**Gavin A. Schmidt** is a climatologist, climate modeler and Direct...
I’ve pondered this ‘possibility’ all my length - that there have be...
One of the most common searched structures that can be seen as an e...
If you are curious here is below is a clip of the first appearance ...
The Negev Desert in Israel has been dated to about 1.8 million year...
It has been estimated that only one bone in a billion gets fossilis...
Carbon has 3 naturally occurring isotopes
- $^{12}C$, about 99% of...
Another element that would be detectable is technetium-98 ($^{98}T$...
The problem is that if the activity of a hypotetical industrial civ...
International Journal of
Astrobiology
cambridge.org/ija
Research Article
Cite this article: Schmidt GA, Frank A (2019).
The Silurian hypothesis: would it be possible
to detect an industrial civilization in the
geological record? International Journal of
Astrobiology 18, 142–150. https://doi.org/
10.1017/S1473550418000095
Received: 11 October 2017
Revised: 21 February 2018
Accepted: 25 February 2018
First published online: 16 April 2018
Key words:
Astrobiology; Drake equation; industrial
civilization; Silurian hypothesis; Anthropocene;
PETM
Author for correspondence:
Gavin A. Schmidt,
E-mail:gavin.a.schmidt@nasa.gov
© Cambridge University Press 2018
The Silurian hypothesis: would it be possible to
detect an industrial civilization in the
geological record?
Gavin A. Schmidt
1
and Adam Frank
2
1
NASA Goddard Institute for Space Studies, 2880 Broadway, New York, NY 10025, USA and
2
Department of Physics
and Astronomy, University of Rochester, Rochester, NY 14620, USA
Abstract
If an industrial civilization had existed on Earth many millions of years prior to our own era,
what traces would it have left and would they be detectable today? We summarize the likely
geological fingerprint of the Anthropocene, and demonstrate that while clear, it will not differ
greatly in many respects from other known events in the geological record. We then propose
tests that could plausibly distinguish an industrial cause from an otherwise naturally occurring
climate event.
Introduction
The search for life elsewhere in the universe is a central occupation of astrobiology and scien-
tists have often looked to Earth analogues for extremophile bacteria, life under varying climate
states and the genesis of life itself. A subset of this search is the prospect for intelligent life, and
then a further subset is the search for civilizations that have the potential to communicate with
us. A common assumption is that any such civilization must have developed industry of some
sort. In particular, the ability to harness those industrial processes to develop radio technolo-
gies capable of sending or receiving messages. In what follows, however, we will define indus-
trial civilizations here as the ability to harness external energy sources at global scales.
One of the key questions in assessing the likelihood of finding such a civilization is an
understanding of how often, given that life has arisen and that some species are intelligent,
does an industrial civilization develop? Humans are the only example we know of, and our
industrial civilization has lasted (so far) roughly 300 years (since, for example, the beginning
of mass production methods). This is a small fraction of the time we have existed as a species,
and a tiny fraction of the time that complex life has existed on the Earth’s land surface (∼400
million years ago, Ma). This short time period raises the obvious question as to whether this
could have happened before. We term this the ‘Silurian hypothesis’
1
.
While much idle speculation and late night chatter has been devoted to this question, we
are unaware of previous serious treatments of the problem of detectability of prior terrestrial
industrial civilizations in the geologic past. Given the vast increase in work surrounding exo-
planets and questions related to detect ion of life, it is worth addressing the question mo re for-
mally and in its astrobiological setting. We note also the recent work of Wright (2017) which
addressed aspects of the problem and previous attempts to assess the likelihood of solar system
non-terrestrial civilization such as Haqq-Misra & Kopparapu (2012). This paper is an attempt
to remedy the gap in a way that also puts our current impact on the planet into a broader per-
spective. We first note the importance of this question to the well-known Drake equation.
Then we address the likely geologic consequences of human industrial civilization and then
compare that fingerprint to potentially similar events in the geologic record. Finally, we
address some possible research directions that might improve the constraints on this question.
Relevance to the Drake equation
The Drake equation is the well-known framework for estimating of the number of active, com-
municative extraterrestrial civilizations in the Milky Way galaxy (Drake, 1961, 1965). The num-
ber of such civilizations, N, is assumed to be equal to the product of; the average rate of star
formation, R*, in our Galaxy; the fraction of formed stars, f
p
, that have planets; the average num-
ber of planets per star, n
e
, that can potentially support life; the fraction of those planets, f
l
,that
1
We name the hypothesis after a 1970 episode of the British science fiction TV series Doctor Who where a long-buried race
of intelligent reptiles ‘Silurians’ are awakened by an experimental nuclear reactor. We are not however suggesting that intelligent
reptiles actually existed in the Silurian age, nor that experimental nuclear physics is liable to wake them from hibernation. Other
authors have dealt with this possibility in various incarnations (for instance, Hogan, 1977), but it is a rarer theme than we ini-
tially assumed.
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actually develop life; the fraction of planets bearing life on which
intelligent, civilized life, f
i
, has developed; the fraction of these civi-
lizations that have developed communications, f
c
, i.e., technologies
that release detectable signs into space, and the length of time, L,
over which such civilizations release detectable signals.
N = R
∗
· f
p
· n
e
· f
l
· f
i
· f
c
· L.
If over the course of a planet’s existence, multiple industrial
civilizations can arise over the span of time that life exists at all,
the value of f
c
may in fact be >1.
This is a particularly cogent issue in light of recent developments
in astrobiology in which the first three terms, which all involve
purely astronomical observations, have now been fully determined.
It is now apparent that most stars harbour families of planets
(Seager, 2013). Indeed, many of those planets will be in the star’s
habitable zones (Dressing & Charbonneau, 2013; Howard, 2013).
These results allow the next three terms to be bracketed in a way
that uses the exoplanet data to establish a constraint on exo-
civilization pessimism. In Frank & Sullivan (2016)sucha‘pessimism
line’ was defined as the maximum ‘biotechnological’ probability
(per habitable zone planets) f
bt
for humans to be the only time a
technological civilization has evolved in cosmic history. Frank &
Sullivan (2016)foundf
bt
in the range ∼10
−24
–10
−22
.
Determination of the ‘pessimism line’ emphasizes the
importance of three Drake equation terms f
l
, f
i
and f
c
.Earth’s
history often serves as a template for discussions of possible
values for these probabilities. For example, there has been con-
siderable discussion of how many times life began on Earth dur-
ing the early A rchean given the ease of abiogenisis (Patel et al.,
2015) including the possibility of a ‘shadow biosphere’ com-
posed o f d escendants of a different origin event from the one
which led to our Last Univers al Common Ancestor (LUCA)
(Cleland & Copley, 2006). In addition, there is a long-standing
debate concerning the number of times intelligence has evolved
in te rms of dolphins and other species (Marino, 2015). Thus,
only the term f
c
has been co mmonly accepted to have a value
on Earth of strictly 1.
Relevance to other solar system planets
Consideration of previous civilizations on other solar system
worlds has been taken on by Wright ( 2017) and Haqq-Misra &
Kopparapu (2012). We note here that abundant evidence exists
of surface water in ancient Martian climates (3.8 Ga) (e.g.
Achille & Hynek, 2010; Arvidson et al., 2014), and speculation
that early Venus (2 Ga to 0.7 Ga) was habitable (due to a dimmer
sun and lower CO
2
atmosphere) has been supported by recent
modelling studies (Way et al., 2016). Conceivably, deep drilling
operations could be carried out on either planet in future to assess
their geological history. This would constrain consideration of
what the fingerprint might be of life, and even organized civiliza-
tion (Haqq-Misra & Kopparapu, 2012). Assessments of prior
Earth events and consideration of Anthropocene markers such
as those we carry out below will likely provide a key context for
those explorations.
Limitations of the geological record
That this paper’s title question is worth posing is a function of the
incompleteness of the geological record. For the Quaternary (the
last 2.5 million years), there is widespread extant physical evi-
dence of, for instance, climate changes, soil horizons and archaeo-
logical evidence of non-Homo Sapiens cultures (Denisovians,
Neanderthals, etc. ) with occasional evidence of bipedal hominids
dating back to at least 3.7 Ma (e.g. the Laetoli footprints) (Leakey
&Hay,1979). The oldest extant large-scale surface is in the Negev
Desert and is ∼1.8 Ma old (Matmon et al., 2009). However,
pre-Quaternary land-evidence is far sparser, existing mainly in
exposed sections, drilling and mining operations. In the ocean
sediments, due to the recycling of ocean crust, there only exists
sediment evidence for periods that post-date the Jurassic
(∼170 Ma) (ODP Leg 801 Team, 2000).
The fraction of l ife that gets fossilized is always extremely
small and varies widely as a funct ion of time, habitat and degree
of soft tissue versus hard shells or bones (Behrensmeyer et al.,
2000). Fossil ization rates are very low in tropical, forested envir-
onments but are higher in arid environments and fluvial systems.
As an example, for all the dinosaurs that ever lived, there are
only a few thousand near-co mplete specimens, or equivalently
only a handful of individual animals across thousands o f taxa
per 100,000 years. Given the rate of new disc overy of taxa of
this age, it is clear that species as short-lived as Homo sapiens
(so far) might not be represented in the existing fossil record
at all.
The likelihood of objects surviving and being discovered is
similarly unlikely. Zalasiewicz (2009) speculates about preserva-
tion of objects or their forms, but the current area of urbanization
is <1% of the Earth’s surface (Schneider et al., 2009), and exposed
sections and drilling sites for pre-Quaternary surfaces are orders
of magnitude less as fractions of the original surface. Note that
even for early human technology, complex objects are very rarely
found. For instance, the Antikythera Mechanism (ca. 205 BCE) is
a unique object until the Renaissance. Despite impressive recent
gains in the ability to detect the wider impacts of civilization on
landscapes and ecosystems (Kidwell, 2015), we conclude that for
potential civilizations older than about 4 Ma, the chances of find-
ing direct evidence of their existence via objects or fossilized
examples of their population is small. We note, however, that
one might ask the indirect question related to antecedents in
the fossil record indicating species that might lead downstream
to the evolution of later civilization-building species. Such argu-
ments, for or against, the Silurian hypothesis would rest on evi-
dence concerning highly social behaviour or high intelligence
based on brain size. The claim would then be that there are
other species in the fossil record which could, or could not,
have evolved into civilization-builders. In this paper, however,
we focus on physicochemical tracers for previous industrial civili-
zations. In this way, there is an opportunity to widen the search to
tracers that are more widespread, even though they may be subject
to more varied interpretations.
Scope of this paper
We will restrict the scope of this paper to geochemical constraints
on the existence of pre-Quaternary industrial civilizations, that
may have existed since the rise of complex life on land. This
rules out societies that might have been highly organized and
potentially sophisticated but that did not develop industry
and probably any purely ocean-based lifeforms. The focus is
thus on the period between the emergence of complex life on
land in the Devonian (∼400 Ma) in the Paleozoic era and the
mid-Pliocene (∼4 Ma).
International Journal of Astrobiology 143
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The geological footprint of the Anthropocene
While an official declaration of the Anthropocene as a unique
geological era is still pending (Crutzen, 2002; Zalasiewicz et al.,
2017), it is already clear that our human efforts will impact the
geologic record being laid down today (Waters et al., 2014).
Some of the discussion of the specific boundary that will define
this new period is not relevant for our purposes because the mar-
kers proposed (ice core gas concentrations, short-half-lived radio-
activity, the Columbian exchange) (e.g. Lewis & Maslin, 2015;
Hamilton, 2016) are not going to be geologically stable or distin-
guishable on multi-million year timescales. However, there are
multiple changes that have already occurred that will persist.
We discuss a number of these below.
There is an interesting paradox in considering the
Anthropogenic footprint on a geological timescale. The longer
human civilization lasts, the larger the signal one would expect
in the record. However, the longer a civilization lasts, the more
sustainable its practices would need to have become in order to
survive. The more sustainable a society (e.g. in energy generation,
manufacturing or agriculture) the smaller the footprint on the rest
of the planet. But the smaller the footprint, the less of a signal will
be embedded in the geological record. Thus, the footprint of civ-
ilization might be self-limiting on a relatively short timescale. To
avoid speculating about the ultimate fate of humanity, we will
consider impacts that are already clear, or that are foreseeable
under plausible trajectories for the next century (e.g. Nazarenko
et al., 2015; Köhler, 2016).
We note that effective sedimentation rates in ocean sediment
for cores with multi-million-year-old sediment are of the order
of a few cm/1000 years at best, and while the degree of bioturb-
ation may smear a short-period signal, the Anthropocene will
likely only appear as a section a few cm thick, and appear almost
instantaneously in the record.
Stable isotope anomalies of carbon, oxygen, hydrogen and
nitrogen
Since the mid-18th century, humans have released over 0.5 trillion
tonnes of fossil carbon via the burning of coal, oil and natural gas
(Le Quéré et al., 2016), at a rate orders of magnitude faster than nat-
ural long-term sources or sinks. In addition, there has been wide-
spread deforestation and addition of carbon dioxide into the air
via biomass burning. All of this carbon is biological in origin and
is thus depleted in
13
C compared with the much larger pool of inor-
ganic carbon (Revelle & Suess, 1957). Thus, the ratio of
13
Cto
12
Cin
the atmosphere, ocean and soils is decreasing (an impact known as
the ‘Suess Effect’ Quay et al., 1992) with a current change of around
− 1‰ δ
13
C since the pre-industrial (Böhm et al., 2002;Eideetal.,
2017) in the surface ocean and atmosphere (Fig. 1(a)).
As a function of the increase of fossil carbon into the system,
augmented by black carbon changes, other non-CO
2
trace green-
house gases (e.g. N
2
O, CH
4
and chloro-fluoro-c arbons (CFCs)),
global industrialization has been accompanied by a warming of
about 1°C so far since the mid-19th century (Bindoff et al.,
2013; GISTEMP Team, 2016). Due to the temperature-related
fractionation in the formation of carbonates (Kim & O’Neil,
1997)(−0.2‰
d
18
O per °C) and strong correlation in the
extra-tropics between temperature and δ
18
O (between 0.4 and
0.7‰ per °C) (and ∼8× as sensitive for deuterium isotopes rela-
tive to hydrogen (δD)), we expect this temperature rise to be
detectable in surface ocean carbo nates (notably foraminifera),
organic biomarkers, cave records (stalactites), lake ostracods and
high-latitude ice cores, though only the first two of these will be
retrievable in the timescales considered here.
The combustion of fossil fuel, the invention of the Haber–
Bosch process, the large-scale application of nitrogenous fertili-
zers and the enhanced nitrogen fixation associated with cultivated
plants, have caused a profound impact on nitrogen cycling
(Canfield et al., 2010), such that δ
15
N anomalies are already
detectable in sediments remote from civilization (Holtgrieve
et al., 2011).
Sedimentological records
There are multiple causes of a greatly increased sediment flow in
rivers and therefore in deposition in coastal environments. The
advent of agriculture and associated deforestation have lead to
large increases in soil erosion (Goudie, 2000; National Research
Council, 2010). Furthermore, canalization of rivers (such as the
Mississippi) have led to much greater oceanic deposition of sedi-
ment than would otherwise have occurred. This tendency is miti-
gated somewhat by concurrent increases in river dams which
reduce sediment flow downstream. Additionally, increasing tem-
peratures and atmospheric water vapour content have led to
greater intensity of precipitation (Kunkel et al., 2013) which, on
its own, would also lead to greater erosion, at least regionally.
Coastal erosion is also on the increase as a function of the rising
sea level, and in polar regions is being enhanced by reductions in
sea ice and thawing permafrost (Overeem et al., 2011).
In addition to changes in the flux of sediment from land to
ocean, the composition of the sediment will also change. Due
to the increased dissolution of CO
2
in the ocean as a function
of anthropogenic CO
2
emissions, the upper ocean is acidifying
(a 26% increase in H
+
or 0.1 pH decrease since the 19th century)
(Orr et al., 2005). This will lead to an increase in CaCO
3
dissol-
ution within the sediment that will last until the ocean can neu-
tralize the increase. There will also be important changes in
mineralogy (Zalasiewicz et al., 2013; Hazen et al., 2017).
Increases in continental weathering are also likely to change ratios
of strontium and osmium (e.g.
87
Sr/
86
Sr and
187
Os/
188
Os ratios)
(Jenkyns, 2010).
As discussed above, nitrogen load in rivers is increasing as a
function of agricultural practices. This in turn is leading to
more microbial activity in the coastal ocean which can deplete
dissolved oxygen in the water column (Diaz & Rosenberg,
2008), and recent syntheses suggests a global decline already of
about 2% (Ito et al., 2017; Schmidtko et al., 2017). This in turn
is leading to an expansion of the oxygen minimum zones, greater
ocean anoxia and the creation of so-called ‘dead-zones’ (Breitburg
et al., 2018). Sediment within these areas will thus have greater
organic content and less bioturbation (Tyrrell, 2011). The ultim-
ate extent of these dead zones is unknown.
Furthermore, anthropogenic fluxes of lead, chromium, antim-
ony, rhenium, platinum group metals, rare earths and gold, are
now much larger than their natural sources (Sen & Peucker-
Ehrenbrink, 2012;Gałuszka et al., 2013), implying that there
will be a spike in fluxes in these metals in river outflow and
hence higher concentrations in coastal sediments.
Faunal radiation and extinctions
The last few centuries have seen significant changes in the abun-
dance and spread of small animals, particularly rats, mice and
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cats, etc. that are associated with human exploration and biotic
exchanges. Isolated populations almost everywhere have now
been superseded in many respects by these invasive species. The
fossil record will likely indicate a large faunal radiation of these
indicator species at this point. Concurrently, many other species
have already, or are likely to become, extinct, and their disappear-
ance from the fossil record will be noticeable. Given the perspec-
tive from many million years ahead, large mammal extinctions
that occurred at the end of the last ice age will also associated
with the onset of the Anthropocene.
Non-naturally occurring synthetics
There are many chemicals that have been (or were) manufactured
industrially that for various reasons can spread and persist in the
environment for a long time (Bernhardt et al., 2017). Most
Fig. 1. Illustrative stable carbon isotopes and temperature (or proxy) profiles across three periods. (a) The modern era (from 1600 CE with projections to 2100).
Carbon isotopes are from sea sponges (Böhm et al., 2002), and projections from Köhler (2016). Temperatures are from Mann et al. (2008) (reconstructions),
GISTEMP (Hansen et al., 2010) (instrumental) and projected to 2100 using results from Nazarenko et al. (2015). Projections assume trajectories of emissions asso-
ciated with RCP8.5 (van Vuuren et al., 2011). (b) The Paleocene–Eocene Thermal Maximum (55.5 Ma). Data from two DSDP cores (589 and 1209B) (Tripati &
Elderfield, 2004) are used to estimate anomalous isotopic changes and a loess smooth with a span of 200 kya is applied to make the trends clearer.
Temperatures changes are estimated from observed δ
18
O
carbonate
using a standard calibration (Kim & O’Neil, 1997). (c) Oceanic Anoxic Event 1a (about
120 Ma). Carbon isotopes are from the La Bédoule and Cau cores from the paleo-Tethys (Kuhnt et al., 2011; Naafs et al., 2016) aligned as in Naafs et al. (2016)
and placed on an approximate age model. Data from Alstätte (Bottini & Mutterlose, 2012) and DSDP Site 398 (Li et al., 2008) are aligned based on coherence
of the δ
13
C anomalies. Temperature change estimates are derived from TEX86 (Mutterlose et al., 2014; Naafs et al., 2016). Note that the y-axis spans the same
range in all three cases, while the timescales vary significantly.
International Journal of Astrobiology 145
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notably, persistent organic pollutants (organic molecules that are
resistant to degradation by chemical, photo-chemical or biological
processes), are known to have spread across the world (even to
otherwise pristine environments) (Beyer et al., 2000). Their per-
sistence is often tied to being halogenated organics since the
bond strength of C–Cl (for instance) is much stronger than C–
C. For instance, polychlorinated biphenyls are known to have life-
times of many hundreds of years in river sediment (Bopp, 1979).
How long a detectable signal would persist in ocean sediment is,
however, unclear.
Other chlorinated compounds may also have the potential for
long-term preservation, specifically CFCs and related compounds.
While there are natural sources for the most stable compound
(CF
4
), there are only anthropogenic sources for C
2
F
6
and SF
6
,
the next most stable compounds. In the atmosphere, their sink
via photolytic destruction in the stratosphere limits their lifetimes
to a few thousand years (Ravishankara et al., 1993). The com-
pounds do dissolve in the the ocean and can be used as tracers
of ocean circulation, but we are unaware of studies indicating
how long these chemicals might survive and/or be detectable in
ocean sediment given some limited evidence for microbial deg-
radation in anaerobic environments (Denovan & Strand, 1992).
Other classes of synthetic biomarkers may also persist in sedi-
ments. For instance, steroids, leaf waxes, alkenones and lipids can
be preserved in sediment for many millions of years (i.e. Pagani
et al., 2006). What might distinguish naturally occurring biomar -
kers from synthetics might be the chirality of the molecules. Most
total synthesis pathways do not discriminate between D- and
L-chirality, while biological processes are almost exclusively
monochiral (Meierhenrich, 2008) (for instance, naturally occur-
ring amino acids are all L-forms, and almost all sugars are
D-forms). Synthetic steroids that do not have natural counterparts
are also now ubiquitous in water bodies.
Plastics
Since 1950, there has been a huge increase in plastics being deliv-
ered into the ocean (Moore, 2008; Eriksen et al., 2014). Although
many common forms of plastic (such as polyethylene and poly-
propylene) are buoyant in sea water, and even those that are nom-
inally heavier than water may be incorporated into flotsam that
remains at the surface, it is already clear that mechanical erosional
processes will lead to the production of large amounts of plastic
micro- and nano-particles (Cozar et al., 2014; Andrady, 2015).
Surveys have shown increasing amounts of plastic ‘marine litter’
on the seafloor from coastal areas to deep basins and the Arctic
(Pham et al., 2014; Tekman et al., 2017). On beaches, novel aggre-
gates ‘plastiglomerates’ have been found where plastic-containing
debris comes into contact with high temperatures (Corcoran et al.,
2014).
The degradation of plastics is mostly by solar ultraviolet radi-
ation and in the oceans occurs mostly in the photic zone
(Andrady, 2015) and is notably temperature dependent (Andrady
et al., 1998) (other mechanisms such as thermo-oxidation or
hydrolysis do not readily occur in the ocean). The densification
of small plastic particles by fouling organisms, ingestion and
incorporation into organic ‘rains’ that sink to the sea floor is an
effective delivery mechanism to the seafloor, leading to increasing
accumulation in ocean sediment where degradation rates are
much slower (Andrady, 2015). Once in the sediment, microbial
activity is a possible degradation pathway (Shah et al., 2008)but
rates are sensitive to oxygen availability and suitable microbial
communities.
As above, the ultimate long-term fate of these plastics in sedi-
ment is unclear, but the potential for very long-term persistence
and detectability is high.
Transuranic elements
Many radioactive isotopes that are related to anthropogenic fission
or nuclear arms, have half-lives that are long, but not long enough
to be relevant here. However, there are two isotopes that are poten-
tially long-lived enough. Specifically, Plutonium-244 (half-life 80.8
million years) and Curium-247 (half-life 15 million years) would be
detectable for a large fraction of the relevant time period if they
were deposited in sufficient quantities, say, as a result of a nuclear
weapon exchange. There are no known natural sources of
244
Pu
outside of supernovae.
Attempts have been made to detect primordial
244
Pu on Earth
with mixed success (Hoffman et al., 1971; Lachner et al., 2012),
indicating the rate of actinide meteorite accretion is small enough
(Wallner et al., 2015) for this to be a valid marker in the event of a
sufficiently large nuclear exchange. Similarly,
247
Cm is present in
nuclear fuel waste and as a consequence of a nuclear explosion.
Anomalous isotopic ratios in elements with long-lived radio-
active isotopes are also possible signatures, for instance, lower
than usual
235
U ratios, and the presence of expected daughter pro-
ducts, in uranium ores in the Franceville Basin in the Gabon have
been traced to naturally occurring nuclear fission in oxygenated,
hydrated rocks ∼2 Ga (Gauthier-Lafaye et al., 1996).
Summary
The Anthropocene layer in ocean sediment will be abrupt and
multi-variate, consisting of seemingly concurrent-specific peaks
in multiple geochemical proxies, biomarkers, elemental compos-
ition and mineralogy. It will likely demarcate a clear transition
of faunal taxa prior to the event compared with afterwards.
Most of the individual markers will not be unique in the context
of Earth history as we demonstrate below, but the combination of
tracers may be. However, we speculate that some specific tracers
that would be unique, specifically persistent synthetic molecules,
plastics and (potentially) very long-lived radioactive fallout in
the event of nuclear catastrophe. Absent those markers, the
uniqueness of the event may well be seen in the multitude of rela-
tively independent fingerprints as opposed to a coherent set of
changes associated with a single geophysical cause.
Abrupt paleozoic, mesozoic and cenozoic events
The summary for the Anthropocene fingerprint above suggests
that similarities might be found in (geologically) abrupt events
with a multi-variate signature. In this section, we review a partial
selection of known events in the paleo-record that have some
similarities to the hypothesized eventual anthropogenic signature.
The clearest class of event with such similarities are the
hyperthermals, most notably the Paleocene–Eocene Thermal
Maximum (56 Ma) (McInerney & Wing, 2011), but this also
includes smaller hype rthermal events, ocean anoxic events in
the Cretaceous and Jurassic and significant (if less well character-
ized) events of the Paleozoic. We do not consider of events (such
as the K–T extinction event or the Eocene–Oligocene boundary)
where there are very clear and distinct causes (asteroid impact
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combined with massive volcanism(Vellekoop et al., 2014), and the
onset of Antarctic glaciation(Zachos et al., 2001) (likely linked to
the opening of Drake Passage Cristini et al., 2012, respectively).
There may be more such events in the record but that are not
included here simply because they may not have been examined
in detail, particularly in the pre-Cenozoic.
The Paleocene–Eocene thermal ma ximum (PETM)
The existence of an abrupt spike in carbon and oxygen isotopes
near the Paleocene/Eocene transition (56 Ma) was first noted by
Kennett & Stott (1991) and demonstrated to be global by Koch
et al. (1992). Since then, more detailed and high-resolution ana-
lyses on land and in the ocean have revealed a fascinating
sequence of events lasting 100–200 kyr and involving a rapid
input (in perhaps <5 kyr Kirtland Turner et al., 2017) of exogen-
ous carbon into the system (see review by McInerney & Wing,
2011), possibly related to the intrusion of the North American
Igneous Province into organic sediments (Storey et al., 2007).
Temperatures rose 5–7°C (derived from multiple proxies Tripati
& Elderfield, 2004), and there was a negative spike in carbon iso-
topes (>3‰), and decreased ocean carbonate preservation in the
upper ocean. There was an increase in kaolinite (clay) in many
sediments (Schmitz et al., 2001), indicating greater erosion,
though evidence for a global increase is mixed. During the
PETM 30–50% of benthic foraminiferal taxa became extinct,
and it marked the time of an important mammalian (Aubry
et al., 1998) and lizard (Smith, 2009) expansion across North
America. Additionally, many metal abundances (including V,
Zn, Mo, Cr) spiked during the event (Soliman et al., 2011).
Eocene events
In the 6 million years following the PETM, there are a number of
smaller, though qualitatively similar, hyperthermal events seen in
the record (Slotnick et al., 2012). Notably, the Eocene Thermal
Maximum 2 event (ETM-2), and at least four other peaks are
characterized by significant negative carbon isotope excursions,
warming and relatively high sedimentation rates driven by
increases in terrigenous input (D’Onofrio et al., 2016). Arctic con-
ditions during ETM-2 show evidence of warming, lower salinity
and greater anoxia (Sluijs et al., 2009). Collectively these events
have been denoted Eocene Layers of Mysterious Origin
(ELMOs)
2
.
Around 40 Ma, another abrupt warming event occurs (the
mid-Eocene Climate Optimum (MECO)), again with an accom-
panying carbon isotope anomaly (Galazzo et al., 2014).
Cretaceous and Jurassic ocean anoxic events
First identified by Schlanger & Jenkyns (1976), ocean anoxic
events (OAEs), identified by periods of greatly increased organic
carbon deposition and laminated black shale deposits, are times
when significant portions of the ocean (either regionally or glo-
bally) became depleted in dissolved oxygen, greatly reducing aer-
obic bacterial activity. There is partial (though not ubiquit ous)
evidence during the larger OAEs for euxinia (when the ocean
column becomes filled with hydrogen sulfide (H
2
S)) (Meyer &
Kump, 2008).
There were three major OAEs in the Cretaceous, the Weissert
event (132 Ma) (Erba et al., 2004), OAE-1a around 120 Ma lasting
about 1 Myr and another OAE-2 around 93 Ma lasting around
0.8 Myr (Kerr, 1998; Li et al., 2008; Malinverno et al., 2010;Li
et al., 2017). At least four other minor episodes of organic black
shale production are noted in the Cretaceous (the Faraoni
event, OAE-1b, 1d and OAE-3) but seem to be restricted to the
proto-Atlantic region (Takashima et al., 2006; Jenkyns, 2010).
At least one similar event occurred in the Jurassic (183 Ma)
(Pearce et al., 2008).
The sequence of events during these events have two distinct
fingerprints possibly associated with the two differing theoretical
mechanisms for the events. For example, during OAE-1b, there is
evidence of strong stratification and a stagnant deep ocean, while
for OAE-2, the evidence suggests an decrease in stratification,
increased upper ocean productivity and an expansion of the oxy-
gen minimum zones (Takashima et al., 2006).
At the onset of the events (Fig. 1(c)), there is often a significant
negative excursion in δ
13
C (as in the PETM), followed by a posi-
tive recovery during the events themselves as the burial of (light)
organic carbon increased and compensated for the initial release
(Jenkyns, 2010 ; Kuhnt et al., 2011; Mutterlose et al., 2014;
Naafs et al., 2016). Causes have been linked to the crustal forma-
tion/tectonic activity and enhanced CO
2
(or possibly CH
4
)
release, causing global warmth (Jenkyns, 2010). Increased sea-
water values of
87
Sr/
86
Sr and
187
Os/
188
Os suggest increased runoff,
greater nutrient supply and consequently higher upper ocean
productivity (Jones, 2001). Possible hiatuses in some OAE 1a sec-
tions are suggestive of an upper ocean dissolution event (Bottini
et al., 2015).
Other important shifts in geochemical tracers during the
OAEs include much lower nitrogen isotope ratios (δ
15
N),
increases in metal concentrations (including As, Bi, Cd, Co, Cr,
Ni, V) (Jenkyns, 2010). Positive shifts in sulph ur isotopes are
seen in most OAEs, with a curious exception in OAE-1a where
the shift is negative (Turchyn et al., 2009).
Early Mesozoic and Late Paleozoic events
Starting from the Devonian period, there have been several major
abrupt events registered in terrestrial sections. The sequences of
changes and the comprehensiveness of geochem ical analyses are
less well known than for later events, partly due to the lack of
existing ocean sediment, but these have been identified in mul-
tiple locations and are presumed to be global in extent.
The Late Devonian extinction around 380–360 Ma, was one of
the big five mass extinctions. It is associated with black shales and
ocean anoxia (Algeo & Scheckler, 1998), stretching from the
Kellwasser events (∼378 Ma) to the Hangenberg event at the
Devonian–Carboniferous boundary (359 Ma) (Brezinski et al.,
2009; Vleeschouwer et al., 2013).
In the late Carboniferous, around 305 Ma the Pangaean trop-
ical rainforests collapsed (Sahney et al., 2010). This was associated
with a shift towards drier and cooler climate, and possibly a
reduction in atmospheric oxygen, leading to extinctions of some
mega-fauna.
Lastly, the end-Permian extinction event (252 Ma) lasted
about 60 kyr was accompanied by an initial decrease in carbon
isotopes (− 5–7‰), significant global warming and extensive
deforestation and wildfires (Krull & Retallack, 2000; Shen et al.,
2
While it is tempting to read something into the nomenclature of these events, it
should be remembered that most things that happened 50 million years ago will forever
remain somewhat mysterious.
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2011; Burgess et al., 2014) associated with widespread ocean
anoxia and euxinia (Wignall & Twitchett, 1996). Pre-event spikes
in nickel (Ni) have also been reported (Rothman et al., 2014).
Discussion and testable hypotheses
There are undoubted similarities between previous abrupt events
in the geological record and the likely Anthropocene signature in
the geological record to come. Negative, abrupt δ
13
C excursions,
warmings and disruptions of the nitrogen cycle are ubiquitous.
More complex changes in biota, sedimentation and mineralogy
are also common. Specifically, compared with the hypothesized
Anthropocene signature, almost all changes found so far for the
PETM are of the same sign and comparable magnitude. Some
similarities would be expected if the main effect during any
event was a significant global warming, however caused.
Furthermore, there is evidence at many of these events that warm-
ing was driven by a massive input of exogeneous (biogenic) car-
bon, either as CO
2
or CH
4
. At least since the Carboniferous
(300–350 Ma), there has been sufficient fossil carbon to fuel an
industrial civilization comparable with our own and any of
these sources could provide the light carbon input. However, in
many cases this input is contemporaneous to significant episodes
of tectonic and/or volcanic activity, for instance, the coincidence
of crustal form ation events with the climate chan ges suggest that
the intrusion of basaltic magmas into organic-rich shales and/or
petroleum-bearing evaporites (Storey et al., 2007; Svensen et al.,
2009; Kravchinsky, 2012) may have released large quantities of
CO
2
or CH
4
to the atmosphere. Impacts to warming and/or car-
bon influx (such as increased runoff, erosion, etc.) appear to be
qualitatively similar whenever in the geological period they
occur. These changes are thus not sufficient evidence for prior
industrial civilizations.
Current changes appear to be significantly faster than the
paleoclimatic events (Fig. 1), but this may be partly due to limita-
tions of chronology in the geological record. Attempts to time the
length of prior events have used constant sedimentation estimates,
or constant-flux markers (e.g.
3
He McGee & Mukhopadhyay,
2012), or orbital chronologies, or suppos ed annual or seasonal
banding in the sediment (Wright & Schaller, 2013). The accuracy
of these methods suffer when there are large changes in sedimen-
tation or hiatuses across these events (which is common), or rely
on the imperfect identification of regularities with specific astro-
nomical features (Pearson & Nicholas, 2014; Pearson & Thomas,
2015). Additionally, bioturbation will often smooth an abrupt
event even in a perfectly preserved sedimentary se tting. Thus,
the ability to detect an event onset of a few centuries (or less)
in the record is questionable, and so direct isolation of an indus-
trial cause based only on apparent timing is also not conclusive.
The speci fic markers of human industrial activity discussed
above (plastics, synthetic pollutants, increased metal concentra-
tions, etc.) are however a consequence of the specific path
human society and technology has taken, and the generality of
that pathway for other industrial species is totally unknown.
Large-scale energy harnessing is potentially a more universal indi-
cator, and given the large energy density in carbon-based fossil
fuel, one might postulate that a light δ
13
C signal might be a com-
mon signal. Conceivably, solar, hydro or geothermal energy
sources could have been tapped preferentially, and that would
greatly reduce any geological footprint (as it would ours).
However, any large release of biogenic carbon whether from
methane hydrate pools or volcanic intrusions into organic-rich
sediments, will have a similar signal. We therefore have a situation
where the known unique markers might not be indicative, while
the (perhaps) more expected markers are not sufficient.
We are aware that raising the possibility of a prior industrial
civilization as a driver for events in the geological record might
lead to rather unconstrained speculation. One would be able to
fit any observations to an imagined civilization in ways that
would be basically unfalsifiable. Thus, care must be taken not to
postulate such a cause until actually positive evidence is available.
The Silurian hypothesis cannot be regarded as likely merely
because no other valid idea presents itself.
We nonetheless find the above analyses intriguing enough to
motivate some additional research. Firstly, despite cop ious exist-
ing work on the likely Anthropocene signature, we recommend
further synthesis and study on the persistence of uniquely indus-
trial byproducts in ocean sediment environments. Are there other
classes of compounds that will leave unique traces in the sediment
geochemistry on multi-million year timescales? In particular, will
the byproducts of common plastics, or organic long-chain syn-
thetics, be detectable?
Secondly, and this is indeed more speculative, we propose that
a deeper exploration of elemental and compositional anomalies in
extant sediments spanning previous events be performed
(although we expect that far more information has been obtained
about these sections than has been referenced here). Oddities in
these sections have been looked for previously as potential signals
of impact events (successfully for the K–T boundary event, not so
for any of the events mentioned above), rang ing from iridium
layers, shocked quartz, micro-tectites, magnetites, etc. But it
may be that a new search and new analyses with the Silurian
hypothesis in mind might reveal more. Anomalous behaviour in
the past might be more clearly detectable in proxies normalized
by weatherin g fluxes or other constant flux proxies in order to
highlight times when productivity or metal production might
have been artificially enhanced. Thirdly, should any unexplained
anomalies be found, the question of whether there are candidate
species in the fossil record may become more relevant, as might
questions about their ultimate fate.
An intriguing hypothesis presents itself should any of the ini-
tial releases of light carbon descri bed above indeed be related to a
prior industrial civilization. As discussed in the section
‘Cretaceous and Jurassic ocean anoxic events’, these releases
often triggered episodes of ocean anoxia (via increased nutrient
supply) causing a massive burial of organic matter, which eventu-
ally became source strata for further fossil fuels. Thus, the prior
industrial activity would have actually given rise to the potential
for future industry via their own demise. Large-scale anoxia, in
effect, might provide a self-limiting but self-perpetuating feedback
of industry on the planet. Alternatively, it may be just be a part of
a long-term episodic natural carbon cycle feedback on tectonically
active planets.
Perhaps unusually, the authors of this paper are not convinced
of the correctness of their proposed hypothesis. Were it to be true
it would have profound implications and not just for astrobiology.
However, most readers do not need to be told that it is always a
bad idea to decide on the truth or falsity of an idea based on
the consequences of it being true. While we strongly doubt that
any previous industrial civilization existed before our own, asking
the question in a formal way that articulates explicitly what evi-
dence for such a civilization might look like raises its own useful
questions related both to astrobiology and to Anthropocene stud-
ies. Thus, we hope that this paper will serve as motivation to
148 Schmidt and Frank
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improve the constraints on the hypothesis so that in future we
may be better placed to answer our title question.
Acknowledgments. No funding has been provided nor sought for this study.
We thank Susan Kidwell for being generous with her time and helpful discus-
sions, David Naafs and Stuart Robinson for help and pointers to data for OAE1a
and Chris Reinhard for his thoughtful review. The GISTEMP data in Fig. 1(a)
were from https://data.giss.nasa.gov/gistemp (accessed Jul 15 2017).
References
Achille GD and Hynek BM (2010) Nature Geoscience 3, 459.
Algeo TJ and Scheckler SE (1998) Philosophical Transactions of the Royal
Society B: Biological Sciences 353, 113.
Andrady A, et al. (1998) Journal of Photochemistry and Photobiology B:
Biology 46, 96.
Andrady AL (2015) Marine Anthropogenic Litter (Springer Nature), 57–72.
Arvidson RE, et al. (2014) Science 343, 1248097.
Aubry M-P, Lucas S and Berggren WA eds (1998) Late Paleocene-early
Eocene Biotic and Climatic Events in the Marine and Terrestrial Records.
Columbia University Press, New York, NY
Behrensmeyer AK, Kidwell SM and Gastaldo RA (2000) Paleobiology 26,
103.
Bernhardt ES, Rosi EJ and Gessner MO (2017) Frontiers in Ecology and the
Environment 15, 84.
Beyer A, et al. (2000) Environmental Science & Technology 34, 699.
Bindoff NL, et al. (2013) Climate change 2013: the physical science basis. In
Stocker TF, Qin D, Plattner G-K, Tignor M , Allen SK , Boschung J ,
Nauels A , Xia Y , Bex V and Midgley P eds. Contribution of Working
Group I to the Fifth Assessment Report of the Intergovernmental Panel on
Climate Change. Cambridge: Cambridge University Press, pp. 867–952.
Böhm F, et al. (2002) Geochemistry, Geophysics, Geosystems 3,1.
Bopp RF (1979) PhD thesis. Columbia University.
Bottini C, et al. (2015) Climate of the Past 11, 383.
Bottini C and Mutterlose J (2012) Newsletters on Stratigraphy 45, 115.
Breitburg D, et al. (2018) Science 359, eaam7240, doi: 10.1126/
science.aam7240
Brezinski DK, et al.
(2009) Palaeogeography, Palaeoclimatology, Palaeoecology
284, 315.
Burgess SD, Bowring S and zhong Shen S (2014) Proceedings of the National
Academy of Sciences USA 111, 3316.
Canfield DE, Glazer AN and Falkowski PG (2010) Science 330, 192.
Cleland CE and Copley SD (2006) International Journal of Astrobiology 4, 165.
Corcoran PL, Moore CJ and Jazvac K (2014) GSA Today, 24,4–8
Cozar A, et al. (2014) Proceedings of the National Academy of Sciences USA
111, 10239.
Cristini L, et al. (2012) Palaeogeography, Palaeoclimatology, Palaeoecology
339–341, 66.
Crutzen PJ (2002) Nature 415, 23.
Denovan BA and Strand SE (1992) Chemosphere 24, 935.
Diaz RJ and Rosenberg R (2008) Science 321, 926.
D’Onofrio R, et al. (2016) Paleoceanography 31, 1225.
Drake FD (1961) Discussion at Space Science Board-National Academy of
Sciences Conference on Extraterrestrial Intelligent Life, Green Bank, West
Virginia, USA.
Drake FD (1965) In Mamikunian G and Briggs MH eds. Current Aspects of
Exobiology. Pergamon, Oxford, pp. 323–345.
Dressing CD and Charbonneau D (2013) The Astrophysical Journal 767, 95.
Eide M, et al. (2017) Global Biogeochemical Cycles 31, 492–514.
Erba E, Bartolini A and Larson RL (2004) Geology 32, 149.
Eriksen M, et al. (2014) PLoS ONE
9, e111913.
Frank A and Sullivan WT (2016) Astrobiology 16, 359.
Galazzo FB, et al. (2014) Paleoceanography 29, 1143.
Gałuszka A, Migaszewski ZM and Zalasiewicz J (2013) Geological Society,
London, Special Publications 395, 221.
Gauthier-Lafaye F, Holliger P and Blanc P-L (1996) Geochimica et
Cosmochimica Acta 60, 4831.
GISTEMP Team (2016) GISS Surface Temperature Analysis (GISTEMP).
Goudie A (2000) The human impact on the soil (The Human Impact on the
Natural Environment. MIT Press), 88.
Hamilton C (2016) Nature 536, 251.
Hansen J, et al. (2010) Reviews in Geophysics 48, EID: RG4004.
Haqq-Misra J and Kopparapu RK (2012) Acta Astronautica 72, 15.
Hazen RM, et al. (2017) American Mineralogist 102, 595.
Hoffman DC, et al. (1971) Nature 234, 132.
Hogan J (1977) Inherit the Stars (Ballantine Books, New York).
Holtgrieve GW, et al. (2011) Science 334, 1545.
Howard AW (2013) Science 340, 572.
Ito T, et al. (2017) Geophysical Research Letters 44, 4214–4223.
Jenkyns HC (2010) Geochemistry, Geophysics, Geosystems 11, Q03004.
Jones CE (2001) American Journal of Science 301, 112.
Kennett J and Stott LD (1991) Nature 353
, 319.
Kerr AC (1998) Journal of the Geological Society 155, 619.
Kidwell SM (2015) Proceedings of the National Academy of Sciences USA 112,
4922.
Kim S-T and O’Neil JR (1997) Geochimica et Cosmochimica Acta 61,
3461.
Kirtland Turner S, et al. (2017) Nature Communications 8, Article ID: 353
Koch PL, Zachos JC and Gingerich P (1992) Nature 358, 319.
Köhler P (2016) Environmental Research Letters 11, 124016.
Kravchinsky VA (2012) Global and Planetary Change 86–87, 31.
Krull ES and Retallack GJ (2000) Geological Society of America Bulletin 112,
1459.
Kuhnt W, Holbourn A and Moullade M (2011) Geology 39, 323.
Kunkel KE, et al. (2013) Bulletin of the American Meteorological Society 94,
499.
Lachner J, et al. (2012) Physical Review C 85, 015801.
Le Quéré C, et al. (2016) Earth System Science Data 8, 605.
Leakey MD and Hay RL (1979) Nature 278, 317.
Lewis SL and Maslin MA (2015) Nature 519, 171.
Li Y-X, et al. (2008) Earth and Planetary Science Letters 271, 88.
Li Y-X, et al. (2017) Earth and Planetary Science Letters 462, 35.
Malinverno A, Erba E and Herbert TD (2010) Paleoceanography 25, EID:
PA2203.
Mann ME, et al. (2008) Proceedings of the National Academy of Sciences USA
105, 13252.
Marino L (2015) Fraction of life-bearing planets on which intelligent life emerges,
fi, 1961 to the present. In Vakoch DA and Dowd MF (eds). The Drake
Equation: Estimating the Prevalence of Extraterrestrial Life through the Ages.
Cambridge University Press, Cambridge, UK, pp. 181–204.
Matmon A, et al. (2009) Geological Society of America Bulletin 121, 688.
McGee D and Mukhopadhyay S (2012) The Noble Gases as Geochemical
Tracers. Berlin, Heidelberg: Springer, pp. 155–176.
McInerney FA and Wing SL (2011) Annual Review of Earth and Planetary
Sciences 39, 489.
Meierhenrich U (2008) Amino Acids and the Asymmetry of Life. Berlin,
Heidelberg: Springer.
Meyer KM and Kump LR (2008) Annual Review of Earth and Planetary
Sciences 36, 251.
Moore CJ (2008) Environmental Research 108, 131.
Mutterlose J, et al. (2014) Geology 42, 439.
National Research Council (2010) Toward Sustainable Agricultural Systems
in the 21st Century. The National Academies Press, Washington, DC.
https://doi.org/10.17226/12832.
Naafs BDA, et al. (2016) Nature Geoscience 9, 135.
Nazarenko L, et al. (2015) Journal of Advances in Modeling Earth Systems 7,
244.
ODP Leg 801 Team (2000) Proceedings of the Ocean Drilling Program
(International Ocean Discovery Program (IODP)).
Orr JC, et al. (2005) Nature 437, 681.
Overeem I, et al. (2011) Geophysical Research Letters 38
, L17503.
Pagani M, et al. (2006) Nature 442, 671.
Patel BH, et al. (2015) Nature Chemistry 7, 301.
Pearce CR, et al. (2008) Geology 36, 231.
Pearson PN and Nicholas CJ (2014) Proceedings of the National Academy of
Sciences USA 111, E1064.
International Journal of Astrobiology 149
https://www.cambridge.org/core/terms. https://doi.org/10.1017/S1473550418000095
Downloaded from https://www.cambridge.org/core. IP address: 67.164.57.139, on 21 Apr 2019 at 20:40:28, subject to the Cambridge Core terms of use, available at
Pearson PN and Thomas E (2015) Climate of the Past 11, 95.
Pham CK, et al. (2014) PLoS ONE 9, e95839.
Quay PD, Tilbrook B and Wong CS (1992) Science 256, 74.
Ravishankara AR, et al. (1993) Science 259, 194.
Revelle R and Suess HE (1957) Tellus 9, 18.
Rothman DH, et al. (2014) Proceedings of the National Academy of Sciences
USA 111, 5462.
Sahney S, Benton MJ and Falcon-Lang HJ (2010) Geology 38, 1079.
Schlanger SO and Jenkyns HC (1976) Geologie en Mijnbouw 55, 179.
Schmidtko S, Stramma L and Visbeck M (2017) Nature 542, 335.
Schmitz B, Pujalte V and Nú nez-Betelu K (2001) Palaeogeography,
Palaeoclimatology, Palaeoecology 165, 299.
Schneider A, Friedl MA and Potere D (2009) Environmental Research Letters
4, 044003.
Seager S (2013) Science 340, 577.
Sen IS and Peucker-Ehrenbrink B (2012) Environmental Science &
Technology 46, 8601.
Shah AA, et al. (2008) Biotechnology Advances 26, 246.
Shen S, et al. (2011) Science 334, 1367.
Slotnick BS, et al. (2012) The Journal of Geology 120, 487.
Sluijs A, et al. (2009) Nature Geoscience 2
, 777.
Smith KT (2009) Journal of Systematic Palaeontology 7, 299.
Soliman MF, et al. (2011) Palaeogeography, Palaeoclimatology, Palaeoecology
310, 365.
Storey M, Duncan RA and Swisher CC (2007) Science 316, 587.
Svensen H, et al. (2009) Earth and Planetary Science Letters 277, 490.
Takashima R, et al. (2006) Oceanography 19,82–92.
Tekman MB, Krumpen T and Bergmann M (2017) Deep Sea Research Part I:
Oceanographic Research Papers 120, 88.
Tripati AK and Elderfield H (2004) Geochemistry, Geophysics, Geosystems
5,EID:Q02006.
Turchyn AV, et al. (2009) Earth and Planetary Science Letters 285, 115.
Tyrrell T (2011) Philosophical Transactions of the Royal Society A:
Mathematical, Physical and Engineering Sciences 369, 887.
van Vuuren DP, et al. (2011) Climatic Change 109,5.
Vellekoop J, et al. (2014) Proceedings of the National Academy of Sciences USA
111, 7537.
Vleeschouwer DD, et al. (2013) Earth and Planetary Science Letters 365, 25.
Wallner A, et al. (2015) Nature Communications 6, 5956.
Waters C, et al. (2014) A Stratigraphical Basis for the Anthropocene, GSL
Special Publications No. 395. London: Geological Society.
Way MJ, et al. (2016) Geophysical Research Letters 43, 8376.
Wignall PB and Twitchett RJ
(1996) Science 272, 1155.
Wright JD and Schaller MF (2013) Proceedings of the National Academy of
Sciences USA 110, 15908.
Wright JT (2017) (submitted), arXiv:1704.07263.
Zachos JC, et al. (2001) Eos Transactions AGU Fall Meeting Supplement 82,
767.
Zalasiewicz J (2009) Earth Without Us. Oxford, UK: Oxford University Press.
Zalasiewicz J, Kryza R and Williams M (2013) Geological Society, London,
Special Publications 395, 109.
Zalasiewicz JN, et al. (2017) Newsletters on Stratigraphy 50, 205–226.
150 Schmidt and Frank
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Other elements listed can be easily present in a planet's atmosphere without an ongoing chemical process generating them; they're stable. Oxygen is rather reactive (see: fires, rust) so it's unlikely to be seen without an ongoing active process like photosynthesis. This reactivity is useful for our own processes.
It's possible to have anaerobic respiration, but it's really inefficient in comparison (the pathway that generates lactic acid is about 1/15th of the efficiency). Or you can use other reducing agents -- Florine is stronger than oxygen, but the options for what you build your body out of that doesn't get set on fire by the florine are somewhat limited.
If you are curious here is below is a clip of the first appearance of Silurians in the Doctor Who series :)
[](https://www.youtube.com/watch?v=9SZEm3uoa5o)
It has been estimated that only one bone in a billion gets fossilised. With those odds, the entire fossil legacy of 320 million people living in the US today would be approximately 60 bones (approximately 25% of a human skeleton).
Another element that would be detectable is technetium-98 ($^{98}T$), which has a 4.2-million-year half-life and no technetium remains from the formation of the Earth.
Technetium is produced each year from spent nuclear fuel rods.
**Gavin A. Schmidt** is a climatologist, climate modeler and Director of the NASA Goddard Institute for Space Studies (GISS) in New York.
**Adam Frank** is an American physicist, astronomer, and writer. His scientific research has focused on computational astrophysics with an emphasis on star formation and late stages of stellar evolution.
The problem is that if the activity of a hypotetical industrial civilization is short-lived we might not be able to see it. The PETM’s spikes mostly show us the Earth’s timescales for responding to whatever caused it, not necessarily the timescale of the cause. We might need new ways to find evidence of a short-lived industrial civilization and we also need to consider that they might not have followed the same trajectory as our own civilization. It seems to be clear though, that if we are not looking for it we might not see it.
I’ve pondered this ‘possibility’ all my length - that there have been civilisation(s) before us on Earth and that for some reason they became extinct and also that any possible archaeological evidence has long degraded and disappeared.
However the paper also looks at possible civilisations from Venus or Mars.
What I’ve NEVER seen discussed is why do we only look for intelligent(?) life like our own that breathes oxygen?
Could there not be lifeforms that evolved that could survive on planets that needed obnoxious elements (cf to us) like Sulphur, CO2, N etc. Lifeforms that could exist under atmospheric pressures that would crush us.
The Negev Desert in Israel has been dated to about 1.8 million years old, the oldest known vast expanse of surface area on the planet. It is more than four times older than the confirmed next oldest desert pavement, in Nevada.
The reason this region was able to be preserved for so long is because it is extremely flat and arid, where tectonic activity is low and rocks are highly resistant to weathering.
One of the most common searched structures that can be seen as an evidence of the existence of an advanced alien civilization is the Dyson Sphere. A Dyson Sphere is a hypothetical structure that an advanced civilization
might build around a star to intercept all of the star’s light
for its energy needs.
Physicist Freeman Dyson wrote a paper in 1960 that attempted to formalize the concept of a Dyson Sphere for the first time. If you are interested in the paper, we've annotated it on Fermat's Library and you can read it [here](https://fermatslibrary.com/s/search-for-artificial-stellar-sources-of-infrared-radiation#email-newsletter).
Dyson argues that a way to look for extraterrestrial life is to search for the presence of Dyson Spheres.
The basic idea is that as a civilization evolves it will have increasingly energy requirements. Once all the resources of the home planet have been explored the next obvious step is to get energy from the closest star.
According to the Kardashev scale there are 3 types of civilizations based on the amount of energy a civilization is able to use:
- **Type I** - can use and store all of the energy which reaches its planet from its parent star.
- **Type II** - can harness the total energy of its planet's parent star.
- **Type III** - can control energy on the scale of its entire host galaxy.
Humans are currently a Type 0 civilization and Michio Kaku predicts that we attain Type I status in 100–200 years, Type II status in a few thousand years, and Type III status in 100,000 to a million years.
As a consequence of large-scale usage of a star's energy, there's going to be a peak in far-infrared radiation which in principle could be detected by radio satellites.
Carbon has 3 naturally occurring isotopes
- $^{12}C$, about 99% of carbon on Earth
- $^{13}C$, about 1% of carbon on Earth
- $^{14}C$, <0.1% of carbon on Earth
The $^{12}C$ and $^{13}C$ isotopes are stable, while $^{14}C$ decays radioactively to nitrogen-{14} with a half life of 5730 years. $^{14}C$ on Earth is produced nearly exclusively by the interaction of cosmic radiation with the upper atmosphere (thus natural $^{14}C$ production and hence atmospheric concentration varies only slightly over time).
Plants take up $^{14}C$ by fixing atmospheric carbon through photosynthesis. Animals then take $^{14}C$ into their bodies when they consume plants (or consume other animals that consume plants). Thus, living plants and animals have the same ratio of $^{14}C$ to $^{12}C$ as the atmospheric CO2. Once organisms die they stop exchanging carbon with the atmosphere, and thus no longer take up new $^{14}C$. Radioactive decay then gradually depletes the $^{14}C$ in the organism. This effect is the basis of radiocarbon dating.
Photosynthetically fixed carbon in terrestrial plants is depleted in $^{13}C$ compared to atmospheric CO2. Fossil fuels such as coal and oil are made primarily of plant material that was deposited millions of years ago. This period of time equates to thousands of half-lives of $^{14}C$, so essentially all of the $^{14}C$ in fossil fuels has decayed. Fossil fuels also are depleted in $^{13}C$ relative to the atmosphere, because they were originally formed from living organisms. Therefore, the carbon from fossil fuels that is returned to the atmosphere through combustion is depleted in both $^{13}C$ and $^{14}C$ compared to atmospheric carbon dioxide.
In geochemistry, paleoclimatology and paleoceanography $δ^{13}C$ (pronounced "delta c thirteen") is an isotopic signature, a measure of the ratio of stable isotopes $^{13}C$ : $^{12}C$, reported in parts per thousand (per mil, ‰).
$$
δ^{13}C = \left( \frac{\frac{^{13}C}{^{12}C}_{sample}}{\frac{^{13}C}{^{12}C}_{standard}}-1 \right) \times 1000‰
$$
where the standard is an established reference material.