AOBPreview originally published online on September 4, 2002
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Annals of Botany 90: 525-536, 2002
© 2002 Annals of Botany Company
Seaweeds in Cold Seas: Evolution and Carbon Acquisition
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,11 Division of Environmental and Applied Biology, School of Life Sciences, University of Dundee, Biological Sciences Institute, Dundee DD1 4HN, UK, 2 Scottish Crop Research Institute, Invergowrie, Dundee DD2 5DA, UK, 3 Department of Botany, University of Western Australia, Crawley, WA 6009, Australia, 4 School of Biological Sciences, Monash University, Clayton, Victoria 3800, Australia, 5 University of Oslo, Department of Biology, Section for Marine Botany, PO Box 1069 Blindern, 0316 Oslo, Norway and 6 University of Texas at Austin, Marine Science Institute, 750 Channel View Drive, Port Aransas, TX 78373, USA
* For correspondence. Fax +44 (0)1382 344275, e-mail j.a.raven{at}dundee.ac.uk Present address: Department of Biology, 18111 Nordhoff Street, California State University, Northridge CA 91330-8363, USA. Present address: British Antarctic Survey, High Cross, Madingley Road, Cambridge CB3 0ET, UK.
Received: 27 November 2001; Returned for revision: 12 February 2002; Accepted: 23 April 2002 Published electronically: 4 September 2002
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONEARTH HISTORY AND ALGAL...THE AQUATIC HABITATFUNCTIONING OF ALGAE IN...CONCLUSIONNOTE ADDED IN PROOFLITERATURE CITED |
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Much evidence suggests that life originated in hydrothermal habitats, and for much of the time since the origin of cyanobacteria (at least 2·5 Ga ago) and of eukaryotic algae (at least 2·1 Ga ago) the average sea surface and land surface temperatures were higher than they are today. However, there have been at least four significant glacial episodes prior to the Pleistocene glaciations. Two of these (approx. 2·1 and 0·7 Ga ago) may have involved a ‘Snowball Earth’ with a very great impact on the algae (sensu lato) of the time (cyanobacteria, Chlorophyta and Rhodophyta) and especially those that were adapted to warm habitats. By contrast, it is possible that heterokont, dinophyte and haptophyte phototrophs only evolved after the Carboniferous–Permian ice age (approx. 250 Ma ago) and so did not encounter low (
5 °C) sea surface temperatures until the Antarctic cooled some 15 Ma ago. Despite this, many of the dominant macroalgae in cooler seas today are (heterokont) brown algae, and many laminarians cannot reproduce at temperatures above 18–25 °C. By contrast to plants in the aerial environment, photosynthetic structures in water are at essentially the same temperature as the fluid medium. The impact of low temperatures on photosynthesis by marine macrophytes is predicted to favour diffusive CO2 entry rather than a CO2-concentrating mechanism. Some evidence favours this suggestion, but more data are needed.
Key words: Review, carbon dioxide, Chlorophyta, glaciations, Heterokontophyta, Phaeophyceae, Rhodophyta.
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONEARTH HISTORY AND ALGAL...THE AQUATIC HABITATFUNCTIONING OF ALGAE IN...CONCLUSIONNOTE ADDED IN PROOFLITERATURE CITED |
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Algae can clearly perform well in cold environments. The Antarctic continent has two native vascular plant species and a substantial number of species of freshwater and terrestrial bryophytes. Most of the primary productivity of Antarctica involves free-living algae in lakes, and algal symbioses with fungi (lichens, and the green alga Prasiola) and, as free-living cyanobacteria, on land. In the surrounding oceans are phytoplankton and sea-ice algae, as well as very large subtidal macroalgae, especially those of the brown algal order Desmarestiales, with kelp (Laminariales)-like organisms such as Himantothallus grandifolius with blades up to 10 m long and 1 m wide (Lüning, 1990; Fogg, 1998; Boyd et al., 2000; Brierley and Thomas, 2002; Thomas and Dieckmann, 2002a, b; Wiencke and Clayton, 2002).
This article points out that the origin of life, and much of Earth’s history over the intervening 3·8 Ga (billion years) or so, has involved average sea-surface and land-surface temperatures greater than those found today, and relates the evolution of marine algae to the temporal and spatial changes in Earth surface temperatures (Falkowski and Raven, 1997). The article also deals with some contrasting effects of air and water as fluid environments in terms of temperature differences between photosynthetic organs and the medium, and with the impact of temperature on CO2 supply for, and assimilation in, photosynthesis (Raven, 1997; Sherlock and Raven, 2001). We do not wish to deny the influence of temperature on the acquisition of resources, other than photons and inorganic carbon, on the transfer of resources within the organism, on growth and development, or on the effectiveness of reproduction.
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EARTH HISTORY AND ALGAL EVOLUTION |
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TOPABSTRACTINTRODUCTIONEARTH HISTORY AND ALGAL...THE AQUATIC HABITATFUNCTIONING OF ALGAE IN...CONCLUSIONNOTE ADDED IN PROOFLITERATURE CITED |
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Earth history
The Earth formed about 4·5 Ga ago, and 13C/12C evidence is consistent with CO2 fixation via Rubisco (ribulose bisphosphate carboxylase-oxygenase) occurring from about 3·8 Ga ago (Falkowski and Raven, 1997). Assuming that the organisms now found on Earth originated on Earth, molecular genetic data and some evidence from biogeochemistry suggest that life originated at hydrothermal vents, possibly as chemolithotrophs (Pace, 1997). This suggests that the earliest life was adapted to warm (relative to most current aquatic habitats) environments. At that time (4·5–3·8 Ga ago), the sun was only emitting 75 % as much energy as it does today, and the maintenance of liquid water on the Earth’s surface (at least after the Hadean meteorite bombardment had decreased) required high partial pressures of greenhouse gases such as CO2, or, in this essentially O2-free atmosphere, CH4.
Greenhouse gases did not always keep the sea-surface and land-surface temperatures high enough to prevent glacial episodes. These occurred in the Palaeoproterozoic (approx. 2·4–2·2 Ga ago) and the Neoproterozoic [approx. 820–550 million years (Ma) ago] and in the Phanerozoic, in the late Ordovician (approx. 450 Ma ago), the late Carboniferous and Permian (approx. 280 Ma ago) and the Pleistocene (the last approx. 2 Ma) (Falkowski and Raven, 1997; Williams et al., 1998). The first two of these involved sea-level glaciation at low latitudes, and it has been suggested that the Palaeoproterozoic and the Neoprotero zoic glaciations corresponded to ‘Snowball Earth’ (Hoffman et al., 1998), with both the land and the sea surface essentially being completely covered in ice. Recovery from this situation would be slow, taking 106–107 years, and would result from volcanic CO2 emissions not balanced by CO2 dissolving in seawater or consumed in weathering of silicate minerals on land because land and sea were covered in ice (Williams et al., 1998).
The occurrence of Snowball Earth has been doubted by some authorities, as have hypotheses based on varying obliquity of the Earth’s axis of rotation relative to the plane of the Earth’s orbit of the Sun, which might explain low palaeolatitude glaciation without requiring world-wide ice cover (Hoffman et al., 1998; Williams et al., 1998; Hoffman and Schrag, 1999; Williams, 1999; Schrag and Hoffman, 2001). There seems to be no evidence of low palaeolatitude glaciation in the three Phanerozoic glacial episodes. It is not certain whether the glacial episodes, and the intervening times when the mean temperature of the Earth’s surface was generally at least as high as it is today, can be related to greenhouse gas changes. Certainly the disposition of the continents and the corresponding equator-to-poles heat flow in the atmosphere and, especially, the sea, was important.
Cyanobacteria alone: 3·5–2·1 Ga ago
In relating variations in the Earth’s temperature to the occurrence of different higher taxa of algae, we start with the cyanobacteria. These were the organisms in which O2 evolution first occurred and they were responsible for the present atmosphere with much more O2 than can be explained by photodissociation of H2O vapour (Falkowski and Raven, 1997; Catling et al., 2001; Hoehler et al., 2001; Kasting, 2001). Cyanobacteria are still biogeochemically very important organisms in their own right, and were evolutionarily important as the source, via endosymbiosis, of O2 evolution in eukaryotes (Falkowski and Raven, 1997). While some microscopic fossil evidence has been said to indicate the occurrence of cyanobacteria 3·5 Ga ago (Golubic and Seong-Joo, 1999; Brasier et al., 2002; Schopf et al., 2002), there are good molecular markers of cyanobacteria in the form of 2-methylhopanoids from 2·7 Ga ago (Summons et al., 1999). Molecular fossils of eukaryotes have also been found from about 2·7 Ga ago (Brocks et al., 1999). The geochemically detectable build up of O2 in the atmosphere began about 2·1 Ga ago (Kasting, 2001).
Cyanobacteria with primarily endosymbiotic eukaryotes
The earliest evidence for eukaryotic algae, based on the size of the fossil Grypanea, is from 2·1 Ga ago (Falkowski and Raven, 1997). The most ancient eukaryotic algal fossil that can be assigned to an extant Division or Phylum of algae is that of the bangiaceous Bangiomorpha pubescens (Rhodophyta), about 1200 Ma old (Butterfield, 2000, 2001), while at least some acritarchs from the Neoproterozoic (1 Ga ago onwards) are cysts of prasinophyte algae (Chlorophyta) (Falkowski and Raven, 1997).
The antiquity of the cyanobacteria means that they must have survived both of the (possible) Snowball Earth episodes, while the two eukaryotic divisions (phyla) must have survived the second Snowball Earth event. If these events were of the global extent and long duration believed by some authors (Hoffman et al., 1998; Hoffman and Schrag, 1999; Kirshvink et al., 2000; Schrag and Hoffman, 2001), then there would have been both cold (1 km of ice over an ocean barely above freezing) and dark (1 km of ice would attenuate essentially all light) conditions over almost all of the Earth’s land and sea surface—with the exception of hot springs and their analogues in the ocean—for at least 106–107 years (cf. McKay, 2000, discussed later). Such events would have involved cyanobacteria and green and red algae living at low CO2 concentrations (low atmospheric CO2) and, perhaps, low temperatures (the hot springs may barely have caused the ice to melt), a combination to which we shall return later. However, McKay (2000) pointed out that sufficient photosynthetically active radiation (PAR) for photolithotrophic growth could penetrate the 10 m of ice which modelling based on Antarctic dry valley lakes suggests would occur in the equatorial ocean of Snowball Earth.
Cyanobacteria with primarily and secondarily endosymbiotic eukaryotes
Cyanobacteria, Chlorophyta and Rhodophyta underwent significant radiation as planktophytes (cyanobacteria, Chlorophyta) and as benthic organisms (all three taxa) in the Neoproterozoic and Palaeozoic (Falkowski and Raven, 1997; Heckman et al., 2001). However, the extant algae have more taxa that evolved by endosymbiosis, involving a green or red algal unicell in association with a variety of heterotrophic eukaryotes. Endosymbiosis of green algal unicells with two different heterotrophs led to euglenoids and to chlororachniophytes. Endosymbiosis involving different red (bangiophycean: Oliveira and Bhattacharya, 2000; Muller et al., 2001) algal unicells with (possibly) four different heterotrophs gave rise to the haptophytes, heterokonts (diatoms, brown algae, chrysophytes), cryptomonads and, probably, dinoflagellates. Whilst the timing of these endosymbiotic events is not certain, evidence from molecular clock analyses is consistent with most of the fossil evidence in suggesting that photosynthetic heterokonts and haptophytes evolved not earlier than the Permian–Triassic boundary about 250 Ma ago (Falkowski and Raven, 1997; cf. Xiao et al., 1998). This was immediately after the Carboniferous–Permian glaciation, and 235 Ma before the Antarctic ice sheet occurred during the cooling which culminated in the Pleistocene glaciations. This means that the photosynthetic marine heterokonts and marine haptophytes had not been through a low sea surface temperature episode until approx. 15 Ma ago. This statement refers to the symbiosis; clearly the partners in the symbiosis (the unicellular red algae and the heterotrophs) had individually experienced the Carboniferous–Permian, late Ordovician and Neoproterozoic glaciations, and the archean and bacterial ancestors of the heterotrophs, and the cyanobacteria, bacteria and archeans that gave rise to the red algae, also experienced the Palaeoproterozoic glaciation.
The case of the Phaeophyceae
Taking the brown algae (Phaeophyceae) as an example of marine algal evolution in the last 250 Ma and its relationship to temperature, it must be acknowledged that molecular phylogenetic approaches to the evolution of these algae have unresolved polychotomies at the ordinal and familial levels (de Reviers and Rousseau, 1999; Draisma et al., 2001; Raven et al., 2001; Rousseau et al., 2001; Yoon et al., 2001). However, robust inferences can be made from a combination of molecular genetic, biogeographical and palaeontological studies (Clayton, 1984, 1994; Lüning, 1990). We concentrate here on the three brown algal orders containing most of the larger representatives of these algae, i.e. the Desmarestiales, Fucales and Laminariales.
On the basis of their present-day natural occurrence, the Desmarestiales and the Fucales are of Gondwanan origin, while the Laminariales are of northern hemisphere, Pacific origin (Clayton, 1984, 1994; Lüning, 1990), although molecular genetic evidence shows a close evolutionary relationship between the Desmarestiales and the Laminariales (Tan and Druehl, 1996). It is likely that all three orders evolved at a time (before 15 Ma ago) when there was no significant low-temperature (<5 °C) surface seawater (Wilson and Norris, 2001; Zachos et al., 2001).
For the Fucales (including Durvillaea and Notheia; de Reviers and Rousseau, 1999), there are 13 genera that are only found in the northern hemisphere compared with 27 that are restricted to the southern hemisphere, and there are many more species in the southern hemisphere, with numerous species in the tropics (Lüning, 1990; Serrão et al., 1999). It has been suggested that the northern hemisphere received its meagre allocation of temperate fucoids during Pleistocene glaciations, when there was a land bridge between Australia and Indonesia, and seawater temperatures in this part of the tropics were 6–8 °C lower than they are today (Clayton, 1984; Lüning, 1990). This suggests a rapid, recent evolution of at least northern hemisphere temperate fucoids (Clayton, 1984, 1994; Lüning, 1990; de Reviers and Rousseau, 1999; Raven et al., 2001). The Fucales are not major elements of the Antarctic algal flora (an exception is Cystosphaera jacquinotii) but are common in cool temperate waters of both the northern and southern hemispheres with some very speciose genera in the tropics (Sargassum, Cystoseira) and in cool and warm temperate southern Australia (Cystophora, Sargassum).
The Desmarestiales are very important elements of the Antarctic marine flora and include the very large Himantothallus grandifolius (10 x 1 m blades). They apparently originated in the southern hemisphere (Peters et al., 1997), if not the Antarctic; an origin in the cold polar oceans would imply an age for the order of less than 15 million years. Members of the Desmarestiales could have moved to the northern hemisphere in the Pleistocene across relatively cool tropical waters in the glacial episodes during that era. Only non-Antarctic species could have survived this still relatively high-temperature passage.
The Laminariales currently have their greatest diversity in the Pacific, and apparently originated there. As the Arctic cooled prior to the Pleistocene glaciations, some laminarians became able to live in Arctic waters as well as cool temperate waters. The southern hemisphere laminarians are not found in the Antarctic; it is likely that only representatives from relatively warm temperate waters could have migrated across the tropics in the glacial episodes of the Pleistocene (Lüning, 1990; Coyer et al., 2001; Wattier and Maggs, 2001). The southern hemisphere laminarians such as Macrocystis pyrifera have not (yet) adapted to Antarctic habitats (Lüning, 1990). Molecular genetic evidence strongly favours a recent (Pleistocene) movement of the genus Macrocystis to the southern hemisphere (Coyer et al., 2001). The laminarians are evolving rapidly, according to molecular genetic and fossil evidence (Lüning, 1990; de Reviers and Rousseau, 1999).
Marine macroalgal biogeography in the Neogene
Turning to more general considerations of marine macroalgal biogeography in relation to the late Tertiary and Quaternary, Lüning (1990), Clayton (1994) and Dunton (1992) have discussed migration and adaptation of marine macroalgae in the Antarctic and the Arctic. As might be expected from the longer existence of ice and very cold seawater in the Antarctic compared with the Arctic, there is a higher degree of endemism of macroalgae in the South Polar than the North Polar flora, and more Antarctic algae that are unable to grow in cool temperate waters than is the case for Arctic algae (Lüning, 1990; Kirst and Wiencke, 1995; Bischoff-Bäsmann and Wiencke, 1996; Voskoboinikov et al., 1996; Molenaar and Breeman, 1997). In addition to the transequatorial migrations in the Pleistocene glaciations of members of the Phaeophyceae mentioned earlier, there are other transequatorial migrations authenticated by molecular genetic evidence (Van Oppen et al., 1993; Bischoff and Wiencke, 1995; Van Oppen et al., 1995a, b; Knowlton, 2000). Migrations within the high northern latitudes in the Pleistocene have also been indicated by molecular genetic analyses (Van Oppen et al., 1995a, b).
The capacity of algae to undertake the scale of migration discussed above is predicated largely on the constraints of temperature and other environmental factors in the physiological performance of the organisms. Below we consider some of the consequences of low temperature for photoautotrophic marine organisms.
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THE AQUATIC HABITAT |
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TOPABSTRACTINTRODUCTIONEARTH HISTORY AND ALGAL...THE AQUATIC HABITATFUNCTIONING OF ALGAE IN...CONCLUSIONNOTE ADDED IN PROOFLITERATURE CITED |
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Background
In considering the aquatic habitat in relation to low environmental temperatures we focus on features that specifically relate to energetics and resource acquisition. An excellent account of the physics and biology of water in comparison with the physics and biology of air can be found in Denny (1993).
Temperature: generalities
Dealing first with the temperature of the environment and of organisms, the specific heat of water is some four times that of air on a mass basis (J kg–1 K–1) or 4000 times that of air on a volume basis (J m–3 K–1) (Denny, 1993). Other things being equal, this means that terrestrial environments show greater diel and seasonal variations than do aquatic environments. Taking these environmental temperatures (air, soil, sea surface) as a constant in a certain habitat at a given time of year and time of day, what temperature difference can be expected between the organism and its environment?
For phototrophs the major energy input that could lead to a higher temperature of the organism than of the fluid environment is the absorption of electromagnetic radiation that is necessary if photosynthesis is to occur. As is discussed in more detail below, the great majority of absorbed radiation is not used in photochemistry but is converted into thermal energy. The absorbed radiation includes infrared as well as PAR, and the infrared absorption properties of water mean that a submerged alga absorbs less infrared per unit PAR absorbed and so is under less of a thermal load than a land plant. Nevertheless, there is a very significant conversion of electromagnetic radiation into thermal energy relative to photosynthetic energy storage in a seaweed. The very high specific heat (on a volume basis) of water suggests that temperature differences between organisms or their organs and their environment should, other things being equal, be much smaller in water than in air. However, an equally important factor is the thermal conductance between the organism and the well-stirred fluid medium. This thermal conductance (W K–1 expressed on a biomass basis) is a function of thermal conductivity (W m–1 K–1) times the diffusion boundary layer thickness (thermal diffusion in this case). The thermal conductivity of water is 23 times that of air, while the diffusion boundary layer thickness around an organism or organ of a given size in their natural flow regime is about ten times greater in air than in water, so that the thermal conductance of the boundary layer in water is at least 200 times that in air (Denny, 1993; Raven, 1997).
These arguments mean that there is perhaps a 200 times greater heat flux through the diffusion boundary layer of aquatic organisms per K temperature difference than for a plant in air. The greater specific heat of water than of air is only of minor significance in determining temperature differences between organisms and their environments when there is significant forced connection (fluid flow not driven by buoyancy differences in the fluid related to heat transfer from the organism).
Temperature of marine photosynthetic organisms relative to that of their environment
Because of their needs for nutrient and photon acquisition, photosynthetic organisms necessarily have a very high surface area per unit volume, so that any volume-based rate of heat production will have a smaller impact on the temperature excess of the organism over its environment than in an organism with a more compact body. Even unitary endothermic heterotrophs have significantly greater restrictions on minimum size (and, for a given shape, surface area per unit volume) for organisms in water than for organisms in air. This is especially the case for organisms whose whole life cycle is completed in water (cetaceans, sirenians) where the minimum size is that of neonates (contrast penguins and pinnipeds whose juveniles live on land) (Downhower and Blumer, 1988, 1989; Innes and Louigne, 1989; Denny, 1993; Gillooly et al., 2001).
Returning to photosynthetic organisms, respiratory rates of photosynthetic structures are unable to alter significantly the temperature of the organism relative to the fluid environment in air and, even more so, in water (Denny, 1993; Breidenbach et al., 1997; Gillooly et al., 2001). An order of magnitude more energy dissipation and heat production occurs in the light. Photosynthetic gas exchange is typically an order of magnitude greater than respiratory gas exchange, while photosynthesis maximally conserves 37 % of the energy absorbed as photosynthetically active radiation and the global average is less than this (Falkowski and Raven, 1997). At light saturation, much less than 37 % of the energy of the absorbed photons is conserved in photosynthetic products, the rest being dissipated as heat (Falkowski and Raven, 1997). Respiration, even with maximal coupling to ADP phosphorylation, again conserves only about 50 % of the energy from organic C oxidation (Breidenbach et al., 1997). Despite this energy dissipation as heat in the light being at least ten times that in the dark, the temperature increment of the photosynthetic organ in the light is negligible in the aquatic organisms, although it can amount to several degrees in photosynthetic organisms in air (Denny, 1993).
A further energy input to marine benthic photosynthetic organisms is wave energy. Leigh et al. (1987) pointed out that waves breaking on 1 m2 of intertidal shore can focus the solar energy absorbed by air and (especially) water over many square metres of open ocean. In certain intertidal habitats the energy dissipated in breaking waves is ten times that from direct solar energy (including PAR) inputs. Some of the energy dissipated by the breaking waves is transferred by friction between moving water and attached algae from the water to the algae; this stretches the algae (Raven, 1989). The energy stored in the stretched algae is dissipated as heat when the algae return to their resting length in the backwash of the wave (which takes approx. 10 s as opposed to approx. 1 s of breaking). Even this large energy input and dissipation as heat does not significantly increase the temperature of the algae relative to that of the seawater medium or, apparently, lead to energy trapping by the reverse of the mechanochemical coupling leading to movement resulting from ATP hydrolysis (Raven, 1989; Bustamante et al., 2001).
The conclusion from this discussion is that marine macroalgae have to operate at essentially the same temperature as that of the surrounding water. Of course, intertidal macroalgae at low tide are subject to the sort of temperature variations found in terrestrial plants. This is a result of the greater variability in air temperature compared with that of the sea, and the lower thermal conductance of the boundary layer in air than in water which is, in part, offset by the loss of heat as the latent heat of evaporation of water from thalli in air (Jones, 1992; Denny, 1993).
Temperature and inorganic C supply and assimilation
Other aspects of the physics and chemistry of water and low temperatures that are relevant to algal ecophysiology include the equilibrium concentrations of O2 and of inorganic C species at low temperature, and diffusion coefficients of O2 and CO2 at low temperatures (Raven, 1984; Falkowski and Raven, 1997).
The effect of temperature on the concentration of CO2, HCO3– and CO32– in seawater can be computed for a realistic carbonate alkalinity ([HCO3–] + 2 x [CO32–]) of 2·2 mol m–3 and atmospheric partial pressure of CO2 (360 µmol CO2 mol–1 with a total atmospheric pressure of 101·3 kPa), using the CO2 solubility coefficient and pKa1 and pKa2 for the inorganic C system values appropriate to the temperatures concerned (Table 1). The equilibrium conditions assumed here give CO2, HCO3– and CO32– concentrations at 5 °C that are, respectively, 1·9, 1·13 and 0·50 those at 25 °C, with pH being reduced by 0·12 units.
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The equilibrium assumptions of the calculations in Table 1 are not all met in the sea, with the thermohaline circulation, biological activity, and slow equilibration of carbon dioxide between the surface ocean and the atmosphere (half time of hours to days) relative to that between CO2 and the dissolved inorganic species (half time of tens of seconds) (Raven and Falkowski, 1999). However, the general conclusion still stands that the CO2 concentration in cooler waters is higher than that in warmer waters (Beardall and Roberts, 1999). The same is true for O2, where the only consideration (in the absence of biological activity) is O2 solubility, with 193·5 mmol O2 m–3 seawater at 25 °C and 310·8 mmol O2 m–3 at 5 °C. Globally, the assumption of equilibrium of CO2 and O2 between surface seawater and the atmosphere has to be slightly modified to accommodate the excess of organic and inorganic C flux down rivers to the ocean (approx. 0·7 Pg C per year) over the sedimentation of particulate organic and inorganic C (approx. 0·2 Pg per year) (Raven and Falkowski, 1999). The approx. 0·5 Pg of C that is lost each year from the ocean to the atmosphere as CO2 in the steady state, with a molecular equivalent flux of O2 from the atmosphere into the ocean (which must occur if ocean composition is to remain constant year on year) is about 1 % of net marine primary productivity (Raven and Falkowski, 1999).
This discussion of the effects of temperature on the concentration of CO2 in seawater cannot be directly used to suggest a 1·9-fold increase in CO2 availability to Rubisco at 5 °C relative to 25 °C. This is because the bulk seawater concentration of CO2 (mol m–3) is only one factor in the relationship of photosynthetic rate to seawater temperature [Co in eqn (1)]:
where J is the photosynthetic CO2 fixation rate on an organism area basis (mol CO2 m–2 s–1), D is the diffusion coefficient for CO2 in the path from turbulently mixed seawater to Rubisco (m2 s–1), Ci is the CO2 concentration at the site of action of Rubisco in steady-state photosynthesis (mol CO2 m–3) and l is the diffusion path length from turbulently stirred seawater to Rubisco (m). We deal next with the effect of temperature on the value of D in eqn (1).
The diffusion coefficient of CO2 in water is 1·92 x 10–9 m2 s–1 at 25 °C and 1·30 x 10–9 m2 s–1 at 5 °C (Denny, 1993). Similar temperature effects are seen for diffusion coefficients for O2, HCO3– and CO32– (Raven, 1984; Denny, 1993). Comparing these temperature effects on diffusion coefficients with those on CO2 concentration in seawater suggests that diffusive CO2 supply to Rubisco should be little affected by temperature, the enhancing effect of low temperature on seawater CO2 concentration being essentially offset by the lower diffusion coefficient of CO2 at low temperatures. The diffusive CO2 supply situation at low temperatures is exacerbated by the larger diffusion boundary layer thickness expected (other things being equal) at lower temperatures; the boundary layer thickness is another component of 1 in eqn (1). However, the temperature effect on the diffusion boundary layer involves the 1/6 power of the ratio of kinematic viscosities of water at the two temperatures (Denny, 1993). Since this ratio is 1·09 for 25 °C relative to 5 °C, the decrease in CO2 supply by the increase in l [eqn (1)] is relatively small. Rather larger effects might occur for the transmembrane components of the diffusive CO2 flux; however, the significance of this component of the pathway in determining the overall conductance for CO2 depends on the occurrence of CO2-permeable proteinaceous aqueous channels in parallel with the lipid bilayer (Raven et al., 2002; cf. Yang et al., 2000; Sun et al., 2001; Tchernev et al., 2001).
The analyses in the three preceding paragraphs suggest that a lower temperature impacts on eqn (1) by having quantitatively similar but opposite effects on D and on Co, and much smaller effects on l. Relating these effects to the value of J, the area-based rate of photosynthesis, requires that we consider the effects of temperature on Ci, the CO2 concentration around Rubisco during steady-state photosynthesis.
Ci is a function of the CO2 supply rate and the capacity of Rubisco to fix CO2. Maintaining the CO2-saturated rate of photosynthesis on an algal surface area basis (with a constant value of Ci) would require a three to four-fold increase in the Rubisco content on a surface area basis, given the effect of a temperature decrease from 25 °C to 5 °C on the specific reaction rate of Rubisco at CO2 saturation (Raven and Geider, 1989). This argument applies when Ci saturates Rubisco or, if Ci is not saturating, when temperature has no effect on the affinity for CO2 or on Srel, the discrimination factor for CO2 relative to O2. If there is no increase in Rubisco content at low temperatures then the ratio of CO2 supply by diffusion to the capacity for CO2 assimilation by Rubisco will increase, so that Ci will also increase. If Ci at 25 °C were saturating Rubisco, the increased Ci at 5 °C would not increase the activity of Rubisco, but would decrease CO2 entry [eqn (1)]. If Ci is not saturating for Rubisco activity at 25 °C, then the increased Ci at 5 °C will increase Rubisco activity, i.e. the achieved Rubisco activity at 5 °C will be more than one-third to one-quarter that at 25 °C. Of course, an iterative procedure is needed to balance quantitatively CO2 diffusion [eqn (1)] and Rubisco activity as Ci increases, but the basic argument stands.
When Ci is not saturating for Rubisco carboxylase activity then Rubisco oxygenase activity becomes significant. Assuming no temperature effect on Srel (but see below) the higher O2 concentration in air-equilibrium seawater at lower temperatures would, in part, offset the higher bulk phase CO2 concentration in dictating carboxylase activity, and in fact the O2:CO2 ratio at air equilibrium is only 15 % lower at 5 °C compared with 25 °C. The lower rate of O2 evolution would offset the lower diffusion coefficient of O2 at low temperatures (Raven, 1984) in determining the increment of O2 concentration at the site of Rubisco over the O2 concentration in bulk seawater during steady-state O2 evolution. Overall, consideration of O2 diffusion effects with temperature-invariant Srel values slightly reduces the stimulation of photosynthesis when Ci is limiting at low temperatures relative to that expected from the temperature dependence of the CO2-saturated specific reaction rate.
In the case of Ci not saturating Rubisco carboxylase activity, two further enzyme kinetic considerations must be taken into account. One is the higher affinity for CO2 at low temperatures that is exhibited by some Rubiscos (Raven and Geider, 1988; Beardall and Roberts, 1999). This would help to offset the lower specific reaction rate per unit Rubisco at CO2 saturation and increase the rate of photosynthesis at lower temperatures, with a rather lower steady-state Ci than in the case of a temperature-invariant CO2 affinity. The second kinetic consideration is the higher Srel at low temperatures shown by higher plant Rubiscos (Sherlock and Raven, 2001). This higher Srel, like the higher CO2 affinity, would give a higher Rubisco carboxylase activity at 5 °C than predicted from invariant CO2 affinity and Srel with temperature. With the higher Ci values at low temperatures even higher Rubisco activity occurs at 5 °C.
These arguments show that, unless there is a very significant increase in Rubisco content in organisms at low temperatures, there is likely to be a higher transport conductance relative to biochemical conductance at low temperatures for organisms relying on diffusive CO2 entry (Raven and Geider, 1988; Beardall and Roberts, 1999). Of course, the arguments have implications for the relative capacities of carboxylation and of light harvesting and transformation in supplying ATP and NADPH, and for the rate at which CO2 can be fixed per unit nitrogen in the alga (Raven and Geider, 1988). However, there is a prima facie case for a greater diffusive supply of CO2 to Rubisco in seawater at low temperatures.
Further factors that influence Rubisco oxygenase activity in vivo at low temperatures are the combination of the increased external O2 concentration in equilibrium with the atmosphere and the greater diffusive conductivity for O2 efflux relative to the rate of O2 evolution, at low rather than at high concentrations. Raven (1997) discusses, in the context of terrestrial C3 plants, the aqueous phase gradients for O2 (higher in chloroplasts than in the cell wall) and for CO2 (lower in choroplasts than in the cell wall) during photosynthesis. Assuming a net photosynthetic quotient of 1·0, i.e. a flux of O2 out of the leaf cells equal to the flux of CO2 into the leaf cells, the rather higher aqueous diffusion coefficient for O2 than for CO2 may be more than offset in terms of Fick’s law [a version of eqn (1) for O2 efflux] by the role of carbonic anhydrase in facilitating inorganic C diffusion in intracellular compartments. While haemoglobin could, in principle, act in facilitating O2 diffusion within the cells in the same way as one of the roles of carbonic anhydrase in facilitating inorganic C diffusion, there is only enough haemoglobin in plants for this to be significant in diazotrophic nodules of rhizobial and some actinorhizal plants and then only at low (well below air equilibrium) O2 concentrations (Couture et al., 1994; Raven et al., 1999; Thumfort et al., 1999; Watts et al., 2001; Weber and Vinogradov, 2001). These arguments suggest that the O2 concentration difference between chloroplasts and the bulk medium is rather greater than that of CO2 during steady-state photosynthesis in macroalgae relying on diffusive C entry, but that a difference in O2 concentration at 25 °C is unlikely to exceed 10 mmol m–3, and less than this at 5 °C. However, the predominant effect on Rubisco oxygenase activity is a result of the higher O2 solubility at low temperatures. The intra-chloroplast O2 concentration may be 204 mmol m–3 at 25 °C and 317 mmol m–3 at 5 °C (cf. Table 1).
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FUNCTIONING OF ALGAE IN THE COLD: INORGANIC C ACQUISITION |
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Diffusive CO2 supply and carbon concentrating mechanisms
In turning to the evidence on the relation of diffusive CO2 entry in marine macroalgae to the environmental temperature, it should be borne in mind that the majority of marine macroalgae have CO2 concentrating mechanisms (CCMs) (Raven, 1997). These partly overcome the limitations on submerged photosynthesis based on diffusive CO2 entry that are imposed by the combination of bulk phase CO2 concentration, CO2 diffusion coefficient and diffusion boundary layer thickness (Raven, 1984, 1997; Raven and Geider, 1988; Raven and Farquhar, 1990; Denny, 1993). A further limitation on diffusive supply of CO2 in submerged photosynthesis comes from the restrictions of O2 loss from the photosynthesizing structure, again due to the combination of O2 diffusion coefficient and the thickness of the diffusion boundary layer, together with the O2 concentration in bulk seawater (Raven, 1984, 1997; see discussion on Rubisco above).
The criteria by which the occurrence of a CCM can be distinguished from the occurrence of diffusive entry of CO2 are discussed by Raven (1984, 1997), Maberly (1990) and Badger et al. (1998). These criteria include the presence or absence of a higher internal than external CO2 concentration during steady-state photosynthesis, and gas exchange characteristics such as CO2 compensation concentration (as well as the related pH compensation value in solution) and O2 sensitivity of inorganic C assimilation (Kübler et al., 1999; Sherlock and Raven, 2001). It must be acknowledged that some of the data do not immediately fit with CO2 diffusion in some cases where the majority of data favour such a mechanism (Kübler et al., 1999).
Diffusive CO2 entry and carbon concentrating mechanisms in marine macroalgae as a function of temperature
Maberly et al. (1992) found that the occurrence of diffusive CO2 entry in macroalgae, as judged from pH compensation points (but see Badger et al., 1998), occurred in those with low (strongly negative) 
13C values, i.e. less than (more negative than) –30 
. By contrast, algae with CCMs had 13C values higher than –30 . The algae examined by Maberly et al. (1992) were from the east coast of Scotland near Dundee and St Andrews, i.e. in the ‘cool temperate’ marine biogeographic zone (Lüning, 1990). Of the algae examined by Maberly (1990) and divided into those with diffusive CO2 entry and those with CCMs, all of the algae dependent on diffusive CO2 entry were members of the Rhodophyta (Maberly et al., 1992). Most of these were from subtidal habitats, and one was from an intertidal but shaded habitat. In later work, more macroalgae from the British Isles and Norway were examined for 13C/12C ratios (Raven et al., 1995b, 2002); all of the algae with 13C values more negative than –30 were red algae from the class Florideophyceae. Geographically much more wide-ranging measurements of 13C/12C ratios in marine macroalgae revealed more organisms with 13C values below –30 . Of these all were red algae apart from two species of brown algae (Heterokontophyta: Phaeophyceae) and five species of green algae (Chlorophyta: Ulvophyceae) (Raven and Johnston, 1991; Fischer and Weincke, 1992; Raven et al., 1995a, b, 1996, 2002; Raven, 1997; Dunton, 2001). While the low 13C value of one of the green algae (Udotea) could be related to C4-like metabolism based on phosphoenolpyruvate carboxykinase as the initial carboxylase (Raven, 1997), in the other algae diffusive CO2 entry with initial fixation by Rubisco seems to occur; even for Udotea inorganic C could enter by CO2 diffusion.
The occurrence of diffusive CO2 entry in some red algae could be related to the relatively high CO2-affinities and the high Srel values of red algal Rubiscos, combined with their rather low CO2-saturated specific reaction rate (Raven, 1997; Badger et al., 1998; Raven et al., 2000; Sherlock and Raven, 2001). This combination of kinetic characteristics means that a high rate of CO2 fixation per unit of Rubisco can occur in an air equilibrated solution, while the CO2 fixation rate per unit of Rubisco at CO2 saturation (as in an organism with a CCM) is relatively low. However, it must be remembered that the majority of red algae examined have CCMs. Furthermore, the kinetics of brown algal Rubiscos are, based on phylogenetic considerations, probably similar to those of red algal Rubiscos (Raven, 1997; Badger et al., 1998), yet few brown algae have characteristics of diffusive CO2 entry (Raven et al., 2002). Green algal Rubiscos have lower Srel maximal values than do the red algal enzymes, and they can have higher CO2-saturated specific reaction rates which seem more appropriate to the presence of a CCM (Badger et al., 1998).
Returning to the red algae, and using 13C as a proxy for diffusive CO2 entry, it is possible to relate the fraction of red seaweeds with low 13C values to the temperature of the habitats from which they originated (Table 2). The tendency is for the fraction of algae examined that have low (less than –30 ) 13C values to be smaller in warmer habitats, although the number of species sampled, in absolute terms and as a fraction of the total flora, is relatively small. More data are needed before definitive conclusions can be drawn (Table 2).
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The red algae with low
13C values are largely epilithic or epiphytic in subtidal habitats; a few are epilithic in shaded (Lomentaria) or exposed (Stictosiphonia) intertidal habitats, while a few are usually epiphytic (and hence in lower light) in high-intertidal environments (Bostrychia, Caloglossa, Catenella) (Dunton, 1992; Maberly et al., 1992; Raven et al., 1995a, b, 2002; Raven, 1997). Many of these algae grow in relatively low photon flux densities (subtidal representatives, and intertidal Lomentaria) with growth rate limited by photon supply (Kain, 1984, 1987; Kain and Norton, 1990); this would permit diffusive supply of CO2 for photosynthesis from seawater due to the low area-based potential rate of photosynthesis (Raven et al., 2002). From the limited data available, the area-based rates of photosynthesis for the high-intertidal algae with very negative 13C values are also rather low, again permitting diffusive CO2 entry at the observed rate (Raven et al., 2002).
The relatively low area-based rate of photosynthesis in the algae with low 13C values suggests that although they can contribute a significant fraction of the number of red algal species tested (Table 1), they probably do not contribute proportionately to the productivity of the red algal component of the benthic macroflora (Kain and Norton, 1990; Gattuso et al., 1998). Certainly there is no indication from the 13C values of consumer organisms (grazers, decomposers and higher trophic levels) that there are food chains supported solely by these red algae with very negative 13C values (Dunton and Schel, 1987; Fischer and Wiencke, 1992; Dunton, 2001; Smit, 2001; Raven et al., 2002).
As for the evolution of these red algae with low 13C values, it is possible that the lack of a CCM is a derived character in the Florideophyceae as is the absence of pyrenoids which occur in some, but not all, algae with CCMs (Raven, 1997; Raven et al., 2002; cf. Harper and Saunders, 2001). It is not clear when any losses of CCMs took place. If it was in the Tertiary (the last 65 Ma), then the general trend has been for a decrease in atmospheric CO2 and a decrease in global mean temperature (Berner and Kothavala, 2001; Zachos et al., 2001). The two factors of decreased atmospheric CO2 and decreased temperature work in opposite directions in terms of CO2 dissolved in seawater but the decrease in temperature would, by increasing the CO2 affinity and Srel and decreasing the CO2-saturated specific reaction rate of Rubisco activity, work in the same direction as increased CO2 in favouring diffusive CO2 entry over CCMs (Raven et al., 1995b, 2002; Sherlock and Raven, 2001). There are also complications in interpreting the selective factors involved in the evolution, and subsequent selective advantages, of C4 photosynthesis in terrestrial plants (Sage and Monson, 1999; Huang et al., 2001; Sage, 2001).
More generally among seaweeds, the apparent ubiquity of CCMs in the Phaeophyceae is consistent with the ancestral nature of CCMs in these algae. The ancestors (Bangiophyceae: Cyanidiales; Oliveira and Bhattacharya, 2000; Muller et al., 2001) of heterokont plastids have a CCM (Zenvirth et al., 1985) but lack pyrenoids, as do the basal brown algae, as indicated by molecular phylogenetic studies (Draisma et al., 2001; Rousseau et al., 2001). These data suggests that the pyrenoid is a derived and perhaps polyphyletic character in brown algae and that it has also been lost in some clades. However, despite exposure of some brown algae to relatively high CO2 conditions during their evolution, they have apparently evolved diffusive CO2 entry.
Regardless of the importance (or lack of it) of diffusive CO2 entry in the ecophysiology of marine macroalgae at low temperatures, there are many other factors involved in their capacity to grow at low temperatures with extremes of day length and such problems as ice scour and ice encasement (Crawford, 1989; Kain and Norton, 1990; Lüning, 1990; Kirst and Wiencke, 1995; Henley and Dunton, 1997; Raven and Scrimgeour, 1997; Fogg, 1998).
The capacity to survive the vicissitudes of life in cold seawater seems to be as pronounced in the brown algae, which have had only this recent experience of cold in their (probably) 250 million years of existence, as in the much more ancient red and green algae which have experienced several major glacial episodes.
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CONCLUSION |
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TOPABSTRACTINTRODUCTIONEARTH HISTORY AND ALGAL...THE AQUATIC HABITATFUNCTIONING OF ALGAE IN...CONCLUSIONNOTE ADDED IN PROOFLITERATURE CITED |
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Over most of their evolutionary history, marine macroalgae have been subjected to higher mean sea surface temperatures than pertain today, and will relatively rarely have been exposed to surface waters cooler than 5 °C. The Pleistocene glaciations and the preceding cooling during the late Tertiary provided these low sea surface temperatures at high latitudes, and permitted trans-tropical algal migrations when tropical surface seawater was relatively cool during glaciations. Low sea surface temperatures are predicted to give diffusive CO2 entry more competitive ability than at higher temperatures in comparison with organisms with CCMs. Some evidence on the occurrence of diffusive CO2 entry is in agreement with this prediction.
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ACKNOWLEDGEMENTS |
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We are grateful to Dr D. N. Thomas for provision of preprints, and to Dr M. C. F. Proctor for helpful comments. Work in J.A.R.’s laboratory on macroalgal photosynthesis has been supported by the Natural Environment Research Council.
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NOTE ADDED IN PROOF |
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TOPABSTRACTINTRODUCTIONEARTH HISTORY AND ALGAL...THE AQUATIC HABITATFUNCTIONING OF ALGAE IN...CONCLUSIONNOTE ADDED IN PROOFLITERATURE CITED |
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The paper by van Zullen et al. (2002) brings into question the evidence cited under Earth History for the occurrence of life using Rubisco to fix CO2 as much as 3·8 Ga ago.
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LITERATURE CITED |
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-
Abbott IA, Hollenberg GJ. 1976. Marine algae of California. Stanford, CA: Stanford University Press.
Adams NM. 1994. Seaweeds of New Zealand – an illustrated guide. Canterbury, NZ: Canterbury University Press.
Badger MR, Andrews TJ, Whitney SM, Ludwig M, Yellowlees DC, Leggat W, Price G. 1998. The diversity and co-evolution of Rubisco, plastids, pyrenoids and chloroplast-based CO2 con centrating mechanisms in the algae. Canadian Journal of Botany 76: 1052–1071.
Beardall J, Roberts S. 1999. Inorganic carbon acquisition by two Antarctic macroalgae, Porphyra endivifolia (Rhodophyta: Bangiales) and Palmaria decipiens (Rhodophyta: Palmariales). Polar Biology 21: 310–315.[CrossRef]
Berner RA, Kothavala Z. 2001. GEOCARBIII: a revised model of atmospheric CO2 over Phanerozoic time. American Journal of Science 301: 182–204.
Bischoff B, Wiencke C. 1995. Temperature adaptation in strains of the amphi-equatorial green alga Urospora penicilliformis (Acro siphonales). Marine Biology 122: 681–688.[CrossRef]
Bischoff-Bäsmann B, Wiencke C. 1996. Temperature requirements for growth and survival of Antarctic Rhodophyta. Journal of Phycology 32: 525–535.[CrossRef][ISI]
Black CC Jr, Bender MM. 1976. 13C values in marine organisms from the Great Barrier Reef. Australian Journal of Plant Physiology 3: 25–32.
Boyd PW, Watson AJ, Law CS, Abraham ER, Trull T, Murdoch R, Bakker R, Bowie AR, Buessler KO, Chang H, et al. 2000. A measurable phytoplankton bloom in the Polar Southern Ocean stimulated by iron fertilization. Nature 407: 695–702.[CrossRef][Medline]
Brasier MD, Green OR, Jephcoat AP, Kleppe AK, van Kranendonk MJ, Lindsay JF, Steele A, Grassineau NV. 2002. Questioning the evidence for Earth’s earliest fossils. Nature 2002: 76–81.
Breidenbach RW, Saxton MJ, Hansen LD, Criddle RS. 1997. Heat generation and dissipation in plants: can the alternative oxidative phosphorylation pathway serve a thermoregulatory role in plant tissue other than specialized organs? Plant Physiology 114: 1137–1140.[CrossRef][ISI][Medline]
Brierley AS, Thomas DN. 2002. On the ecology of Southern Ocean pack ice. Advances in Marine Biology (in press).
Brocks JJ, Logan GA, Buick R, Summons RE. 1999. Archean molecular fossils and the early rise of eukaryotes. Science 285: 1033–1036.
Brouwer PEM, Geilen EFM, Gremer NJM, van Lent F. 1995. Biomass, cover and zonation pattern of sublittoral macroalgae at Signy Island, South Orkney Islands, Antarctica. Botanica Marina 38: 259–270.
Bustamante C, Keller D, Oster G. 2001. The physics of molecular motors. Accounts of Chemical Research 34: 412–420.[CrossRef][ISI][Medline]
Butterfield NJ. 2000. Bangiomorpha pubescens n.gen, n.sp.: implications for the evolution of sex, multicellularity, and the Mesoproterozoic/Neoproterozoic radiation of eukaryotes. Palaeobiology 26: 386–404.
Butterfield NJ. 2001. Palaeobiology of the late Mesoproterozoic (ca. 1200 Ma) Hunting Formation, Somerset Island, Arctic Canada. Precambrian Research 111: 235–256.[CrossRef][ISI]
Catling DC, Zahnle KJ, McKay CP. 2001. Biogenic methane, hydrogen escape, and the irreversible oxidation of early earth. Science 293: 839–843.
Clayton MN. 1984. Evolution of the Phaeophyta with particular reference to the Fucales. Progress in Phycological Research 3: 11–46.
Clayton MN. 1994. Evolution of the Antarctic benthic algal flora. Journal of Phycology 30: 897–904.[CrossRef][ISI]
Couture M, Chamberland H, St Pierre B, Lafontaine J, Guertin M. 1994. Nuclear genes encoding chloroplast hemoglobins in the unicellular green alga Chlamydomonas eugametos. Molecular and General Genetics 243: 185–197.
Coyer JA, Smith GJ, Andersen RA. 2001. Evolution of Macrocystis spp. (Phaeophyceae) as determined by ITS1 and ITS2 sequences. Journal of Phycology 37: 574–585.[CrossRef][ISI]
Crawford RMM. 1989. Studies in plant survival. Ecological case histories of plant adaptation to adversity. Oxford: Blackwell Scientific Publications.
Cribb AB. 1996. Seaweeds of Queensland. Handbook No. 2 of the Queensland Naturalist Club. Brisbane: The Queensland Naturalists’ Club Inc.
Denny MW. 1993. Air and water. The biology and physics of life’s media. Princeton, NJ: Princeton University Press.
de Reviers B, Rousseau F. 1999. Towards a new classification of the brown algae. Progress in Phycological Research 13: 107–201.
Downhower JF, Blumer LS. 1988. Calculating just how small a whale can be. Nature 335: 675.
Downhower JF, Blumer LS. 1989. Size of aquatic endotherms. Nature 341: 192.
Draisma SGA, Van Reine WFP, Stam WT, Olsen JL. 2001. A reassessment of phylogenetic relationships within the Phaeophyceae based on RUBISCO large subunit and ribosomal DNA sequences. Journal of Phycology 37: 586–603.[CrossRef][ISI]
Dunton KH. 1992. Arctic biogeography: the paradox of the marine benthic fauna and flora. Trends in Ecology and Evolution 7: 183–189.[CrossRef]
Dunton KH. 2001. 15N and 13C measurements of Antarctic peninsula fauna: trophic relations and assimilation of benthic seaweeds. American Zoologist 41: 99–112.[CrossRef]
Dunton KH, Schel DM. 1987. Dependence of consumers on macroalgal (Laminaria solidungula) carbon in an Arctic kelp community: 13C evidence. Marine Biology 93: 615–625.[CrossRef]
Falkowski PG, Raven JA. 1997. Aquatic photosynthesis. Malden, MA: Blackwell Science.
Fischer G, Wiencke C. 1992. Stable carbon isotope composition, depth distribution and fate of macroalgae from the Antarctic Peninsula region. Polar Biology 12: 341–348.
Fogg GE. 1998. The biology of polar habitats. Oxford: Oxford University Press.
Gattuso J-P, Frankignoulle M, Wollast R. 1998. Carbon and carbonate metabolism in coastal aquatic ecosystems. Annual Review of Ecology and Systematics 29: 405–434.[CrossRef][ISI]
Gillooly JF, Brown JH, West GB, Savage VM, Charnov EL. 2001. Effects of size and temperature on metabolic rates. Science 293: 2248–2251.
Golubic S, Seong-Joo L. 1999. Early cyanobacterial fossil record: preservation, palaeoenvironments and identification. European Journal of Phycology 34: 339–348.[CrossRef]
Harper JT, Saunders GW. 2001. Molecular systematics of the Florideophyceae (Rhodophyta) using nuclear large and small subunit rDNA sequence data. Journal of Phycology 37: 1073–1082.[CrossRef][ISI]
Heckman DS, Geiser DM, Eldel BR, Stauffer RL, Kardos NL, Hedges SB. 2001. Molecular evidence for the early colonization of land by fungi and plants. Science 293: 1129–1133.
Henley WJ, Dunton KH. 1997. Effects of nitrogen supply and continuous darkness on growth and photosynthesis of the Arctic kelp Laminaria solidungula. Limnology and Oceanography 42: 209–216.
Hoehler TM, Bebout BM, Des Marais DJ. 2001. The role of microbial mats in the production of reduced gases on the early Earth. Nature 412: 324–327.[CrossRef][Medline]
Hoffman PF, Schrag DP. 1999. ‘Snowball Earth’ theory still stands. Nature 400: 708.
Hoffman PF, Kaufman AJ, Halverson GP, Schrag DP. 1998. A Neoproterozoic Snowball Earth. Science 281: 1342–1346.
Huang Y, Street-Perrott FA, Metcalfe SE, Brenner M, Moreland M, Freeman KH. 2001. Climate change as the dominant control on glacial-interglacial variations in C3 and C4 plant abundance. Science 293: 1647–1651
Huisman J, Walker DI. 1990. A catalogue of the marine plants of Rotnest Island, Western Australia, with notes on their distribution and biogeography. Kingia 1: 349–549.
Innes S, Louigne DM. 1989. Size of aquatic endotherms. Nature 341: 192.
Jones HG. 1992. Plants and microclimate. A quantitative approach to environmental plant physiology. 2nd edn. Cambridge: Cambridge University Press.
Kain JM. 1984. Seasonal growth of two sublittoral species of Rhodophyta off the Isle of Man. Journal of Experimental Marine Biology and Ecology 82: 207–220.[CrossRef]
Kain JM. 1987. Seasonal growth and photoinhibition of Plocamium cartilagineum (Rhodophyta) off the Isle of Man. Phycologia 26: 88–99.
Kain JM, Norton TA. 1990. Marine ecology. In: Cole KM, Sheath RG, eds. Biology of the red algae. Cambridge: Cambridge University Press, 377–422.
Kasting JF. 2001. The rise of atmospheric oxygen. Science 293: 819–820.
Kirshvink J, Gaidos EJ, Bertani LE, Beukes NJ, Gutzher J, Maepa LN, Steinberger RE. 2000. Paleoproterozoic Snowball Earth: extreme climatic and geochemical global change and its biological consequences. Proceedings of the National Academy of Sciences of the USA 97: 1400–1405.
Kirst GO, Wiencke C. 1995. Ecophysiology of polar algae. Journal of Phycology 31: 181–199.[CrossRef][ISI]
Knowlton N. 2000. Molecular genetic analyses of species boundaries in the sea. Hydrobiologia 420: 73–90.[CrossRef][ISI]
Kübler JE, Johnston AM, Raven JA. 1999. The effects of reduced and elevated CO2 on the seaweed Lomentaria articulata. Plant, Cell and Environment 22: 1303–1310.