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Background
Section 1
Limnol. Oceanogr., 36(8), 1991, 1507-1526
0 1991, by the American
Society
of Limnology
and Oceanography,
Inc.
Role of the marine biosphere in the global carbon cycle
Alan R. Longhurst
Biological Oceanography Division, Bedford Institute
P.O. Box 1006, Dartmouth, Nova Scotia B2Y 4A2
of Oceanography,
Department
of Fisheries and Oceans,
Abstract
The geochemical disequilibrium
of our planet is due mainly to carbon sequestration by marine
organisms over geological time. Changes in atmospheric CO, during interglacial-glacial
transitions
require biological sequestration of carbon in the oceans. Nutrient-limited
export flux from new
production in surface waters is the key process in this sequestration. The most common model for
export flux ignores potentially important nutrient sources and export mechanisms. Export flux
occurs as a result of biological processes whose complexity appears not to bc accommodated by
the principal classes of simulation models, this being especially true for food webs dominated by
single-celled protists whose trophic function is more dispersed than among the multicelled metazoa.
The fashionable question concerning a hypothetical
“missing sink” for CO2 emissions is unanswerable because of imprecision in our knowledge of critical flux rates. This question also diverts
attention from more relevant studies of how the biological pump may be perturbed by climatic
consequences of CO, emissions. Under available scenarios for climate change, such responses may
seem more likely to reinforce, rather than mitigate, the rate of increase of atmospheric CO,.
There are other ways of saying it, but really the most urgent question for biological
oceanography is whether the marine biosphere will mitigate or reinforce the increase
of anthropogenic greenhouse gases in the
atmosphere. This increase, as Fig. 1 shows,
is an instantaneous step function on the time
scale of natural changes of similar magnitude.
A simple probe of the gross chemistry of
the atmosphere, ocean, and sediments would
reveal a state of geochemical disequilibrium. A reactive gas, oxygen, is abundant in
both atmosphere and ocean, and carbon
(with which it would combine on a lifeless
but watery planet) is largely sequestered in
Acknowledgments
Though they won’t all agree with what I have written,
I am particularly grateful to many colleagues and friends
for spirited discussion of this problem and for providing me with charts for my voyage into unknown waters:
especially Karl Banse, David Bird, Steve Calvert, John
Cullen, Glen Harrison, Erica Head, Peter Jones, Paul
Kepkay, Trevor Platt, and Jorge Sarmiento.
biogenic carbonate formed below or along
the margins of oceans. This geochemical
disequilibrium
must be maintained by biological processes against a background of
geological episodes of subduction and vulcanism, during which C is cycled between
biogenic carbonate rocks and the atmosphere (Owen and Rea 1985; Volk 1989b).
Our probe might not sense the critical fact
that the biologically
mediated flux that
maintains the geochemical disequilibrium
is small compared with the transfer of CO2
across the ocean surface driven simply by
diffusion and solubility. Figure 2 shows a
new version of a familiar figure to make this
point. We would surely see that the surface
layer of the oceans is in equilibrium
with
the atmosphere, while the deeper water
masses are oversaturated with dissolved
CO,. From this we draw the basic concept
of the “biological pump”: downward flux of
C is maintained by gravitational settling of
organic and carbonate C, fixed autotrophitally in the surface layer by plant cells. We
1507
Longhurst
1508
600
1
CO2
2050 +
(atmosphere)
1990,
500
Glacial transition periods
g 400
CL
..I
lA 300
3
200
Most recent glacial period
100
Years
-150000
Fig. 1. Estimates ofpCO,atm
Pole data, with a projection to
(-150,000 to +lOO)
100
for the last 150,000 yr from ice-core data and from modern Mauna Loa-South
A.D. 2050 based on the present-day
incremental rate.
a.ssume that this flux is constrained by the
supply of a limiting nutrient, usually supposed to be inorganic N as NO+
To predict the response of the ocean (and
its biological pump) to a step function in
the radiatively
active gases in the atmosphere is difficult or even impossible, given
our uncertainties about relevant processes
and rates. Paradoxically, we are more confident about the role of marine biota in the
remote past than in the present perturbation
and, by constraining our predictions, this
knowledge may help us to predict the role
of marine biota in the immediate future
(Berger et al. 1989).
Riogeochemical evolution of the
primitive atmosphere
The CO,-rich atmosphere of the Proterozoic was progressively modified by the sequestration of C in carbonate rocks and organic sediments by stromatolite-building
and other cyanobacteria (Rothschild
and
Mancinelli
1990) in shallow seas. By the
early Cambrian, sufficient free oxygen was
available both to oxidize reactive elements
in the rocks (e.g. Fe to Fe oxides) and to
produce a surplus sufficient for the evolution of protists and metazoa.
The geochemical consequences of marine
productivity
dominate the geological record: terrestrial plants have endowed us with
onl:y 3-7 x 1O3Gt C of coal, compared with
40 x lo6 Gt C of carbonate rocks from
marine biota. However, the pCO,atm decline over the 60 x lo6 yr from the Cretaccous to the Pleistocene may be due to the
activity of land plants. If the Cretaceous atmosphere resulted from active sea-floor
spreading and vulcanism, and if Cretaceous
oceanic circulation was as active as today
(Manabe and Wetherald 1980), the subsequent decrease in pCO,atm may have been
due to the evolution of the angiosperm forests (Volk 1989~). If, on the other hand, the
C&-rich Cretaceous atmosphere originated
in weak oceanic circulation and low plankton productivity,
strengthening circulation
and increasing marine productivity
may
have progressively reduced pCO,atm to
Pleistocene levels.
Role of the biosphere in Pleistocene glacial
transitions-Because the rate of change in
pCO,atm during Pleistocene glacial-interglacial transitions approaches our scale of
interest, it may be profitable to examine how
small changes in solar forcing (Berger et al.
1990) could be so amplified by changes in
pCO,atm as to mediate transitions between
glacial and interglacial climates (Genthon et
al. 1987). Perhaps such a process might also
amplify or reduce the effects of anthropogenic forcing of climate change in the coming centuries.
1509
Carbon cycle--&cdthe biota
Flows as Gt C yr-1
Stocks as Gt C
0.3 f 0.3
0.3 - 11.0
&
ii
ATMOSPHERE
747 ( = 355 ppm)
(annual
1
1
change
+3.4 f 0.2 ppm)
54.3 f 9.5
LAND
BIOTA
I 639
(net 1.5 - 2.5)
80 f 23
1
+ 186
FLUX
L
I
DIC
DOC - 0.2-0.5
DIC 1000
DOC 313
25.3-27.3
h
f
3
PYCNOCLINE
POC 7-9
PIG 0.75 f 0.3
OCEAN
PRODUC ‘TIVITY
OPTIONS
INTERMEDIATE
DEEP WATER
AND
DIC 37
Doc ??
DIC 36700
POC 4.7
DOC 1198
P/C 0.15
POC 0.04
Fig. 2. Summary of the global carbon budget, excluding carbonate and other rocks, modified
and Bolin 1986. The levels of uncertainty indicated are mostly from the tabulations of Sundquist
are explained in the text. Export flux is shown for different values for the global$ratio.
A cornucopia of scenarios, having a rapid
replacement rate, purports to explain such
changes in pCO,atm and also the concomitant changes in terrestrial biosphere C: the
sum of these two fluxes must have been
absorbed by (or released from) the oceans
during a glacial transition. It is generally
supposed that the marine biosphere was in
some way implicated (e.g. Barnola et al.
1987; Berger et al. 1987).
At glacial onset, a flux of 93- 1,3 50 Gt C
from terrestrial plants, litter, and soil oc-
from Moore
1985; others
curred (Faure 1990), to which must be added 2 11 Gt C from the atmosphere (to reduce
it from 280 to 180 ppm), thus requiring a
sink in the ocean of 304-l ,56 1 Gt C to
maintain global balance. This amount is
within
observational
constraints,
for
Sarnthein et al. (1988) computed a return
of 580-720 Gt C from the ocean to the atmosphere during the transition to our own
interglacial period.
In recent years, thinking on these fluxes
has evolved rapidly. From reliance on the
1510
Longhurzt
relationship between ocean temperature and
pCO,ocean (e.g. Broecker 1982), recent scenarios are more likely to invoke increased
ocean productivity
during glaciations to
draw down pCO,atm. Candidate mechanisms have ranged from increased production of biogenic carbonate by corals (Berger
1982), through changes in the Redfield ratio
of sinking material (Broecker and Pcng 1982;
Broecker and Takahashi 1984) to increased
nutrient supply from continental shelves exposed during glaciations (McElroy
19 8 3;
Broecker and Peng 1984; Walsh 1988). Current thinking, however, converges on the
effects of altered ocean circulation on the
supply of nutrients to the photic zone, especially in low latitudes where most new
production
occurs (Barber and Chavez
1986), but perhaps not in the Arabian Sea
monsoon region (Sarkar et al. 1990). Sediment cores suggest that during Pleistocene
glaciations the ocean was characterized by
stronger upwelling centers (Muller and Suess
1979; Boyle 1986; Mix 1989a), while subtropical gyres were more stratified and less
productive (Sarnthein et al. 1987; Sarnthein
and Winn 1990). Such data suggest 38%
more productivity
at low latitudes in the
Atlantic (Mix 19893), as well as an increase
in the tropical Pacific (Pedersen et al. 1988),
for a significant increase globally (Sarnthein
and Fenner 1988).
Other scenarios have emphasized highlatitude processes. In the North Atlantic, as
the polar front approached the mcteorological intertropical
convergence zone, increased wind stress would favor pulsed production and export flux, resulting in a
drawdown of pCO,atm. At the same time,
formation of d.eep water masses may have
ceased at high latitudes in the North Atlantic (Oppo and Fairbanks 1987), with consequences for Southern Ocean alkalinity and
photic zone NO3 and Si03 (Sarmiento and
Toggweiler 1984; Broecker and Peng 1989;
Knox and McElroy 1984).
It has also been suggested that plant
growth in our present interglacial Southern
Ocean may be Fe-limited because of low
aeolian input of this element, a flux thought
to have been greater during glaciations
(Martin and Fitzwater 1988; Martin 1990)
though sediment cores suggest reduced,
rather than enhanced, export flux in Ant-
arctic seas during glaciations (Boyle 1988;
Mortlock et al. 199 1). The question of the
role of Fe in the Southern Ocean is, of course,
one of the central subjects of this symposium.
Mechanism of the biological pump-For
’
simplicity, we might characterize the biological pump as having three components,
each representing a different ecological process: a rotary pump which circulates material in the microplankton
food web of the
photic zone, an Archimedian pump by which
flux of fecal and aggregated material occurs
continually under gravity, and a reciprocating pump by which diel migrants actively
carry material down at dawn, to rise again
at dusk to feed.
The usual concept of the biological pump
does not accommodate the consequences of
the complex interactions
between plant
growth and the alkalinity of seawater for
flux of CO, across the sea surface. Further,
it usually simply assumes that sinking flux
of POC is a measure of the rate of N03d:riven production in the photic zone. Although this balance may be confirmed by
15N uptake measurements (e.g. Knauer et
al. 1990), the simple model ignores active
flux of DON and respiratory C by diel migrants, vertical advection and diffusion of
JDOM, and variable organic-inorganic C and
C : N : P ratios of the flux.
These omissions have had an important
impact on our thinking. Many (most?) biological oceanographers would probably still
accept the simple paradigm that weakened
equatorial upwelling should reduce the biological uptake of C and thus enhance the
normal rate of CO, outgassing there. Conversely, most might be surprised at the concomitant concept that reduced upwelling of
CO,-rich water would actually result in reduced outgassing of CO, from a warmerthan-normal surface water (e.g. Siegenthaler
and Wenk 1984). Actually, it must be the
balance between these two processes that
cletermines flux across the sea surface.
Nutrient sources and carbon sinks--In
simple models, photosynthesis is the only
autochthonous source of particulate C and
the limiting nutrient is usually supplied in
our models only by eddy flux from below,
although diverse intermittent
mechanisms
(Gargett 199 1) should also be accommo-
Carbon cvcle and the biota
dated: wind- or buoyancy-driven
winter
overturn, seasonal upwelling driven by nonlinear local or remote forcing, cross-shelf
eddy transfer (e.g. Houghton et al. 1987),
event-scale wind mixing (Eppley and Renger 1988), vorticity induced by eddy shear
(Pollard and Regier 1990), wind-stress curl,
and isopycnal upwelling. Only by invoking
all such mechanisms can we explain the spatial variability of sea-surface chlorophyll.
The concept of a single source 0” fixed N
may not be sustained much longer. The supply of fixed N by the cyanobacteria Trichodesmium is likely to be an important source
of export flux, regionally perhaps comparable to NO,-driven
production, and representing true net sequestration of CO, (Bird
pers. comm.). Dense blooms of Trichodesmium are widespread in moderate latitudes
and supply fixed N at rates significant relative to algal demand and sometimes more
rapidly than NO3 is supplied by vertical
mixing (Carpenter et al. unpubl.; Karl unpubl.). Further, the endosymbiotic
cyanophyte Richelia releases fixed N to sustain
diatom blooms in central gyral waters (Venrick 1974) and may also be a significant
source of N.
Also usually ignored is atmospheric deposition, especially leeward of continental
masses. On the eastern continental shelf of
North America, this may contribute 20-30%
of biologically available N (Paerl et al. 1990).
Although Legendre and Gosselin (1989) recently calculated that aeolian input of NH3
in the North Atlantic could exceed upward
transport of NO, by a factor of three or
more, this has now been shown to be incorrect, and - 10% is more likely (Laws
199 1; Galloway et al. 1985). Nor can we
ignore isopycnal advection of properties,
since variability
of plankton abundance in
the California Current (and perhaps other
eastern boundary currents) is principally determined by the variable transport of nutrient-rich water from higher latitudes, not
by local coastal upwelling
(Bernal and
McGowan 198 1). The simple concept of the
biological pump does not acknowledge vertical advection of dissolved organic matter.
If the inverse, linear relationship between
NO3 and DON in the upper kilometer observed by Sugimura and Suzuki (1988) is
correct, then downward mixing of DON (and
1511
DOC) must be an important component of
the export flux from the photic zone
(Toggweiler 1989).
We also need to know how much of the
total annual particle flux for each ocean basin comprises seasonal or intermittent
blooms which may sink almost directly to
the sediments. Such flux is caused by rapid
growth of diatoms outstripping the capacity
of herbivore consumption (Riemann 19 8 9;
Lampitt 19 8 5; Prezelin and Alldredge 19 8 3;
Takahashi 1986; Nair et al. 1989) or else
by the rapid development of salp populations capable of consuming the standing crop
of small cells and producing rapidly sinking
fecal material (Bathmann 1988; Wiebe et
al. 1979).
Finally, we must separate flux from micro- and mesoplanktonic
food webs (Michaels and Silver 1988). Microplankton
produce relatively small fecal material, including abundant minipellets from protists
(Gowing and Silver 1985) with low rates of
sinking, and even more abundant submicron particles from flagellates (Koike et al.
1990). On the other hand, mesoplankton
produce larger, faster sinking fecal material
and also interzonal diel migrants may actively transport material downward (Longhurst and Harrison 1989).
Balance between micro- and mesoplankton consumption of primary production Classically, mesoplankton were thought to
dominate grazing of phytoplankton,
but actually they rarely consume more than a small
fraction of daily primary production.
Since the relative size of phytoplankton
cells is progressively smaller offshore and
toward low latitudes, grazers must follow
the same trend. Mesoplankton
consumption in oceanic situations is almost always
~20% of daily production
(Tsuda et al.
1989). In intermittent upwelling systems, as
off Peru, open-ocean consumption by mesozooplankton may be that low (3-8%, Dagg
et al. 1980). During the JGOFS 1989 North
Atlantic spring bloom study, mesoplankton
removed only l-2% of the daily primary
production (Morales et al. 199 1). Of course,
at the end of any algal bloom, consumption
may briefly exceed production.
In shallow shelf seas, mesozooplankton
consumption
is somewhat higher than in
the open ocean, usually lo-40% of daily
1512
Longhu&
production (Dagg et al. 1982; Chervin et al.
198 1; Pace et al. 1987; Longhurst 1983) but
may approach a balance with primary production, as in Puget Sound (60-70% daily,
Welschmeyer and Lorenzen 1985a, b), or
even exceed it, as in summer in the Celtic
Sea (Joint and Williams 198 5). Oceanic microplankton, on the other hand, frequently
consume 70-90% of daily primary production (Welschmeyer and Lorenzen 1985a).
In coastal and high latitudes, experiments
return quantities usually between 10 and
50% of standing stock or 25-100% of potential production daily (Gifford 1988; Paranjape 1987, 1990; Burkill et al. 1987).
Role of mesoplankton in particulate exportflux- If mesoplankton are not the principal consumers we thought in the past, what
is the importance of their apparently fastsinking fecal material in oceanic export flux?
Can we relate this material directly to new
production as the biological pump concept
requires (e.g. Michaels and Silver 1988),
since mesoplankton fecal pellets will usually
not represent more than a small fraction of
the total consumption in the overlying photic zone?
The concept that sinking fecal pellets contribute to vertical flux dates from Lohmann
(1902) and is confirmed by many trap observations (e.g. Lorenzen and Welschmeyer
1983; Sasaki et al. unpubl.), but some
anomalous observations compel our attention: crustacean fecal pellets having sinking
rates that should take them to the bottom
may instead be consumed or broken up by
other organisms while held in suspension
by vertical turbulence in coastal areas (Smetacek 1980); there is also a remarkably close
similarity between open-ocean profiles of
pellets and of copepods, suggesting little or
no vertical pellet flux (Krause 198 1; Bathmann et al. 1987). A similar case can be
made for profiles of crustacean exuviae (Sameoto 1986; Longhurst et al. 1984).
Depth proJiles ofPOCproduction and consumption -The biological pump is usually
modeled as flux between two boxes: a surface photosynthetic and a deep aphotic box
where sinking C is respired and NO3 regenerated. Whether this geochemical shorthand is supportable can be determined only
from a knowledge of the profiles of light,
temperature,
nutrients,
primary production, heterotrophic respiration, and sources
of POC. Because profile characteristics are
determined principally by physical processes, we can generalize plankton profiles for
regions of the ocean having characteristic
profiles for density (Fig. 3).
Depths of maximum algal and growth rate
coincide when these are shallow but are progressively separated in profiles where the
maxima are deeper. In such cases, the depth
of maximum productivity
is shallower than
the chlorophyll maximum, which lies at or
just above the nutricline
(Cullen 1982;
Longhurst and Harrison 1989). Heterotrophic bacterial biomass follows algal biomass
and productivity,
reflecting the utilization
of DOM released directly from plant cells
and heterotrophs (Sherr and Sherr 1988).
:New observations of oceanic DOM profiles (Sugimura and Suzuki 1988) show steep,
biologically
driven gradients to -500 m,
raising the possibility of a flux-driven connection between the “zooplankton
remineralization loop” discussed by Banse (1990a)
and the microbial loop of the photic zone.
Fr,om the depth distribution
of trophic
groups of mesoplankton we can conclude at
what depths fecal POC, excretory DOC, and
respiratory C are released and at what depths
diRerent fractions of POC are consumed
(Fig. 4); results are consistent with most partic:le flux from the photic zone being consumed and remineralized within the subsequent 200-300 m (e.g. Martin et al. 1987;
Bishop et al. 1986).
Active export flux by vertical migrantsDiel migrations, both within the photic zone
and between the photic zone and the mesopelagial, have been well described so we
can predict which genera will participate in
each kind of migration (Fig. 5). Diel migration in coastal and seasonal situations can
be: considered modifications
of the openocean paradigm. Why diel migration should
occur was a Holy Grail for planktologists
for many decades. Only recently have we
begun to ask, rather, what are its consequences for vertical flux of material? (e.g.
Longhurst et al. 1989, 1990; Longhurst and
Harrison 1988, 1989). Interzonal migrants
feed in the photic zone at night and rest in
colder, dark water by day at 300-800 m.
1513
Carbon cycle and the biota
BIDSTAT - E. Trop. Pacific
1.0
Chl a ,
INO3
R
I
2.5
zoo
2.0 ,F
EASTROPAC
,Chla
,
,NO3
,
- 20”s
,
,
,NO3
zoo
It-L3
2.0
,
0
zw
DAY % GROUPS
50
NIGHT
10
0
% GROUPS
50
100
40
1.0,
Southern Ocean Chl
02
21
4.0
Costa Rica Dome
1 .o
,
,Chla
,NO3
a
45”s
,
0.0
,
2.8 200
II
I
4.0
I
Y
BIOMASS,
NIGHT
Fig. 4. Relative contribution
of functional feeding
groups to the day and night zooplankton biomass profile in the eastern tropical Pacific Ocean (GP-gelatinous predators; RP-raptorial
predators; H-herbivores; O-omnivores;
D and DETR-detritivores;
FILT - filter-feeders; PRED - predators).
Fig. 3. Cartoons to show characteristic profiles of
chlorophyll,
NO3 primary production (Pt), and zooplankton biomass in relation to the structure of the
water column, particularly
comparing several HNLC
regions having significant mixed-layer nitrate (Chl a as
mg m--j; NO, as pg-atms liter-l; Pt as mg C m-3 d-l.
Aphotic zone, shaded; thermocline, striped).
The guts of copepods, which are usually
-70% of these migrants, are emptied prior
to, or early in, their descent. The gut contents of only the larger nekton are carried
down with them. Both plankton and nekton
migrants release respiratory C and excretory
NH4 at depth, usually at 400-600 m. Their
feeding pattern ensures that this dissolved
material originates in the surface layers.
At seven stations in low latitudes (Longhurst et al. 1990), the mean flux of respiratory C from diel migration was equivalent
to -30% of sinking flux across the thermocline. From the global biomass of zooplankton in regions where diel migration is
significant, respiratory flux is 0.14-0.27 Gt
Cyr-I, whichisg-17%ofSundquist’s(1985)
estimate for total sinking flux of C at 200
m between these latitudes, extrapolated from
Betzer et al. (1984). To this must be added
the flux from gut contents of nekton and the
loss of C (by respiration and predation) from
populations of deep over-wintering
oceanic
copepods. Work in progress in the North
Atlantic suggests that two-thirds of the migrating biomass do not survive the winter
and represent active transport from the
photic zone to depths of -600-800
m. The
comparable flux of metabolic nitrogen as
NH4 downward out of the photic zone as a
result of diel migration requires upward revision of the f-ratio (relating NO,-driven to
Longhurst
25
g-175
c
zQ)
200
P
Cd
225
275
300
325
r
3
STYLOUiElFlON
THYSANOESSA
NEMATOBRACHION
NEMATOSCELIS
~cmJsocALANus
CALOCAIANUS
MK=YI\IDCERA
PARACALANUS
TEMoRoPlA
ACARTIA
ACROCAIANUS
PSEUDOCALANUS
OITHONA
NANNOCAIANUS
UNDINUIA, ETC
’
350
375
Fig. 5. Depth distribution of trophic groups, day and night, excluding interzonal migrants. From the left, in
sequence, the groups are: cpipelagic euphausiids, small Calanoid copepods, a large herbivorous Calanoid and
pteropods, large epipelagic Calanoid copepods, large mesopelagic Calanoid copepods, an interzonal migrant
Calanoid and, on the right, ostracods. (Dl, D2-top,
bottom of thermocline; SCM and shading-subsurface
chlorophyll maximum.)
4)
total production) of from 5 to 65%; the relation between N flux due to diel migrants
and sinking particle flux is positively correlated with the depth of the mixed layer
(Longhurst et al. 1989), the ratio being 10%
for mixed layers 40-50 m deep and 20% at
1 lo-120 m.
Balance between production-consurnption: Excess nutrients-If we are to understand high N03, low chlorophyll
(HNLC)
regions, we must be able to quantify the
balance between photosynthesis and consumption, both regionally and seasonally.
.Seasonally pulsed production
of phyto-
plankton, exceeding the consuming capacity
of herbivores (e.g. Billet et al. 1983), is easily
observed, but herbivore constraint on the
numerical growth of plant cells is harder to
demonstrate. In grazing-constrained
HNLC
situations, we might expect to find unusually high pheopigment : chlorophyll
ratios
(Thomas 1979), especially where metazoan
grazers are abundant. Although the dynamits of pheopigments in the water column
are: complex (reingestion,
photodegradation), these processes should not differ signiftcantly in normal and in HNLC regions.
The HNLC region of the eastern tropical
Carbon cycle and the biota
1515
Pacific lies consistently to the south of the must ask to what extent their output is really
equator in the EASTROPAC data, but me- reliable. The basic assumption for sizescaled models is that material passes in a
soplankton biomass, chlorophyll, and pheopigments all take their highest values along linear fashion from small to large biota and
or to the north of the equator. The Pheo: *for network analysis that biota can be arranged into a few compartments, comprisChl a ratio is close to unity over the whole
ing subsets of zooplankton, phytoplankton,
region (Thomas 1979; Beers and Stewart
197 l), suggesting that the NO3 anomaly does and bacteria. The consequences of modeling
not result from mesoplankton grazing pres- all metazoan biota as just a few interacting
sure. Further, a meridional section across compartments are difficult enough to prethe region (Chavez et al. 1990) shows no dict, even though functional feeding groups
can be identified for mesoplankton, as we
correlation
between microheterotroph
abundance and surface NO,; an earlier me- have discussed above. But what of the proridional section of Protist abundance (Beers tist plankton, the principal photic zone consumers? Can we treat them as reducible to
and Stewart 197 1) does show higher abuna few feeding groups like metazoa?
dance at some stations south of the equator,
Network models: Dispersed autotrophy and
but this is not a clear result. Thus, neither
meso- nor microheterotroph
abundance is heterotrophy among protists -Although
bacteria and protists are the major comunequivocally
associated with high nearsurface NO3 in available observational data, ponent of respiration in the ocean, we are
so suggestions that the anomaly is due to only now learning that to compartmentalize
grazing constraint (e.g. Toggweiler 1990) are their trophic functions involves uncertainnot supported by the EASTROPAC
data. ties that are much greater than for metazoa
The same conclusion
was reached by because the functions of photosynthesis,
Thomas (1979) for the region where superpredation, and heterotrophy are dispersed
ficial HNLC water is advected from the disamong all Protist groups, and many individual cells frequently perform more than
tant coastal upwelling region, rather than
originating in equatorial upwelling.
one of these functions. It is now apparent
that all groups of functionally
photosynModels for C flux analysis
thetic “algae” include forms (sometimes
Two generalizations
are available for abundant) that also attack and consume
models of the marine biosphere: by simple
other protists: diatoms appear to be the only
network analysis to a few interacting biotic
exceptional group. Conversely, most groups
compartments (e.g. Volk and Hoffert 1985; of “heterotrophic”
protists, with the posFasham et al. 1990), and by allometric ar- sible exception of tintinnids,
also include
guments to size-scaled linear biomass spec- abundant forms which harbor photosyntra (e.g. Platt 1985; Echevarria et al. 1990). thetic symbiotic algal cells or photosynSuch models are capable of delivering outthetic plastids.
put within observational
constraints, alThus, the classical distinction
between
though it is difficult to be sure that they do protozoa and algae is no longer tenable
so for the right reasons, and the arguments
(Sleigh 199 l), and we should abandon this
that follow suggest that most currently used misleading
classification.
However, our
models actually simulate trophic functions
ecological interpretation
of function among
dispersed among taxa, rather than the purprotists must be sensitive to the Great Diported function-specific
taxonomic units.
vide between procaryotes (metabolically diBecause such models seem to work, we must
verse and adaptable to changing ambient
assume that our generalizations about biochemistry through DNA recombination
by
logical processes are sufficiently robust to plasmid exchange) and eucaryotes (strucdeliver realistic output despite inappropriturally diverse but metabolically less flexiate assumptions concerning the nature of ble through DNA exchange at gametogenthe simulated compartments.
esis).
Nevertheless, these two main classes of
The autotrophic
symbionts accommomodels require simplifications
that cannot
dated by “protozoan”
taxa whose dimenbe supported by ecological analysis and we sions permit such a relationship are diverse:
1516
Longhurst
cyanobacteria,
chrysomonads,
prymnesiomonads, cryptomonads, and dinoflagellates (Taylor 1987). To further complicate
the issue, heterotrophic dinoflagellates may
themselves harbor symbionts, using a variety of mechanisms to do so. Symbionts
may or may not dominate the physiology
of the host by an exchange of molecules
partly satisfying the nutrient demands of
both symbiont and host. The opposite case,
of chlorophyll-bearing,
“phytoplanktonic”
protists that regularly consume other living
cells, is also widespread. The dinoflagellate
genus Ceratium is a classical component of
autumnal blooms and of the oceanic phytoplankton, yet includes species capable of
consuming diatoms larger than themselves,
with a variety of engulfing mechanisms
(Taylor 1989). The same is true for the small,
abundant genus Protoperidinium, which is
capable of daily consuming > 20 large diatom cells to support a growth rate of 1.7
divisions daily (Jacobsen 1987).
Many genera of nanophytoflagellates
(including coccolithophores) ingest bacteria by
a variety of techniques. These chlorophyllbearing flagellates are among the principal
consumers of bacteria in lakes (Bird and
Kalff 1986) and in the marine pelagial (D.
F. Bird pers. comm.). In the Sargasso Sea,
phagotrophic (mixotrophic)
nanoflagellates
may constitute 50% of all photosynthetic
nanoplankton
and contribute significantly
to consumption of bacteria (Caron and Lim
unpubl.).
Our knowledge
of predatory
“protophytes” and photosynthetic
“protozoa”
is
quite insufficient to generalize their effect
on our measurements of (for example) heterotroph respiration or size-fractioned photosynthesis. Plastid-bearing ciliates in both
neritic and oceanic environments may represent significant proportions of large cell
photosynthesis,
reaching -25% in both
temperate and Arctic environments.
Even
so, because of their fragility on polycarbonate filters, the ciliate fraction of the largecell component of size-fractioned photosynthesis may be underestimated (Putt 1990).
The functional complexity of the pelagic
protistan ecosystem approaches the spatial
and functional complexity of coral reef organisms within their carbonate matrix. The
Protist ecosystem exists in a high viscosity
matrix which permits spatial complexity
through the development
of microzones
within which single Protist cells interact
chemically with ambient organic and inorganic molecules. Their symbionts and the
bacteria associated with their cell wall share
th.eir microzone and participate in its chemistry. The complexity of this part of the biosphere stems also from the interaction within. it between the metabolically
labile
procaryotes, the metabolically complex eucaryotes, and the structurally diverse small
metazoa (Longhurst 1990).
Attempts to simplify and compartmentalize even the major trophic relationships
in this protistan ecosystem result in complex networks (Fig. 6), which recall the unstructured food-web model of Isaacs (1972).
As discussed by Banse (1990b), the complexity of metazoan food webs may be more
apparent than real, but it is perhaps in the
protistan food web, if anywhere, that we can
expect to find a lack of biomagnification
because of the complexity of trophic links.
Biomass spectrum models: Variability of
predator-prey ratios -Though
size-scaled
plankton biomass spectra usually have a
slope close to that predicted by allometric
models, it is only recently that they have
been made to accommodate events such as
loss from the linear biomass spectrum,
which occurs during a spring bloom when
autotrophic production overwhelms the capacity for consumption of the relevant larger size classes and algal cells sink unconsumed (e.g. Rodriguez et al. 1987). Further
assumptions concerning the relative sizes of
predators and prey are inherent in linear
biomass models. A prey-predator linear ratio approaching what is often (but probably
incorrectly) assumed for terrestrial ecosystems is also proposed for marine biota (e.g.
Fenchel 19 8 8). Available
generalizations
suffer the same defect as classical terrestrial
models, only more acutely: this ratio fits
chase-and-grab predators (hawks-sparrows,
fish-shrimps) but other important, feeding
relationships (caterpillars-trees,
salps-picoplankton) fall far outside the 0.05-O. 15 range
usually assumed, with many B 1.O.
Extending -an earlier analysis (Longhurst
11390) a review of > 100 pairs of prey and
ISa
V
‘_._.._ __..:._.:_.
._
_‘....
:
.:.., ::.:.y
j :
v
VIRUSES
2.0 pm
0.2 mm
PARTICULATE ORGANIC MATTER
20 pm
2.0 mm
Fig. 6. Size-structured compartment flux formulation based on discussions at a NATO Advanced Studies Institute in 1988. To the original has been
added the flux of submicron particles from flagellate feeding (Koike et al. 1990) and also the role of marine viruses in lysing bacteria during algal blooms
(Bratbak et al. 1990; Proctor and Fuhrman 1990).
HEIEFIOTROPHS
.:.
_:__:
:.
.>
::....
I .:_
MIXOTFWHS
+
LOMOLwr
DOM
HIMOLWT
DOM
lEzz!l
AUTOTROPHS
(
H _
/ j
1
Longhurst
1518
61
4
PROTISTA
CETACEA
o
2
0
-2
-4
-6
-6
10
-10-c
-8
1
-4
-6
-2
y = 0.626x - 2.13, r2 = 0.551
2
0
6
4
6
10
In(x) OF PREDATOR LENGTH (mm)
loo
g2
E
E
E
E
5
:
PROTISTA
CETACEA
10.
0
0
O0
1 -0
*,.O
0 0
O
0
.Ol -
&”
0
0
0
0
0
0
0
f8’
0
-0 00
ocp ’
O:
Q ““Og
O o.
0
0
4 0”
00
0
0:o”Q?a 0 OeO
0
-0
.OOl -
0
0
0
0
0
l.OOOE-4 -8
.
-6
0
-4
-2
0
7
4
6
8 -----TO
In(x) OF PREDATOR LEiNGTH
ratio for 100 cases from the
Fig. 7. The feeding ratio (in longest dimensions) shown as a prey/predator
marine ecosystem, including all major taxa. Standard regression of ratio and predator length-scaled ratio scaled
by predator length shown.
predators (from picoplankton to whales) in
the pelagic ecosystem, selected to cover all
kinds of feeding mechanisms, gives the results shown in Fig. 7. Prey really are usually
smaller than predators, but the linear ratio
is 0.65 (+ 1.62!) in these selected samples.
The high variance suggests that we should
place little reliance on the actual value of
this ratio, and it is unlikely that increasing
the data base would reduce the scatter. For
metazoa, the ratio covers five orders of magni tude; high values ( 1.2-9.09) occur among
protists, copepods, ctenophores,
heteropods, fish, and cetacea, while the lowest values (0.0001-0.001) occur in fish, salps, and
pteropods. Taxonomic position (except for
protists) and relative abundance do not appear capable of predicting the ratio-even
in the most approximate manner.
Su.ch results suggest that strict linearity in
biomass spectra must be to some extent an
illusion. Because such models can have realistic output, we can only assume that there
is internal
balance between the consequences of shortcuts up the size-structured
food chain and reverse flow back along it.
Response of the biological pump to
forcing at d@erent time scales
After this short excursion into our uncertainty about the mechanism of the biological pump and how to model it, we can
return to the question of how the ocean,
including its biota, may mitigate or reinforce the rate of increase of greenhouse gases
in the atmosphere. More fashionably (e.g.
Moore and Bolin 19&6), this question is of-
Carbon cycle and the biota
1519
geochemical budgets require, most are of an
ten posed as “what part of the CO, emissions has the ocean-already assimilated?” or accuracy that would cause them to be re“where is the missing CO, sink?” It is not jected by an experimental scientist. From
Fig. 2 we could select values to support consurprising that answers are contradictory
tradictory
conclusions
concerning
the
given the low accuracy and precision of our
amount of the imbalance (or even its exisestimates of the relevant fluxes, although
tence) and the sink(s) to which the missing
this problem often seems not to be recogCO, must have gone. As it happens, Detnized.
The quest for the missing CO, sink de- wiler and Hall ( 198 8) have already done
this: using high values for oceanic uptake
rives from our inability to balance presentand low values for fossil fuel and forest
day sources and sinks. Increases in the airclearance, they showed that it is possible to
borne fraction of CO, emissions and estimates of oceanic uptake fall short by - 1.2 balance the budget within 0.3 Gt C yr-‘,
Gt C yr-’ of the canonical estimates of 6.3 which is a negligible quantity. If we reverse
their argument and use high emissions and
Gt C yr-’ emitted from all anthropogenic
sources (Broecker et al. 1979; Detwiler and low ocean uptake, the missing CO, sink can
Hall 1988). The atmospheric increase is, of become 2.3 Gt C yr- l. Updating their argument and using the range of estimates now
course, based on measurements with rather
small error terms, but this is not the case available for flux from the terrestrial biosphere (e.g. Houghton et al. 1987; Woodwell
for ocean uptake. Computation of flux into
et al. 1978), we can obtain an excess of
the ocean is based on box diffusion models,
sources over sinks ranging from 0.2 to 13.1
parameterized by uptake rates for passive
Gt C yr-l . In view of such great uncertaintracers into the ocean (radon, tritium, bombties, strong justification
is needed to work
14C), assumptions concerning ocean ventilation rates, transfer velocities of CO, across with computations based on values chosen
the surface, and rates of vertical mixing (e.g. simply because they lie centrally in the scatBroecker et al. 1979; Oeschger et al. 1975); ter of estimates.
A recent study (Tans et al. 1990) illuserror estimates are usually not given for such
trates this difficulty. A global data base of
models, which converge on oceanic uptake
of 1.5-2.5 Gt C yr- l. With median values,
oceanic mixed-layer pC0, is used together
with assumptions of transfer velocity of CO2
then, - 1.2 Gt C yr- 1 cannot be accounted
for and represents the missing COZ.
across the sea surface to suggest that observed differences in atmospheric pC0, beBalance could be obtained if the terrestrial biosphere was a net sink of C (Tans et tween the northern and southern hemial. 1990; Siegenthaler and Oeschger 1987). sphere require a mid-northern
latitude sink
However, terrestrial ecologists concur that
of 1.5-2.0 Gt C yr-’ for which it is suggested
there is presently a net flux of CO, from the that the terrestrial biosphere is the best canterrestrial biosphere to the atmosphere at a didate. No statistics are provided for the
median rate of 3.25 Gt C yr-’ (Woodwell
pCO,ocean data base, but it is known that
et al. 1978; Houghton et al. 1987; Detwiler
this is highly variable on the scale of days
and Hall 1988). That this result is contrary
to weeks; in fact, because this variability is
to what is apparently required to balance
driven by growth of phytoplankton,
pC0,
global geochemical models is acknowledged
is probably at least as patchy as the chlo(e.g. Houghton et al. 1987; Mellilo et al. rophyll field seen in satellite imagery. Re1990), although no convincing solution to gional computations of pC0, in the mixed
this paradox has been advanced. These cal- layer, therefore, can be no better than the
culations accommodate
soil weathering,
notoriously unreliable regional estimates of
forest clearance and regrowth, and growth
chlorophyll or primary production based on
stimulation
by increasing
atmospheric
ship data. Even supposing that the transfer
pCO,, for which there is little direct evivelocity assumptions used in this study are
dence at present (Detwiler and Hall 1988; correct, the level of uncertainty of the asKohlmaier et al. 1987).
sociated pCO,ocean data base seems quite
Of the many estimates for flux terms that
inadequate with which to challenge the can-
1520
Longhurst
on of terrestrial ecology concerning the sign
of C flux with the atmosphere.
What are we to make of all this? Probably
we should admit that the fashionable question (where has the apparently unaccounted-for CO2 gone to?) is not very profitable,
as are global geochemical models in which
interplay across the range of estimates for
each flux is not permitted. Unless the flux
from the terrestrial biosphere is really at the
very lowest end of all current estimates, we
should admit the possibility that our estimate of ocean uptake may be far too low.
New thinking on this flux is appropriate.
Since most scenarios for change in pCO,atm
during glacial-interglacial
transitions
require some perturbation
of the biological
processes of the photic zone, the first point
that we ought to explore is whether the biological pump is presently in steady state
(if, indeed, a steady state can be envisaged
given the lack of concordance of the time
scales relevant to the rotary, the arch.imedean, and the reciprocating pumps). And,
secondly, how do we expect it to be perturbed as climate changes induced by greenhouse gases begin to manifest themselves?
The significance of these to points is easily
overlooked. Unless we overturn our basic
concept of phytoplankton
limitation by inorganic nutrients and accept some form of
C limitation,
the biological pump cannot
respond to increasing pCOZ in surface waters unless nutrient limitation in the photic
zone is relaxed. Only if we think that this
(or the opposite case) has already occurred,
because of changing global weather and circulation patterns, could we claim that the
marine biosphere has already modified the
rate of increase of atmospheric COz.
Is the biological pump already perturbed?- Fisheries scientists need knowledge of interannual variation in fish stocks,
so information on decadal-scale variability
in ocean circulation, and the distribution of
biota is available for many parts of the ocean.
It is certainly premature to conclude that
any of the trends seen in these data reflect
the effects of increased greenhouse gases,
but they do indicate how rapidly ocean biota can respond to changes in wind patterns
resembling those likely to occur as radiative
effects strengthen.
The most profound quasi-periodic event
in ocean circulation
is El Niiio/Southem
Oscillation (ENSO). During the 1982-l 983
event, pCO,atm ceased to increase, only to
rebound immediately afterward to above the
accustomed 1.5 ppm yr-l (Volk 1989a;
Keeling and Revelle 1985). This global perturbation must have involved transfers between anomalous
sources and sinkschanges which interacted to produce the observed signal in the atmosphere.
The reduced upwelling and deeper mixed
layer over the eastern tropical Pacific Ocean
during the 1982-1983 ENS0 caused a reduction of primary production to 29% in
the equatorial upwelling, normally responsible for 20-50% of global new production,
and a reduction to 1O”h of the normal area
of the Peru-Chile upwelling region (Barber
and Chavez 1986; Walsh 1988). Such a decrease in export flux might be expected to
result in increased outgassing of CO, at the
anomalously warm sea surface (Siegenthaler
1990). However, measurements (Feeley et
al. 1987) and simulations (Volk 1989a) suggest that during 1982 the equatorial ocean
ceased to be a net source for CO2 because
weaker upwelling reduced the near-surface
supply of CO,-rich deeper water. Weaker
upwelling and reduced plant growth interacted to determine transfer rates of CO2
across the sea surface. Since the effect of
ENS0 drought conditions
reduces plant
growth in monsoon regions, the observed
changes in the atmospheric pC0, during the
19 82- 19 8 3 ENS0 requires major intermittent sinks during an ENS0 event in higher
latitude oceans or in the nontropical terrestrial biosphere.
Events on a longer time scale may also
have important effects on biological production. Decadal increases in wind strength
and wave height in the northern hemisphere
ma.y explain decadal increase in chlorophyll
values in the North Pacific gyre through increased vertical transport of nutrients (Venriclk et al. 1987; Carter and Draper 1988).
Anomalous and sustained northerly winds
in the northeast Atlantic resulted in a decadal-scale surface salinity anomaly that was
advected oceanwide in the thermohaline
circulation. This event has been invoked to
explain changes in the abundance and tim-
Carbon cycle and the biota
ing of diatom blooms and a reduction by
half of high latitude primary production due
to increased mixed-layer stability during its
passage (Cushing 1990).
Potential responsesof the biological pump
to an anthropogenic greenhouse climateOur predictions for the response of the marine biosphere to a changed climate can be
no better than the available predictions for
how ocean circulation will change with the
climate. The main outlines, though not regional details, of changes in circulation have
been predicted (Manabe and Stouffer 1980):
weaker meridional thermal gradients will
decrease polar ice cover, decrease the
strength of the tradewinds and-at
least
during a period of transition-strengthen
the
flow of low-salinity
surface water from the
polar regions. In some situations, increased
thermal gradients may develop between
continental mass and adjacent ocean so that
coastal upwelling may be enhanced (e.g.
Bakun 1990). It is difficult to predict the
balance of the various biological responses
to these changes, which may be of opposite
sign in low and high latitudes.
Can we get any help by reference to the
response of the marine biosphere to changes
in wind-stress patterns so as to mediate glacial transitions? Can we assume that after a
period of rapid and perhaps chaotic response to the present anomalous rate of increase of atmospheric pCO,, the biogeochemical balance will attain a quasi-steady
state resembling the much slower rate of
change at a glacial transition? Output from
a two-dimensional
model of Baes and Killough (1986) appears to confirm, in fact,
that biological uptake of C by the marine
biosphere is capable of important modification of atmospheric CO2 levels: an ocean
with no biological uptake induces a doubling of atmospheric COZ, and a doubling
of nutrient supply to the phytoplankton
almost halves atmospheric COz. Much smaller changes in atmospheric CO, are induced
by halved or doubled circulation rates in
the absence of biological processes.
There are also suggestions concerning
other potential changes likely to occur if the
anthropogenic greenhouse effect strengthens. Ice cover in the Arctic decreased - 5%
between 1979 and 1987 (recalling the con-
1521
current global retreat of glaciers) and a further decrease of the magnitude suggested by
climate scenarios is computed to increase
levels of primary production over presently
ice-covered Arctic shelf areas by an order
of magnitude, or to about - 1 Gt C yr-’
(Walsh 1989). This process, by sequestering
C to the sediments, would reduce the rate
of increase in atmospheric pC0, levels and
may already have started to take effect, if
observations of reduced Arctic ice cover in
recent years are correct.
On the other hand, any oceanic primary
production processes presently driven by
wind stress at the sea surface and consequent renewal of photic zone nutrients
(equatorial upwelling, event-scale mixing,
and perhaps spring blooms driven by winter
overturn) will become weaker and reduce
the present rate of C sequestration: the socalled phytoplankton
accelerator effect in
relation to increasing pCO,atm. However,
as noted above, there is evidence that in
recent decades, when the first effects of increased greenhouse gases may be expected
to have taken effect, these processes may
have become stronger rather than weaker in
midhigh latitudes of the northern hemisphere.
Conclusions
Several conclusions can now be drawn
concerning our original question: will the
marine biosphere mitigate or reinforce the
rate of increase of anthropogenic pCO,atm?
Our first conclusion must be that the output from formal biogeochemical models deployed for this purpose should be accorded
less deference than has become customary,
given our extreme uncertainty
about the
rates of almost all critical fluxes and even
of the size of the stocks between which flux
occurs. Where such models do not accommodate reasonable error terms for each flux
and stock, their out.put should be regarded
with great caution.
The second conclusion concerns the paradox that highly aggregated simulation
models of biochemical processes in the pelagial should give output within observational constraints, although their formulation departs so strikingly from our emerging
understanding of flow within protistan food
1522
Longhurst
webs. Perhaps the complexity of simulated
processes is, in reality, lost as noise in the
overall signal representing the fluxes simulated, or perhaps, since we know approximately what constraints to place on outputs, we may unconsciously parameterize
the models accordingly. Output may be correct, but if network flow is incorrect, predicted responses to novel conditions may
be fallacious. Thus, a way must be found to
subsume the dispersion of function in procaryotes and eucaryotes (still very largely
undescribed, but becoming more complex
with each new finding) into a few interacting
compartments, representing dispersed fLmclions rather than groups of taxa, as is the
current practice.
If we do not defer to geochemical models
to answer our question, can we obtain any
help from scenarios of the role of biota in
pCO,atm
changes during glacial transitions? Can these scenarios be a guide to biotic response to other global perturbations?
ILf this is so, and if the changed greenhouse
climate is to be dominated by a reduced
meridional thermal gradient (reduced meridional wind stress), then we can at least
guess what the sign of the biotic response
will be. It is more likely that C uptake by
lthemarine biosphere will weaken rather than
strengthen and consequently will enhance
the rate of increase of pCO,atm rather than
mitigating the effect. Hopefully, in the foreseeable future, it will be possible to upgrade
this guess (through the step of being a plausible scenario) to a predictive model on
which we can place some confidence.
References
BAES, C. F., JR., AND G. G. KILLOUGH. 1986. Changes
and biological processes in CO, ocean models, p.
329-347. In J. R. Trabalka and D. E. Reichle [eds.],
The changing carbon cycle. Springer.
BAKUN, A. 1990. Global climate change and intensification of coastal ocean upwelling. Science 247:
198-201.
BANSE, K. 1990a. New views on the degradation and
disposition of organic particles as collected by sediment traps in the open sea. Deep-Sea Rcs. 37:
1177-l 195.
.-. 1990b. On pelagic food web interactions in
large water bodies, p. 556-579. In M. M. Tilzer
and C. Serruya [eds.], Large lakes-ecological
structure and function. Springer.
BARBER, R. T., AND F. P. CHAVEZ. 1986. Ocean variability in relation to living resources during the
1982-83 El Nifio. Nature 319: 279-285.
BARNOLA, J. M., D. RAYNAIJD, Y. S. KOROTKEVICH,
AND C. LORIUS. 1987. Vostok ice core provides
160,000-year record of atmospheric CO,. Nature
329: 408-4 13.
BATHMANN, U. V. 1988. Mass occurrence of Salpa
fusiformis in the spring of 1984 off Ireland: Implications for sedimentation processes. Mar. Biol.
97: 127-135.
--,
T. T. NOJI, M. Voss, AND R. PEINERT. 1987.
Copepod faecal pellets: Abundance, sedimentation and content at a permanent station in the
Norwegian Sea. Mar. Ecol. Prog. Ser. 38: 45-5 1.
BEERS,J. R., AND G. L. STEWART. 197 1. Micro-zooplankters in the plankton communities of the upper waters of the eastern tropical Pacific. DeepSea Res. 18: 861-883.
BERGER, A., H. GALLI?E, T. FICHEFET, I. MARSIAT, AND
C. TRICOT. 1990. Testing the astronomical theory with a coupled climate-ice sheet model. Palaeogeogr. Palaeoclimatol. Palaeoecol. 89: 125-14 1.
BERGER, W. H. 1982. Increase of carbon dioxide in
the atmosphere during deglaciation: The coral reef
hypothesis. Naturwissenschaften
69: 87-88.
--,
K. FISCHER, C. LAI, AND G. WV. 1987. Ocean
productivity and organic carbon flux. Part 1. Overview and maps of primary production and export
production. Univ. Calif. San Diego. SIO Ref. 8730.
--,
V. S. SMETACEK, AND G. WEFER. 1989. Ocean
productivity
and palaeoproductivity-an
overview, p. 1-31. In W. H. Berger et al. teds.], Productivity of the oceans. Wiley.
BER.NAL, P. A., AND J. A. MCGOWAN. 1981. Advection and upwelling in the California Current, p.
381-399. Zn F. A. Richards [ed.], Coastal upwelling. Am. Geophys. Union.
BETZER, P. R., AND OTHERS. 1984. Primary productivity and particle fluxes on a transect of the equator at 153W in the Pacific Ocean. Deep-Sea Res.
31: l-11.
BILLET, D. S. M., R. S. LAMPITT, A. L. RIUCE, AND R.
F. C. MANTOURA. 1983. Seasonal sedimentation
of phytoplankton
to the deep-sea benthos. Nature
302: 520-522.
BIRD, D. F., AND J. KALFF. 1986. Bacterial grazing
by planktonic lake algae. Science 231: 493-495.
BISHOP, J. K. B., J. C. STEPIEN,AND P. H. WIEBE. 1986.
Particulate matter distributions, chemistry and flux
in the Panama Basin: Response to environmental
forcing. Prog. Oceanogr. 17: l-59.
BO’U’LE,E. A. 1986. Deep circulation, preformed nutrients and atmospheric CO,: Theories and evidence from oceanic sediments, p. 49-59. In K. J.
Hsii [ed.], Mesozoic and Cenozoic oceans. Am.
Geophys. Union.
1988. Cadmium: Chemical tracer of deep--.
water paleoceanography. Paleoceanography 3: 47 l489.
BR~TBAK, (i., M. HELDAL, S. NORLAND, AND T. F.
THINGSTAD. 1990. Viruses as partners in spring
bloom microbial trophodynamics.
Appl. Environ.
Microbial. 56: 1400-1405.
Carbon cycle amndthe biota
BROECKER,W. S. 1982. Glacial to interglacial changes
in ocean chemistry. Prog. Oceanogr. 11: 15 l-l 97.
-,
AND T.-H. PENG. 1982. Tracers in the sea.
Eldigio.
-,
AND 1984. The climate-chemistry
connection. Gedphys. Monogr. 29, p. 327-336.
-,
AND . 1989. The cause of the glacial
to interglacial atmospheric CO, change: A polar
alkalinity hypothesis. Global Biogeochem. Cycles
3: 2 15-239.
-,
AND T. TAKAHASHI. 1984. Is there a tie between atmospheric CO, content and ocean circulation? Geophys. Monogr. 29, p. 314-326.
-,
H. J. SIMPSON,AND T.-H. PENG. 1979.
Fate of fossil fuel carbon dioxide and the global
carbon budget. Science 206: 409-4 18.
BURKILL, P. H., R. F. C. MANTOURA, C. A. LLEWELLYN,
AND N. J. P. OWENS. 1987. Microplankton
grazing and selectivity of phytoplankton
in coastal waters. Mar. Biol. 93: 581-590.
CARTER, D. J. T., AND L. DRAPER. 1988. Has the
north-east Atlantic become rougher? Nature 332:
494.
CHAVEZ, F. P., K. R. BUCK, AND R. T. BARBER. 1990.
Phytoplankton
taxa in relation to primary production in the equatorial Pacific. Deep-Sea Res.
37: 1733-1752.
CHERVIN, B., T. C. MALONE, AND P. J. NEALE. 198 1.
Interactions between suspended organic matter and
copepod grazing in the plume of the Hudson River. Estuarine Coastal Shelf Sci. 13: 169-183.
CULLEN, J. J. 1982. The deep chlorophyll maximum:
Comparing profiles of chlorophyll-a.
Can. J. Fish.
Aquat. Sci. 39: 791-803.
CUSHING, D. H. 1990. Plankton production and yearclass strength in fish populations: An update of the
match/mismatch
hypothesis. Adv. Mar. Biol. 26:
249-293.
DAGG, M., AND OTHERS. 1980. Grazing and excretion
by zooplankton in the Peru upwelling system during April 1977. Deep-Sea Res. 27: 43-59.
-,
J. VIDAL, T. E. WHITLEDGE, R, L. IVERSON,
AND J. J. GOERING. 1982. The feeding, respiration and excretion of zooplankton in the Bering
Sea during a spring bloom. Deep-Sea Res. 29: 4563.
DETWILER, R. P., AND C. A. S. HALL. 1988. Tropical
forests and the global carbon cycle. Science 239:
42-47.
ECHEVARR~A, F., AND OTHERS. 1990. The size-abundance distribution and taxonomic composition of
plankton in an oligotrophic,
high mountain lake
(La Caldera, Sierra Nevada, Spain). J. Plankton
Res. 12: 4 15-422.
EPPLEY, R. W., AND E. H. RENGER. 1988. Nanomolar
increase in surface layer nitrate concentration following a small wind event. Deep-Sea Res. 35: 11191125.
FASHAM, M. J. R., H. W. DUCKLOW, AND S. M.
MCKELVIE. 1990. A nitrogen-based
model of
plankton dynamics in the oceanic mixed layer. J.
Mar. Res. 48: 591-639.
FAURE, H. 1990. Changes in the global continental
reservoir of carbon. Palaeogeogr. Palaeoclimatol.
Palaeoecol. 82: 47-52.
1523
FEELEY, R. A., AND OTHERS. 1987. Distribution
of
chemical tracers in the eastern equatorial Pacific
Ocean during and after the 1982-83 El Nifiol
Southern Oscillation event. J. Geophys. Res. 92:
654516558.
FENCHEL, T. 1988. Ecology: Concepts and limitations. Ecol. Inst.
GALLOWAY, J. N., R. J. CHARLSON, M. 0. ANDRE&
AND H. RODHE [EDs.]. 1985. The biogeochemical
cycling of sulfur and nitrogen in the remote atmosphere. Reidel.
GARGETT, A. E. 199 1. Physical processes and the
maintenance of nutrient-rich euphotic zones. Limnol. Oceanogr. 36: 1527-l 545.
GENTHON, C., AND OTHERS. 1987. Vostok ice core:
Climatic response to CO, and orbital forcing
changes over the last climatic cycle. Nature 329:
414-418.
GIFFORD, D. J. 1988. Impact of grazing by microzooplankton in the northwest arm of Halifax Harbour, Nova Scotia. Mar. Ecol. Prog. Ser. 47: 249258.
GOWING, M. M., AND M. W. SILVER. 1985. Minipellets: A new and abundant size class of marine
fecal pellets. J. Mar. Res. 43: 395-418.
HOUGHTON, R. A., AND OTHERS. 1987. The flux of
carbon from terrestrial ecosystems to the atmosphere in 1980 due to changes in land use: Geographical distribution and global flux. Tellus 39B:
122-139.
ISAACS, J. 1972. Unstructured food webs, and “pollutant analogies.” Fish. Bull. 70: 1053-1059.
JACOBSEN, D. M. 1987. The ecology and feeding of
thecate dinoflagellates. Ph.D. thesis, Woods Hole
Oceanogr. Inst. WHOI-87- 10.
JOINT, I. R., AND R. WILLIAMS. 1985. Demands of
the herbivore community on phytoplankton
production in the Celtic Sea in August. Mar. Biol. 87:
297-306.
KEELING, C. D., AND R. REVELLE. 1985. Effects of El
Niiio/Southern
Oscillation
on the atmospheric
content of carbon dioxide. Meteoritics
20: 437450.
KNAUER, G. A., D. G. REDALJE, W. G. HARRISON, AND
D. M. KARL. 1990. New production at the VERTEX time-series site. Deep-Sea Res. 37: 11211134.
KNOX, F., AND M. B. MCELROY. 1984. Changes in
atmospheric CO,: Influence of the marine biota at
high latitude. J. Geophys. Res. 89: 4629-4637.
KOHLMAIER, G. H., H. BRQL, E. 0. SIR& M. PL~CHL,
AND R. REVELLE. 1987. Modelling stimulation
of plants and ecosystem response to present levels
of excess atmospheric CO,. Tellus 39B: 155-l 62.
KOIKE, I., S. HARA, K. TERAUCHI, AND K. KOGURE.
1990. Role of sub-micrometre
particles in the
ocean. Nature 345: 242-244.
KRAUSE, M. 198 1. Vertical distribution of faecal pellets during FLEX ‘76. Helgol. Meeresunters. 34:
313-327.
LAMPITT, R. S. 1985. Evidence for the seasonal deposition of detritus to the deep-sea floor and its
subsequent resuspension. Deep-Sea Res. 32: 885897.
1524
Longhurst
LAWS, E. A. 1991. Photosynthetic
quotients in the
open ocean. Deep-Sea Res. 38: 143-167.
LEGENDRE, L., AND M. GOSSELIN. 1989. New production and export of organic matter to the deep
ocean: Consequences of some recent discoveries.
Limnol. Oceanogr. 34: 1374-l 380.
LOHMANN, H. 1902. Die Coccolithophoridae.
Arch.
Protistenk. 1: 89-l 65.
LONGHURST, A. R. 1983. Benthic-pelagic
coupling
and export of organic carbon from a tropical Atlantic continental shelf-Sierra
Leone. Estuarine
Coastal Shelf Sci. 17: 26 l-285.
-1990. Pelagic ecology: Definition of pathways
for material and energy flux, p. 263-288. Zn M.
M. Denis [ed.], Oceanologie, actualite et prospective. Proc. Centennial Sta. Endoume, Marseille.
-, AND OTHERS. 1989. NFLUX:
A test of vertical nitrogen flux by diel migrant biota. Deep-Sea
Res. 36: 1705-1719.
---,
A. W. BEDO, W. G. HARRISON, E. J. H. HEAD,
AND D. D. SAMEOTO. 1990. Vertical flux of respiratory carbon by oceanic diel migrant biota.
Deep-Sea Res. 37: 685-694.
--, AND W. G. HARRISON. 1988. Vertical nitrogen flux from the oceanic photic zone by diel migrant zooplankton and nekton. Deep-Sea Res. 35:
88 l-889.
-‘-,
AND -.
1989. The biological pump: Profiles of plankton production and consumption in
the upper ocean. Prog. Oceanogr. 22: 47-123.
---,
D. SAMEOTO, AND A. HERMAN. 1984. Vertical
distribution
of Arctic zooplankton
in summer:
Eastern Canadian archipelago. J. Plankton Res. 6:
137-168.
LORENZEN, C. J., AND N. A. WELSCHMEYER. 1983.
The in situ sinking rates of herbivore fecal pellets.
J. Plankton Res. 5: 929-933.
MCELROY, M. B. 1983. Marine biological controls
on atmospheric CO, and climate. Nature 302: 328329.
MANABE, S., AND R. 1. STOUFFER. 1980. Sensitivity
of a global climate change model to an increase of
CO, in the atmosphere. J. Geophys. Res. 85: 55295554.
--,
AND R. T. WETHERALD. 1980. On the distribution of climate change resulting from an increase in CO, content ofthe atmosphere. J. Atmos.
Sci. 37: 99-l 18.
MARTIN, J. H. 1990. Glacial-interglacial
CO, change:
The iron hypothesis. Palcoceanography
5: l-l 3.
-., AND S. E. FITZWATER. 1988. Iron deficiency
limits phytoplankton
growth in the north-east Pacific subarctic. Nature 331: 34 l-343.
----,
G. A. KNAUER, D. M. KARL, AND W. W.
BROENKOW. 1987. VERTEX: Carbon cycling in
the northeast Pacific. Deep-Sea Res. 34: 267-285.
MELLILO, J. M., T. V. CALLAGHAN, F. I. WOODWARD,
E. SALATI, AND S. K. SINHA. 1990. Effects on
ecosystems, p. 283-310. Zn J. T. Houghton et al.
[eds.], Climate change-the
IPCC scientific assessment. Cambridge.
MICHAEL& A. F., AND M. SILVER. 1988. Primary production, sinking fluxes and the microbial food web.
Deep-Sea Res. 35: 473-490.
MIX, A. C. 1989a. Pleistocene productivity:
Evi-
dence from organic carbon and foraminifera1 species, p. 313-340. Zn W. H. Berger et al. [eds.],
Productivity
of the oceans-past
and present. Wiley.
--,
1989b. Influence of productivity
variations
on long-term atmospheric CO,. Nature 337: 54 l544.
MOC)RE, B., AND B. BOLIN. 1986. The oceans, carbon
dioxide and global climate change. Oceanus 30:
9-l 5.
MORALES, C., A. BEDO, R. P. HARRIS, AND P. R. G.
‘TRANTER. 1991. Grazing of copepod assemblages in the north-east Atlantic: The importance
lofthe small size fraction. J. Plankton Res. 13: 455,472.
MOR.TLOCK, R. A., AND OTHERS. 199 1. Evidence for
lower productivity
in the Antarctic Ocean during
the last glaciation. Nature 351: 220-223.
MUL.LER, P. J., AND E. SUESS. 1979. Productivity,
sedimentation rate and sedimentary organic matter in the oceans. 1. Organic carbon preservation.
Deep-Sea Res. 26: 1347-1362.
NAIF!, R. R., AND OTHERS. 1989. Increased particle
flux to the deep ocean related to monsoons. Nature
,338: 749-75 1.
OESC:HGER,H., U. SIEGENTHALER, U. SCHO’ITERER,AND
.4. GUGELMANN. 1975. A box diffusion model to
study the carbon dioxide exchange in nature. Telhs 27: 168-192.
OPP~, D. W., AND R. G. FAIRBANKS. 1987. Variability in the deep and intermediate water circulation ofthe Atlantic Ocean during the past 25,000
Iyears: Northern hemisphere modulation
of the
Southern Ocean. Earth Planet. Sci. Lett. 86: l-l 5.
OWEN, R. W., AND D. K. REA. 1985. Sea-floor hydrothermal activity links climate to tectonics: The
‘Eocene carbon dioxide greenhouse. Science 227:
166-169.
PACE), M. L., G. A. KNAUER, D. M. KARL, AND J. H.
MARTIN. 1987. Primary production, new production and vertical flux in the eastern Pacific
Ocean. Nature 325: 803-804.
PAERL, H. W., J. RUDEK, AND M. A. MULLIN. 1990.
Stimulation ofphytoplankton
production in coastal
-waters by natural rainfall inputs. Mar. Biol. 107:
247-254.
PARANJAPE, M. A. 1987. Grazing by microzooplankton in the eastern Canadian arctic in summer 1983.
Mar. Ecol. Prog. Ser. 40: 239-246.
--.
1990. Microzooplankton
herbivory on the
IGrand Bank: A seasonal study. Mar. Biol. 107:
321-328.
PEDEIRSEN,T. F., M. PICKERING, J. S. VOGEL, J. N.
SOUTHON, AND D. E. NELSON. 1988. The re:sponse of benthic foraminifera to productivity
cycles in the eastern equatorial Pacific: Faunal and
geochemical constraints on glacial bottom water
Ioxygen levels. Paleoceanography 3: 15 7- 168.
PLAIT, T. 1985. Structure of the marine ecosystem:
its allometric basis. Can. Bull. Fish. Aquat. Sci.
213, p. 55-64.
POLI,ARD, R. T., AND L. REGIER. 1990. Large variations in potential vorticity at small spatial scales
in the upper ocean. Nature 348: 227-229.
PRÉZELIN, B. A., AND A. ALLDREDGE. 1983. Primary
Carbon cycle and the biota
production of marine snow during and after an
upwelling event. Limnol. Oceanogr. 28: 11561167.
PROCTOR, L. M., AND J. A. FUHRMAN. 1990. Viral
mortality of marine bacteria and cyanobacteria.
Nature 343: 60-62.
PUTT, M. 1990. Abundance, chlorophyll content and
photosynthetic
rates of ciliates in the Nordic seas
during summer. Deep-Sea Res. 37: 17 13-1732.
RIEMANN, F. 1989. Gelatinous phytoplankton
detritus aggregates on the Atlantic deep-sea bed. Structure and mode of formation. Mar. Biol. 100: 533539.
RODRIGUEZ, J., F. JIMBNEZ, B. BAUTISTA, AND V.
RODRIGUEZ. 1987. Planktonic biomass spectra
dynamics during a winter production
pulse in
Mediterranean
coastal waters. J. Plankton Res. 9:
1183-l 194.
ROTHSCHILD, L. J., AND R. L. MANCINELLI. 1990.
Model of carbon fixation in microbial mats from
3,500 Myr ago to the present. Nature 345: 7 lo712.
SAMEOTO, D. D. 1986. Influence of the biological and
physical environments on the vertical distribution
of mesoplankton and micronekton in the tropical
Pacific. Mar. Biol. 93: 263-279.
SARKAR, A., R. RAMESH, S. K. BHATTACHARYA, AND
G. RAJAGOPALAN. 1990. Oxygen isotope evidence for a stronger winter monsoon current during the last glaciation. Nature 343: 549-55 1.
SARMIENTO, J. L., AND J. R. TOGGWEILER. 1984. A
new model for the role of the oceans in determining atmospheric PCoz. Nature 308: 621-624.
SARNTHEIN, M., AND J. FENNER. 1988. Global windinduced change of deep-sea sediment budgets, new
ocean production and CO, reservoirs. Phil. Trans.
R. Sot. Lond. Ser. B 318: 487-504.
-,
AND K. WINN. 1990. Reconstruction
of low
and middle latitude export productivity,
30,000
years BP to present: Implications for global carbon
reservoirs, p. 3 19-342. Zn M. E. Schlesinger [ed.],
Climate-ocean interaction. Kluwer.
-,
J.-C. DUPLESSY, AND M. R. FONTUGNE.
1988. Global variations of surface ocean productivity in low and mid latitudes: Influence on CO,
reservoirs of the deep ocean and atmosphere during the last 2 1,000 years. Paleoceanography 3: 36 l399.
SHERR, E., AND B. SHERR. 1988. Role of microbes in
pelagic food webs: A revised concept. Limnol.
Oceanogr. 33: 1225-l 227.
SIEGENTHALER, U. 1990. El Niiio and atmospheric
CO,. Nature 345: 295-296.
-,
AND H. OESCHGER. 1987. Biospheric CO,
emissions during the past 200 years by deconvolution of ice-core data. Tellus 39B: 140-l 54.
AND T. WENK. 1984. Rapid atmospheric CO,
vaiiations and ocean circulation, Nature 308: 624626.
SLEIGH, M. A. 1991. A taxonomic review of heterotrophic protists important in marine ecology, p.
9-38. Zn P. C. Reid et al. [eds.], Protozoa and their
role in marine processes. Springer.
SMETACEK, V. S. 1980. Zooplankton
standing stock,
copepod faecal pellets and particulate detritus in
1525
Kiel Bight. Estuarine Coastal Mar. Sci. 11: 477490.
SUGIMURA, Y., AND Y. SUZUKI. 1988. A high temperature catalytic oxidation method for the determination of non-volatile dissolved organic carbon
in seawater by direct injection of a liquid sample.
Mar. Chem. 24: 105-131.
SUNDQUIST, E. T. 1985. Geological perspectives on
carbon dioxide and the carbon cycle. Geophys.
Monogr. 32, p. 5-59.
TAKAHASHI, K. 1986. Seasonal fluxes of pelagic diatoms in the subarctic Pacific, 1982-1983. DeepSea Res. 33: 1225-125 1.
TANS, P. P., I. Y. FUNG, AND T. TAKAHASHI. 1990.
Observational
constraints on the global atmospheric CO, budget. Science 247: 143 1-1438.
TAYLOR, F. J. R. 1987. Heterotrophic
nutrition, p.
225-267. Zn F. J. R. Taylor [ed.], The biology of
dinoflagellates. Bot. Monogr. 2 1. Blackwell.
TAYLOR, G. T. 1989. Variability
in the vertical flux
of micro-organisms
and biogenic material in the
epipelagic zone of a North Pacific central gyre station. Deep-Sea Res. 36: 1287-l 308.
THOMAS, W. H. 1979. Anomalous nutrient-chlorophyll interrelationships
in the offshore eastern
tropical Pacific Ocean, J. Mar. Res. 37: 327-335.
TOGGWEILER, J. R. 1989. Is the downward DOM flux
important in carbon transport?, p. 84-107. Zn W.
H. Berger et al. [eds.], Productivity
of the oceans:
Past and present. Wiley.
1990. Modeling workshop offers first look at
ne’w simulation of equatorial Pacific. U.S. JGOFS
News 2(2): 1, 11.
TSUDA, A., K. FURUYA, AND T. NEMOTO. 1989. Feeding of micro- and macrozooplankton
at the subsurface chlorophyll
maximum in the subtropical
North Pacific. J. Exp. Mar. Biol. Ecol. 132: 41-52.
VENRICK, E. L. 1974. The distribution
and significance of Richelia intracelIularis
in the North Pacific central gyre. Limnol. Oceanogr. 19: 437-445.
-,
J. A. MCGOWAN, D. R. CAYAN, AND T. L.
HAYWARD. 1987. Climate and chlorophyll
a:
Long-term trends in the central North Pacific
Ocean. Science 238: 70-72.
VOLK, T. 1989a. Effect of the equatorial Pacific upwelling on atmospheric CO, during the 1982-83
El Niiio. Global Biogeochem. Cycles 3: 267-279.
-.
19893. Sensitivity of climate and atmospheric CO, to deep-ocean and shallow-ocean carbonate
burial. Nature 337: 637-640.
1989~. Rise of angiosperms as a factor in
long-term climatic cooling. Geology 17: 107-l 10.
-,
AND M. I. HOFFERT. 1985. Ocean carbon
pumps: Analysis of relative strengths and efficiencies in ocean-driven
atmospheric CO, changes.
Geophys. Monogr. 32, p. 99-l 10.
WALSH, J. J. 1988. On the nature of continental
shelves. Academic.
1989. Arctic carbon sinks: Present and future. Global Biogeochem. Cycles 3: 393-4 11.
WELSCHMEYER, N. A., AND C. J. LORENZEN. 1985a.
Chlorophyll
budgets: Zooplankton
grazing and
phytoplankton
growth in a temperate fjord and
the central Pacific gyres. Limnol. Oceanogr. 30: l21.
1526
-,
Longhurst
AND -.
198%. Role of herbivory in
controlling
phytoplankton
abundance: Annual
pigment budget for a temperate marine fjord. Mar.
Biol. 90: 75-86.
WIEBE, P. H., L. P. MADIN, L. R. HAURY, AND G. R.
HARBISON. 1979. Diel vertical migration by Snl-
pa aspersa
and
its potential for large-scale particulate organic matter transport to the deep-sea. Mar.
Biol. 53: 249-255.
WOODWELL, G. M., AND OTHERS. 1978. The biota
and the world carbon budget. Science 199: 14 l146.