Algal Calcification and
Silification
Secondary article
Article Contents
. Calcification
Colin Brownlee, Marine Biological Association of the UK, Plymouth, UK
Alison R Taylor, Marine Biological Association of the UK, Plymouth, UK
. Silification
The algae represent major producers of calcium carbonate and silica among the world’s
biota. Calcification involves the precipitation of CaCO3 from Ca2 1 and CO32 ions. Algal
calcification may account for up to half of global oceanic CaCO3 production. Silicification is
less widespread among algal groups, which transform dissolved silicate to skeletal material.
Diatoms play a key role in marine silica cycling. Diatomaceous deposits have long been
exploited for building and filling materials, and the low-temperature, low-pressure
biogenic formation of silica has potential for biotechnological application in novel
industrial processes.
Calcification
Overview
Calcification is widespread among plant, animal, algal and
prokaryotic groups and occurs with varying degrees of
sophistication and complexity. Calcification involves the
precipitation of CaCO3 from Ca2 1 and CO23 2 ions in
solution. In most cases this involves the generation of
microenvironments that allow supersaturation of CaCO3.
Although many terrestrial and freshwater organisms are
able to produce CaCO3, most of the world’s CaCO3 is
produced in the oceans. Calcification has been widespread
among the biota since at least the Cambrian era and
probably evolved considerably earlier than this. Various
explanations exist for the origin of calcification. As Ca2 1
levels increased in the early oceans, the precipitation of
phosphate by Ca2 1 inside cells would become problematic, leading to selection of organisms that were able to
extrude Ca2 1 from inside their cells to the external
medium (Degens and Itterkkot, 1986). Grazing pressure
and competition would also provide evolutionary driving
forces for the development of biomineralized structures for
defence and structural support of increasingly complex
multicellular organisms.
The most striking feature of calcification in the oceans is
that it occurs almost exclusively by biogenic processes.
Ca2 1 and CO23 2 inputs into the oceans occur through
weathering of rocks, geothermal activity and hydrothermal seepage (Westbroek et al., 1993). The ocean is in fact
supersaturated with Ca2 1 and CO23 2 . However, spontaneous chemical precipitation is largely absent owing to the
high concentrations of inhibitory ions in seawater and the
presence of crystal poisons produced by the biota.
Among the algae, calcification can be found in both
freshwater and marine species. The brackish-water giantcelled Charophyte alga Chara produces bands of CaCO3
along the length of its surface and has provided a good
model for the study of the transport processes involved in
external calcification (see Figure 1). A variety of marine
multicellular macrophyte algae also produce CaCO3,
including Corallina spp. and Halimeda spp. However, the
most abundant calcifying algae are the free-living unicellular members of the Haptophyte division collectively
known as the coccolithophores (see Figure 2). These are
primarily marine and occur in all of the world’s oceans,
sometimes forming vast monospecific blooms under
appropriate conditions.
Geological and economic importance
Marine biogenic calcification and its connection to
photosynthesis is an important though little understood
process in the global carbon cycle. Three major classes of
organisms are responsible for the bulk of oceanic
calcification. These are the corals, foraminifera and the
coccolithophorid phytoplankton. Many species of coral
and forams form symbiotic associations with photosynthetic dinoflagellate algae (zooxanthellae). Although these
symbiotic algae have been shown in many studies to
enhance the process of calcification, they do not have a
direct role in CaCO3 production, which is carried out by
specialized cells of the host species. In terms of total CaCO3
production, the pelagic foraminifera and coccolithophores
together probably represent the bulk of modern-day global
CaCO3 production. While accurate estimates of the
relative contributions of these organisms are lacking, it is
likely that the coccolithophores represent up to half of all
current global oceanic CaCO3 production.
The oceans represent a significant sink for atmospheric
CO2, removing approximately 30% of anthropogenic CO2
emissions. In this context oceanic calcification plays a
significant role in the formation of a sink for inorganic
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1
Algal Calcification and Silification
Ca2+
(a)
(b)
CaCO3
CaCO3
CO2
Ca2+
Alkaline
H+
CO2 1 H2OKH2CO3KHCO32 1 H 1 KCO23 2 1 2H 1
[I]
H+
A significant, and perhaps surprising, feature of oceanic
calcification is that it is potentially a net producer of CO2,
whereas dissolution of CaCO3 during rock weathering
comprises a sink for atmospheric CO2 levels (eqn [II])
HCO3
H+
H+
HCO3
Acid
CO2
CO2
HCO3
Chloroplast
C
Out
carbon. Bicarbonate comprises more than 90% of the total
dissolved inorganic carbon (DIC) in the oceans and most
evidence suggests that this is the main carbon source for
calcification. Marine photosynthesis and respiration are
linked to calcification and CaCO3 dissolution, respectively,
with significant effects on the alkalinity and pH of the
oceans. DIC speciation occurs according to the reaction [I].
V
In
C
Out
Figure 1 Alternative models for the mechanism of external calcifying
band production in the giant-celled alga Chara. In (a), calcification is driven
by the diffusion of H 1 into the cell in localized regions, producing localized
alkalinization of the cell surface and precipitation of CaCO3. In an adjacent
region, H 1 is actively pumped out of the cell producing localized
acidification of the cell surface. In (b), calcification is driven by the extrusion
of Ca2 1 in exchange for H 1 in the alkaline zone. CO2 diffusion from the
cell’s interior provides the carbon source for CO23 2 formation and CaCO3
precipitation. In both models, H 1 extrusion in the acidic zone facilitates the
production of CO2 from HCO32 which can be used by photosynthesis in the
chloroplasts (green).
Ca2+ + 2HCO3–
Calcification
CaCO3 + H2O + CO2
CaCO3 dissolution
[II]
Photosynthesis requires CO2 and it is not surprising that the
major calcifiers have direct associations with photosynthesis.
CaCO3 production is largely limited to the upper photic zone
of the ocean, while particulate inorganic carbon in the form
of CaCO3 sinks to the deep ocean. The sinking is facilitated
by zooplankton grazing and the formation of faecal pellets.
In deeper waters, below the lysocline up to 90% of this
CaCO3 gradually redissolves. The remainder contributes to
vast calcareous sediments. Over much longer geological time
scales ( 100 My) tectonic activity results in the subduction
of sediments, and release of CO2 during volcanic activity or
uplifting of sediments and weathering and dissolution of
carbonates and recycling of DIC back to the oceans
(Varekamp et al., 1992).
The preservation of coccoliths in the sedimentary record
has led to their use as palaeoclimate proxies. Coccolithophores also produce a wide range of long-chain alkene and
alkenone hydrocarbons that are also preserved in the
Figure 2 Model for fluxes of Ca2 1 , HCO32 and H 1 during calcification in coccolithophores. (a) Scanning electron micrograph of heterococcoliths on
the surface of Coccolithus pelagicus cells (cell diameter=20 mm). (b) Ca2 1 and HCO32 uptake into a Golgi-derived compartment leads to the production
of CaCO3 and H 1 during calcite precipitation. H 1 production can be used to counter the alkalinizing effect of CO2 production from HCO32 in the
chloroplast (green) and removal of CO2 by photosynthesis.
2
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Algal Calcification and Silification
sedimentary record (Grossi et al., 2000). Good evidence
exists that the alkenone composition of coccolithophore
populations varies with sea surface temperature, allowing
studies of past coccolithophore abundance and palaeotemperature. Despite the obvious widespread uses of
chalk, limestone and quartz in construction, chemical
and other industries, the biotechnological applications of
coccolithophores have been largely unexplored. Knowledge of the molecular processes involved in the production
of highly ordered crystal structures such as coccoliths will
undoubtedly improve understanding of the control of
microscale ordered crystal growth.
Mechanisms
Different groups of algae form CaCO3 in strikingly
different ways, and the cellular sites of calcification can
be either internal or external, suggesting that calcification
mechanisms evolved independently in different groups.
The partial reaction of calcification (eqn [III]) indicates a
potential physiological role for calcification in providing a
source of H 1 that may be used in the production of CO2
from HCO32 (McConnaughey and Whelan, 1997) or for
any other H 1 -requiring process.
HCO32 KH 1 1 CO23 2
[III]
In the marine environment, CO2 is present in micromolar
concentrations, which may be lower than the dissociation
constant (Km) for the photosynthetic carbon-fixing enzyme
ribulose bisphosphate carboxylase oxygenase (Rubisco).
Phytoplankton species that rely on CO2 diffusion to supply
photosynthesis may be rate-limited by the CO2 concentration and diffusion to the site of Rubisco in the chloroplast.
Several species of aquatic algae appear to have evolved
mechanisms that allow utilization of external HCO32 as
the external carbon substrate for photosynthesis (Raven,
1997). If taken in to the cell, this can be used as a source of
CO2, facilitated by the action of carbonic anhydrase (CA)
in the chloroplast (eqn [IV]).
HCO32 KCO2 1 OH
[IV]
However, this process may become limited by the
availability of H 1 , which will be needed to maintain
cellular pH homeostasis.
The giant-celled brackish-water alga Chara deposits
distinct bands of external CaCO3 along its length. These
correspond to extracellular alkaline regions that are
separated from one another by acidic bands (Figure 1).
Two hypotheses have been forwarded for the mechanism
of banding and calcification in Chara (McConnaughey and
Whelan, 1997). Each mechanism clearly involves the
spatial separation of H 1 extrusion regions from those
involved in H 1 uptake. In both models H 1 ions produced
during calcification are ultimately used to generate CO2
from HCO32 . Calcifying macroalgae such as Halimeda and
Corallina also produce CaCO3 in alkaline extracellular
spaces. However, the spatial separation of H 1 uptake and
efflux into different zones is less clear in these organisms.
Coccolithophorid phytoplankton produce CaCO3 in the
form of elaborate crystalline structures known as coccoliths (Figure 2). Two distinct classes of coccoliths have been
identified. Holococcoliths are simple single-crystalline
structures, while heterococcoliths are more complex multicomponent structures (Young et al., 1999). Current
evidence suggests that holococcoliths are formed on the
external cell surface while heterococcoliths are produced in
intracellular compartments (coccolith vesicle) (Young
et al., 1999). Certain species (e.g. Coccolithus pelagicus)
have both hetrococcolith-and holococcolith-producing
phases, while others (e.g. Emiliania huxleyi) produce only
heterococcoliths. Various functions have been assigned to
coccoliths, ranging from protection against grazing,
regulation of depth in the water column, or modification
of optical properties to favour photosynthesis. However,
the frequent correlation of calcification rates with photosynthesis and the often observed continual production and
shedding of coccoliths suggests that coccolithophore
calcification production may have at least a partial
metabolic role.
Heterococcoliths develop inside intracellular vesicles
derived from the Golgi cisternae. Calcification occurs in a
highly ordered manner from a protein base plate (Young
et al., 1999). As the coccolith matures, the vesicle moves
towards the cell periphery and the mature coccolith is
extruded in a single exocytotic event and attaches to the
external cell surface. Estimates of coccolith formation in E.
huxleyi have shown that approximately one coccolith can
be formed per hour (Paasche, 1964). Under conditions that
promote high rates of calcification, carbon is fixed into
CaCO3 at the same molar rate as in photosynthetic carbon
fixation, i.e. the ratio of calcification to photosynthesis is
unity (Paasche, 1964). Figure 2 summarizes current knowledge of the main fluxes of carbon and Ca2 1 in
coccolithophore calcification. The precise routes of uptake
of DIC and Ca2 1 are still unresolved. Uptake of HCO32 to
the coccolith vesicle and chloroplast can potentially occur
passively without expenditure of cellular energy. However,
the uptake of Ca2 1 into the coccolith vesicle may be
subject to considerable energetic barriers, particularly if
significant fluxes of Ca2 1 occur through the cytoplasm en
route to the coccolith vesicle, since the cytoplasmic Ca2 1
concentration is kept extremely low (around
100 nmol L 2 1) in all eukaryotic cells. This energetic barrier
could be overcome by utilization of alternative transport
routes and/or by binding of Ca2 1 to polysaccharides or
proteins. While Ca2 1 -binding proteins and polysaccharides have been characterized in some detail in E. huxleyi
and Pleurochrysis carterii (e.g. Corstjens et al., 1998), these
so far appear to be involved in the regulation of CaCO3
precipitation within the coccolith vesicle rather than the
influx of Ca2 1 to this organelle. So far no transport protein
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3
Algal Calcification and Silification
has been specifically identified that is exclusively involved
in calcification in coccolithophores. Characterization of
the pathways and transporters for DIC and Ca2 1 is
eagerly awaited.
Silification
Overview
Silica is the second most abundant element in the earth’s
crust. It is surprising, therefore, that few algal groups
utilize silicon to any great extent. Of these, all transform
dissolved silicate to particulate skeletal material. Cell wall
silica deposition has been recorded in some marine
macroalgae. However, unicellular algae are responsible
for the majority of biogenic silicification. Only four groups
of these phytoplankton produce silicified skeletal material:
silicoflagellates, chrysophyta, xanthrophyta and – by far
the most significant producer of biogenic silica – the
bacillariophyta or diatoms. Diatoms contribute approximately 25% of world net primary productivity and up to
50% of marine primary productivity, reflected by their
dominance in productive ocean upwelling and shelf areas.
This in turn dominates the marine silicon cycle. The archive
of diatom frustules found in sedimentary deposits has
important applications in palaeogeology and hydrocarbon
exploration, and the sedimentary deposits themselves are
an important raw material for industrial use. The study of
the mechanisms of biogenic silica deposition is likely to
reveal biotechnological applications in the future and is
therefore a critical area of future research.
Geological and economic importance
Diatoms play a key role in marine silicon cycling (Tréguer
et al., 1995) and, because diatom growth has an absolute
requirement for silicate, Si(OH)24 2 availability is critical in
determining primary productivity and therefore CO2
fluxes in large areas of ocean (Dugdale and Wilkerson,
1998). It is estimated that over 30 million km2 of ocean
floor are covered with sedimentary deposits of diatom
shells. Diatomite is a low-density and highly porous
opaline silica sedimentary rock formed by the compaction
of diatom frustule deposits between 50 and 80 million years
ago in both marine and freshwater locations. Some
commercial deposits of diatomite may contain up to 90%
SiO2 and have been utilized in many industrial applications. The high porosity, low permeability, high stability
and chemical inertness of diatomite have been exploited for
centuries in its use as a building material, filler and additive
to mortar. Diatomite continues to be used today as a fine
filler in ceramics and construction materials. Alfred Nobel
implemented one of the earliest commercial applications of
diatomite with his discovery that highly explosive nitro4
glycerine could be stabilized by absorption into diatomite,
thereby producing dynamite. While the discovery of
dynamite played a key role in the industrial revolution, it
has largely been superseded by modern explosives.
Given that industrial synthesis of silica-based materials
requires high temperature and pressure, the biogenic
formation of silica under ambient conditions has great
potential for biotechnological applications such as controlled nanoscale crystal formation, large-scale silica
harvesting and novel means of industrial silica production.
While research is gathering momentum in these applications, naturally occurring diatomite and diatomaceous
earth continue to be an important raw material for
biotechnology and filtration applications, particularly in
the water purification and brewing industry. A more recent
application has been in pest control during post-harvest
storage of cereals, when treatment of cereals with nontoxic
diatomaceous powder appears to cause a high mortality in
a number of food-spoilage beetles.
A robust taxonomic key of the diatoms is possible
because of the unique and stable silica frustule morphotypes of diatom species. This is particularly useful in the
analysis of cores from diatomaceous deposits, which
provide palaeoenvironmental and geochronological information. The silica microfossils within sediment samples
can be assessed for the species composition and diversity.
Such profiles can be used to reconstruct the environmental
conditions prevailing at the time of deposition. Particular
examples include the use of diatom microfossils in
extrapolation of quaternary climate and hydrology from
lake cores in East Africa (Gasse et al., 1996). Species
composition is directly related to lakes’ patterns of level
and salinity and consequently reflects long-term climate
variation. While the majority of these studies deal with
long-term changes, there is scope for shorter, decadal,
reconstructions of diatom-inferred salinity to provide a
detailed view of drought conditions; an example is Moon
Lake in North Dakota (Laird et al., 1996). Stable-isotope
fractionation by diatoms during biogenic silica formation
has been exploited to reconstruct primary productivity
from ocean sediment cores (DeLa Rocha et al., 1998).
Core profiles from diatomaceous deposits are also
crucial for predictive hydrocarbon exploration. Characteristic diatoms from different cores can be correlated with
biostratigraphy. Heavy oil and gas deposits are also
frequently associated with such deposits.
Mechanism
Silicon is an essential element for diatom growth. Silicon
availability in aquatic habitats varies considerably, largely
as a product of rock weathering, river inputs and
dissolution of biogenic silica sediments. The most dominant soluble form for biogenic silification is monomeric
undissociated silicic acid, Si(OH)4, although the less
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Algal Calcification and Silification
abundant anion (SiO(OH)32 ) may be utilized by some
diatom species. In the euphotic zone of aquatic ecosystems,
silicic acid is present in concentrations ranging from
submicromolar up to 200 mmol L 2 1 with 1–20 mmol L 2 1
being common.
While uptake of Si(OH)4 can occur passively across the
lipid membrane, active transport is necessary to satisfy the
demands of silification. Active uptake of silicic acid is
implied by studies of tracer uptake and effects of metabolic
inhibitors. Carrier-mediated silicon uptake is indicated for
a wide range of diatom species fitting Michelis–Menten
kinetics. Typically the Km for Si(OH)4 lies in the range 0.3
to 4 mmol L 2 1 for marine diatoms. The observation that
dissipation of the Na 1 gradient across the plasma
membrane inhibits Si(OH)4 uptake supports a model for
secondary active transport in marine diatoms whereby
silicon uptake is coupled to a favourable Na 1 gradient.
Silicon uptake is closely governed by metabolic demand
and cell cycle, with maximal rates associated with periods
of cell wall deposition. Uptake rates are also variable
depending on the availability of external silicon, with
maximal surge uptake occurring on reintroduction of a
silicon source after a period of silicon starvation. While
silica deposition and cell division can occur close to
maximal rates at silicon concentrations that are limiting for
silicon uptake, diatoms can accumulate soluble silicon
between 20 and 400 mmol L 2 1, exceeding the solubility for
silicon and in some cases enough to synthesize an entire
frustule without further silicon input. While storage in
specialized vesicles or by association with intracellular
organic silicon-binding components has been proposed,
the mechanism for maintaining these supersaturated
internal pools of silicon is unclear. Likewise, the internal
silicon pool size is known to be influenced by a number of
metabolic and environmental variables, yet the precise
mechanism underlying this regulation remains unknown.
Recent progress has been made in the cloning and
functional expression of cDNAs coding diatom silicic acid
transporters (Hildebrand et al., 1997). The clones derived
from Cylindrotheca fusiformis can be heterologously
expressed in Xenopus oocytes, where they enable Na 1 dependent uptake of germanium-68 (a silicon analogue
tracer). These clones encode the diatom SIT gene family of
unique conserved integral membrane domains and less
conserved carboxyl-terminal portion, indicating that they
most likely encode silicic acid transporters with different
cellular locations and/or affinities. This is supported by the
fact that different SITs exhibit different levels of mRNA
abundance during silification and cell division. Overall,
however, SIT gene expression patterns during silification
and cell division in Cylindrotheca fusiformis correlate well
with transport activity characterized in other diatoms.
Removal of silicon allows diatom cells to reach maturity,
but they fail to undergo cell division. Synthesis of both
protein and DNA is inhibited within a few hours of silicon
starvation, whereas photosynthesis and glycolysis are only
slightly reduced during this time. Silicon-starved cells
increase production of lipid, which appears as oil droplets.
Resupply of Si(OH)4 triggers a rapid and specific increase
in DNA polymerase and cells resume cell division. Silicon
starvation has been used as a tool to synchronize diatom
cultures.
Polycondensation of Si(OH)4 to form a cell wall is
energetically favourable compared to synthesis of cell wall
cellulose or to calcification. It has been estimated that 1
ATP is required for every silicon taken up and deposited,
Figure 3 (a) Scanning electron micrograph showing complete silica frustules of the diatom Thallassiosira eccentrica (cell diameter 5 50 mm). (b)
Working model of silica biogenesis in a diatom. Na 1 -dependent silicate uptake is mediated by specialized silica transporters (SIT) at the plasma
membrane. Once inside the cell, the soluble silica is delivered to the silica deposition vesicle (SDV), possibly involving isoforms of the SITs. Silica
polycondensation occurs under acidic conditions within the SDV to form insoluble amorphous silica. Soluble silica may be stored in intracellular pools,
the size of which is sensitive to silica demand and metabolism. Silifin proteins are thought to play a key role in regulating the formation of insoluble
silica within the SDV. Frustulins are a separate group of proteins that together with polysaccharides and lipids are likely to play a role in stabilization of the
mature silica wall structure. On maturity, the complete valve is released onto the cell surface by exocytosis.
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5
Algal Calcification and Silification
which represents only 2% of the cellular energy budget to
produce 20% of cell dry weight (Raven, 1983). This
compares very favourably to calcification (see above) and
is likely to have been a major evolutionary advantage.
While the morphology and ultrastructure of silica
deposition have been described in detail for a number of
diatom species, little is known about the mechanisms of
deposition at the molecular level. In an analogous way to
coccolithophorid calcification (see above), silica polymerization occurs intracellularly within a specialized membrane-bound compartment, known as the silica deposition
vesicle (SDV), that lies close to the plasma membrane. The
membrane of the SDV, the silicalemma, extends with, and
remains closely attached to, the mineralized structure
within. The SDV is an acidic compartment, which favours
mineralization and prevents dissolution of the newly
formed frustule (Figure 3).
The most promising recent work has focused on the
organic components of the amorphous silica walls of
diatoms. Some of these appear to have a regulatory role in
silica deposition and wall patterning and others promote
stability of the mature wall. Diatom cell walls are rich in
hydroxylated amino acids that may play a role in
promoting silica polymerization. A group of polycationic
peptides (silaffins), in tandem with long-chain polyamines,
that rapidly induce silica nanosphere formation in vitro are
thought to play a key role in directing silica deposition
inside the acidic SDV (Kroger et al., 1999). Additionally, a
group of Ca2 1 -binding glycoproteins (frustulins) that is
associated with mature valves may act as a protective layer
to retard dissolution in the slightly basic, undersaturated
extracellular environment.
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ENCYCLOPEDIA OF LIFE SCIENCES / & 2002 Macmillan Publishers Ltd, Nature Publishing Group / www.els.net