efforts, the magnitude of changes is well constrained, suggesting glacial SSTs which were on average 4°C lower than today, and amplified glacial cooling and associated faunal shifts at high latitudes. These results are essentially in agreement with modeling studies, but substantial uncertainties remain in tropical areas and on regional scales (Annan & Hargreaves, 2015).
Microfossil-based temperature reconstructions are increasingly complemented by the application of sophisticated geochemical proxies, biomarkers (e.g., alkenone UK’37, TEX 86), and the Mg/Ca value of foraminiferal test calcite (e.g., Rosell-Melé et al., 2004; Barker, Cacho, Benway, & Tachikawa, 2005; de Vernal et al., 2006; Wade et al., 2012). The Mg/Ca proxy is based on the observation that the Mg content in foraminifera test calcite increases proportionally to the calcification temperature (Lea, 2014). Various temperature calibrations have been established for a number of planktic and benthic foraminifera (for an example see Fig. 14). Since Ca and Mg have comparatively long oceanic residence times, late Quaternary Mg/Ca changes can be directly related to temperature changes, although accurate temperature estimates need to consider potential alteration by dissolution (Rosenthal & Lohmann, 2002). Over longer timescales (>1 Ma), the reconstruction of absolute temperatures has to consider changes in seawater Mg/Ca (Lear et al., 2015).
Changes in sea-surface salinity (SSS) are traditionally estimated by the combination of planktic foraminiferal δ18O and independent temperature proxies (e.g., Sarnthein et al., 2004), although the relations between residual δ18O values and SSS contain considerable regional uncertainties (Rohling, 2000; Ravelo & Hillaire-Marcel, 2007). Independent SSS estimates can be retrieved from the process length of specific dinoflagellate cysts (Verleye et al., 2012; Mertens et al., 2012). However, these studies suggested that cyst morphology responds to surface water density as a function of SSS and SST rather than to salinity alone. So far the only reliable independent salinity proxy was proposed by Bollmann, Herrle, Cortés, and Fielding (2009), who found a significant linear correlation between the size of placoliths of the cosmopolitan coccolithophorid species Emiliania huxleyi from plankton samples and in-situ SSS. This function was successfully applied to the reconstruction of the early Holocene water-mass exchange between the Aegean and Black seas (Herrle et al., 2018).
Quantitative Reconstruction of Surface Primary and Export Productivity, Organic Matter Fluxes, and Oxygen
Surface-water productivity, organic matter fluxes, and oxygen concentrations characterize the marine organic carbon pump and are commonly linked to each other, complicating their separation (see “Benthic-Pelagic Coupling and Ecology of Benthic Foraminifera”). Qualitative information on export production can be derived from diatom and radiolarian fluxes and the stable Si isotope composition of their skeletal opal (e.g., Crosta & Koç, 2007; Abelmann et al., 2015). Export productivity has been quantified on the basis of the accumulation rate of benthic foraminifera in sediments from the western equatorial Pacific Ocean (Herguera & Berger, 1991). However, the transfer of this relation to other regions failed, suggesting a non-linear response of the foraminiferal number to organic matter fluxes and the interference of other parameters, such as oxygen levels (Schmiedl & Mackensen, 1997; Naidu & Malmgren, 1995). Also, minor changes in the transfer efficiency of exported organic matter (e.g., due to temperature changes and related changes in metabolic rates) could significantly change the food fluxes at the seafloor (Laws, Falkowski, Smith, Ducklow, & McCarthy, 2000; John et al., 2013).
A number of semi-quantitative to quantitative approaches have been developed on the benthic foraminiferal fauna, either based on species-specific flux regimes (Schönfeld & Altenbach, 2005) or on multivariate statistics (e.g., Kuhnt, Hess, & Jian, 1999; Wollenburg, Kuhnt, & Mackensen, 2001), but none of these approaches proved widely applicable as a straightforward quantitative paleoproductivity proxy (summary in Jorissen et al., 2007).
The δ13C difference between shallow infaunal and epifaunal benthic foraminifera varies proportional to the organic matter flux rate (e.g., Zahn et al., 1986; McCorkle et al., 1990; Schilman et al., 2003). Accordingly, Theodor, Schmiedl, Jorissen, and Mackensen (2016) used the δ13C signals of epifaunal taxa and the shallow infaunal Uvigerina mediterranea to develop a transfer function for organic carbon flux rate in the Mediterranean Sea. A comprehensive testing of this transfer function is still missing, but its applicability is likely restricted to open-ocean settings since isotope data from marginal settings with substantial lateral organic carbon fluxes lead to an overestimation of vertical fluxes and related surface water productivity (Theodor et al., 2016).
Oxygen exerts a strong control on deep-sea benthic diversity and microhabitat partitioning, facilitating the development of oxygen proxies on the basis of the benthic foraminiferal fauna (summary in Jorissen et al., 2007). Semi-quantitative oxygen indices use the ratio of various morphological groups (Kaiho, 1994), or a combination of the ratio between high- and low-oxygen tolerant taxa, and faunal diversity (Schmiedl et al., 2003). The latter approach was applied to the characterization of deep-water oxygen changes across the early Holocene sapropel S1 in the Mediterranean Sea (Schmiedl et al., 2010), illustrating the response of deep-water formation to low- and high-latitude climate forcing (see “Resilience and Recovery Potential of Marine Ecosystems with Respect to Perturbations”).
Field and laboratory studies demonstrated that various benthic foraminifera taxa increase their pore density and size in response to reduced oxygen and enhanced nitrate concentrations in the bottom water (Perez-Cruz and Machain-Castillo, 1990; Moodley & Hess, 1992; Glock et al., 2011). This morphological adaptation is likely related to the respiration of the foraminifera as demonstrated by the clustering of mitochondria in the vicinity of the pores (Leutenegger & Hansen, 1979). In the meantime, transfer functions exist for a variety of taxa and regions (Glock et al., 2011; Kuhnt et al., 2013; Rathburn, Willingham, Ziebis, Burkett, & Corliss, 2018) and allowed estimating changes in the late glacial and Holocene nitrogen inventory of the Peruvian upwelling region (Glock et al., 2018).
Changes in bottom water oxygen concentrations are also recorded in the isotope and trace element geochemistry of benthic foraminiferal test calcite. The δ13C of dissolved inorganic carbon in the pore water at sediment depth where oxygen approaches zero is directly related to the oxygen concentration of the bottom water (McCorkle & Emerson, 1988) (Fig. 15 left). Accordingly, the δ13C difference between epifaunal benthic foraminifera, such as Cibicidoides wuellerstorfi, and deep infaunal taxa, such as Globobulimina affinis, can be used as proxy for bottom water oxygen (McCorkle et al., 1990). Two calibration data sets were generated on the basis of modern foraminifera and successfully applied to late Quaternary successions from the Arabian Sea, North Atlantic Ocean, and the equatorial Pacific Ocean (Schmiedl & Mackensen, 2006; Hoogakker et al., 2015, 2018). The applicability of this proxy is restricted to oxygen concentrations below 235 µmol kg−1 and lacks a clear relation above this threshold (Fig. 15 right).
In the 21st century, analytical progress, such as the application of secondary ion mass-spectrometry, fostered the development of novel geochemical proxies for the redox state of ambient water. Field studies demonstrated that, in the absence of diagenetic alteration, Mn/Ca and I/Ca ratios of foraminiferal test calcite are highly redox-sensitive and increase proportionally to oxygen concentration (e.g., Glock et al., 2012; Glock, Liebetrau, Eisenhauer, & Rocholl, 2016). Accordingly, the I/Ca ratios in planktic foraminifera from a sediment core of the Southern Ocean were analyzed and used to document glacial decrease in dissolved oxygen concentration in the near-surface ocean (Lu et al., 2016). Similarly, planktic foraminiferal I/Ca gradients indicate lateral expansion of oxygen minimum zones in the Atlantic, Indian and Pacific oceans during the Paleocene–Eocene thermal maximum (Zhou, Thomas, Rockaby, Winguth, & Lu, 2014).
Quantitative Reconstruction of Bottom Current Strength
Bottom currents shape benthic ecosystems because they raise the energy at the benthic boundary layer, modify substrate at the sea floor, and transport suspended food particles. Bottom currents are particularly relevant in shallow-water ecosystems, on submarine elevations, and in areas influenced by oceanic gateways. In the South Atlantic Ocean, benthic foraminiferal faunas with a dominance of Angulogerina angulosa were assigned to sandy substrates in high-energy environments on submarine highs, the shelf edge, and upper slope (Mackensen, Schmiedl, Harloff, & Giese, 1995). The first quantitative relation between near-bottom current velocities and abundance of elevated epifaunal benthic foraminifera was established for the pathway of the Mediterranean Outflow Water (MOW) undercurrent in the Gulf of Cadiz, northeastern Atlantic Ocean (Schönfeld, 2002). The applicability of this function to the reconstruction of changes in MOW strength was evaluated based on early Pliocene benthic foraminiferal faunas from a sediment core of the Gulf of Cadiz (García-Gallardo, Grunert, Voelker, Mendes, & Piller, 2017). The reliability of this proxy is biased by the downslope transport of epifaunal foraminifera from the shelf, but appears a suitable indicator for current velocity after removal of allochthonous tests.
Response of Marine Ecosystems to Climate Changes and Abrupt Perturbations
As outlined in “Quantitative Reconstruction of Marine Environmental Parameters,” marine microfossils provide a variety of proxies for quantitative environmental reconstructions. Marine microfossils also prove useful for the evaluation of ecological responses to both long-term and abrupt climate change. Such information is relevant in order to assess the resilience and recovery potential of marine ecosystems with respect to past and future climate perturbations.
Response of Marine Ecosystems to Orbital Climate Changes
The excellent applicability of marine protists in quantitative environmental reconstructions is based on their immediate response to climate forcing on various timescales. Some prominent examples are documented for the marine ecological impacts of glacial–interglacial climate variability during the late Quaternary.
Planktic microfossils document the impact of glacial–interglacial climate changes on sea surface temperature, sea-ice cover and other environmental parameters of the surface ocean. Joint programs dedicated to the reconstruction of the last glacial ocean (see “Quantitative Reconstruction of Surface-Water Temperature and Salinity”) documented changes in the zonal distribution of certain plankton groups. For example, the distribution area of transitional and subpolar planktic foraminifera in the northern North Atlantic and Arctic oceans shifted further south during the last glacial maximum (LGM) (Kucera et al., 2005; Kucera, 2007) (Fig. 16). At high northern latitudes, shifts in the zonation of plankton associations reveal immediate responses to both orbital and suborbital climate variability, which retrace the northward inflow of warm surface waters (e.g., Kandiano, Bauch, & Müller, 2004; Barker et al., 2015). Similarly, the distribution of siliceous microfossils in the Southern Ocean responded to orbital changes in sea-ice extent, which was displaced northward by 7–10° during the LGM (e.g., Gersonde et al., 2005; Studer et al., 2015). Shifts of the austral frontal systems during the past five glacial–interglacial cycles also affected the intensity of the Agulhas Current as documented by changes in the abundance of subtropical planktic foraminiferal species in a sediment core off the Cape of Good Hope (Peeters et al., 2004).
At low latitudes and in upwelling areas, the orbital-scale variability of planktic ecosystems is linked to changes in nutrient availability and surface productivity. The temporal variability is strongly coherent on the obliquity and precession bands because the position of the Intertropical Convergence Zone, and the intensity of the monsoon circulation, exhibit substantial seasonal variations (e.g., Clemens, Prell, Murray, Shimmield, & Weedon, 1991; Beaufort et al., 1997). Tropical plankton communities also mirror changes in surface-water stratification and thermocline depth, and respond to climate oscillations, such as the El Niño Southern Oscillation (e.g., Beaufort, de Garidel-Thoron, Mix & Pisias, 2001; Wara, Ravelo, & Delaney, 2005).
Deep-sea ecosystems reflect orbital-scale climate changes through the immediate processes of benthic-pelagic coupling, including the export of organic matter from the photic zone (Fig. 7). In addition, the deep-sea is ventilated by the advection of intermediate, deep, or bottom-water masses, which may result in a lagged response of the deep-sea benthos to climate forcing, depending on the residence time of the water mass. A good example for the interaction between organic matter fluxes and ventilation rate is the deep Arabian Sea. The relative abundances of oxygen-tolerant infaunal benthic foraminifera and faunal diversity reveal periodic changes in the deepening of the oxygen minimum zone (Fig. 17). Reconstructed oxygen values vary between approximately 50 and 105 µmol kg-1, applying the δ13C-based transfer function of Hoogakker et al. (2015) (Fig. 15). The estimated changes in oxygen concentration lag the coherent changes in SW monsoon strength and related organic matter fluxes by several thousand years (Fig. 17). The data suggest that the deep-sea benthic ecosystems of the Arabian Sea are forced by the combined influence of regional organic matter fluxes and the entrainment of oxygen-enriched deep-water from the Atlantic Ocean (Schmiedl & Leuschner, 2005; Schmiedl & Mackensen, 2006).
Resilience and Recovery Potential of Marine Ecosystems With Respect to Perturbations
The ecological impact of past ocean perturbations can provide valuable information for the assessment of marine ecosystem response to future anthropogenic changes. Relevant examples include the mass extinction at the Cretaceous–Paleogene boundary (KPg boundary), the ecosystem crisis during the carbon cycle disturbance at the Paleocene–Eocene thermal maximum (PETM), and the deep-water anoxia during Neogene sapropel formation in the Mediterranean Sea.
The mass extinction at the KPg boundary, around 66 million years ago, has been associated with the impact of a large asteroid on the Yucatan carbonate platform in the southern Gulf of Mexico (Alvarez, Alvarez, Asaro, & Michel, 1980; Hildebrand et al., 1991). Approximately 76% of all species became extinct globally, of which the marine planktic ecosystems were most severely affected (Schulte et al., 2010). The role of Deccan volcanism in the mass extinction is highly disputed. Enhanced volcanic CO2 emission before and during the KPg event may have contributed to ocean acidification and stress for marine calcifiers (Punekar et al., 2016), but the resulting climate effects were probably only moderate (Schulte et al., 2010), and the environmental impacts cannot account for the observed extinction patterns of planktic foraminifera (Molina, 2015).
The breakdown of stable carbon isotope gradients between surface ocean and deep-sea of ~500,000 years duration (Zachos, Arthur, & Dean, 1989) was associated with a global collapse of pelagic marine primary productivity (“Strangelove” Ocean; Hsü & McKenzie, 1985) or export productivity (“Living” Ocean; D’Hondt, Donaghay, Zachos, Luttenberg, & Lindinger, 1998). However, benthic foraminifera were but slightly affected by the mass extinction, suggesting regional and only moderate decrease in export productivity (Thomas, 2007; Culver, 2003; Alegret et al., 2012). The interpretation of sustained export productivity across the KPg event is supported by biomarker data suggesting only a short decline of eukaryotic algal and continuation of cyanobacterial primary productivity (Sepúlveda, Wendler, Summons, & Hinrichs, 2009).
In the Chixculub crater basin first life returned within years and a productive ecosystem re-established within 30,000 years after the impact, implying a high recovery potential of planktic communities (Lowery et al., 2018). The evolutionary recovery of planktic foraminifera peaked a few million years after the KPg boundary, concurrent to the full recovery of the marine carbon cycle (Coxall et al., 2006) and the evolution of foraminiferal photosymbiosis around 63.5 million years ago (Birch, Coxall, & Pearson, 2012).
The marine ecological crisis of the PETM, around 56 million years ago, was associated with a negative carbon isotope excursion, which was likely caused by rapid emission of a large volume of greenhouse gasses resulting in a transient temperature increase of 5–8°C (Zachos et al., 2003; McInerney & Wing, 2011). The period of carbon release has likely lasted for less than 20,000 years and the duration of the whole PETM is estimated to around 200,000 years (McInerney & Wing, 2011; Zeebe, Dickens, Ridgwell, Sluijs, & Thomas, 2014). The carbon sources remain controversial, and may have included the dissociation of methane hydrates (Dickens, O’Neil, Rea, & Owen, 1995), volcanic carbon from the North Atlantic Igneous Province (Gutjahr et al., 2017), or both.
Deep-sea benthic foraminifera experienced a massive extinction, concerning 30–50% of all species during a few thousand years (Thomas & Shackleton, 1996; Thomas, 2007). By contrast, planktic organisms, including dinoflagellates, calcareous nannofossils, and planktic foraminifera, exhibit rapid evolutionary turnover, distributional range shifts, and species-specific growth response, but lack major extinctions (Speijer, Scheibner, Stassen, & Morsi, 2012; Self-Trail, Powars, Watkins, & Wandless, 2012; Gibbs et al., 2006, 2013). The mass extinction of deep-sea benthic foraminifera and its biogeographic pattern is complex, and has been attributed to the combined effects of ocean warming, deep-water circulation changes, ocean acidification, oxygen depletion, and reduced food supply (Thomas, 1998; Winguth, Thomas, & Winguth, 2012). This combination was confirmed by the dwarfing of some surviving benthic foraminiferal taxa at deeper sites (Schmidt et al., 2018), implying bathymetric gradients in the resilience of deep-sea benthic ecosystems, depending on the magnitude of perturbation.
The marine ecosystems of marginal basins, such as the Mediterranean Sea and the Red Sea, react very sensitively to global and regional climate changes and have experienced substantial regime shifts in their marine ecosystems during the past (Hemleben et al., 1996; Rohling, Marino, & Grant, 2015). Basin-wide compilations of deep-sea benthic foraminiferal oxygen index values and epibenthic δ13C data delivered a detailed history of deep-water stagnation and reventilation across the past 25,000 years of the eastern Mediterranean Sea, including the last glacial termination and the early Holocene sapropel S1 interval (Schmiedl et al., 2010; Grimm et al., 2015) (Fig. 18). The observed rapid deep-sea benthic ecosystem collapse at the onset of sapropel deposition reflects a lagged response to the insolation-driven intensification of the African monsoon system, and associated hydrological changes. Abrupt high-latitude hydrological perturbations and associated cooling events are superimposed on the long-term evolution, which is highlighted by a transient reventilation of benthic ecosystems during the 8.2 ka cold event (Rohling, Jorissen, & de Stigter, 1997; Schmiedl et al., 2010) (Fig. 18). Under the oligotrophic boundary conditions of the late Holocene Mediterranean Sea, the recovery of deep-sea faunas strongly depended on the duration of the anoxic phase. While deep-sea ecosystems exhibited a rapid recolonization by opportunistic taxa (Jorissen, 1999), the full recovery of abyssal benthic ecosystems under the influence of ultra-oligotrophic conditions may have taken up to several millennia (Schmiedl, Hemleben, Keller, & Segl, 1998).
The contemporary rapid global warming affects the biogeography of marine protists as reflected by latitudinal shifts of distribution belts and colonization of new marine ecosystems by temperature-sensitive taxa (e.g., Weinmann et al., 2013; Schmidt et al., 2015). Enhanced greenhouse gas emissions and the associated temperature rise impose thermal stress and acidification of surface waters, particularly affecting the life of calcifying organisms. Culture experiments suggest that some coccolithophorids respond to ocean acidification by reduced growth and calcification rates, but other species or strains seem to be able to maintain their survival and functionality by rapid adaptive evolution (e.g., Langer, Nehrke, Probert, Ly, & Ziveri, 2009; Lohbeck, Riebesell, & Reusch, 2012; Schlüter et al., 2014). Similarly, shallow-water benthic foraminifera exhibit specific tolerance levels in terms of acidification and thermal stress (Haynert, Schönfeld, Schiebel, Wilson, & Thomsen, 2014; Schmidt et al., 2016). In the geological record, phases of ocean acidification were commonly associated with extinction and evolutionary turnover of marine calcifying organisms (Hönisch et al., 2012). The rapidity of ongoing anthropogenic warming and CO2 emission rates are probably unprecedented during the past 66 million years (Zeebe, Ridgwell, & Zachos, 2016) but the abrupt climate perturbation, acidification, and ecological and evolutionary responses at the KPg boundary event may probably serve as an analogue for the anticipated future changes.
Marine micropaleontology investigates the diversity, biostratigraphy, ecology, and geochemistry of planktic and benthic microfossil groups. The paleo-environmental applications of marine microfossils are manifold, and deliver a wealth of information on past ocean circulation and climate, and evolution of oceanic biota. Specifically, microfossil-based proxies have been developed for the quantitative reconstruction of past changes in water depth, sea-surface temperature and salinity, surface productivity and organic matter fluxes, oxygen concentration, and current strength. The majority of these proxies use transfer functions, which are based on modern training data sets and a variety of statistical methods.
Contemporary micropaleontological research is challenged by technical innovations, such as the availability of sophisticated analytical techniques, which are used for the establishment of novel geochemical proxies. For many proxies the use of single taxon material is essential, which requires a careful taxonomy of the selected specimens.
Regardless the analytical progress, a well-grounded understanding of plankton and benthos systematics and ecology forms the basis for micropaleontological research involving appropriate biological and ecological field studies, laboratory experiments, and compilation of existing data. Microfossil-based paleoclimate research creates added value in the frame of interdisciplinary research, for example through combining proxy studies with experiments from earth system models. Last but not least, marine micropaleontological research is invoked to address important societal future challenges, such as pollution monitoring or assessment of coastal ecosystem resilience. In this context, the profound understanding of past ecosystem dynamics appears invaluable to assess the future impacts of global climate change and biodiversity loss.
The author thanks Dania Achermann and Simone Rödder for organizing the workshop Towards a History of Paleoclimatology: Changing Roles and Shifting Scales in Climate Research at Hamburg University, September 2017, which motivated this contribution. The author is grateful to the members of the micropaleontology group at the Institute for Geology for discussion, and to Silke Schmiedl for constructive comments on the manuscript. The author greatly appreciates the thorough comments of an anonymous reviewer, which helped to improve the manuscript.
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