Marine Ecosystems E-Book

• Herausgeber: John Wiley & Sons
• Kategorie: Wissenschaft und neue Technologien
• Sprache: Englisch

Seitenzahl: 512

Veröffentlichungsjahr: 2015

Beschreibung

In a perspective of sustainable management, the balance between ecological dynamics, social and economic are now at the heart of ecological modeling and environmental strategies screenwriting. Diversity and marine ecosystems function illustrates biodiversity, habitat diversity, structures and food webs in various oceans of the world and systems: pelagic and benthic ecosystems, coral reefs and seagrass beds, oasis of hydrothermal vents ridges or areas rich upwelling. Appropriate observation methods, long-term monitoring and modeling reveal the complexity of systems, trophic interactions and spatiotemporal dynamics. The ecosystem approach is a prerequisite to assess the state of these systems, their living resources and ecological services involved in local and global environmental changes.

Leseprobe

Table of Contents

Cover

Title

Copyright

Foreword

1: Marine Biosphere, Carbonate Systems and the Carbon Cycle

1.1. Introduction

1.2. Marine organisms and carbon

1.3. Variability in the production of organic matter

1.4. From the biosphere to the atmosphere to climate

1.5. Carbonate production

1.6. The coupling of carbonaceous and organic productions

1.7. Modification of equilibria and consequences on marine life

1.8. Conclusion

1.9. Bibliography

2: Biodiversity of Phytoplankton: Responses to Environmental Changes in Coastal Zones

2.1. Introduction

2.2. Phytoplankton ecology

2.3. Phytoplankton responses to anthropogenic pressures

2.4. Observation systems for the identification of phytoplankton

2.5. Conclusion

2.6. Bibliography

3: Marine Seagrasses (Magnoliophyta) in the Intertropical Zone

3.1. From plant to habitat

3.2. Role of seagrass beds in the coastal environment

3.3. Functioning of seagrass beds

3.4. Challenges in the conservation of seagrass beds

3.5. Pressures on and threats to seagrasses

3.6. Restoration of seagrass beds

3.7. The functional role of seagrasses in the lagoon ecosystem

3.8. Conclusion

3.9. Bibliography

4: Biocomplexity of Coral Ecosystems: Diversity in All its States

4.1. Introduction

4.2. Diversity in the coral world

4.3. Links between diversities

4.4. Conclusion

4.5. Bibliography

5: Man and Diversity in the Coral Environment

5.1. Introduction

5.2.

Diversity and ecological services

5.3. Local versus global threats: what are local threats and what are their effects?

5.4. What are the combined effects of local and global threats on corals?

5.5. Functions and diversity

5.6. Conclusion

5.7. Bibliography

6: Hydrothermal Vents: Oases at Depth

6.1. Introduction to deep-sea ecosystems

6.2. Discovery of hydrothermal sources

6.3. Geology and geochemistry of hydrothermal systems

6.4. Microbial chemosynthesis

6.5. Symbioses and trophic chains

6.6. Distribution of fauna at different spatial scales

6.7. Faunal microdistribution and interactions

6.8. Temporal dynamics of hydrothermal ecosystems

6.9. Mineral resources and exploitation

6.10. Bibliography

List of Authors

Index

End User License Agreement

List of Tables

2: Biodiversity of Phytoplankton: Responses to Environmental Changes in Coastal Zones

Table 2.1.

Systematics of several groups within the marine phytoplankton. The classification is based on the taxonomic database WoRMS [WOR 14]

Table 2.2.

Phytoplankton classification according to different size fractions

Table 2.3.

Examples of studies on temporal variations of biomass and the abundance of phytoplankton

,

based on observations in different marine regions

Table 2.4.

Examples of studies on the phenological variations in phytoplankton. The results were obtained using data from observations on land

,

satellites and mesocosm experiments

Table 2.5.

Main observation networks of phytoplankton and the network characteristics

Table 2.6.

Examples of analyses applied to time series (TS) in ecology

List of Illustrations

1: Marine Biosphere, Carbonate Systems and the Carbon Cycle

Figure 1.1.

Parameters of the Earth’s orbit [LAS 04] molding the global climate (CO

2

concentration in ice core d), oxygen stable isotopes in marine sediments e), precession b) and obliquity c) influence the seasonal contrasts, eccentricity a) modulates (amplifies or not) precession effects (Beaufort, compilation of [LAS 04, LUE 08, MAR 87])

Figure 1.2.

Example of the evolution of organic production in the tropical zone over 150,000 years

Figure 1.3.

Photographs taken using a scanning electron microscope of several representatives of the two main groups of pelagic calcifying organisms

Figure 1.4.

Images of the deep sea in the Southern islands at a depth of 2,500 m (images from the oceanographic campaign SO2O4 Pacenpal by Beaufort, 2010)

Figure 1.5.

Correlation between mass and calcified plates of coccolithophores (coccoliths) and the concentration of dissolved carbonate in seawater, in different regions of the current and past oceans (0 – 40,000 years ago) [BEA 11]

Figure 1.6.

Two images taken using a scanning electron microscope of the coccolithophore Emiliania huxleyi. Left: collected in April 2014 offshore of Pisco (Peru) in water with a pH of 7.66. Right: this coccolithophore was collected in February 2009 between Cap Horn and Antarctica in water with a pH of 8.05

2: Biodiversity of Phytoplankton: Responses to Environmental Changes in Coastal Zones

Figure 2.1.

Representative species of certain groups within the marine phytoplankton. a) Dictyocha (Dictyochophyceae)

,

b) Eutreptiella (Euglenoidea)

,

c) Scenedesmus (Chlorophyceae)

,

d) Phaeocystis (colonial form

,

Prymnesiophyceae)

,

e) Chattonella (Raphidophyceae)

,

f) Prorocentrum (Dinophyceae) and g) Guinardia (chain

,

Bacillariophyceae) (photo credits a)

,

b)

,

c)

,

e)

,

f) and g) Ifremer LER Pertuis Charentais

,

d) Robert Andersen)

Figure 2.2.

Different size scales of phytoplankton relative to macroscopic organisms and objects (from [FIN 09])

Figure 2.3.

Diagram showing the physical and biological factors that influence the phytoplankton

Figure 2.4.

Representation of seasonal cycles of phytoplankton (continuous line) and herbivores (dotted line) at different latitudes according to Cushing [CUS 96]

Figure 2.5.

Models of phytoplankton organization

Figure 2.6.

Representation of the CSR organization with regard to turbulence gradients and nutrient concentration. The loops represent temporal successions over one year in a) oligotrophic lakes

,

b) eutrophic lakes

,

c) and d) small and rich systems and e) autogenic succession (figure according to Reynolds [REY 06]])

Figure 2.7.

Typology of phytoplankton traits according to Litchman and Klausmeier [LIT 08])

Figure 2.8.

Schematic representation of an ecological niche in a two-dimensional environment. The grey zone represents the space where growth (r) is positive (from Hirzel and Le Lay [HIR 08])

Figure 2.9.

Possible responses of three species of phytoplankton to changes in their environment. Each species is represented by a shape (circle

,

triangle or rectangle) and each phenotype (and potentially a genotype) by a given color

Figure 2.10.

Conceptual diagram of responses and impacts on ecosystems under multiple stress factors. The effects these factors have on ecosystems depend on the inherent properties of systems

,

which act as filters

,

and mitigate or exacerbate these effects

Figure 2.11.

Predicted response of phytoplankton as the ocean temperature increases (according to Doney [DON 06])

Figure 2.12.

Time series of the winter mean (from January to March) of the climate index north Atlantic oscillation (NAO) and the annual mean of the index Atlantic multidecadal oscillation (AMO)

3: Marine Seagrasses (Magnoliophyta) in the Intertropical Zone

Figure 3.1.

Morphology of marine magnoliophytes – example of the species Thalassodendron ciliatum (photo C. Hily).

Figure 3.2.

Variability in the morphology of three genera of tropical magnoliophytes: a) Halophila

,

b) Syringodium

,

c) Cymodocea

Figure 3.3.

Mitigation of the effect of coastal surges: seagrass mat (Thalassia testudinum) eroded by surges (photo C. Hily)

Figure 3.4.

Seagrass beds

,

biodiversity hotspots: epibiosis under a canopy of Talassodendron ciliatum (photo C. Hily).

Figure 3.5.

Seagrasses are widely used by local populations and

,

under the impact of trampling

,

sampling and uprooting

,

weaken the biocenosis of seagrasses and the erosion of plant biomass. a) Subsistence fishing in Madagascar; b) recreational water activities in New Caledonia (photos C. Hily)

Figure 3.6.

Functional continuity between a) seagrass and mangrove and b) seagrass and coral reef (photos C. Hily)

4: Biocomplexity of Coral Ecosystems: Diversity in All its States

Figure 4.1.

Conceptual diagram of biocomplexity, adapted from [CAD 06]. vs: versus. The vertical dotted arrows indicate interaction between levels. Gray arrows indicate the actions of the factors (either global, therefore external to the system, or local, therefore internal to the system)

Figure 4.2.

Map showing the main coral reefs in the world. Source UNEP, WCMC-008. Black lines represent the areas with high freshwater input forming a barrier for corals and associated ecosystems. In black are other natural barriers, essentially linked to cold currents or upwellings. The dots represent the western limit of the oceanic barrier between the Indo-Pacific and East Pacific

Figure 4.3.

Diagram of a coral reef showing both the gradients of depth and the gradients of exposure in the spatial organization of a reef

Figure 4.4.

Cross-section of a reef system. Note that the lagoon system may be absent or not include an intermediate reef

Figure 4.5.

Diagram showing the structure and main elements of a polyp (soft part) and the polypierite (skeleton) of scleractinian reef-building corals (illustration: P. Bosserelle)

Figure 4.6.

Main growth forms of coral colonies (illustration: P. Bosserelle)

Figure 4.7.

Diagram showing the coral life cycle and their main modes of reproduction (illustration: M. Adjeroud)

Figure 4.8.

Diversity gradient of corals by number of species and genera. Data taken from the IUCN database on corals

Figure 4.9.

Proportion of families according to number of species, in the four oceans and globally

Figure 4.10.

Biogeographic regions defined by the species composition of reef fish [KUL 13]. The R numbers indicate the number of known species for the geographic area and the E numbers indicate the number of endemic species.

Figure 4.11.

Grouping of sites according to the relationship between abundance and size of geographical distribution range of reef fish species in the Pacific (unpublished analysis on 118 islands and 8,000 transects – GASPAR program of FRB)

Figure 4.12.

Species distribution of endemic reef fish according to their size and diet in the Pacific

Figure 4.13.

Model (SEM: structural equation model) representing the relative role of different factors to explain the level of reef fish diversity throughout the world [PAR 13]. The maps show, on the one hand, the distribution observed in the diversity of these fish, and on the other hand the diversity predicted by the model. The adjustment is optimal in the zones where the diversity is at its maximum; the most significant differences are observed around the periphery (south-west Indian Ocean; oceanic islands in the East Pacific).

Figure 4.14.

Deviation from a neutral model for the spatial distribution of reef fish for taxonomic composition (at the family level), trophic composition (seven feeding classes) and size (six size groups).

Figure 4.15.

Relationship between local (alpha) and regional (gamma) diversity. The dotted like indicates equality between local and regional diversity (all known species in the region are locally present). The beta diversity represents the difference between a mean local diversity and the available regional diversity (according to [LOR 00])

Figure 4.16.

Relationship between alpha diversity and sampled area for twelve islands on French Polynesia [KUL 00]

Figure 4.17.

Size and trophic structure in diversity according to atoll size (small, medium, large and very large), the two top histograms are for a local scale (500 m²) and the two bottom histograms are for the integrality of atolls. The notable values or trends are indicated by black arrows. The bar for the region is constant on both series and represents the proportions of all species in the region

Figure 4.18.

Theoretical forms of the relationship between species diversity (number of species) and functional diversity (number of functions performed by the species) (adapted from [MIC 05])

Figure 4.19.

Relationship between species and functional diversity for coral reef fish in New Caledonia (Voh-Koné-Pouembout region, Northern Province) for different functional classification diagrams, DS, DSH and DSHG(D: diet; S: size; H: Home range: G; schooling). The dotted line represents a straight line with a slope of 1 (adapted from [GUI 11])

Figure 4.20.

Relationship between species and functional diversity on a global scale. Each point corresponds to a geographic cell of 5° × 5° (data from GASPAR database (FRB))

Figure 4.21.

Relationship between species and function diversity of coral reef fish in New Caledonia [GUI 11]

Figure 4.22.

Space of functional volume occupied by coral fish species from six biogeographical provinces of the world [KUL 13] (adapted from [MOU 14]).

Figure 4.23.

Redundancy of functional groups by biogeographic region (see Figure 4.6 for the regions), organized by increasing diversity. EA: East Atlantic ; EP: East Pacific; WA: West Atlantic; WI: West Indian; CP: Central Pacific; CIP: Central Indo-Pacific. The figures on the abscissa represent the number of species per region

5: Man and Diversity in the Coral Environment

Figure 5.1.

Expected evolution in biomass of reef fish from t

0

(no fishing) to t

1

(low fishing), t

2

(moderate fishing) and t

3

(very intense fishing)

Figure 5.2.

a) Let us assume that two islands 1 and 2 are different in size. Around each island, stations of the same size are regularly sampled (small circles). b) The number of species on island 1 will be lower than on island 2 and the mean size of species on island 1 will be greater than on island 2. c) Both islands belong to the same region; the local diversity (α) observed at a station on one of these islands will be proportional to the regional diversity and number of species known on the island; the number of species observed at a station in this region cannot exceed the α

max

of the region. d) The abundance (or density) of a species is associated with size; the abundance is also associated with the number of resources (fewer on island 1 due to its smaller size), we will, therefore, observe a lower abundance for the same fish size on island 1 and a general decrease in abundance as fish size increases. e) The weight of a fish is a function of its size, according to an exponential law. f) By combining graphs B, C, D and E, the biomass on station i on one of these two islands will be a function of the alpha diversity observed; with species being larger on average on island 1 their biomass will be greater at an equal diversity compared with island 2; the maximal local diversity observed is bounded on each of these two islands (α

max1

; α

max2

)

Figure 5.3.

a) Moving from a low level of fishing to a high level of fishing causes a loss of biomass (loss of individuals removed by fishing) (1), but also a depletion of individuals, which causes a decrease in species density (2), this loss of diversity is greater when the initial diversity is low (3). These losses of diversity (2 and 3) in turn cause a decrease in resource partitioning and therefore the amount of biomass attainable, as there is a dual effect of fishing on biomass. b) On island 1 (small), the loss of biomass will be greater at an equal fishing effort, whereas the loss of diversity will be greater on island 2 (large)

Figure 5.4.

Relationship between the growth coefficient (k) and size (top graph) or log of weight (bottom graph)

Figure 5.5.

Relationships between biomass, size and biogeographical origin. Three regions, Raja Ampat (RA circle), Fiji (FJ circle) and Tuamotu (TU circle), have declined diversities of reef fish (sources: Raja Ampat: A. Green (pers. com.), Fiji and Tuamotu: IRD database)

Figure 5.6.

Relationships between size, diversity, biomass and production. Example taken from Tuamotu. a) Ten atolls were classed into two size classes (small < 100 km²; large > 100 km²), each circle represents the biomass for a given size class (abscissa) the level of the circle being given by the diversity of the size class (ordinates); the largest circle represents 50.3 g/m². b) Based on biomass-production relationships according to size, this diagram shows a higher expected production in large atolls

Figure 5.7.

Erosion of species, phylogenetic and functional diversity of parrotfish from coral reefs in 17 insular countries in the tropical Pacific (adapted from [DAG 14])

Figure 5.8.

Photos of the outer slope reef of the island of Moorea (French Polynesia), before (photo from November 2007) and after an outbreak of the corallivorous starfish Acanthaster planci (photo from October 2009), and after cyclone Oli in February 2010 (photo from May 2010) (photo credits: M. Kayal)

Figure 5.9.

Conceptual diagram illustrating the concept of ecological resilience (according to [HOL 73])

Figure 5.10.

Variation in the percentage of coral cover (across all genera) and algal cover (macroalgae and microalgal turf) on the outer slope of Tiahura (Island of Moorea, French Polynesia), impacted by five coral bleaching events, two cyclones and an outbreak of the corallivorous starfish Acanthaster planci

Figure 5.11.

Variation in percentage of coral cover of four main coral genera (which represent more than 70% of total cover) on the outer slope of Tiahura, showing a change in their relative contributions within assemblages

Figure 5.12.

World maps a) of the levels of threats (according to [HAL 08]); b) the proportion of protected habitats (source: IUCN); c) taxonomic vulnerability and d) functional vulnerability of reef fish throughout the world. The graphs to the right indicate the frequency (according to [PAR14])

Figure 5.13.

Production cycle for Scylla serrata illustrating the different stages of breeding (adapted from [QUI 10]) – see [SHE 11] for a detailed description of each stage

Figure 5.14.

Diagram showing the principle of post-larval capture and culture (PCC))

6: Hydrothermal Vents: Oases at Depth

Figure 6.1.

Cross-section of the ocean illustrating the different ecosystems of the deep sea (Capsule Graphik 2014)

Figure 6.2.

Distribution of active hydrothermal sites (in red) and sites where indices of activity have been identified (in yellow). Continuous lines represent ridges and faults whereas dotted lines represent trenches. Exclusive economic zones are indicated in pale blue (according to [BEA 10]).

Figure 6.3.

Diagram illustrating the circulation of fluids and the formation of hydrothermal vents (Capsule Graphik 2014)

Figure 6.4.

Photosynthesis and chemosynthesis: two processes producing an organic matter that uses different energy sources (Capsule Graphik 2014)

Figure 6.5.

Endosymbiosis in the siboglinid worm Riftia pachyptila (Capsule Graphik 2014)

Figure 6.6.

Simplification of the hydrothermal trophic network (Capsule Graphik 2014)

Figure 6.7.

The EMSO-Azores observatory has been in operation since 2010 on the Lucky Strike vent field (1,700 m) on the Mid-Atlantic Ridge. It was implemented by French interdisciplinary teams – Institut de Physique du Globe, University of Toulouse and Ifremer (Capsule Graphik 2014)

Figure 6.8.

Diagram showing the potential exploitation of sulfide deposits from the deep sea (Capsule Graphik 2014)

Figure 6.9.

The main engineer species of the hydrothermal ecosystem. a) Riftia pachyptila, EPR; b) Alvinella pompejana, EPR; c) Bathymodiolus azoricus, MAR; d) Rimicaris exoculata, MAR; e) Alviniconcha spp., Lau basin; f) Vulcanolepas sp., Lau basin. (Photos a, b, c and d belong to Ifremer and photos eandfarea courtesy of C. Fisher, Penn State University).

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Marine Ecosystems

Diversity and Functions

André Monaco

Patrick Prouzet

From the Seas and Oceans Set coordinated by André Mariotti and Jean-Charles Pomerol

First published 2015 in Great Britain and the United States by ISTE Ltd and John Wiley & Sons, Inc.

Apart from any fair dealing for the purposes of research or private study, or criticism or review, as permitted under the Copyright, Designs and Patents Act 1988, this publication may only be reproduced, stored or transmitted, in any form or by any means, with the prior permission in writing of the publishers, or in the case of reprographic reproduction in accordance with the terms and licenses issued by the CLA. Enquiries concerning reproduction outside these terms should be sent to the publishers at the undermentioned address:

ISTE Ltd27-37 St George’s RoadLondon SW19 4EUUK

www.iste.co.uk

John Wiley & Sons, Inc.111 River StreetHoboken, NJ 07030USA

www.wiley.com

© ISTE Ltd 2015

The rights of André Monaco and Patrick Prouzet to be identified as the authors of this work have been asserted by them in accordance with the Copyright, Designs and Patents Act 1988.

Library of Congress Control Number: 2015948076

British Library Cataloguing-in-Publication Data

A CIP record for this book is available from the British Library

ISBN 978-1-84821-782-9

Foreword

We have been asked by ISTE to stimulate work in the area of the environment. Therefore, we are proud to present the “Seas and Oceans” set of books, edited by André Monaco and Patrick Prouzet.

Both the content and the organization of this collection have largely been inspired by the reflection, initiatives and prospective works of a wide variety of national, European and international organizations in the field of the environment.

The “oceanographic” community, in France and internationally – which is recognized for the academic quality of the work it produces, and is determined that its research should be founded on a solid effort in the area of training and knowledge dissemination – was quick to respond to our call, and now offers this set of books, compiled under the skilled supervision of the two editing authors.

Within this community, there is a consensus about the need to promote an interdisciplinary “science of systems” – specifically in reference to the Earth’s own “system” – in an all-encompassing approach, with the aim of providing answers about the planet’s state, the way it works and the threats it faces, before going on to construct scenarios and lay down the elementary foundations needed for long-term, sustainable environment management, and for societies to adapt as required. This approach facilitates the shift of attention from this fundamental science of systems (based on the analysis of the processes at play, and the way in which they interact at all levels and between all the constituent parts making up the global system) to a “public” type of science, which is finalizable and participative, open to decisionmakers, managers and all those who are interested in the future of our planet.

In this community, terms such as “vulnerability”, “adaptation” and “sustainability” are commonly employed. We speak of various concepts, approaches or technologies, such as the value of ecosystems, heritage, “green” technologies, “blue” chemistry and renewable energies. Another foray into the field of civilian science lies in the adaptation of research to scales which are compatible with the societal, economic and legal issues, from global to regional to local.

All these aspects contribute to an in-depth understanding of the concept of an ecosystemic approach, the aim of which is the sustainable usage of natural resources, without affecting the quality, the structure or the function of the ecosystems involved. This concept is akin to the “socio-ecosystem approach” as defined by the Millennium Assessment (http://millenniumassessment.org).

In this context, where the complexity of natural systems is compounded with the complexity of societies, it has been difficult (if only because of how specialized the experts are in fairly reduced fields) to take into account the whole of the terrestrial system. Hence, in this editorial domain, the works in the “Seas and Oceans” set are limited to fluid envelopes and their interfaces. In that context, “sea” must be understood in the generic sense, as a general definition of bodies of salt water, as an environment. This includes epicontinental seas, semi-enclosed seas, enclosed seas, or coastal lakes, all of which are home to significant biodiversity and are highly susceptible to environmental impacts. “Ocean”, on the other hand, denotes the environmental system, which has a crucial impact on the physical and biological operation of the terrestrial system – particularly in terms of climate regulation, but also in terms of the enormous reservoir of resources they constitute, covering 71% of the planet’s surface, with a volume of 1,370 million km3 of water.

This set of books covers all of these areas, examined from various aspects by specialists in the field: biological, physical or chemical function, biodiversity, vulnerability to climatic impacts, various uses, etc. The systemic approach and the emphasis placed on the available resources will guide readers to aspects of value-creation, governance and public policy. The long-term observation techniques used, new techniques and modeling are also taken into account; they are indispensable tools for the understanding of the dynamics and the integral functioning of the systems.

Finally, treatises will be included which are devoted to methodological or technical aspects.

The project thus conceived has been well received by numerous scientists renowned for their expertise. They belong to a wide variety of French national and international organizations, focusing on the environment.

These experts deserve our heartfelt thanks for committing to this effort in terms of putting their knowledge across and making it accessible, thus providing current students with the fundaments of knowledge which will help open the door to the broad range of careers that the area of the environment holds. These books are also addressed to a wider audience, including local or national governors, players in the decision-making authorities, or indeed “ordinary” citizens looking to be informed by the most authoritative sources.

Our warmest thanks go to André Monaco and Patrick Prouzet for their devotion and perseverance in service of the success of this enterprise.

Finally, we must thank the CNRS and Ifremer for the interest they have shown in this collection and for their financial aid, and we are very grateful to the numerous universities and other organizations which, through their researchers and engineers, have made the results of their reflections and activities available to this instructional corpus.

André MARIOTTIProfessor Emeritus at University Pierre and Marie CurieHonorary Member of the Institut Universitaire de FranceFrance

Jean-Charles POMEROLProfessor Emeritus at University Pierre and Marie CurieFrance

1Marine Biosphere, Carbonate Systems and the Carbon Cycle

1.1. Introduction

It is now accepted that the recent increase in the concentration of carbon dioxide (CO2) in the atmosphere is a consequence of human activities, in particular the combustion of fossil fuels and the production of cement. Approximately 30% of CO2 emissions are absorbed by the ocean, which clearly indicates the importance of the sea in the regulation of the level of atmospheric CO2. Once absorbed by the ocean, this carbon becomes an important parameter in biogeochemical cycles.

For several million years, the concentration of CO2 in the atmosphere has shown cyclic variations directly linked to global changes in volume. These changes occur at regular intervals of 100,000, 40,000 and 23,000 years (Figure 1.1) which represent the Earth’s orbit about the Sun; these phases control the insolation on the surface of the Earth (solar forcing) and are responsible for the natural variability of the climate.

COMMENTS ON FIGURE 1.1.– These parameters, each with their own frequency, combine together and cause climatic variations, particularly glacial/interglacial variations. The two curves to the bottom of the figure show: the variations in isotopes of oxygen in the plankton shells (SPECMAP) [MAR 87] which are determined by water temperature and the global volume of ice over the past 400,000 years. Ice cores entrap bubbles containing past atmospheres. The concentration of CO2 in these bubbles (EPICA) [LUE 08] is very similar to that in the ocean, by identifying glacial (gray areas) and interglacial areas.

Figure 1.1.Parameters of the Earth’s orbit [LAS 04] molding the global climate (CO2 concentration in ice core d), oxygen stable isotopes in marine sediments e), precession b) and obliquity c) influence the seasonal contrasts, eccentricity a) modulates (amplifies or not) precession effects (Beaufort, compilation of [LAS 04, LUE 08, MAR 87])

The reasons why the atmospheric concentration of CO2 follows climatic variations are still not fully understood, but the ocean could be responsible for their long term natural variations. In fact, the ocean surface layer contains an enormous reservoir of carbon than can react with the atmosphere over these orbital time scales. Marine organisms are of particular importance in these mechanisms as they help incorporate (pump) and transfer a large amount of carbon from the surface of the ocean to the deep sea and into sediments (sinks). Their role becomes crucial in the global carbon cycle.

If we know that human activities, by emitting large quantities of CO2 into the atmosphere, disturb the biological fluxes of carbon in the ocean, it is extremely difficult to predict the response of marine ecosystems and what the future holds for natural carbon in the biological cycle. Studying marine sediments, which can be considered as historical archives of ocean ecology, allows us to better understand how the ocean participates in the carbon cycle and how marine biodiversity adapts to global changes. In the past, these changes in biodiversity have sometimes had very significant retroactive effects on the environment and climate.

This contribution does not intend to describe the chemistry of oceans or biogeochemical cycles, widely developed in the first volume of this set of books [BER 14, LEG 14]. As stated by Legendre [LEG 14], despite their low biomass, pelagic ecosystems are the driver of biogeochemical cycles in the oceans. In this context, we will focus in particular on calcareous phytoplankton in the dynamics of carbon and its role in the evolution of past and future climates. The complexity of these processes is why this system is often disregarded in experimental research, models and projections.

1.2. Marine organisms and carbon

Marine organisms are clearly adapted to the ocean properties, in which they live, but they also actively contribute to its composition; since organismal biology is based on the chemistry of carbon, this is particularly true for the concentration of carbon dissolved in the sea. In fact, marine organisms use carbon to build their tissues (organic form of carbon), many of which also form solid skeletons, particularly in the form of calcium carbonate (inorganic form of carbon). The concentrations of dissolved and particulate carbon therefore change according to the mass and activity of these organisms. Two equations can express the dual effect of this activity on dissolved carbon: on the one hand, carbon sequestration by photosynthesis into their tissues [1.1] and, on the other hand, the release of CO2 upon the construction of their carbonaceous skeleton from bicarbonate and dissolved calcium [1.2]:

[1.1]

[1.2]

These processes can be reversed depending on the environmental conditions. In addition, if the organic matter produced is used by another organism (grazing, predation), this will release carbon dioxide; however, if it is buried in the sediments; the carbon will be trapped, sometimes for very long periods of time (sequestration). With regard to the carbonates that form the skeletons (calcification), they will be either buried in sediment or dissolved in the ocean.

Calcification and biomineralization will depend on the biodiversity of the organisms; different species will not produce the same quantity of matter, skeletons do not have the same degree of calcification as they are produced at different paces. We will therefore see how marine biodiversity impacts the global carbon cycle and therefore climate.

1.3. Variability in the production of organic matter

During photosynthesis, marine algae absorb dissolved CO2 to produce their biomass, thereby releasing oxygen (equation [1.1]). The algal production is not distributed uniformly throughout the ocean. Most multicellular algae live on the seabed and on a substrate; they are therefore known as benthic; and their distribution is limited to the shallowest depths of the ocean where sunlight can penetrate. On the contrary, algae that constitute the phytoplankton are almost all unicellular and floating; they are widely distributed along oceanic margins, and beyond, over a maximum of approximately 200 m wide in the water column.

Continental margins are generally much more productive in terms of organic matter than zones situated in the center of oceanic basins. In fact, rivers feed the coastal zones with nutritive salts (nitrogen, phosphorus, etc.), required for the growth of phytoplankton. Offshore, in the pelagic zone, this lateral input is increasingly less as one strays from the continents. The wind sometimes brings dust rich in nutritive salts, but these fluxes rarely compensate for this deficit.

Pelagic phytoplankton that lives exclusively in the photic zone (zone exposed to light) depletes the nutrient stores during photosynthesis; the only way to regenerate this stock at the surface is by the vertical pumping of nutritive salts in the deep sea (>200 m). These vertical nutrient transfers, required for pelagic life, are only produced by updrafts caused by gusts of wind in favorable directions (upwelling) or by the deep mixing of surface layers during episodes of strong winds.

One example of this mechanism can be found in the Indian Ocean where monsoon winds are powerful enough to break down the vertical stratification and allow high phytoplanktonic production by fertilization at depth. This mixing acts over a certain depth and is called the mixed layer; when it reaches the thermocline that separates warm surface waters from cold deep waters, the nutritive salts then diffuse towards the surface. We can then easily understand that during periods favored by strong Indian monsoons, or strong trade winds, offshore of West Africa, oceanic primary production (produced by phytoplankton or PP) is reinforced.

Studying sediment cores taken from these zones has revealed that highly biologically productive periods alternate with depletions, and that these changes follow the rhythms of the Earth’s orbit. Thus, primary production, expressed in grams of carbon per meter squared per year, has increased from 120 to 200 gC/m2/yr, since the previous precession cycle of the equinoxes in the center of the Indian Ocean [BEA 97] or in the Banda Sea (Figure 1.2). By causing seasonal variations and by heating different tropical zones, these cycles cause the increase or decrease in winds above the Indian Ocean, beginning with oceanic production.

COMMENTS ON FIGURE 1.2.– Top: variations in organic productivity, common (calculated by EOF) to the entire tropical band in the Indian and Pacific Ocean [BEA 01] (continuous line and circles). The organic production varies according to the concentration of CO2 measured in the core of Antarctic ice [LUE 08] (dotted line and full squares). Bottom: recording of the organic production (dotted line, full circles and vertical error bars) in the Banda Sea. It is highly dependent on the intensity of the Australian monsoon [BEA 10]. The monsoon activity deduced follows the rhythms of insolation which are parameters of the Earth’s orbit (continuous line) [LAS 04]. This highlights the difference that can exist between the local and global variations.

Figure 1.2.Example of the evolution of organic production in the tropical zone over 150,000 years

In this example, it is also clear that these changes in primary productivity are accompanied by significant variations in biodiversity. In the sediments deposited in the center of the Indian Ocean 20,000 years ago, 80% of fossil micro flora observed (Figure 1.3) is represented by species known to currently inhabit depths of 100 to 200 m beneath the ocean surface, at the lower limit of the photic zone, in conditions of very low levels of light. However, 10,000 years ago, this microflora represents only 10% of the assembly (assembly of a fossilized community); it has also been replaced by more diverse forms known to live between 0 and 100 m.

We can compare this location in the ocean 20,000 years ago to where we currently find the poorest zones of the central South Pacific (around Easter Island or North Pacific (Hawaii)). These zones, called gyres, represent oceanic deserts (except for areas close to islands) where planktonic organisms live at their deepest, to the extreme limits of light able to access nutritive salts. These zones are occupied by sparse ecological communities, especially at the surface and there are significant vertical differences in the composition of phytoplankton. In the central South Pacific, the specific composition at depth (example 150 m) is more similar to that of the north of the central North Atlantic, at the same depth, than those that live just below the surface. Longhurst [LON 98] defines the base of the mixing zone as the most significant ecological barrier that exists in the ocean. Biogeographical regions or biomes can be defined based on variations in some physicochemical parameters such as the depth of the thermocline and the photic zone, or seasonality.

1.4. From the biosphere to the atmosphere to climate

If climate changes greatly influence phytoplanktonic communities in the sea, in turn, these ecological changes can alter the CO2 composition in the atmosphere and hence the climate. Similarly, global changes to the wind regime have significant consequences on the carbon cycle and marine biodiversity.

Currently, the global annual primary production is approximately 40–45 gigatons of carbon (45 × 109 g). Most of this biomass is consumed by heterotrophic organisms (planktivores, carnivores and decomposers) and the mineralization of this organic matter will produce CO2. Ultimately, only a small part falls to the oceanic depths where it is either consumed by heterotrophic benthic organisms, or buried in the sediments. In the former case only, carbon is removed from biogeochemical cycles, often for very long periods of time. We are therefore referring to exported primary production. Although this represents a very small fraction (0.02%) of what is produced at the ocean surface, this fraction of deposited carbon actually represents a significant carbon sink due to the immensity of the ocean surface; it is currently estimated to be more than 10 Mt (107 g) of carbon per year. Changes in climate will induce changes in the deposition rate of organic matter via changes to the general circulation speed or degree of oxygenation of the deep ocean which depends on the activity of the benthos; the absence of benthic organisms favors carbon sinks.

The issue that now deserves to be discussed is how a change to the oceanic biological production and its deposition rate as sediments alter the concentration of CO2 in the atmosphere. This was the first hypothesis proposed by Broecker [BRO 82] to explain the discovery, in 1980, of lower levels of CO2 in the atmosphere during glacial periods than during interglacial periods. This author proposes an increase in concentration of nutritive salts (essentially phosphate) which would favor organic production. In fact, the decrease in sea levels during the ice age would expose marine sediments rich in phosphate, deposited during the previous interglacial period, and transfer them into the ocean via rivers. This fertilization of the glacial ocean would have induced a greater sequestration of carbon.

Other reasons have been put forth to explain the decrease in concentration of atmospheric CO2 during the ice age. Note, a better use of nutritive salts by phytoplankton in the Antarctic Ocean [FRA 97]; an increase in strong dustladen winds bringing iron into zones where this micronutrient is currently not found [MAR 90]; an increase in productivity resulting from a transfer of dissolved silica from the Antarctic Ocean towards the lower latitudes [MAT 02]; a decrease in the general depth of the thermocline at low latitudes which would have fertilized the ocean surface [BEA 01]; or even, an increase in biological production parallel with a decrease in carbonate production [BRO 82]. The latter point will be discussed later on.

Most of these hypotheses themselves cannot completely explain the decrease in concentration of atmospheric CO2 during ice ages. Moreover, reconstructions of the primary production during these episodes do not always confirm the increases, and large local variations are observed.

At low latitudes, organic production seems to vary independently of global changes in climate [BEA 97] (Figure 1.2). At these latitudes, changes in organic productivity occur every 20,000 years which corresponds to local variations in seasonal insolation, whereas variations in global ice volume occur less frequently (every 41,000 and 100,000 years). Numerical simulations also indicate that local factors, for example wind strength and that of upwelling, are responsible for distribution structures and changes in production of organic matter. Sediments in the tropical zone of the Pacific and Indian Ocean also reveal that the organic production during the last ice age was 50% greater than that produced today (Figure 1.2).

However, many records of organic productivity reveal variations with a dynamic opposite to variations in concentrations of atmospheric CO2 (Figure 1.2 top); but locally the variations may obey different rules (Figure 1.2 bottom).

At high latitudes (above 40°) and in many tropical regions (for example in the monsoon region), numerical simulations show an increase in PP during the ice age which implies greater sequestration of CO2 during these periods. A compilation of many data from studies of sediments of the world’s oceans indicates that the exported organic production (including sediments) during the last maximum ice age is slightly higher than that of today. However, this variation does not appear to be enough to explain the differences in CO2 between the glacial and interglacial period. As a result, even if the mechanism proposed by Broecker [BRO 82] cannot fully account for these variations, we know that the primary production has varied greatly in the past and contributed significantly to the atmospheric CO2 dynamics. So, there is a feedback loop between climate and the marine biosphere, both influencing one another.

1.5. Carbonate production

Calcium carbonate forms the greatest reservoir of carbon on Earth [LEG 14] and many marine species produce shells based on this bio mineral. These secretions, just like organic production, play an important role in the chemistry of the oceans. Just as Broecker stated for organic production, Berger [BER 82] also mentioned changes in the production and conservation of biogenic carbonates to explain variations in the concentration of CO2 in the atmosphere during glacial and interglacial periods.

The phenomena responsible for this relationship are more difficult to explain since they involve complex chemical processes. So, if we conventionally assume that the dissolution of limestone by acid produces carbon dioxide, in the ocean, this is the opposite reaction: the dissolution of limestone sequesters CO2 and secretes CO2. As expressed in equation [1.2], calcification (precipitation of CaCO3) is accompanied by the release of CO2 from the ocean into the atmosphere. This is what is called the carbonate counter-pump [WES 93].

1.5.1. Importance of biological carbonate production in the evolution of the planet

In the ocean, most calcification is of biological origin. Non-biological precipitations of carbonate are only produced in the presence of extreme concentrations of carbonate or bicarbonate ions. These conditions only currently exist in rare environments, at warm shallow depths with high evaporation, but they are responsible for the formation of carbonaceous rocks in the primitive ocean.

1.5.1.1. The Neoproterozoic

In the primitive ocean and before the first secretions of biological origin, chemical precipitations only operated in conditions where there was an over saturation of carbonate ions which therefore meant atmospheric CO2 was not controlled by the ocean. Carbonate ions from the weathering of aluminosiliceous continental rocks are transported towards the ocean. It was shown that the carbon input into the ocean from erosion was greater than that released into the atmosphere by marine calcification; the content of CO2 in the atmosphere would have progressively declined. For example, if the flow of carbon into the ocean is greater than 25% of the flow of carbon into the atmosphere, in only one million years, the atmosphere would be depleted of CO2 [BER 97]; the Earth’s surface would therefore have been covered in ice, without a greenhouse effect.

The biological carbonate production began 3.5 billion years ago, with stromatholiths from the activity of photosynthetic cyanobacteria. These would have dominated the production of carbonates for a long time but insufficiently so to decrease the saturation of oceans with carbonate ions from the erosion of continental rocks. An increase in continental erosion, sequestration of atmospheric CO2 and a decrease in the greenhouse effect [DON 04] arose as a result of extreme glacial events. The oceans were certainly completely covered in ice; between 0.8 and 0.6 billion years ago, we find traces of glaciers up to the equator, which was called Snowball Earth by Kirschivink [KIR 92].

It therefore appears that carbonate ions deposited into the ocean were not used quickly enough by stromatoliths and the biological release of oceanic carbon into the atmosphere was too limited to prevent these extreme conditions from occurring.

1.5.1.2. The Paleozoic

At the beginning of the Paleozoic (primary era), approximately 550,000,000 years ago, several different groups of organisms, mollusks, sponges, corals, arthropods, etc., began to produce carbonaceous skeletons in shallow environments. This revolution of marine biodiversity would have had consequences on atmospheric CO2 concentrations.

The rate of carbonate production depended on the relative sea level height and the usable area for calcification, during periods of high sea levels (submerged continental shelves), the total carbonate production will be significant and will cause an increase in CO2 in the atmosphere; this relates to relatively warm periods. During periods of low sea levels, there are fewer calcifying organisms in shallow zones, which will cause a decrease in the total marine carbonate production which will not be able to compensate for the carbonate input from the continents. In particular, this is the case of periods of tectonic activity and uplift of mountain ranges where erosion is great. If the glaciations were severe, they would not have had the same extensions and consequences as before this first revolution, as during Snowball Earth.

1.5.1.3. Mesozoic revolution

During the Triassic (220 million years ago), a group of unicellular algae, the coccolithophores, began to produce a carbonaceous skeleton made of microscopic plates (coccoliths) external to their cells. These microalgae can develop offshore due to their floating ability and rapidly colonize the oceans. From the Jurassic onwards, they are accompanied by planktonic for aminifera, which, despite also being unicellular, belong to the zooplankton. Most current coccolithophores have a diameter of between 2 and 15 μm whereas for aminifera usually have a diameter of up to several hundred micrometers (Figure 1.3). Despite their tiny size and their low weight (less than a nanogram), coccolithophores can produce very dense blooms extending over thousands of kilometers squared. The density of calcareous scales is so great that they can be detected by satellite images. Found at the bottom of oceans, remains of coccolithophores can contribute to the formation of huge quantities of light carbonate-rich sediments; the extreme example is chalk from the Cretaceous (from the Latin Creta) uniquely composed of coccolithophore remains (Figure 1.4).

Figure 1.3.Photographs taken using a scanning electron microscope of several representatives of the two main groups of pelagic calcifying organisms

COMMENTS ON FIGURE 1.3.– To the left, planktonic for aminifera Globorotalia menardii; to the right, five coccolithophores (from left to right and top to bottom): Discosphaera tubifera, Helicosphaera carteri, Emiliania huxleyi, Solisphaera galbula and Florisphaera profunda. The two last species at the bottom left are typical of the deepest parts of the photic zone. Note the difference in scale (horizontal bars) for the foraminifera and for coccolithophores (author’s photos).

The colonization of the entire surface of the ocean by calcite-producing organisms was a major evolutionary event in the history of the climate. It was called the “mid-Mesozoic revolution” [RID 05] due to its high impact on marine chemistry. Pelagic calcification helped better stabilize the climate; even during glacial episodes of the quaternary period, it helped maintain an equilibrium between the production of marine carbonates and atmospheric carbon.

Numerical models have shown that carbonate concentrations have remained relatively stable over the past 100 million years, whereas they were two to three times greater and highly variable before the coccolithophores appeared (more than four times greater before the Paleozoic [RID 05]). This shows how important the role the diversity of the biosphere plays in the stability of the Earth’s climate.

1.5.2. Carbonate compensation depth

Currently, coccolithophores and planktonic foraminifera produce the largest portion of calcium carbonate in the pelagic domain which represents 90% of the sea surface. However, if the open ocean contributes to more than half of the marine production, only a third approximately is accumulated in the form of CaCO3. This difference is due to the fact that bio-mineral particles produced in shallow zones have a much greater chance of being buried than those produced offshore. So, aragonite crystals (one of two forms of CaCO3) secreted by coralin tahiti certainly contribute to reef growth whereas calcite crystals (the other form of calcium carbonate), secreted at the surface by a coccolithophore tens of kilometers offshore of the reef, will not be buried in the ocean floor if they are too deep. In fact, beyond a depth of 4,000 m, calcite dissolves. It is estimated that 50–80% of carbonates that are secreted on continental margins are stored in coastal sediments, against 45% of carbonates produced in the open ocean. In comparison, each year, 11 × 1012 moles of CaCO3 accumulate in the pelagic zone and 7 × 1012 moles of CaCO3 in coral reefs [MIL 93]. Despite their better efficiency in storing calcium carbonate, coral reefs still accumulate less due to a relatively low useable area 0.6 million km2 versus 106 million km2 for the pelagic zone whose depth is less than 4,000 m. Carbonate particles dissolve below a certain depth mainly due to the high pressure; the solubilities of calcite and aragonite strongly depend on pressure, and at great depths, the pressure is great enough that CaCO3 dissolves; this limited depth is called the lysocline.

First note, the transfer of particulate carbon towards the sea floor due to the aggregation of diverse particles produced on the surface and the ballast effect of coccoliths; since this falling of particles is referred to as “marine snow”, the deep sea is often represented as an alpine landscape with snowcapped peaks (white calcite and aragonite), and the deeper snow-free parts (without carbonates) (Figure 1.4). This physiographic limit corresponds to the base of the lysocline, carbonate compensation depth (CCD) below which there are no carbonate deposits. On the map of the central Pacific (Figure 1.4(c)) the white/black limit which corresponds to the mean depth of the lysocline (4 000 m +/– 900) represents the contrast between oceanic basalt (black) and the carbonates (light). A photo of the depths (Figure 1.3(b)), in this zone, at –2,000 m, shows that the sediments are whitish; a core sampled at this depth (Figure 1.4(c)) allows us to verify, from assessing samples using an electronic microscope, that the sediment is almost entirely composed of planktonic foraminifera (Figure 1.4(d)) and calcareous scales from coccolithophores (Figure 1.4(e)). In the same zone, but at greater depths, basalt occurs without being covered by sedimentary rock.

Figure 1.4.Images of the deep sea in the Southern islands at a depth of 2,500 m (images from the oceanographic campaign SO2O4 Pacenpal by Beaufort, 2010)

COMMENTS ON FIGURE 1.4.– (a) General map. (b) Detailed map; light areas are depths of less than 4,000 m occupied by carbonaceous sediments and the shaded areas are depths greater than 4,000 m where basalticrocks are found. (c) Photograph of the deep sea during the sampling of white calcareous sediment cores. (d) Photograph of the sediment viewed under a scanning electron microscope, showing the foraminifera (the largest particles) and their fragments (smaller identifiable elements). (e) Detailed image of (d) showing that the cement surrounding the foraminifera is exclusively composed of coccolithophore skeletons.

Carbonate compensation does not always occur at the same depth since the solubilities of calcite and aragonite also depend on the concentration of carbonate and calcium. So, a high productivity of CaCO3 at the surface, during a coccolithophore bloom, will cause a high vertical flux of calcareous skeletons which, upon arrival at the CCD, will dissolve and release calcium and carbonate ions into the water; the increase in concentration of these ions will tend to displace the lysocline at depth. However, a zone of low carbonate production at the surface will have alysocline situated at middle depth. As long as there is a pelagic production of carbonate, the overall chemical concentration of Ca and carbonate ions in the ocean remains fairly constant and is redistributed by the currents.

The overall mechanism that transfers the sum of all chemical species of inorganic carbon or ΣCO2 from the surface towards the sea floor and that maintains the high concentrations at depth is called the “carbonate-counter pump” and plays a key role in the regulation of the climate; Legendre [LEG 14] describes four types: physical, organic (or biological pump), carbonate-counter pump and the microbial pump.

1.5.3. Carbonates and climate

Very soon after the discovery of low concentrations of atmospheric CO2 during ice ages, Berger [BER 82] stated that the possession of continental margins by coral reefs during deglaciation could explain a substantial part of the post-glacial increase in CO2. During the last major ice age which occurred between 70,000 and 15,000 years ago, the volume of ice distributed over the continents was so large that it caused a decrease in the sea level by 120 m. Organic and carbonate production was limited to a small coastal fringe and the global production of CaCO3 decreased to almost half [BER 82]. This decrease was slightly compensated for by the more constant production of pelagic carbonate following a deepening of the lysocline by approximately 1,000 m, as in the Pacific [FAR 89]. During deglaciation, the continental margins became gradually submerged and benthic calcification caused a rapid increase in reefs and therefore a gradual increase in CO2 in the atmosphere.

The numerical modeling of this phenomenon shows that this amount of carbon accounts for a large part of the difference in concentration of atmospheric CO2 between a glacial and an interglacial period. However, this scenario is often overlooked in current studies on the marine carbon cycle. The reason for the infrequent consideration of this phenomenon is due to the fact that it occurs as a result of climate change and therefore cannot initiate it. In fact, it requires that the sea levels rise as a result of warming. This initial warming cannot therefore be attributed to this phenomenon. Nevertheless, changes in the habitat area of marine organisms and ecological systems have certainly had a non-negligible retroactive effect on the climate variations during deglaciation.

Various experimental data show that during ice ages, the concentration of carbonate ions is greater, and hence the pelagic carbonate production is probably greater:

1) the planktonic foraminifera shells are then thicker and coccolithophores produce more calcified plates;

2) the pelagic production at low latitudes and particularly that of coccolithophorids is greater than that during interglacial periods;

3) as we have seen, the lysocline retreats.

So, a greater production of carbonate and a better conservation of carbonates will cause an increase in the concentration of atmospheric CO2 which will mitigate the effects of glaciation. However, a precise quantification of the pelagic carbonate production and its conservation has rarely been carried out because of the complexity of its implementation.

1.6. The coupling of carbonaceous and organic productions

We have seen that the marine organic production as well as the carbonate production interfere, at different scales of time and space, with the concentration of CO2. Some models play with this combination.

In this category, the most elegant is definitely that of the rain ratio by Archer and Maier-Reimer [ARC 94]. During the ice age, in the pelagic zone, there appears to have been an increase in organic productivity coupled with a decrease in carbonate productivity. So, carbon has a better chance of being sequestered in the sediments causing the observed decrease in carbon during this period. For these authors, a 40% decrease in the ratio of CaCO3 in organic matter in sediment particles would be sufficient to explain the decrease in concentration of atmospheric CO2 recorded during ice ages.

Several reasons could explain the decrease in this ratio; they are all linked to a major ecological change in the pelagic zone: the replacement of coccolithophores by diatoms, phytoplanktonic organisms by a siliceous skeleton. Diatoms would outcompete coccolithophores due to an influx of silicic acid whose sources, according to hypotheses, would be: wind (silicarich dust), rivers from exposed margins, or even an oceanic source of silica from the Antarctic.

This theory, still a very topical issue, has been faced with observations that contradict it:

1) as we saw above, the coccolithophores do not appear to have decreased in abundance nor in calcification; rather the opposite is observed;

2) in the cores from low and middle latitudes, we do not find a synchronous or large increase in the number of diatom skeletons.

We know that the process for storing marine organic matter is much more efficient than the aggregation of matter in suspension (marine snow). In this sense, calcareous coccolithophorid scales are the best vectors of carbon towards the sea floor and sedimentary sinks. In these conditions, if the coccolithophores were outcompeted by diatoms, the hypothesis of organic carbon sequestration in sediments would not be valid.

1.7. Modification of equilibria and consequences on marine life

Time constants play an important role in the carbon cycle; as described above, the marine carbon cycle follows the rhythms of global climate variations (Figure 1.1