Article
Volume 11, Number 8
3 August 2010
Q08003, doi:10.1029/2009GC002968
ISSN: 1525‐2027
Extent of the microbial biosphere in the oceanic crust
Cara Heberling and Robert P. Lowell
Department of Geosciences, Virginia Polytechnic Institute and State University, Blacksburg, Virginia 24061, USA
(rlowell@vt.edu)
Lei Liu
School of Earth and Atmospheric Sciences, Georgia Institute of Technology, Atlanta, Georgia 30332, USA
Martin R. Fisk
College of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, Oregon 97331, USA
[1] We estimate the depth of the 120°C isotherm by constructing crustal thermal gradients based on theoretical
and observed conductive heat flux as a function of lithospheric age. We chose the 120°C isotherm because it is
close to the upper limit for prokaryotic life, and therefore, the isotherm approximates the maximum depth
at which life can persist in the ocean crust. The depth of the potential microbial biosphere increases with
lithospheric age from approximately 0.5 to 1 km for 1 Ma lithosphere to as much as 5 km at a subduction
age of 180 Ma. We use global models of oceanic plate creation to estimate the volume of crust occupied
by the biosphere today and throughout geologic time. Presently, the volume of the ocean crust that is capable
of containing life is similar to the volume of the oceans (∼1018 m3). Depending on the model used for the
growth of continental crust, the volume of rock available to house the subsurface biosphere may have
remained constant or doubled since the Archean. Although the thermal models presented here provide
estimates for the potential depth and volume of rock in which microbes may live, the biomass in this volume
is not well constrained. Using a previously published model, the prokaryotic biomass in the igneous ocean
crust is estimated to exceed that in all aquatic and soil environments and is similar to that in the continental
subsurface and in marine sediment. Most of the crustal biomass beneath the sediments is likely contained
within the extrusive layer, and this has probably been the case since microbes first colonized the oceanic crust
in the Archean.
Components: 8600 words, 8 figures.
Keywords: microbial biosphere; oceanic crust; thermal structure.
Index Terms: 0456 Biogeosciences: Life in extreme environments; 0463 Biogeosciences: Microbe/mineral interactions;
3015 Marine Geology and Geophysics: Heat flow (benthic).
Received 18 November 2009; Revised 30 March 2010; Accepted 9 April 2010; Published 3 August 2010.
Heberling, C., R. P. Lowell, L. Liu, and M. R. Fisk (2010), Extent of the microbial biosphere in the oceanic crust, Geochem.
Geophys. Geosyst., 11, Q08003, doi:10.1029/2009GC002968.
Copyright 2010 by the American Geophysical Union
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1. Introduction
[2] The extent of the subsurface microbial biosphere
is of considerable scientific interest as current
research suggests that it contributes significantly to
global biomass and diversity [Amend and Teske,
2005]. Whitman et al. [1998] suggest that there may
be as much living carbon in the subsurface biosphere
as there is in terrestrial plants and they estimate that
the subsurface biosphere in marine sediments contains 3 × 1017g C, approximately two orders of
magnitude greater than in the open ocean.
[3] Although a significant amount of the subseafloor microbial biosphere is contained in marine
sediments, microbial activity occurs within the
basaltic substrate as well. Microchannels and intricate chemical alteration patterns associated with
DNA in marine basalt glass (Figure 1) recovered
from depths of hundreds of meters below the seafloor [e.g., Furnes et al., 1996; Giovannoni et al.,
1996; Fisk et al., 1998; Furnes and Staudigel,
1999; Staudigel et al., 2004, 2008] provide evidence for biologically mediated chemical weathering. Evidence of active biogenic alteration of basalt
is found in basalts from ridge crests [Mason, 2008]
to the oldest cored crust [Fisk et al., 1999; Staudigel
et al., 2004].
[4] This biogenic alteration decreases with depth in
the oceanic crust. In the upper 250 m of oceanic
crust, 60%–85% of total alteration of basaltic glass is
biogenic; at depths of 500 m, biogenic alteration
decreases to 10%–20% of the total [Furnes and
Staudigel, 1999]. Most of the biogenic alteration
occurs where subsurface temperatures are between
15°C and 80°C [Furnes and Staudigel, 1999; Walton
and Schiffman, 2003] and decreases significantly
until it disappears at about 120°C [Staudigel et al.,
2004]. The 120°C limit is consistent with estimates
for the upper limit for biological life [e.g., Gold,
1992; Holden and Daniel, 2004; Amend and Teske,
2005; Takai et al., 2008].
[5] Although temperature and other environmental
stresses will limit biological activity and biomass in
the igneous ocean crust [D’Hondt et al., 2002; Fisk
et al., 1998], the microbial subsurface biosphere
could occupy all regions of the crust that are less
than approximately 120°C, provided that there is
enough interconnected pore space to allow water to
circulate, and that the other requirements for life
are met.
[6] In this paper we calculate the depth of the sub-
surface oceanic biosphere by estimating the depth of
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the 120°C isotherm as a function of marine heat
flow, which in turn is dependent on lithospheric age.
To do this, we first use standard lithospheric cooling
models between the ages of 1 Ma and 180 Ma, the
estimated maximum age of subduction. Second, we
use the observed conductive heat flux means where
hydrothermal circulation depresses the observed
heat flux below the theoretical cooling curve to
calculate the thermal gradient. We then incorporate
the effects of hydrothermal circulation on expected
thermal gradients; both near the ridge axis and as the
lithosphere ages and hydrothermal effects lessen. By
multiplying the estimate for the maximum depth of
the biosphere by the rate of plate production and
integrating over time, we obtain estimates of the
volume of the biosphere in the present oceanic crust.
Finally, we extend our results back through geologic
time by considering the effects of greater rates of
plate production, younger lithospheric subduction
ages, and models of continental growth. By combining our estimates of the volume of rock available
for biological activity with existing calculations of
subsurface microbial abundance, one can formulate
estimates of the present biomass in the igneous
ocean crust.
[7] For this study the 120°C isotherm was chosen
because it approximates the maximum temperature
at which microorganisms survive [Kashefi and
Lovley, 2003; Takai et al., 2008]. The maximum
temperature for prokaryotic life may be higher than
120°C but it is not likely to exceed 140°C because
key organic molecules denature at or below this
temperature [Daniel et al., 2004]. If it is verified
that life continues to a significant extent above
120°C, then the depth and volume of the oceanic
crustal biosphere modeled in this paper can be
adjusted accordingly.
2. Depth of the Submarine Microbial
Biosphere
[8] In this section we determine the depth to the
120°C isotherm and the potential depth of the
microbial biosphere in several ways. In sections 2.1
and 2.2 we use expressions for theoretical and
observed conductive heat flux and determine the
depth to the 120°C isotherm from a conductive
thermal gradient. Then in section 2.3 we discuss the
possible effects of hydrothermal circulation.
2.1. Theoretical and Observed Heat Flux
[9] The first step in determining the depth to the
120°C isotherm is to consider conductive heat
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somewhat. Following Stein and Stein [1994] we
write:
H0 ðtÞ ¼ 510t 1=2 mW=m2 f or 1 Ma < t < 55 Ma
ð1Þ
and
H0 ðtÞ ¼ 48 þ 96 expðt=36Þ mW=m2 f or t > 55 Ma: ð2Þ
[ 10] Time t is in millions of years. Equation (1)
Figure 1. Basalt glass from ODP Leg 197 with microchannels that suggest microbial activity. Abbreviations
are as follows: mbsf, meters below seafloor; mib, meters
into basement (basalt).
flux H0(t) as a function of lithospheric age t. One
of the signature results of seafloor spreading is
the observed correlation of seafloor topography,
heat flow, and lithospheric plate thickness with
seafloor age. Currently the best fit lithospheric
cooling model, GDH1 [Stein and Stein, 1992],
suggests that H0(t) is best approximated by a half‐
space cooling model [e.g., Davis and Lister, 1974]
to an age of approximately 55 Ma; whereas at
greater ages, the heat flux versus age curve flattens
does not hold for lithosphere younger than 1 Ma
because hydrothermal circulation and advective heat
transfer appears to dominate the thermal structure of
young lithosphere [e.g., Elderfield and Schultz,
1996].
[11] Although equations (1) and (2) provide theo-
retical estimates for the heat flux from spreading
oceanic lithosphere, conductive heat flow measurements are typically much less than the theoretical
value for lithospheric ages younger than 65 ± 10 Ma
[Stein and Stein, 1992, 1994]. Figure 2 shows the
theoretical heat flux from equations (1) and (2) along
with the observed heat flux. Stein and Stein [1994]
show that the ratio of the observed heat flux to the
theoretical heat flux increases monotonically from
about 0.4 at 1 Ma to 1 at 65 Ma (Figure 2).
[12] The difference between observed and theoret-
ical conductive heat fluxes out to a lithospheric age
Figure 2. Heat flow (conductive heat loss) as a function of lithospheric age. Theoretical fit using equations (1) and
(2) is shown by the solid line. The dots are observed averaged in 2 Ma bins, and the dotted lines show the standard
deviation (provided courtesy of Carol Stein). The dot‐dashed line data show the fit to the binned data using equations (3)
and (4).
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of ∼65 Ma is largely attributed to hydrothermal
cooling and the transport of heat to the seafloor by
warm springs [Elder, 1965; Talwani et al., 1971;
Stein and Stein, 1994; Elderfield and Schultz,
1996]. Beyond 65 Ma, advective heat transfer resulting from hydrothermal circulation appears to be
negligible and the crust is thought to be “sealed.”
This does not necessarily mean hydrothermal circulation is absent; rather hydrothermal processes in
the basement rocks may redistribute heat, resulting
in conductive heat flow anomalies on old ocean
floor [Von Herzen, 2004].
[13] Although several sites of high‐temperature
hydrothermal venting at ocean ridge crests have
been intensely studied over the past two decades,
patterns of hydrothermal circulation in young lithosphere away from ridge crests are poorly known
[e.g., Mottl, 2003]. Heat flow studies on the young
(<3.6 Ma), heavily sedimented eastern flank of
the Juan de Fuca Ridge suggest that basaltic outcrops that penetrate the sediment cover act as either
hydrothermal recharge or discharge sites [e.g., Mottl
et al., 1998; Hutnak et al., 2006] with the result that
fluid is transported nearly isothermally across large
lateral distances beneath the sediment blanket
[Fisher et al., 2003]. Similar processes occur on the
Cocos Plate where fluid flow between widely
spaced outcrops removes more 60% of the heat
expected from lithospheric cooling models [Hutnak
et al., 2008]. Fisher and Von Herzen [2005] show
that even in fully sediment‐covered areas of low
basal heat flux, hydrothermal circulation controlled
by basement topography can redistribute heat. In
addition to hydrothermal cooling, rapid sedimentation can suppress conductive heat flow, even as fluid
seeps upward as a result of sediment compaction. In
areas of rapid sedimentation, conductive rebound
can take millions of years after hydrothermal circulation ceases [Hutnak and Fisher, 2007].
[14] The investigation of the thermal regimes asso-
ciated with these local and regional hydrothermal
patterns is beyond the scope of this paper. However,
as a first step toward approximating the effect of
hydrothermal circulation of the global thermal
regime, we construct a simple fit to the observed heat
flux between 1 Ma and the sealing age of 65 Ma.
Equations (3) and (4) provide a good fit to the
observed heat flux and produce the observed ratio
between observed and theoretical heat flux as a
function of age.
Hobs ¼ ½
0:6
0:6 * 65
t þ ð1 ÞH0 ðtÞ
64
64
ð3Þ
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from 1 Ma to 25 Ma and
Hobs ¼ 63:75 mW=m2
ð4Þ
from 25 Ma to 65 Ma. We find that equations (3) and
(4) give a better fit to the observed data than those of
Pelayo et al. [1994]. Stein [2003] provides a three‐
part linear fit to the global data. The fit from
equations (3) and (4) is superimposed on the
observed heat fluxes in Figure 2.
2.2. Potential Depth of the Biosphere
[15] The depth to the 120°C isotherm, can be found
from the relationship between conductive heat flow
H and the thermal gradient dT/dz
HðtÞ ¼ ðz; tÞ
dT
dzðtÞ
ð5Þ
where the thermal conductivity l could be considered a function of depth and lithospheric age.
Although the temperature distribution with depth
in the oceanic lithosphere can be approximately
represented by an error function, especially in
regions where the half‐space cooling model holds,
we will assume a linear temperature distribution
with depth. In young lithosphere, sediment thickness is negligible, but in older lithosphere, the low
thermal conductivity of marine sediments impacts
the temperature distribution as a function of depth.
Although the seafloor temperature is typically ≈ 1–
3°C in the ocean basins, for simplicity we assume
the seafloor temperature is 0°C. The error in depth
introduced by this simplification is small compared to the uncertainties in other parameters such
as local sediment thickness and thermal conductivity. The depth zb to the 120°C isotherm can be
found from equation (5) as:
zb ðtÞ ¼
120r
r
þ ð1 Þzs ðtÞ
HðtÞ
s
ð6Þ
where lr and ls are the thermal conductivities of
the rock and sediment, respectively, and zs is the
sediment thickness. For rock we assume a typical
thermal conductivity of basalt lr = 2.0 W/m°C, and
for sediment we assume a typical values of thermal
conductivity ls = 0.8 W/m°C [Clark, 1966; Sclater
et al., 1969]. Sedimentation rate and consequently
sediment thickness vary greatly from region to
region in the ocean (e.g., D. L. Divins, NGDC Total
Sediment Thickness of the World’s Oceans and
Marginal Seas, retrieved 19 September 2009, http://
www.ngdc.noaa.gov/mgg/sedthick/sedthick.html);
however, to obtain an approximate global perspec4 of 15
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Figure 3. Calculated depth to the 120°C isotherm based for several situations as shown in the legend. The solid
curves are calculated from the theoretical conductive heat flux considering the presence and absence of sediment
thickness increasing with age, whereas the dot‐dashed curves are calculated from the observed heat flux. These curves
join at 65 Ma. The dotted curves are calculated assuming uniform downflow of fluid from equations (10) and (11) for
two different depths of downflow.
tive we assume that sedimentation rate is constant at
3 m/Ma [e.g., Fowler, 2005]; and sediment thickness
is a linear function of age. Hence,
zs ðtÞ ¼ 3t
ð7Þ
Equation (7) yields a sediment thickness of ≈ 200 m
at 65 Ma and 540 m at an age of 180 Ma. In some
regions, the sedimentation rate may be considerably
higher than given by equation (7) because of the
proximity to the sediment source and sediment
transport.
[16] The cessation of advective hydrothermal heat
loss at ∼65 Ma is partially a result of the sealing of
fluid flow across the seafloor by low‐permeability
sediment and partially a result of permeability loss
within the crust itself by mineral precipitation
[Anderson et al., 1977; Stein and Stein, 1994].
[17] One estimate of the depth to the 120°C iso-
therm can be obtained by using the theoretical heat
flux for cooling lithosphere from equations (1) and
(2). Another estimate can be obtained by recognizing that hydrothermal cooling reduces the conductive heat flux in lithosphere less than 65 Ma. Then
H(t) in equation (6) is the measured heat flow given
by equations (3) and (4) for lithosphere less than
65 Ma and equations (1) and (2) for greater ages.
The results of these calculations are shown in
Figure 3.
2.3. Effects of Hydrothermal Circulation
2.3.1. Near Axis and Off Axis
[18] Very near the ridge axis, where vigorous
high‐temperature circulation occurs, the depth to the
120°C isotherm could very widely. In some places
it could be within ∼10 cm of the seafloor [e.g., Rona
et al., 1990]. Diffuse flow discharge with temperatures typically less than 50°C often occurs adjacent
to high‐temperature vents [e.g., Schultz et al., 1992;
Von Damm and Lilley, 2004]. These fluids represent
a mixture between high‐temperature vents and seawater that may occur at depths of ∼100 m [Crowell
et al., 2008; Ramondenc et al., 2008], suggesting
that the microbial biosphere could exist locally near
this depth, even at the ridge axis. Further off axis, as
high‐temperature discharge become less common,
the depth of the 120°C isotherm should continue to
deepen. The base of layer 2A is a plausible estimate
for the maximum depth of the microbial biosphere
in lithosphere younger than 1 Ma.
[19] There are relatively few studies of off‐axis
hydrothermal circulation, and it is difficult in a
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general way to incorporate the effects of hydrothermal circulation on observed heat flux. One
simple way is to suppose that the globally low
observed heat flux as a function of age (Figure 2) is a
result of uniform vertical hydrothermal flow. Either
uniformly downward or upward moving fluid could
result in decreased conductive heat flux at the surface. Downward moving fluid, by transporting cold
seawater into the crust could lower the surface heat
flux from the theoretical value to the observed value.
Alternatively, ascending warm seawater could
transport heat advectively across the seafloor. In this
scenario, the theoretical heat flux would be a combination of advective and conductive heat, and the
measured conductive heat flux would be reduced.
To estimate the depth of the microbial biosphere in
oceanic crust locally in areas that are affected by
hydrothermal circulation data on near surface thermal
gradients and warm spring temperatures can be used
to construct crustal geotherms. In the absence of local
data, the estimates given in sections 2.3.2–2.3.4
concerning the effect of hydrothermal circulation
on the depth of the biosphere are only suggestive.
2.3.2. Hydrothermal Downflow at Constant
Speed
[20] To consider the effect of cold seawater flowing
vertically into the crust at constant speed, we construct the steady state geotherm resulting from the
combination of advective and conductive heat transfer
across a layer of thickness h. The thermal problem is
given by [e.g., Bredehoeft and Papadopulos, 1965;
Lowell and Yao, 2002]:
d 2 Td
u dTd
¼0
a* dz
dz2
ð8Þ
Td ð0Þ ¼ 0; Td ðhÞ ¼ T0 ðtÞ ¼ H0 ðtÞh=r
ð9Þ
where Td is the temperature of the downflowing
fluid, u is the Darcian velocity (specific discharge),
a* is the effective thermal diffusivity. The basal
boundary condition, reflects the assumption that in
the absence of hydrothermal flow, the temperature at
the base of the layer is given by the temperature that
would be obtained if heat were transported only by
conduction through lithosphere of a given age t. This
means that for hydrothermal downflow to affect the
depth of the 120°C isotherm, the depth h must be
such that T0 exceeds 120° at age t. For example, if
we assume h = 4 km at all lithospheric ages, hydrothermal downflow would affect the estimated depth
of the biosphere up to 65 Ma. On the other hand, if we
assume h = 2 km, hydrothermal downflow would only
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affect the estimated depth of the biosphere out
to about 17 my. For simplicity, we have ignored
the effect of sediment cover in constructing these
geotherms.
[21] The solution to the differential equation (8)
with boundary conditions (9) is
Pe
Td ðzÞ ¼ T0
ðe h z 1Þ
ðePe 1Þ
ð10Þ
where Pe = uh/a* is the Peclet number. From (10),
the heat flow at the surface is
dTd Pe
¼
H
¼
H
obs
0
dz z¼0
ePe 1
ð11Þ
The observed heat flow Hobs at a given lithospheric
age is the left hand side of (11), H0 is the theoretical heat flow. Since both of these are known
from equations (1)–(4), equation (11) yields a value
of Pe as a function of lithospheric age for different
values of the depth h. Once Pe is known for a given
assumed depth h and age t, equation (10) is used to
determine the depth z to the 120°C isotherm. Figure 3
shows the results for h = 2 km, and 4 km. As can be
seen in Figure 3 hydrothermal downflow results in
depth estimates that are intermediate between those
given by using the observed heat flow data and those
estimated from the theoretical cooling curve.
2.3.3. Hydrothermal Upflow at Constant Speed
[22] One can also construct a steady state thermal
model for hydrothermal upflow at constant speed.
In this case a mixed boundary condition is used at
the seafloor to allow for advective heat loss to the
surface. In this case, equation (8) is used with the
Darcian velocity u being replaced by −u. The boundary condition at the seafloor is given by
dTu
bTu ¼ 0
dz
z¼0
ð12Þ
and the boundary condition at the base of the upflow
at depth h is
Tu ðhÞ ¼ T0 at z ¼ h
ð13Þ
where Tu is the temperature of the upflowing fluid,
T0 is the temperature at the base of the upflow zone,
and b is a heat transfer coefficient. The solution to
equation (8) with boundary conditions (12) and (13) is
3
2
Pe
Pe z
h 1þ
e
6
hb 7
7
Tu ðz; tÞ ¼ T0 6
4
5
Pe
Pe
1þ
e
hb
ð14Þ
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The temperature of springs at the seafloor is found by
setting z = 0 in equation (14). Hence
Pe=bh
Tu ð0; tÞ ¼ T0
1 þ Pe=bh ePe
ð15Þ
The observed conductive heat flux is found by taking
the derivative of equation (14) and evaluating the
result at z = 0. That is,
dTu T0 Pe=h
¼ Hobs ðtÞ ¼
1 þ Pe=hb ePe
dz z¼0
ð16Þ
Substituting equation (15) into equation (16) one
obtains the relation
Hobs ¼ bTu ð0; tÞ
ð17Þ
Another relationship comes from recognizing that the
advective heat flux is the difference between the theoretical heat flux H0(t) and the Hobs(t). That is,
f cf uTu ð0; tÞ ¼ H0 Hobs
ð18Þ
This leads to the relationship
Pe
H0
¼
1
hb Hobs
ð19Þ
[23] For an observed spring temperature Tu(0,t),
equations (17) and (19) constrain the heat transfer
coefficient b and the fluid velocity u, respectively.
These values can then be put into equation (14) to
calculate the depth of the 120°C isotherm. Without
constraints on the depth of hydrothermal upflow
and basal fluid temperature, however, equation (14)
cannot be solved uniquely. Some simplification
occurs if it is assumed the fluid comes from great
depth so the Pe 1, but one would still need an
estimate of the basal temperature. Therefore we do
not carry this analysis further.
2.3.4. Uniform Basement Temperature
[24] As discussed previously, basaltic rock out-
crops tend to control the hydrology of the shallow
crust, resulting in quasi‐isothermal layer 2A [Fisher
et al., 2003; Hutnak et al., 2006]. The depth to the
120°C isotherm can then be estimated by calculating
the conductive geotherm from the base of layer 2A.
The basement temperature would serve as the temperature at the top of the geotherm and the slope of
the geotherm would be constrained by H0. For
example, suppose the sediment thickness is zs =
100 m, and the layer 2A thickness z2A = 500 m.
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Suppose the basement temperature in layer 2A as a
result of lateral fluid flow is 50°C and that the theoretical heat flux at the given lithospheric age t is
H0 = 400 mW/m2. Then the depth z to the 120°C
isotherm from the base of layer 2A is
z¼
r T 2:0ð70Þ
x103 ¼ 350m
¼
H0
400
ð20Þ
and the depth to the 120°C isotherm would be zb =
zs + z2A + z = 950 m. By assuming a constant
temperature throughout layer 2A, this estimate is a
maximum.
3. Volume of the Suboceanic Microbial
Biosphere
3.1. Present‐Day Volume
[25] To determine the volume of the biosphere,
both at present and over geologic time t, we integrate the depth zb(t) with the rate of change in area
of the lithosphere A(t,t). Sclater et al. [1980] show
that the seafloor area per unit age decreases nearly
linearly with age. We assume that this formula
holds throughout geologic time, so that
dAðt; Þ ¼ C ð Þ½1 t=tm ð21Þ
where C(t) is the rate of plate creation at a given
geologic time t, and tm is the maximum age of
plate subduction. To obtain C(t), equation (21) is
integrated from 0 ≤ t ≤ tm to obtain
C ð Þ ¼ 2Aðtm ; Þ=tm
ð22Þ
[26] At present, tm = 180 Ma, yielding a value C(0) =
3.43 km2/yr of crustal production for an area of
3.09 × 108 km2 of present‐day seafloor area [Fowler,
2005]. Various estimates have been made for the
maximum age of plate subduction in the past. For
example, Bickle [1978] suggests that C0 was at least
18 km2/yr at t = 2.8 Ga, corresponding to a maximum subduction age of 40 Ma. This estimate is
consistent with lithosphere old enough to become
denser than the mantle [Oxburgh and Parmentier,
1977]. Abbott and Hoffman [1984], however, suggest that the occurrence of moderate abundances of
andesite in Archean greenstone belts, beginning at
3.4 Ga [Condie, 1982], points to a change in subduction style. They further argue that this implies
that lithosphere older than 50 Ma began to be subducted around 3.4 Ga. If we assume this value and
accept Abbott and Hoffman’s [1984] supposition of
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Figure 4. Estimated volume of lithosphere that may be occupied by the suboceanic microbial biosphere under various conductive heat loss scenarios as indicated.
a linear increase of subduction age over geologic
time, we can write
tm ¼ 180 38:2
ð23Þ
where tm is in Ma and t is in Ga.
[27] To estimate the volume of the present‐day
suboceanic biosphere, we set t = 0 in equation (23)
and insert the resulting value of tm = 180 Ma in
equations (21) and (22). We use the present surface
total area of oceanic crust of 3.09 × 1014 m2 [Fowler,
2005] to obtain C(0) in equation (21). The volume
Vb(t) is then obtained by integrating
Zt
Vb ðtÞ ¼
1
0
dAðt ; 0Þ 0 0
zb ðt Þdt
dt0
ð24Þ
where zb(t) is given by equation (6). The error
incurred by neglecting biosphere <1 Ma is less than
1%. The results are shown in Figure 4.
3.2. Potential Volume of the Suboceanic
Biosphere Over Geologic Time
[28] There is considerable uncertainty regarding the
style and rate of heat loss from Earth’s interior
during the Archean [Bickle, 1978; Abbott and
Hoffman, 1984; Windley, 1984; Sleep, 1992; Isley
and Abbott, 1999], and the mechanisms for creation of oceanic lithosphere and lithospheric heat loss
may have changed over geologic time. Because of a
higher rate of radioactive heat generation and more
vigorous mantle convection, heat loss from Earth’s
interior may have been a factor of 3 or more greater
than at present [e.g., Bickle, 1978; Turcotte, 1980].
For simplicity, we will assume that creation of
oceanic lithosphere and lithospheric heat loss
occurred in a manner similar to the present, and that
the higher total heat loss from the lithosphere in
the geologic past is simply a function of the rate of
plate creation and the maximum age of the subduction [e.g., Lowell and Keller, 2003]. With this
assumption, we can assume that the theoretical heat
flux from the lithosphere as a function of lithospheric age is still given by equations (1) and (2). We
will further assume that the depth to the base of
the biosphere can be represented by equation (6),
neglecting the effect of sediment cover and changes
in bottom water temperature. Neglecting sediment
cover in the Archean is not critical because a younger
subduction age implies less sediment accumulation.
Equations (21)–(23) will be used to determine the
function dA(t, t)/dt and different geologic times, but
the total area of oceanic lithosphere A(tm, t) depends
upon the model used for growth of continental
lithosphere.
[29] To determine A(tm,t) we use three different
growth models for continental crust [McLennan and
Taylor, 1983; Lowe, 1992]: (1) constant continental
area of 40% of Earth’s surface; (2) linear growth of
continental crust from 0% to 40% between 4.0 Ga
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Figure 5. The estimated volume of the lithosphere that may be occupied by the suboceanic microbial biosphere as a
function of lithospheric age at a number of geologic ages t in Ga before present assuming that the continental and
oceanic surface areas are constant and that oceanic lithosphere occupies 60% of Earth’s surface.
and present; and (3) segmented linear continental
growth such that continental area grows from 0% to
10% of its present value between 4.0 and 3.2 Ga,
from 10% to 80% of its present value between 3.2
and 2.5 Ga, and from 80% to 100% of its present
value between 2.5 Ga and the present. The total
surface area of the earth is assumed to be constant.
Figure 5 shows the volume of the oceanic biosphere
as a function of geologic age t for case 1. Figure 6
shows the ratio of the biosphere volume to the
Figure 6. Ratio of the volume of the oceanic biosphere to the present volume as a function of geologic ages t in Ga.
Model 1 assumes constant continental surface area. Model 2 assumes linear continental growth. Model 3 assumes
episodic continental growth.
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Figure 7. Generalized vertical structure of oceanic crust. Sediment, if present, is typically less than 1 km thick.
Extrusive pillow basalts and sheet flows are also usually less than 1 km thick. Sheeted dikes are thought to be 1–
2 km thick based on their occurrence in exposed and drilled ocean sections and ophiolites. The gabbroic layer includes
shallow massive gabbros and deeper layered cumulative peridotite and is thought to be 3–4 km thick. The total ocean
crust thickness is typically 5–6 km based on the depth of the Moho. Modified after Bratt and Purdy [1984] with
additional data from Kennett [1982] and Kearey et al. [2009].
present volume as a function of t for the 3 different
continental growth scenarios.
4. Discussion
4.1. Geological Context of the Depth of the
120°C Isotherm
[30] Figure 3 shows that the depth of the 120°C
isotherm increases with lithospheric age from
approximately 0.5 to 1 km for 1 Ma lithosphere to
as much as 5 km at the maximum subduction age
of 180 Ma. For crust less than 1 Ma, where vigorous hydrothermal circulation dominates the thermal regime, the depth to the 120°C isotherm may be
highly variable. It may range between a few centimeters to perhaps ∼100 m near the ridge axis, to
perhaps near the base of layer 2A as the lithosphere
ages. By comparing depth of the 120°C isotherm
with an ocean crust model (Figure 7) one can see that
the microbial biosphere is confined to the crust.
This will be true except where the volcanic crust is
anomalously thin or absent, such as near fracture
zones or along ultraslow spreading ridges, in which
case the isotherm will be in the mantle, or where
sediments are anomalously thick and the isotherm is
in the sediment.
[31] The stratigraphic sequence of oceanic crust
upward from the crust‐mantle boundary can be
generalized as cumulate ultramafic rocks, gabbro,
intrusive basalt dikes, extrusive basalt, and sediment
(Figure 7). The total thickness of this sequence
averages 5 to 7 km, with the lower cumulates and
gabbro making up 3 to 4 km, intrusive basalt making
up 1 to 2 km, extrusive basalt (pillow and sheet
flows) making up 0.5 to 1 km, and sediment up to
1 km thick. Although the seafloor stratigraphy presented here does not apply to all seafloor, we will use
it for discussing our models of biosphere depth and
volume. Using this stratigraphy and the “observed
with sediment” curve in Figure 3, the maximum
potential depth of the biosphere typically occurs in
the pillows and sheet flows (0.5 to 1.0 km) in the
youngest seafloor (approximately <1 Ma), typically
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occurs in the dikes (1.0 to 3.0 km) in seafloor that is
∼1 to 10 Ma, and occurs in the gabbro/layered
peridotites in seafloor older than ∼10 Ma.
[32] In crust less than about 65 Ma, hydrothermal
circulation affects lithospheric heat loss and thus the
depth of the 120°C isotherm. In regions of hydrothermal recharge, the resulting lower geothermal
gradient could lead to a deepening of the subsurface
biosphere by as much as several hundred meters in
some places. In the absence of better constraining
data, the effects of hydrothermal discharge through
warm springs cannot be easily quantified. Depending upon the warm spring temperature and the depth
of hydrothermal circulation, the depth of the biosphere could be affected only slightly. If warm fluid
is rising relatively rapidly, however, local sediment
temperatures of 120C° could occur just below the
seafloor.
4.2. Volume of the Subsurface Microbial
Biosphere Over Geologic Time
[33] Estimates of the volume of the oceanic crust
biosphere are obtained on a global basis by convolving the depth estimate with estimates of the
rate of plate creation as a function of lithospheric
age. The results are shown in Figure 4. The volume
of the ocean crust biosphere that is capable of containing life today (∼1018 m3) is similar to the volume
of the oceans (∼1.3 × 1018 m3). The volume of the
continental crust that is capable of containing life is
probably less than half that of the ocean crust,
mainly because of the smaller area of continental
crust. More detailed quantification of continental
biosphere is more difficult than for the oceanic crust
because thermal gradients in the continental crust are
highly variable and do not obey a simple mathematical expression with age. Figure 5 provides estimates
of the biosphere volume over geologic time for a
single model of continental growth in which the ratio
of continental to oceanic surface area has remained
constant over geologic time. With this constraint the
volume of the potential ocean crust biosphere has
increased from 0.4 × 1018 m3 at t = 4 Ga to 1.0 ×
1018 m3 today. Comparison of young ocean crust
(<75 Ma), however, shows that in the Archean (t =
3 Ga) there is more potential crustal biosphere in this
young crust than today. This reflects the greater rate
of production of crust in the Archean and that the
<75 Ma crust represents the whole of the ocean
crust in the Archean. Today, however, crust less than
75 Ma only composes about half of the ocean crust.
[34] Figure 6 gives the ratio of the estimated poten-
tial biosphere volume to the present estimate as a
10.1029/2009GC002968
function of geologic time for three different continental growth models. The lower “green” curve, in
which continental and oceanic surface areas were
assumed to be constant, follows directly from
Figure 5. The upper curves show that growth of
continental surface area over geologic time could
have a significant impact on the volume of the
suboceanic biosphere. The upper curves in Figure 6,
which are based on the assumptions of episodic or
linear continental growth suggest that the volume
of biosphere may have been relatively stable, and
nearly equal to the present‐day volume, throughout
much of geologic history.
4.3. Potential Biomass of the Current
Ocean Crust Biosphere
[35] Even though the oceanic igneous crust is suit-
able for life does not mean that life is present and
viable. In addition to high temperature, life can be
inhibited by the absence of nutrients, available space,
or energy. Igneous rocks and the water present in
pore space provide all the elements needed for life
(C, H, N, O, P, and S, and minor and trace nutrients).
Available space is also not likely to be a limitation.
Porosity measurements of oceanic igneous rocks
from the surface to depths of 2400 m have been
made as part of the Ocean Drilling Program, Integrated Ocean Drilling Program, and Deep Sea Drilling Program. The porosity of the extrusive layer
(pillows and sheet flows) in relatively young ocean
crust (6 Ma and 15 Ma) is 3% to 15% [Alt et al., 1993;
Expedition 309 and 312 Scientists, 2006] and there
appears to be little difference between young and old
crust. In the oldest sampled ocean crust (165 Ma) the
porosity typically ranges from 2% to 15% [Plank
et al., 2000]. The porosity in the intrusive sheeted
dikes is lower than in the extrusive lavas. In 6 Ma
and 15 Ma crust the porosity of dikes is typically
0.2% to 3%, based on measurements from two sites
where these layers have been sampled [Alt et al.,
1993; Expedition 309 and 312 Scientists, 2006]. In
the gabbro layer (the deepest layer in which we
expect to encounter the 120°C isotherm) the porosity
is typically between 0.1% and 2.5% [Robinson et al.,
1989]. Thus space is available at all depths in the
ocean crust.
[36] The flux of chemical energy that is needed to
maintain the subsurface biosphere comes from
reductants and oxidants in water that is advected
or convected through the crust and reductants and
oxidants in the rocks. Igneous rocks contain both
Fe2+ and Fe3+ in the proportion of approximately
6 to 1, so that the rocks can act as either electron
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Figure 8. Estimated prokaryotic mass of the present biosphere in the major reservoirs on Earth. Units are in 1015g.
Aquatic reservoir includes the oceans and fresh water. Land subsurface does not include soil. Solid circles, Whitman
et al. [1998]; open circle, Lipp et al. [2008]; open square, Parkes et al. [1994]; triangle, this study.
donors or electron acceptors in biological energy
producing reactions. The crust is also a source
of S and Mn that can provide energy to microorganisms. In addition Fe2+ bearing silicate
minerals in the lower crust and upper mantle can
react with water to produce H2, another source of
metabolic energy [McCollom and Bach, 2009].
The temperature of maximum H2 production is
about 300°C [McCollom and Bach, 2009], so H2
generated at higher temperatures beneath the biosphere is expected to diffuse upward where it could
fuel microbial activity in the overlying ocean crust.
[37] The combined impact of increasing tempera-
ture, reduced porosity, and therefore decreased fluid
flow and decreased flux of oxidants and reductants
with increasing depth in the crust, will result in a
decrease in biomass from the surface downward. At
some depth the crust will become sterile. Temperature is the only measure that can confidently indicate
sterility. To estimate the biomass of the ocean crust
we applied the calculation of Whitman et al. [1998]
to the three igneous layers of the ocean crust. This
calculation is based on the porosity of the layer and
the amount of this porosity that is occupied by
prokaryotes. The porosities for our calculation are
typical for pillow and sheet flows (5%), dikes (2%),
and gabbros (1%) The amount of the pore space
occupied by prokaryotes was taken as 0.016%
for the pillows and sheet flows [Gold, 1992]. To
account for the environmental stresses of higher
temperatures and lower fluid fluxes in the deeper
layers of the crust we assumed the fraction of space
occupied by microorganisms decreased logarithmically to 0.0016% in dikes and to 0.00016% in
gabbros. This logarithmic decrease is consistent
with the logarithmic decreases in cell counts with
depth that are observed in marine sediments [Parkes
et al., 1994]. With these model parameters the pillow
and sheet flows contain over 90% of the biomass of
the igneous ocean crust. The gabbros, which are at
the highest temperature and have lowest porosity
contain less than 1% of the biomass. The average
thickness of Layer 2A of 1 km that is chosen for the
model has a significant impact on the total biomass
of the igneous crust. In this model the igneous ocean
crust contains an amount of biomass that is intermediate between the estimates of the ocean sediment
biomass of Whitman et al. [1998] and those of Lipp
et al. [2008] and Parkes et al. [1994] (Figure 8).
5. Conclusions
[38] In this paper we have estimated the potential
depth and volume of the microbial biosphere in the
oceanic crust. To calculate the depth we have used
relatively simple thermal models and assumed the
limit of life is controlled the depth of the 120°C
isotherm. Our results suggest that the biosphere is
limited to crustal depths. To calculate the volume we
assume that plate tectonics has occurred throughout
geologic time and have used relatively simple models
of plate creation, a maximum subduction age that
increases linearly with geologic time, and various
continental growth scenarios. These results show
that if continental area has remained constant, the
volume of the biosphere has grown by approximately a factor of two over geologic time. If continental area has grown over geologic time, the
potential volume of the biosphere may have remained
relatively constant throughout geologic history.
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Based on the present‐day potential volume of the
biosphere in oceanic crust, we estimate that the
biomass is approximately 2 × 1017 g. Approximately 90% of this biomass likely occurs in the
extrusive layer. Since the depth to the base of
the extrusive layer is considerably shallower than
the potential depth defined by the 120°C isotherm,
this result may be true from the time the oceanic
crust was first colonized, probably before 3.5 Ga
[Furnes et al., 2004].
Acknowledgments
[39] We thank Carol Stein and an anonymous reviewer for
their insightful comments on the original version of this manuscript. We acknowledge the Future of Marine Heat Flow workshop organized by Rob Harris, Andrew Fisher, Fernando
Martinez, and Carolyn Ruppel and funded by NSF for making
it possible to begin the collaboration that resulted in this paper.
This work was supported in part by NSF grant OCE 0527208 to
R.P.L.
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