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Earth and Planetary Science Letters 227 (2004) 427 – 439
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Temporal variation of oceanic spreading and crustal production
rates during the last 180 My
Jean-Pascal Cognéa,*, Eric Humlerb,1
a
Laboratoire de Paléomagnétisme, UMR CNRS 7577, Institut de Physique du Globe de Paris, 4 place Jussieu, 75252 Paris cedex 05, France
b
Laboratoire des Géosciences Marines, UMR CNRS 7097, Institut de Physique du Globe de Paris, 4 place Jussieu,
75252 Paris cedex 05, France
Received 3 October 2003; received in revised form 8 August 2004; accepted 1 September 2004
Available online 7 October 2004
Editor: E.Bard
Abstract
We present a re-evaluation of seafloor spreading and generation rates, mainly based on a direct measurement of the
remaining surfaces of oceanic crust and isochron lengths defined in the most recent isochron maps [J.Y. Royer, R.D. Mqller,
L.M. Gahagan, L.A. Lawyer, C.L. Mayes, D. Nqrnberg, J.G. Sclater, A global isochron chart, Tech. Rep. 117, Austin, Univ. of
Tex. Inst. for Geophys., 1992; R.D. Mqller, W.R. Roest, J.Y. Royer, L.M. Gahagan, J.G. Sclater, Digital isochrons of the world’s
ocean floor, J. Geophys. Res., 102 (1997), 3211–3214]. Our evaluation of the amount of oceanic crust per unit age {dA/dt} as a
function of age, which can be expressed as dA/dt=C o(1t/t m), is in fairly good agreement with previous determinations [J.G.
Sclater, B. Parsons, C. Jaupart, Oceans and continents: similarities and differences in the mechanisms of heat loss, J. Geophys.
Res., 86 (1981) 11,535–11,552; D.B. Rowley, Rate of plate creation and destruction: 180 Ma to present, Geol. Soc. Amer. Bull.,
114 (2002) 927–933], with C o=2.850F0.119 km2 year1 and t m=180.2F9.7 Ma. Dividing these dA/dt by the ridge lengths L,
defined as the isochron length at each epoch allowed us to compute the evolution of global half-spreading rates. These have
been roughly constant at 25.9F3.3 mm year1 for at least the last 150 Ma. We propose that the global seafloor surface
generation rate is roughly constant as well, with a mean half-value of 1.298F0.284 km2 year1 and varying F20% with time.
This study corroborates the recent conclusion of Rowley [D.B. Rowley, Rate of plate creation and destruction: 180 Ma to
present, Geol. Soc. Amer. Bull., 114 (2002) 927–933], of a constant generation rate since 180 Ma, and completely contradicts
the commonly accepted idea of high seafloor spreading and surface generation rates during a large part of the Cretaceous.
Combining the oceanic surface generation rates derived here with crustal thicknesses deduced from the chemical composition of
old oceanic crusts and seismic measurements [E. Humler, C.H. Langmuir, V. Daux, Depth versus age: new perspectives from
the chemical compositions of ancient crust, Earth Planet Sci. Lett., 173 (1999) 7–23], the magmatic flux at young (0–80 Ma)
oceanic ridges appears to be about 18.1F3.4 km3 year1 and was possibly 15% to 30% higher during the Mesozoic. We
* Corresponding author. Tel.: +33 1 44 27 60 93.
E-mail addresses: [email protected] (J.-P. Cogné)8 [email protected] (E. Humler).
1
Tel.: 33 1 44 27 50 88.
0012-821X/$ - see front matter D 2004 Elsevier B.V. All rights reserved.
doi:10.1016/j.epsl.2004.09.002
428
J.-P. Cogné, E. Humler / Earth and Planetary Science Letters 227 (2004) 427–439
propose that mantle temperature variation provides an alternative mechanism to spreading rate for the Cretaceous highstand in
sea-level and atmospheric CO2 generation.
D 2004 Elsevier B.V. All rights reserved.
Keywords: seafloor spreading rate; ridge length; magmatic flux; sea-level; atmospheric CO2; Cretaceous; mantle temperature
1. Introduction
Most of the Earth’s volcanism occurs at midocean ridges and forms the oceanic crust. Until
recently, spreading rates and mean crustal thicknesses were difficult to evaluate and the amount of
crust generated at ridges was poorly estimated.
Menard [6], Deffreyes [7] and Dickinson and Luth
[8] estimated the rates of basalt production at
spreading centers at about 5–15 km3 year1 since
the late Mesozoic. In the last 20 years, considerable
advances in marine geosciences have allowed a
better estimate of the production at oceanic ridges,
which is now accepted to be about 20 km3 year1
(e.g. [9–11]). This value and its temporal evolution
are important because they are used in various
models of the Earth dynamics such as: heat flux
balances [3,12], geochemical mass balance calculations (e.g. [9,10]), sea-level changes [13,14] and the
CO2 evolution of the atmosphere (e.g. [15]).
It is often assumed that the Cretaceous sea-level
highstand (e.g. [16,17]) is the result of increased ridge
volumes linked to rapid sea-floor spreading during
that period, following the mechanism proposed by
Hays and Pitman [13] (e.g. [18,19]). This idea of rapid
mid-Cretaceous seafloor spreading, however, has been
strongly debated, and is judged unnecessary, and
perhaps wrong by Heller et al. [20], who believe that
it is an artifact of poorly defined timescales. It should
be noted, as Hardebeck and Anderson [21] did, that
(1) the inferred high rates took place during the
Cretaceous Long Normal Superchron (LNS), where
there are no reversals to precisely determine the
seafloor generation history; (2) the original derivation
of these rates [14] was based mainly on reconstructions of the Pacific Ocean assuming symmetric oceans
of which a large part has now been subducted; (3)
ridge jumps were not taken into account, which can
result in a large overestimation of seafloor spreading;
(4) the influence of other oceans (so-called
TGondwanar oceans such as Atlantic and Indian
oceans) was underestimated.
Rowley [4] recently reappraised seafloor spreading
rates since 180 Ma, using the oceanic age grid
proposed by Royer et al. [1] and Mqller et al. [2].
Integrating the differential surface/age {dA/dt} distribution proposed by Sclater et al. [3] and Parsons
[27], which is of the form {dA/dt}=C o(1t/t m), where
C o is the present-day oceanic crust production in km2
year1, and t m the maximum age of the crust, the
author proposed that plate production remained
roughly constant since 180 Ma at 3.4 km2 year1.
This conclusion, however, is based on two main
assumptions, which are (1) the destruction of oceanic
floor at subduction zones is uniformly distributed with
age, and (2) from present to 180 Ma, the oldest age of
oceanic crust remained constant at 180 Ma. Both
hypotheses may be debated, but their validity while
commonly accepted is difficult to prove (e.g. Parsons
[27], Rowley [4]).
For these reasons, we decided to revisit the
seafloor production rates, using a different method:
the direct measurement of currently visible seafloor
surfaces encompassed between pairs of isochrons in
each of the main oceanic basins. Our work follows
the method of Sclater et al. [3], but uses the more
recent, and accurate, database of Royer et al. [1] and
Mqller et al. [2]. Dividing surfaces by the isochron
length, which are the ridge remnants which produced
these surfaces, we thus obtain a figure of spreading
rates on each of the main oceanic ridges, which we
then average to evaluate the mean global spreading
rate since 180 Ma. In a second and more speculative
step, we compute the global seafloor surface
generation rates since 180 Ma by applying our
derived ridge spreading rates to hypothesized global
ridge lengths including subducted ridges of the
Pacific and Tethys oceans [22,23] (see Section
3.3). Finally, in Section 3.4, we estimate the temporal
variation of oceanic crust production, in volume, in
J.-P. Cogné, E. Humler / Earth and Planetary Science Letters 227 (2004) 427–439
429
the light of the recently proposed temperature
variations in the upper mantle [5].
2. Method
2.1. General
We have based all of the present study on the
synthetic isochron database published by Royer et al.
[1] and Mqller et al. [2]. This database (which may
be downloaded at ftp://ftp-sdt.univ-brest.fr/jyroyer/
Agegrid/) reconstructs magnetic anomalies 5, 6, 13,
18, 21, 25, 31, 34, M0, M4, M10, M16, M21 and M25,
calibrated using geomagnetic timescales of Cande and
Kent [24] for anomalies younger than chron 34 (83
Ma), and of Gradstein et al. [25] for older periods. An
error gridmap is also provided, computed as a function
of (1) errors on ocean floor ages, (2) distances of each
grid cell to the nearest magnetic anomaly, and (3) the
gradient of the age grid. Except in the vicinity of large
fracture zones, most part of this age map is estimated
to be defined with an accuracy better than F3 Ma.
We first computed the area of oceanic crust {dA}
comprised between each successive pair of isochrons,
in each of the main oceanic basins (Atlantic and
Indian Oceans, Antarctic, Pacific, Nazca, and Cocos
plates), using the PaleoMac application of Cogné [26]
(Fig. 1). We imposed an error of F3% on {dA} from
the computations of the spherical shells defined by the
isochron outlines.
Second, we estimate the area of ocean floor per
unit age {dA/dt} as a function of age by dividing each
{dA} by the time {dt} between each isochron pair
(Fig. 2; Table 1, and Table A in the online appendix).
A comparison with the estimates of Sclater et al. [3]
will be discussed below. The errors on {dt} have been
largely discussed in the literature, and may be the
cause of significant variations in {dA/dt} (e.g.
[2,14,19]). Following the error map proposed by
Royer et al. [1] and Mqller et al. [2], we assume an
average constant error of F1 Ma on each isochron,
thus on each area boundary, resulting in a FM2 Ma
error on the {dt} parameter.
Finally, we calculated spreading rates by dividing
the {dA/dt} values by the isochron length {L} (Table
A in the online appendix) at each epoch. L is
computed by summing up the angular distance
Fig. 1. Illustration of the method used to compute (a) surfaces and
(b) length of productive segments of isochrons. The surfaces are
integrated by summing up the area e of spherical triangles {a,b,c}
made up of each consecutive pair of points defining the outline, the
third point being the vector sum of all the points (star); this
necessitates continuous, closed outlines, in order to be able to
remove overlapping parts of triangles. The ridge lengths are
computed by summing up the angular distances h i between each
pair of points defining the productive segments of the isochrons.
between each pair of points defining the bproductiveQ
segments of each isochron (Fig. 1). We quantified the
uncertainty on {L} by comparing the measurements
of symmetric pairs of isochrons in both (symmetric)
Atlantic and Indian basins. It appeared to vary
between 5% and 10% of the total length. We therefore
assumed a 7% error on these determinations. The {dA/
dt/L} parameter (Fig. 3; Table A in the online
appendix) obtained for each of the Atlantic, Indian,
Pacific, Nazca, Cocos and Antarctic basins expresses
the area of oceanic crust per unit age and unit ridge
length as a function of age. This therefore represents
an average spreading rate (in mm year1) for each
basin. The error on these computations was obtained
using the classical error propagations calculations on
{dA}, {dt} and {L}.
2.2. Atlantic Ocean
Because the Atlantic Ocean is symmetric and
bounded by passive margins, the {dA} parameters
(Table A(a)) were simple to measure. We arbitrarily
cut the Atlantic Ocean from the Indian Ocean at about
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J.-P. Cogné, E. Humler / Earth and Planetary Science Letters 227 (2004) 427–439
Fig. 2. Area of the ocean floor per unit age {dA/dt} as a function of age for (a) Atlantic, (b) Indian, (c) Antarctic, (d) West Pacific, Nazca, and
Cocos and (e) Global ocean floors. The straight dotted line in (e) is the best-fit line over points given by Eq. (1) with C o=2.850 km2 year1 and
t m=180.2 Ma.
J.-P. Cogné, E. Humler / Earth and Planetary Science Letters 227 (2004) 427–439
Table 1
Seafloor spreading data as computed from Isochrons 0 to M29, after
data from Table A in the online appendix
Chron
t
(Ma)
Dt
(Ma)
dA
Million
km2
dA/dt a
km2/yr
Lb
km
0
5
6
13
18
21
25
31
34
M0
M4
M10
M16
M21
M25
M29
NM29
0
9.7
20.1
33.1
40.1
47.9
55.9
67.7
83.5
120.4
126.7
131.9
139.6
147.7
154.3
160.5
180
9.7
10.4
13
7
7.8
8
11.8
15.8
36.9
6.3
5.2
7.7
8.1
6.6
6.2
19.5
31.230
25.587
35.502
14.099
14.113
15.819
20.623
25.266
56.559
4.563
4.111
4.089
4.191
3.918
5.550
2.488
0.116
3.220F0.244
2.460F0.178
2.731F0.180
2.014F0.220
1.809F0.201
1.977F0.211
1.748F0.140
1.599F0.102
1.533F0.059
0.724F0.111
0.791F0.140
0.531F0.083
0.517F0.075
0.578F0.091
0.641F0.146
0.128F0.013
51747.1F1999.8
51421.3F1986.8
51098.3F2003.6
47393.1F1898.2
45623.8F1857.6
45383.5F1871.6
42401.4F1730.9
39412.2F1616.4
33449.9F1380.7
16122.2F663.2
17320.4F718.5
11636.1F524.0
12721.2F603.4
12499.7F593.1
10301.4F550.6
4155.8F290.9
t: isochron age; dt: time interval between 2 isochrons; dA: area of
crust generated between 2 isochrons; dA/dt: area per unit age; L:
isochron (i.e. ridge) length.
a
Total area using the Indian ocean without its Antarctic part
(see Section 2.3).
b
Total isochron length excluding Antarctic, Nazca and Cocos
Isochrons (see Section 2.6).
108E longitude, south of Africa. The ridge lengths
{L} given in Table A(a) were evaluated by averaging
the measure of symmetric pairs of isochrons from
both sides of the Atlantic Ridge.
2.3. Indian Ocean
The Indian Ocean is a little bit more complicated. It
is composed of symmetric crust segments around
three main ridges: the Central Indian Ridge (CIR), the
South East Indian Ridge (SEIR) and the South West
Indian Ridge (SWIR). We calculated the ridge lengths
{L} by averaging the values measured on symmetric
isochron pairs from both sides of these ridges as in the
Atlantic Ocean (Table A(b)). Similarly, we computed
the spreading rates {dA/dt/L} using {dA} values from
both sides, including the Antarctic part of this ocean,
south of SEIR and SWIR. However, to compute the
Global {dA} of Table 1, we had to exclude the
Antarctic part of the Indian Ocean, to avoid adding
this crust segment twice. The {dA} and {dA/dt}
431
parameters of the Indian Ocean excluding the
Antarctic part are given in Table A(c).
2.4. Antarctic plate
The surface {dA} and isochron lengths {L} of
Antarctic plate are given in Table A(d). They allowed
to compute spreading rates {dA/dt/L} which, compared to Atlantic and Indian Oceans, represent halfspreading rates. It should be noticed that the global
data given in Table 1 do include {dA} from the
Antarctic, but exclude ridge lengths {L} from this
plate because they are already included in the Indian,
Pacific and Nazca estimates (see below).
2.5. Pacific, Nazca and Cocos plates
The Pacific plate is bounded to the east by the East
Pacific Rise (EPR), and to the south by the Pacific–
Antarctic Rise. The Nazca plate is bounded by the
Chile Rise, the EPR and the Cocos–Nazca Ridge, and
the Cocos plate is bounded by the EPR and the
Cocos–Nazca Ridge. The spreading rates {dA/dt/L}
(Table A(e, f and g)) were determined independently
within these three plates and represent half-spreading
rates. However, the total ridge lengths given in Table
1 were obtained by summing the Pacific {L} with
only the Cocos–Nazca and Chile Rise parts of the
Nazca plate, which are given in Table A(h).
2.6. Summary
The Global surfaces {dA} of Table 1 were
obtained by summing the Atlantic, Indian without
Antarctic, Antarctic, Pacific, Nazca and Cocos surfaces quoted in Table A. The Global ridge lengths {L}
(Table 1) were calculated by summing the Atlantic,
Indian, Pacific, Chile and Cocos–Nazca {L} of Table
A, excluding other Nazca and Cocos isochrons as well
as Antarctic isochrons.
3. Results
3.1. Surfaces
The area of oceanic crust per unit time {dA/dt} as a
function of time (Table 1, and Table A in the online
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J.-P. Cogné, E. Humler / Earth and Planetary Science Letters 227 (2004) 427–439
Fig. 3. Half-spreading rates on (a) Atlantic, (b) Indian, (c) Antarctic, (d) Pacific, Nazca, and Cocos ridges as a function of age. (e) Average halfspreading rates without (thin line) and with (bold line) the Pacific, Nazca and Cocos data. The bold dotted lines with dark grey areas are the
averages computed for the 0–33.1, 33.1–120.4, and 120.4–147.7 Ma periods (see Section 3.2).
J.-P. Cogné, E. Humler / Earth and Planetary Science Letters 227 (2004) 427–439
appendix) is illustrated in Fig. 2 for each of the main
ocean basins. It is clear from this plot that a large
variability exists between them. Although seafloor
production of the Atlantic has remained fairly stable
since the opening of the South Atlantic Ridge, the
Indian ocean exhibits a jump in production between
chrons 21 and 31 (47.9 to 67.7 Ma), the period during
which the India plate rapidly drifted to the north,
before colliding with Eurasia. The Antarctic plate
displays a regularly increasing spreading rate since
chron M0, which is due to the fact that it is entirely
bounded by ridges. The West Pacific plate exhibits a
more sporadic production rate, with period of constant
production between chrons M21 and M0 (147.7 to
120.4 Ma), and between M0 and isochron 13 (120.4
to 33.1 Ma).
Early methods to evaluate the evolution of the
(observable) total amount of oceanic crust per unit age
{dA/dt} as a function of time {t} employed the
relationship:
dA
t
¼ Co 1 ð1Þ
dt
tm
where C o is the present-day oceanic crust production
in km2 year1, and t m the maximum age of the crust
(Sclater et al. [3] and Parsons [27]). They found
roughly linear rate on global scale with C o of 3.45
km2 year1 and t m=180 Ma, for a total measured
surface of 310 106 km2. The more recent evaluations of these parameters by Rowley [4] give t m=182
Ma and C o=2.960 km2 year1.
Our results are similar to those one proposed by
Sclater et al. [3] and Rowley [4]. However, the best fit
line over our points defines a C o parameter of
2.850F0.119 km2 year1 which is significantly lower
than the value reported by Sclater et al. [3]. Our t m is
strictly identical (t m=180.2F9.7 Ma) to that of Sclater
et al. [3] but the surface either measured (268106
km2), or obtained by integrating the above equation (1/
2C ot m=257106 km2) is significantly (~15%) lower
than their value. However, our estimate of C o is
reasonably close to the values proposed by Rowley
([4]; C o=2.96 km2 year1) and Parsons ([27]; C o=2.51
km2 year1), with the latter being based on an
evaluation of consumption rates at subduction zones.
We follow Rowley [4] in concluding that discrepancies between Sclater et al. [3] and other studies
433
probably result from the neglected back-arc spreading
and marginal basins which may contribute ~0.38 km2
year1 to the accretion rate estimates [27]. Adding this
to our evaluation leads to C o=3.230 km2 year1, which
is almost identical to previous estimates [3,4,12,27].
3.2. Spreading rates
In order to evaluate the variations in spreading rate
through geological time, we normalized the measured
{dA/dt} quantities by the ridge lengths {L} at the time
of production (Table 1). Not surprisingly, because
most of the remaining ridge traces result from the
breakup of Pangea, the ridges tend to lengthen with
time, except for the Pacific ridges which are currently
subducting. The spreading rates for each of the six
oceanic areas considered, computed as {dA/dt/L}, in
mm year1, are listed in Tables A and B and in the
online appendix and are illustrated in Fig. 3. Because
Pacific, Nazca, Cocos and Antarctic data are halfspreading rates, the spreading rates of the Atlantic and
Indian oceans were divided by two in Fig. 3 (Table A
in the online appendix).
Here again, a great variability between oceans is
obvious (Fig. 3). During the Cretaceous (between
chrons M0 and 31; 120.4 to 67.7 Ma), the Atlantic
ocean is characterized by an half-spreading rates of
~22 mm year1 followed by a decrease to the presentday half-rate of ~14 mm year1. In contrast, the
Indian ocean rates first decrease from ~22 mm year1
at chron M16 (139.6 Ma) to ~10 mm year1 at M0
(120.4 Ma), and then remain stable at 18–20 mm
year1 except for the high rates of ~30 mm year1
during the 67.7–47.9 Ma period (chrons 31 to 21),
which corresponds to the time of rapid northward drift
of the India plate. The Antarctic plate shows a
different trend, with a slight decrease from ~20 to
~10 mm year1 between M21 (147.7 Ma) and M0
(120.4 Ma), followed by a regular increase to present
value of ~30 mm year1. On average, the Pacific rates
are twice as large as rates from the other oceans. Apart
from the first two points (M29 and M25), which may
be questionable in terms of accuracy, the Pacific rates
tend to increase from ca. ~40 to ~75 mm year1 from
the end of the Jurassic to the beginning of the
Cretaceous, between M21 (147.7 Ma) and M0
(120.4 Ma). Production rates decrease to ~30 mm
year 1 at chron 25 (55.9 Ma) followed by a
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J.-P. Cogné, E. Humler / Earth and Planetary Science Letters 227 (2004) 427–439
significant, albeit scattered, increase to present-day
values of 60–70 mm year1. Such high values in
recent times are corroborated by the Nazca and Cocos
data.
From the above analysis, global spreading rates are
dominated by the Pacific ocean. We however attempt
to define a mean global spreading rate by averaging
the values of Table A. This global mean (Table 2) was
obtained by first averaging the Pacific, Nazca, and
Cocos rates, in order to have an average spreading rate
for the Pacific system. This value was then averaged
with the Antarctic, half Indian and half Atlantic rates.
The temporal evolution of spreading rates (Table 2,
Fig. 3), computed only from currently observable
elements, shows no significant variation since at least
155 Ma. There are slight variations in the mean halfspreading rates (Table 2). The mean before 83.5 Ma
(chron 34) is: 25.0F1.3 mm year1, 23.3F1.9 mm
Table 2
Mean half-spreading rates computed after the data of Table 1, using
the average rates of Pacific, Nazca and Cocos plates, half-rates for
Indian and Atlantic oceans, and full rates for Antarctic plate
Chron
t
(Ma)
Dt
(Ma)
Mean
(mm
yr1)
0
5
6
13
18
21
25
31
34
M0
M4
M10
M16
M21
M25
M29
NM29
Averages
0.0
9.7
20.1
33.1
40.1
47.9
55.9
67.7
83.5
120.4
126.7
131.9
139.6
147.7
154.3
160.5
180.0
9.7
10.4
13.0
7.0
7.8
8.0
11.8
15.8
36.9
6.3
5.2
7.7
8.1
6.6
6.2
19.5
33.2
27.1
30.8
29.2
21.4
25.9
22.9
22.8
24.5
26.5
22.9
25.9
24.9
25.3
45.2a
30.7a
–
25.9
30.1
23.3
25.0
All:
0.0–40.1 Ma:
40.1–83.5 Ma:
N83.5 Ma:
n
4
4
4
4
4
4
4
4
4
4
4
4
4
4
2
1
14
4
4
6
r
(mm
yr1)
20.6
15.7
21.1
21.8
11.6
11.3
10.0
7.4
15.4
32.2
13.6
21.4
13.2
12.9
45.2
–
3.3
2.6
1.9
1.3
n: number of entries in the statistics, r: standard deviation of the
means.
a
Excluded from the final means.
year1 between 40.1 and 83.5 Ma (chrons 34 to 18),
and 30.1F2.6 mm year1 since 40.1 Ma. However,
even if we consider these differences as significant
within the 2r error bars, the most striking and
surprising feature is that the average spreading rates
do not show a significant increase during the Cretaceous, as generally believed, but remain fairly stable
at 25.9F3.3 mm year1.
3.3. Seafloor generation rates
Beyond knowing the spreading rates themselves,
the question of whether seafloor production underwent significant variations in surface is relevant to
which length of ridges these rates applied to.
Following Kominz [14], we assume that the total
ridge length remained roughly constant with time
(e.g.Tables 1a and b of Kominz [14]), and multiplying
our constant spreading rate by a constant ridge length
leads to a constant seafloor generation rate. However,
we note that the higher spreading rates are biased by
the larger production centers in the past, in particular,
the Pacific system of ridges. This bias makes it
necessary to re-evaluate the time-dependent production rate in another way, which implies, as discussed
in Kominz [14], a number of hypotheses and
assumptions. The main problem is to recover disappeared (subducted) ridges. From late Jurassic to
present, two main oceanic ridges have undergone
significant changes: the Tethys ridge(s), and the paleoPacific ridges.
The geologic evolution of the Tethyan domain was
synthesized by Ricou [23], who also calculated the
rotation parameters between the main plates. From
these data, paleogeographic maps were drawn to
reconstruct plate position through time, which
included hypothesized ridges and subduction zones.
We digitized these ridges from the maps [23], in order
to obtain the Tethys ridge length over time (Table 3).
For the paleo-Pacific ocean, we followed Engebretson
et al. [22], and digitized the ridges separating the
Izanagi (Iz), Farallon (Fa), Kula (Ku) and Pacific (Pa)
plates, for the relevant time (Table 3). Fig. 4a
illustrates the temporal evolution of total ridge length,
following the above assumptions. Our results suggest
global ridge lengths were slightly longer during the
Mesozoic, which could also be the result of a poor
determination of ridges lost by subduction. Despite
J.-P. Cogné, E. Humler / Earth and Planetary Science Letters 227 (2004) 427–439
435
Table 3
Hypothesized subducted Ridges (km); interpolated values are in italics
Chron
t
(Ma)
Dt
(Ma)
0
5
6
13
0.0
9.7
20.1
33.1
37.0
40.1
45.0
47.9
55.9
65.0
67.7
80.0
83.5
95.0
110.0
120.4
126.7
131.9
135.0
139.6
145.0
147.7
154.3
160.5
180.0
9.7
10.4
13.0
3.9
3.1
4.9
2.9
8.0
9.1
2.7
12.3
3.5
11.5
15.0
10.4
6.3
5.2
3.1
4.6
5.4
2.7
6.6
6.2
19.5
18
21
25
31
34
M0
M4
M10
M16
M21
M25
M29
NM29
Tethysa
Fa/Izb
Iz/Pab
Fa/Pab
Fa/Kub
(c)
780.3
967.9
555.0
6939.7
6870.9
6681.1
6464.6
6400.3
7151.7
7364.9
8067.5
12599.6
9428.3
8925.5
8510.4
8264.0
8531.5
8845.6
8806.7
8693.5
8585.9
9479.4
1852.6
2750.6
3295.6
3745.1
4013.8
4411.1
4440.0
4440.0
4440.0
4440.0
0.0
2455.3
4910.6
7366.0
7018.5
6816.5
6648.9
6550.1
6402.5
5550.0
5550.0
5550.0
5550.0
3765.1
3705.2
3669.7
3640.8
3624.2
3598.6
3663.0
3663.0
3663.0
3663.0
444.0
444.0
777.0
999.0
1521.8
3313.4
3495.4
4325.7
4440.0
4440.0
4440.0
4440.0
4440.0
4440.0
4440.0
4440.0
4440.0
4440.0
4440.0
4440.0
Total Pacific ridges
0.0
780.3F78.0
967.9F96.8
555.0F55.5
444.0F66.6
444.0F66.6
777.0F116.6
999.0F149.9
1521.8F228.3
3313.4F497.0
3495.4F524.3
4325.7F648.9
6895.3F961.3
9350.6F1153.7
17423.7F1551.0
17914.3F1541.4
18221.8F1543.2
18474.8F1548.2
18628.0F1553.0
18852.2F1562.1
18093.0F1829.2
18093.0F1829.2
18093.0F1829.2
18093.0F1829.2
Fa: Farallon, Iz: Izanagi, Pa: Pacific, Ku: Kula.
Uncertainties on the total are evaluated by assuming a 20% uncertainty on Tethys estimates, and a 15% uncertainty for the other disappeared
ridges.
a
Following Ricou [23].
b
Following Engebretson et al. [22].
c
Subducted Nazca isochrons after Royer et al. [1].
this, our evaluation is in quite good agreement with
Kominz [14].
The next step in evaluating the surficial global
seafloor production rates is to assign spreading rates
to these ridges. It seems appropriate to assign the
Pacific spreading rates, as we determined above, to
the subducted paleo-Pacific ridges, especially because
the older rates were already determined on parts of the
segments that are almost entirely subducted. In
contrast, the spreading rates of the Tethys sea are far
less constrained, as this ocean has entirely disappeared. We thus decided to model seafloor productions in two ways: (1) the more conservative, and less
favoured option applied the average global spreading
rates we have determined above to the consumed
ridges of the Tethysian and paleo-Pacific oceans; (2)
the more realistic, and preferred option assigns the
Pacific spreading rates to the paleo-Pacific ridges, and
the Indian ocean rates to the Tethys ridges. The
production rates {dA/dt} of the subducted ridges are
then added to the measured {dA/dt} of the presentday oceanic crust.
The seafloor half-production rates from the present
measurements, and from Tables 3 and 4 are illustrated
in Fig. 4b. The total seafloor generation rates do not
show large variations with time, in either option 1 or
2, and generally vary within F20% about the mean.
We note again that the average half-production rate of
option 2 in the last 180 Ma, established at
1.298F0.284 km2 year 1 (full rate=2.596 km 2
year1), compares well with the current-day production rate of 2.850 km2 year1 evaluated above.
436
J.-P. Cogné, E. Humler / Earth and Planetary Science Letters 227 (2004) 427–439
Fig. 4. (a) Ridge lengths as measured on present-day isochrons (in blue), and as hypothesized from Engebretson et al. [22] for paleo-Pacific and
Ricou [23] for Tethys (in red), and total ridge lengths (in bold black lines) as a function of age; the data of Kominz [14] are given in thin black
lines. (b) Half-production rates in km2 year1 as a function of age following our two options 1 (dotted blue line) and 2 (bold black line); red line:
average half seafloor generation rate from option 2 (see Section 3.3); blue continuous line: actual {dA/dt}, as measured on present-day surfaces;
thin continuous black curve is the 1st order sea-level variations curve (right scale) following Hardenbol et al. [17], the blue curve is a polynomial
smoothing of this curve. Colored horizontal bars in (a) and (b) are the magnetic timescales as proposed by Royer et al. [1] and Mqller et al. [2].
However, it seems that option 2 coincides with an
increase in production rate, which reaches 42F20% of
the mean rate, between M4 (126.7 Ma) and M0 (120.4
Ma), at the beginning of the Cretaceous. Because the
period between M0 and M4 is short, a slight shift in
time of one or both of these chrons could lead to a
drastic change in the generation rate. Furthermore, the
rates then rapidly decrease during the Cretaceous,
falling down to minimum values (~1 km2 year1)
around chron 31 (67.7 Ma) and chron 18 (40.1 Ma). A
similar evolution is obvious using option 1, yet with
lower amplitudes, in particular during the Cretaceous.
Another quite surprising feature, obvious from both
options 1 and 2, is that the production rates have
tended to increase since 40.1 Ma.
3.4. Oceanic crust production rates
It is straightforward to evaluate magmatic fluxes at
oceanic ridges provided surface production rates and
crustal thicknesses are known. The chemical composition (major and trace elements) of oceanic basalt is
important because it can be quantitatively linked to
mantle temperature [28,29,31]. Humler et al. [5]
investigated ancient crustal compositions from the
Atlantic, Pacific and Indian oceans. They found that
the chemical composition of oceanic crust older than
80 Ma was statistically different from the modern one.
In order to explain this observation, they proposed
that the mantle temperature was 50 8C hotter prior to
80 Ma, whereas the mean mantle temperature of the
mantle after 80 Ma is identical to the temperature
beneath the modern oceanic system [32].
On the other hand, the amount of melt in the
mantle beneath an oceanic ridge is controlled by the
amount of pressure release that each individual
parcel of mantle has experienced (e.g. [28,29]). The
higher the mantle temperature, the more melt
produced, and hence the greater the thickness of
the oceanic crust (e.g. [30]). Regardless of the
origin of mantle temperature variations [32,33], the
chemical data suggest that crustal thicknesses were
1–2 km thicker before 80 Ma (~9 km) than at
present (~7 km) [5] . Klein and Langmuir [28]
found a positive global correlation between the
extent of melting, derived from the major-element
chemistry of basalts, and seismically determined
crustal thicknesses.
J.-P. Cogné, E. Humler / Earth and Planetary Science Letters 227 (2004) 427–439
Table 4
Average seafloor half-production rates
Chron
t
(Ma)
dt
(Ma)
Option 1
(km2 yr1)
Option 2
(km2 yr1)
0
5
6
13
0.0
9.7
20.1
33.1
37.0
40.1
45.0
47.9
55.9
65.0
67.7
80.0
83.5
95.0
110.0
120.4
126.7
131.9
135.0
139.6
145.0
147.7
154.3
160.5
180.0
9.7
10.4
13.0
3.9
3.1
4.9
2.9
8.0
9.1
2.7
12.3
3.5
11.5
15.0
10.4
6.3
5.2
3.1
4.6
5.4
2.7
6.6
6.2
19.5
1.610F0.119
1.251F0.088
1.395F0.090
1.023F0.096
1.020F0.095
0.914F0.087
1.070F0.148
1.193F0.158
1.062F0.133
1.098F0.147
1.026F0.114
1.062F0.129
1.117F0.264
1.194F0.318
1.503F0.535
1.087F0.946
1.017F0.429
0.964F0.638
0.962F0.635
0.941F0.420
0.929F0.420
0.970F0.414
1.378F1.327
0.923F0.077
1.610F0.119
1.260F0.087
1.427F0.089
1.025F0.095
1.021F0.095
0.923F0.088
1.054F0.105
1.228F0.138
1.139F0.119
1.186F0.121
1.040F0.078
1.082F0.086
1.218F0.103
1.356F0.137
1.841F0.239
1.808F0.571
1.331F0.393
1.497F0.405
1.501F0.407
1.263F0.312
1.235F0.316
1.238F0.336
1.957F0.766
0.923F0.161
1.113F0.190
1.298F0.284
18
21
25
31
34
M0
M4
M10
M16
M21
M25
M29
NM29
Averages
In effect, using the average seismic crustal thickness of oceanic crust from White et al. [35] and
Nagihara et al. [36], one obtains an average of
6.6F0.8 km (n=24) for 0–15 Ma oceanic crust, and
of 7.4F1.0 km (n=34) for crust older than 80 Ma [5].
Albeit ~1 km smaller, these estimates are in good
agreement with those deduced from the chemical
composition of oceanic basalts. Moreover, the geochemical results suggest that the crust was not grown
thicker [37,38] but rather that old crust was thicker at
the time of its formation.
As a consequence, despite the constant surface
generation rates reported here at oceanic ridges for
the last 180 Ma, production rates changed due to
crustal thickness variations linked to mantle temperature variations. Thus, for 0–80 Ma, we propose
that the magmatic flux at oceanic ridges was
18.1F3.9 km3 year1, being the product of the
average surface generation rate 2.596F0.568 km2
year1 multiplied by a 7-km-thick crust. In contrast, during the Mesozoic, the magmatic flux was
437
higher by about 30% (23.3F5.1 km3 year1 for a
9-km-thick crust). Accounting for seismically determined thicknesses, this difference could be reduced
to ~15%.
Note that our analysis has neglected the volumes of
oceanic plateaus. An evaluation of oceanic plateaus
production was given by Larson [19], following the
work of Schubert and Sandwell [34]. The mean
oceanic plateau production rate was about 3 km3
year1 since 150 Ma [19], some 60% lower than the
excess production rate at ocean spreading ridges
during the Mesozoic. We therefore conclude that,
although oceanic plateau production is an important
factor, the effect of mantle temperature variations on
ridge production is far from being negligible, and may
be the major cause in the balance of mass production
within ocean basins.
4. Discussion and conclusion
We have calculated/measured the area of oceanic
crust per unit age {dA/dt} as a function of age, from
the isochron database of Royer et al. [1] and Mqller et
al. [2]. Our results are in fairly good agreement with
the previous estimates of Sclater et al. [3], even
though we used a different and probably more precise
method of computing surface areas on the Earth, as
well as an updated data set of Isochron positions on
the Earth, and magnetic timescales.
We then computed the spreading rates within
each oceanic basin by dividing these surfaces by the
ridge lengths as measured from the isochron
lengths. These calculations show that the global
average half-spreading rates have remained constant
since at least 150 Ma, at 25.9F3.3 mm year1.
Making some hypotheses on subducted paleoPacific ridges (following [22]) and Tethys ridges
(following [23]), and on the velocity on these
ridges, we finally propose a roughly constant
seafloor surface generation rate since 180 Ma,
within F20% variation around a mean (half) value
at 1.298F0.284 km2 year1 (Table 4). There may
exist a slightly higher production rate at the
beginning of the Cretaceous, but this observation
relies on hypothetical ridges, with unknown velocities, and, as a consequence, should be considered
questionable.
438
J.-P. Cogné, E. Humler / Earth and Planetary Science Letters 227 (2004) 427–439
Our evaluation of the seafloor surface generation
half rate is notably lower than the full production of 3.4
km2 year1 proposed by Rowley [4], because we did
not take into account small ocean basins. However, in
agreement with Rowley [4], but based on a different
method, we conclude that spreading rates remained
constant since at least 180 Ma. Importantly, the fact
that both studies reach similar conclusions using quite
different methods clearly shows that the initial
assumptions used in the integration of Rowley [4]
are valid: (1) the destruction of oceanic floor at
subduction zones is uniformly distributed with age,
and (2) from present to 180 Ma, the oldest age of
oceanic crust remained constant at 180 Ma.
If significant, the evolution of seafloor surface
generation rates during the last 180 Ma is in
significant disagreement with the results of Kominz
[14], who suggested spreading rates in the Cretaceous, between 120 and 80 Ma, attained two to four
times the present-day value, followed by a regular
decrease since then. Apart from the fact that we
found an increase in production rates since 40 My,
which is not assumed by Kominz [14] model, the
Cretaceous appears to us as a period characterized by
a particularly quiet, and fairly slow, seafloor generation rate. We point out, however, that the
determinations of Kominz [14] are primarily based
on Pacific plate reconstruction models, involving a
significant amount of subducted oceanic crust. The
Kominz model [14] also assumes symmetric spreading and neglects the effects of ridge jumps during the
Cretaceous LNS. It is also bothersome that the
inferred high spreading rates take place during the
LNS, a period for which, by definition, no precise
description of seafloor spreading history can be
obtained. Another limiting factor of this work is that
the timescales were poorly defined at that time. We
therefore conclude that our evaluation, which is
based on much more conservative methods and
assumptions, and probably more accurate timescales,
provides a more accurate description of seafloor
generation rates.
From our results, no clear correlation appears
between spreading or surface generation rates and
the 1st order variations in sea-level ([17]; Fig. 4b).
Following Heller et al. [20] and Hardebeck and
Anderson [21], who proposed Pangea breakup as an
alternative mechanism, we also conclude that the 1st
order sea-level changes are not primarily controlled by
the average seafloor spreading rates, as proposed by
Hays and Pitman [13]. However, instead of Pangea
breakup, we suggest that variations in oceanic crust
thickness related to mantle temperature variations
since the Mesozoic could account for the sea-level
changes.
Finally, our analysis shows that despite a constant
spreading history during the past 180 Ma, the crust
production at ridges could actually be higher during
the Mesozoic because the mantle was hotter. Although
this mechanism differs from previous suggestions, it
provides a consistent alternative hypothesis for the
high CO2 concentration in the Earth’s atmosphere
during the Cretaceous.
Acknowledgements
We thank L.E. Ricou, J. Dyment and V. Courtillot
for their thoughtful insights and discussions all along
this work. J.Y. Royer and J. Francheteau are thanked
for their helpful discussions. W. Crawford, S. Gilder
and J. Ludden have greatly improved the final version
of this manuscript. We also thank D. Mqller, D.
Rowley and an anonymous reviewer for their comments and corrections.
Appendix A. Supplementary data
Supplementary data associated with this article can
be found, in the online version, at doi:10.1016/
j.epsl.2004.09.002.
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